Science, Faculty of
Earth, Ocean and Atmospheric Sciences, Department of
DSpace
UBCV
Rasmussen, Kirsten Louise
Doctor of Philosophy - PhD
University of British Columbia
Syn- to post-accretionary magmatic episodes of Cretaceous to Eocene age comprises the most voluminous magmatism in the northern Cordillera. Our understanding of the age, composition, and petrogenesis of this magmatism thus far, however, has been severely limited by a lack of robust geochronological, geochemical, and isotopic datasets for these rock units. This study fills the need for better temporal and compositional constraints on mainly Early to ‘mid’-Cretaceous (~115–90 Ma) magmatism in the Cassiar Terrane and eastern Selwyn Basin regions of the northern Cordilleran miogeocline. An extensive set of new U–Pb and ⁴⁰Ar/³⁹Ar age data, together with major and trace element geochemical analyses, radiogenic (Nd, Sr, Pb, and Hf) and stable (S and O) isotopic compositions, and detailed magmatic fluid compositional studies of intrusive rocks in the northern Cordillera are presented here. These results, and the integration of compiled datasets, have led to the identification of several new plutonic suites and sub-suites, as well as modifications of previously defined plutonic suites. Inferences regarding the age and composition of the basement to the northern Cordilleran miogeocline have also been made based on mineralogical, geochemical, and isotopic interpretations of magma sources for crustally derived plutonic suites from across this region. These extensive datasets have also allowed a re-evaluation of previously proposed models for syn- to post-accretionary magmatic episodes, and have resulted in the proposal of modified tectono–magmatic models for the particularly voluminous Early to ‘mid’-Cretaceous magmatic episode. Finally, estimates of magmatic fluid compositions determined from apatite and biotite compositions (with respect to HF, HCl, and H₂O) and extended cooling histories indicated by U–Pb and ⁴⁰Ar/³⁹Ar geochronology, have been used to identify intrusive phases (locally and regionally) that are associated with reduced tungsten skarn mineralization, which comprises some of the most economically important intrusion-related mineralization in the northern Cordillera. Collectively, this work has substantially advanced the current knowledge of syn- to post-accretionary magmatism in the northern Cordillera, and it has broad implications for similar magmatic provinces world-wide.
THE TIMING, COMPOSITION, AND PETROGENESIS OF SYN- TO POST- ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERAN MIOGEOCLINE, EASTERN YUKON AND SOUTHWESTERN NORTHWEST TERRITORIES by Kirsten Louise Rasmussen B.Sc., The University of Calgary, 2004 A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY in THE FACULTY OF GRADUATE STUDIES (Geological Sciences) THE UNIVERSITY OF BRITISH COLUMBIA (Vancouver) April 2013 © Kirsten Louise Rasmussen, 2013 ii ABSTRACT Syn- to post-accretionary magmatic episodes of Cretaceous to Eocene age comprises the most voluminous magmatism in the northern Cordillera. Our understanding of the age, composition, and petrogenesis of this magmatism thus far, however, has been severely limited by a lack of robust geochronological, geochemical, and isotopic datasets for these rock units. This study fills the need for better temporal and compositional constraints on mainly Early to ‘mid’-Cretaceous (~115–90 Ma) magmatism in the Cassiar Terrane and eastern Selwyn Basin regions of the northern Cordilleran miogeocline. An extensive set of new U–Pb and 40Ar/39Ar age data, together with major and trace element geochemical analyses, radiogenic (Nd, Sr, Pb, and Hf) and stable (S and O) isotopic compositions, and detailed magmatic fluid compositional studies of intrusive rocks in the northern Cordillera are presented here. These results, and the integration of compiled datasets, have led to the identification of several new plutonic suites and sub-suites, as well as modifications of previously defined plutonic suites. Inferences regarding the age and composition of the basement to the northern Cordilleran miogeocline have also been made based on mineralogical, geochemical, and isotopic interpretations of magma sources for crustally derived plutonic suites from across this region. These extensive datasets have also allowed a re-evaluation of previously proposed models for syn- to post-accretionary magmatic episodes, and have resulted in the proposal of modified tectono–magmatic models for the particularly voluminous Early to ‘mid’-Cretaceous magmatic episode. Finally, estimates of magmatic fluid compositions determined from apatite and biotite compositions (with respect to HF, HCl, and H2O) and extended cooling histories indicated by U–Pb and 40Ar/39Ar geochronology, have been used to identify intrusive phases (locally and regionally) that are associated with reduced tungsten skarn mineralization, which comprises some of the most economically important intrusion-related mineralization in the northern Cordillera. Collectively, this work has substantially advanced the current knowledge of syn- to post-accretionary magmatism in the northern Cordillera, and it has broad implications for similar magmatic provinces world-wide. iii PREFACE The research program was initially designed as an MSc project (focusing on the Selwyn Basin field area) by, J. Mortensen, and was initiated in January, 2005. Chapters 4 and 7, and a large proportion of Chapter 2, are the results of this earlier work. Upon my transfer to the PhD program in January, 2007, J. Mortensen and I together designed a research programme that added several new aspects to the original project, including the addition of an extensive stable isotopic dataset. I proposed the reconnaissance-scale zircon composition study presented in Chapter 5, and the detailed magmatic fluid composition study presented in Chapter 6. J. Mortensen suggested adding an additional field area, the Cassiar Terrane, to the scope of the project, which resulted in the data presented in Chapter 3 and a large proportion of Chapter 2. I have been the principal investigator for all field, U–Pb (with the exception of the ID–TIMS samples), and magmatic fluid composition studies. Other data collection was out-sourced and (or) resulted from collaboration with colleagues at the University of British Columbia (UBC) and the research facilities detailed below. I am the primary author of all eight chapters of this dissertation, although each chapter has benefited significantly from editorial and scientific comments from J. Mortensen (and others listed below). All interpretations of the data, however, are mine, as are any errors. CHAPTER 2 – GEOCHRONOLOGICAL AND GEOCHEMICAL CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY PLUTONIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE All samples collected for petrographic examination, geochemical analysis, and U–Pb geochronology were obtained over the course of several field seasons, from 2005 to 2007 (samples beginning with: “KR-05”, “CL-06”, and “KR- 07”). Field assistance was provided by H. Falck and J. Mortensen. I am responsible for collecting the “KR-05” and “KR-07” samples, and preparing and submitting these samples for geochemical work at ALS Chemex (Vancouver), preparing these samples for U–Pb geochronological work. I was present when most of the “CL-06” samples were collected, and was responsible for preparing these samples for geochronological work, although C. Leslie prepared and submitted these samples for geochemical analysis at ALS Chemex. J. Mortensen, R. Friedman, and H. Lin assisted with zircon separation and sample preparation for the above-listed samples by Laser Ablation – Inductively Coupled Plasma – Mass Spectrometry (LA–ICP–MS) geochronological analysis, which I performed between 2006– 2009 at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at UBC. Some of the LA–ICP–MS geochronological data (analyzed in June 2010) were collected by J. Cross at the PCIGR, although I oversaw the analytical work. All sample numbers not beginning with “KR-05”, “CL-06”, and “KR-07” were collected and prepared for geochronological ± geochemical analysis by J. Mortensen, and were analyzed by Isotope Dilution – Thermal Ionization Mass Spectrometry (ID–TIMS) geochronological methods by J. Mortensen at the PCIGR. 40Ar/39Ar geochronological analysis for “KR-05” mafic dyke samples were conducted by Tom Ullrich at the PCIGR. This chapter integrates the geochemical and geochronological data produced by this study with a large compilation of geochronological and geochemical data tabulated in Appendices F.1 and F.2, respectively. Dr. T. Liverton read an early version of this chapter and provided several constructive comments mainly pertaining (but not limited) to the iv Seagull plutonic suite. CHAPTER 3 – ISOTOPIC CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY, EARLY CRETACEOUS TO EOCENE MAGMATISM IN THE CASSIAR TERRANE AND CHAPTER 4 – ISOTOPIC CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY, EARLY TO ‘MID’-CRETACEOUS MAGMATISM IN THE SELWYN BASIN B. Kieffer undertook all aspects of the Sm–Nd and Rb–Sr isotopic analysis at the PCIGR. J. Gabites undertook the lead isotopic analysis at the PCIGR. S. Heffernan allowed me access to his sample material for stable isotope analysis. I prepared samples for sulfur and oxygen isotopic analysis, and travelled to the Nevada Stable Isotope Laboratory in Reno, Nevada, to do the whole rock sulfur separations under the supervision of G. Arehart, S. Poulson, and C. Sladek. S. Poulson performed the sulfur isotope analyses, and G. Arehart performed the oxygen isotope analyses. This chapter integrates the data produced by this study with a large compilation of magmatic and basement isotopic data tabulated in Appendices F.3 and F.4. G. Arehart read a near-final version of Chapter 3 and provided input that clarified and improved some of the isotopic interpretations. T. Liverton also read a near-final version of Chapter 3 and provided several corrections and comments pertaining to the Seagull plutonic suite. CHAPTER 5 – A RECONNAISSANCE-SCALE STUDY OF ZIRCON TRACE ELEMENT AND HAFNIUM ISOTOPIC COMPOSITIONS IN THE EASTERN SELWYN BASIN I selected several representative samples of the plutonic suites in the eastern Selwyn Basin, did a detailed morphological classification of zircon textures, and chose the analysis spots. Zircon compositional analytical work and data reduction was completed at CREAIT Network, Memorial University, by R. Lam and M. Tubrett. R. Lam contributed a draft of the analytical methods for the hafnium isotopic analysis of zircon (Section 5.1.3) and M. Tubrett contributed a draft of the analytical methods for the trace element analysis of zircon (Section 5.1.2). CHAPTER 6 – APPROXIMATING MAGMATIC FLUID COMPOSITIONS OF BARREN AND TUNGSTEN SKARN-ASSOCIATED PLUTONS WITH APATITE AND BIOTITE: CASE STUDIES FROM THE NORTHERN CANADIAN CORDILLERA AND CHAPTER 7 – PETROGENETIC, THERMAL, AND FLUID EVOLUTION OF MAGMATISM IN THE EASTERN SELWYN BASIN: INTEGRATING GEOCHRONOLOGICAL AND MINERAL COMPOSITIONAL DATASETS A near-identical version of Chapter 6 has been published as: Rasmussen, K.L., and Mortensen, J.K., 2013. Magmatic petrogenesis and the evolution of (F:Cl:OH) fluid composition in barren and tungsten skarn-associated plutons using apatite and biotite compositions: Case studies from the northern Canadian Cordillera. Ore Geology Reviews, 50, 118–142. v A very early draft of the data presented in Chapter 7 was published as: Rasmussen, K.L., Mortensen, J.K., Falck, H., and Ullrich T.D., 2007. The potential for intrusion-related mineralization within the South Nahanni River MERA area, Selwyn and Mackenzie Mountains, Northwest Territories. In: Mineral and energy resource assessment of the Greater Nahanni Ecosystem under consideration for the expansion of the Nahanni National Park Reserve, Northwest Territories. Edited by: D.F. Wright, D. Lemkow, J.R. Harris. Geological Survey of Canada, Open File 5344. I selected and prepared apatite and biotite samples for analysis on the Electron-probe Microanalyzer (EPM). I. Coulson and J. Lang offered advice regarding the analysis of apatite. Analysis was done with the assistance and supervision of M. Raudsepp, the Director of the Electron Microbeam/X-ray Diffraction Facility at UBC. M. Raudsepp also contributed drafts of the EPM analytical details in Sections 6.3 and 7.3. I. Coulson and G. Dipple provided assistance with deciphering the thermodynamic calculations. The 40Ar/39Ar geochronological analysis for “KR-05” plutonic ’cooling’ ages presented in Chapter 7 was conducted by Tom Ullrich at the PCIGR. Chapter 6 benefited substantially from comments and recommendations made by the Ore Geology Reviews reviewers (I. Coulson and D. Kontak) and the managing editor (S. Rowins). Various versions of the material presented in Chapter 7 also benefited from reviews by J. Kerswill, L. Ootes, J. Walker, S. McClenaghan, and H. Falck. vi TABLE OF CONTENTS ABSTRACT ................................................................................................................................................................. ii PREFACE ................................................................................................................................................................... iii TABLE OF CONTENTS ........................................................................................................................................... vi LIST OF TABLES .................................................................................................................................................... xvi LIST OF FIGURES ................................................................................................................................................. xvii ACKNOWLEDGEMENTS ..................................................................................................................................... xxi 1 CHAPTER: INTRODUCTION ......................................................................................................................... 23 1.1 NORTHERN CORDILLERAN GEOLOGY ............................................................................................ 24 1.1.1 Nomenclature.. ............................................................................................................................ 24 1.1.2 The Amalgamation and Composition of the Northern Cordillera ............................................... 25 1.1.2.1 Crystalline Basement ................................................................................................. 26 1.1.2.2 Proterozoic Metasedimentary Assemblages .............................................................. 27 1.1.2.3 Paleozoic Sedimentation, Rifting, and Magmatism ................................................... 28 1.1.2.4 Mesozoic to Cenozoic Terrane Accretion ................................................................. 29 1.1.3 Deformation in the Northern Cordilleran Miogeocline ............................................................... 30 1.1.3.1 Neoproterozoic Rifting and Basin Formation ............................................................ 30 1.1.3.2 Mesozoic Thin-Skinned Deformation ....................................................................... 31 1.1.3.3 Cretaceous–Cenozoic Dextral Strike-slip Deformation............................................. 32 1.2 AN OVERVIEW OF SYN- TO POST-ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERA ..................................................................................................................................................... 33 1.2.1 The Early to ‘Mid’-Cretaceous Magmatic Event ........................................................................ 34 1.2.1.1 Early Cretaceous Arc Magmatism ............................................................................. 34 1.2.1.2 Early Cretaceous Back-Arc Magmatism ................................................................... 35 1.2.1.3 ‘Mid’-Cretaceous Post-Arc Magmatism .................................................................... 35 1.2.1.4 Previously Proposed Petrogenetic Models for Early to ‘Mid’-Cretaceous Magmatism .................................................................................................................................. 36 1.2.2 The Late Cretaceous Magmatic Event......................................................................................... 41 1.2.3 The Latest Cretaceous to Eocene Magmatic Event ..................................................................... 43 1.3 RESEARCH OBJECTIVE AND METHODOLOGIES ............................................................................ 44 1.3.1.1 U–Pb and 40Ar/39Ar Geochronology .......................................................................... 45 1.3.1.2 Whole Rock Geochemistry ........................................................................................ 45 1.3.1.3 Radiogenic Isotopes................................................................................................... 46 1.3.1.4 Stable Isotopes ........................................................................................................... 46 1.3.1.5 Zircon Compositions ................................................................................................. 47 1.3.1.6 Magmatic Fluid Compositions .................................................................................. 47 vii 1.4 OUTLINE OF THE DISSERTATION ...................................................................................................... 48 2 CHAPTER: GEOCHRONOLOGICAL AND GEOCHEMICAL CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY MAGMATIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE ....................................................................................................................................................... 39 2.1 INTRODUCTION ..................................................................................................................................... 39 2.2 GEOLOGICAL BACKGROUND ............................................................................................................ 41 2.2.1 Neoproterozoic to Cenozoic Structural Evolution of the Northern Cordilleran Miogeocline ..... 42 2.2.2 Syn- to Post-accretionary Magmatism in the Northern Cordilleran Miogeocline ....................... 43 2.2.2.1 Estimated Depths of Emplacement for Syn- to Post-accretionary Magmatism ......... 44 2.2.2.2 Previously Defined Plutonic Suites in the Cassiar Terrane and Selwyn Basin .......... 45 2.3 METHODS AND RESULTS .................................................................................................................... 45 2.3.1 Geochronology ............................................................................................................................ 45 2.3.1.1 Summary of Geochronological Results ..................................................................... 46 2.3.2 Major and Trace Element Chemistry .......................................................................................... 49 2.3.2.1 Basis of Geochemical Composition and Classification of the Plutonic Suites .......... 49 2.4 CHARACTERIZATION OF PLUTONIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE .................................................................................................................................................. 52 2.4.1 Cassiar Plutonic Suite .................................................................................................................. 53 2.4.2 Anvil Plutonic Suite .................................................................................................................... 55 2.4.3 Hyland Plutonic Suite .................................................................................................................. 56 2.4.4 Seagull Plutonic Suite ................................................................................................................. 57 2.4.5 Tay River Plutonic Suite ............................................................................................................. 59 2.4.5.1 Western Selwyn Basin ............................................................................................... 60 2.4.5.2 Eastern Selwyn Basin ................................................................................................ 61 2.4.6 Tungsten Plutonic Suite............................................................................................................... 52 2.4.7 Mayo Plutonic Suite .................................................................................................................... 64 2.4.8 Tombstone Plutonic Suite ........................................................................................................... 65 2.4.9 Rancheria Plutonic Suite ............................................................................................................. 67 2.4.10 Black River Plutonic Suite .......................................................................................................... 67 2.5 PETROGENESIS OF SYN- TO POST-ACCRETIONARY PLUTONIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE............................................................................................... 68 2.5.1 A Review of the High-K Calk-alkaline Magma Series ............................................................... 68 2.5.1.1 I-Type and S-type Variants of High-K Calc-alkaline Magmas ................................. 70 2.5.1.2 Temperatures and Mechanisms of Crustal Anatexis ................................................. 71 2.5.2 High-K Calc-Alkaline Plutonic Suites in the Northern Cordilleran Miogeocline ....................... 72 2.5.2.1 Cassiar Plutonic Suite: high-K calc-alkaline (to alkaline), S > I-type ....................... 72 2.5.2.2 Anvil and Hyland Plutonic Suites: (medium-K to) high-K calc-alkaline, S > I-type……. .................................................................................................................................. 74 viii 2.5.2.3 Tay River Plutonic Suite: high-K calc-alkaline, I-type ............................................. 75 2.5.2.4 Mayo and Tungsten Plutonic Suites: high-K calc-alkaline to alkali-calcic (with late shoshonitic), I- (± S-) type .......................................................................................................... 77 2.5.3 A Review of the Alkali-Calcic to Alkaline Magma Series .......................................................... 78 2.5.3.1 Mantle and Crustal Melt Sources of Alkali-Calcic to Alkaline Magmas .................. 79 2.5.4 Alkali-Calcic to Alkaline Plutonic Suites in the Northern Cordilleran Miogeocline .................. 80 2.5.4.1 Seagull, Rancheria, and Black River Plutonic Suites: aluminous alkali-calcic (high-K calc-alkaline) ± alkaline, A-type .................................................................................... 81 2.5.4.2 Tombstone Plutonic Suite: aluminous alkali-calcic (high-K calc-alkaline) > peralkaline, mixed A-/I-type ....................................................................................................... 82 2.6 METALLOGENIC ASSOCIATIONS FOR MAGMATISM IN THE NORTHERN CORDILLERAN MIOGEOCLINE .................................................................................................................................................. 83 2.6.1 High-K Calk-Alkaline “S-type” Suites........................................................................................ 84 2.6.2 High-K Calk-Alkaline “I-type” Suites ........................................................................................ 84 2.6.3 Alkali-Calcic to Alkaline “A-type” Suites .................................................................................. 85 2.7 TEMPORAL AND COMPOSITIONAL PATTERNS OF SYN- TO POST-ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERA ...................................................................................... 87 2.7.1 The Early to ‘mid’-Cretaceous (~115-100 Ma) ........................................................................... 87 2.7.1.1 ‘Outboard’ Magmatism ............................................................................................. 87 2.7.1.2 ‘Inboard’ Magmatism ................................................................................................ 88 2.7.2 The Late Cretaceous (~84–67 Ma) .............................................................................................. 89 2.7.2.1 ‘Outboard’ Magmatism ............................................................................................. 89 2.7.2.2 ‘Inboard’ Magmatism ................................................................................................ 90 2.7.3 The Latest Cretaceous to Eocene (~67–47 Ma) .......................................................................... 90 2.7.3.1 ‘Outboard’ Magmatism ............................................................................................. 90 2.7.3.2 ‘Inboard’ Magmatism ................................................................................................ 91 3 CHAPTER: ISOTOPIC CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY, EARLY CRETACEOUS TO EOCENE MAGMATISM IN THE CASSIAR TERRANE .............................................. 152 3.1 INTRODUCTION ................................................................................................................................... 152 3.1.1 Isotopic Datasets as Petrogenetic Indicators ............................................................................. 153 3.2 GEOLOGICAL BACKGROUND .......................................................................................................... 153 3.2.1 Evolution and Composition of Northern Cordilleran Basement ............................................... 153 3.2.1.1 Precambrian Basement to the Cassiar Terrane ........................................................ 154 3.2.1.2 The Assembly of the Cassiar Terrane ...................................................................... 156 3.2.2 Regional and Contact Metamorphism in the Cassiar Terrane ................................................... 158 3.2.2.1 Burial Depths and Estimates of Cretaceous Crustal Thickness in the Cassiar Terrane….. ................................................................................................................................ 159 3.2.3 Syn- to Post-Accretionary Magmatism ..................................................................................... 159 ix 3.2.3.1 Magmatism ‘Outboard’ (west) of the Cassiar Terrane ............................................ 160 3.2.3.2 ‘Inboard’ Magmatism within the Cassiar Terrane ................................................... 161 3.2.4 Relevant Regional Studies on Igneous Rocks in the Cassiar Terrane ....................................... 164 3.2.4.1 The Cassiar Batholith .............................................................................................. 164 3.2.4.2 A Regional Transect ................................................................................................ 165 3.2.4.3 Southwest to South–Central Yukon ........................................................................ 166 3.2.4.4 The Mesozoic Mantle .............................................................................................. 166 3.3 METHODS AND RESULTS .................................................................................................................. 167 3.3.1 Radiogenic Isotopes .................................................................................................................. 167 3.3.1.1 Sm–Nd and Rb–Sr ................................................................................................... 168 3.3.1.2 Pb–Pb ...................................................................................................................... 169 3.3.2 Stable Isotopes........................................................................................................................... 169 3.3.2.1 Oxygen .................................................................................................................... 170 3.3.2.2 Sulfur ....................................................................................................................... 171 3.4 ISOTOPIC COMPOSITIONS AND COMPARISONS .......................................................................... 171 3.4.1 Radiogenic Isotopic Datasets .................................................................................................... 171 3.4.1.1 87Sr/86Sr100 versus εNd100 ......................................................................................... 172 3.4.1.2 147Sm/144Nd versus εNd100 ....................................................................................... 174 3.4.1.3 TDM versus εNd100 .................................................................................................... 175 3.4.1.4 206Pb/204Pb versus 207Pb/204Pb and 208Pb/204Pb .......................................................... 176 3.4.1.5 207Pb/206Pb versus 208Pb/206Pb .................................................................................. 178 3.4.2 Stable Isotopic Datasets ............................................................................................................. 178 3.4.2.1 δ18O ......................................................................................................................... 179 3.4.2.2 δ34S .......................................................................................................................... 181 3.4.3 Isotopic Comparisons to Coeval Intermediate–Felsic Intrusive Rocks in the Northern Cordillera……. .......................................................................................................................... 184 3.4.3.1 The Northern Cordillera .......................................................................................... 184 3.4.3.2 The Southern Cordillera .......................................................................................... 185 3.5 DISCUSSION .......................................................................................................................................... 187 3.5.1 Isotopic and Geochemical Variation ......................................................................................... 187 3.5.1.1 Sm/Nd, La/Sm, and Ba/Nb ...................................................................................... 187 3.5.1.2 Sm/Yb ...................................................................................................................... 190 3.5.1.3 Rb/Sr, Eu/Eu*, and A/CNK ..................................................................................... 191 3.5.2 Sources of Isotopic Variation in the Back-Arc Cassiar Plutonic Suite ...................................... 195 3.5.2.1 Possible Isotopically Juvenile Protoliths ................................................................. 196 3.5.2.2 Possible Isotopically Highly Evolved Protoliths ..................................................... 198 3.5.2.3 Mineralogical ± Isotopic Variation with Time ........................................................ 199 3.5.3 Isotopic Comparisons between the Cassiar Plutonic Suite and Northwest Laurentian Basement……………… ......................................................................................................................... 200 x 3.5.3.1 An Overview of the Isotopic Composition of Northwest Laurentian Basement ..... 201 3.5.3.2 Possible Crustal Magma Sources for the Cassiar Plutonic Suite ............................. 203 3.5.3.3 Implications Regarding the Age and Composition of Cordilleran Basement .......... 205 3.5.4 Sources of Isotopic Variation in the Post-Arc A-type Plutonic Suites ...................................... 206 3.5.4.1 An Evaluation of Possible A-type Magma Sources ................................................. 207 3.5.5 Petrogenesis of Syn- to Post-accretionary Magmatism in the Cassiar Terrane ......................... 210 3.5.5.1 Early Cretaceous (115–104 Ma) .............................................................................. 210 3.5.5.2 ‘Mid’ Cretaceous (103–94 Ma) ............................................................................... 213 3.5.5.3 Late Cretaceous to Eocene (87–75 Ma and 65–47 Ma) ........................................... 214 3.6 CONCLUSIONS ..................................................................................................................................... 215 3.6.1 The Back-Arc Cassiar Plutonic Suite ........................................................................................ 215 3.6.2 The Post-Arc A-type Plutonic Suites ........................................................................................ 217 4 CHAPTER: ISOTOPIC CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY, EARLY TO ‘MID’-CRETACEOUS MAGMATISM IN THE SELWYN BASIN ........................................................... 254 4.1 BACKGROUND ON SYN- TO POST-ACCRETIONARY MAGMATISM ........................................ 254 4.1.1 Magmatism ‘Outboard’ (southwest) of the Selwyn Basin ........................................................ 254 4.1.2 ‘Inboard’ Magmatism within the Selwyn Basin ........................................................................ 255 4.2 METHODS AND RESULTS .................................................................................................................. 258 4.2.1 Radiogenic Isotopes .................................................................................................................. 258 4.2.1.1 Sm–Nd and Rb–Sr ................................................................................................... 259 4.2.1.2 Pb–Pb ...................................................................................................................... 259 4.2.2 Stable Isotopes........................................................................................................................... 260 4.2.2.1 Oxygen .................................................................................................................... 260 4.2.2.2 Sulfur ....................................................................................................................... 261 4.3 DISCUSSION AND INTERPRETATION OF THE ISOTOPIC DATASETS ...................................... 261 4.3.1 Anvil, Hyland, and Tay River Plutonic Suites .......................................................................... 261 4.3.1.1 Radiogenic Isotopic Datasets ................................................................................... 261 4.3.1.2 Stable Isotopic Datasets ........................................................................................... 263 4.3.1.3 Possible Sources of the Anvil, Hyland, and Tay River Suite Magmas .................... 263 4.3.2 Tungsten, Mayo, and Tombstone (TTB) Plutonic Suites .......................................................... 265 4.3.2.1 Radiogenic Isotopic Datasets ................................................................................... 265 4.3.2.2 Stable Isotopic Datasets ........................................................................................... 268 4.3.2.3 Possible Sources of the Tungsten, Mayo, and Tombstone Suite Magmas............... 269 4.4 SUMMARY OF PETROGENETIC INTERPRETATIONS FOR SYN- TO POST-ACCRETIONARY MAGMATISM IN THE SELWYN BASIN ...................................................................................................... 272 4.4.1 A Tectonic Model for Syn- to Post-Accretionary Magmatism in the Selwyn Basin ................. 274 xi 5 CHAPTER: A RECONNAISSANCE-SCALE STUDY OF ZIRCON TRACE ELEMENT AND HAFNIUM ISOTOPIC COMPOSITIONS, EASTERN SELWYN BASIN ....................................................... 302 5.1 INTRODUCTION ................................................................................................................................... 302 5.2 ANALYTICAL METHODS ................................................................................................................... 303 5.2.1 Trace Element Analysis ............................................................................................................. 304 5.2.2 Lu–Hf Isotopic Analysis and Data Reduction ........................................................................... 305 5.3 DISCUSSION OF RESULTS ................................................................................................................. 306 5.3.1 Texture and Morphology of Zircons from Representative Samples .......................................... 306 5.3.2 Characterization of Zircon Trace Element Compositions ......................................................... 310 5.3.2.1 Markers of Magmatic Differentiation in Zircon ...................................................... 312 5.3.2.2 Magmatic and Inherited Zircon Compositional Comparisons ................................. 314 5.3.2.3 Anomalously Light REE Enriched Zircon .............................................................. 314 5.3.3 εHf100 and TDM.. ......................................................................................................................... 315 5.3.3.1 Integration of Lu–Hf isotopic Compositions in Zircon with Other Isotopic Datasets… ................................................................................................................................. 316 5.4 PETROGENETIC IMPLICATIONS ...................................................................................................... 318 5.4.1 Crustal Protolith Compositions ................................................................................................. 318 5.4.2 Magma Mixing .......................................................................................................................... 318 5.5 A COMPARISON OF ZIRCON CRYSTALLIZATION TEMPERATURES ........................................ 319 5.6 SUMMARY OF PRELIMINARY OBSERVATIONS AND POTENTIAL DIRECTIONS FOR FUTURE ZIRCON COMPOSITION WORK ................................................................................................... 322 5.6.1 Recommended Future Zircon Compositional Work .................................................................. 324 6 CHAPTER: APPROXIMATING MAGMATIC FLUID COMPOSITIONS OF BARREN AND TUNGSTEN SKARN-ASSOCIATED PLUTONS WITH APATITE AND BIOTITE: CASE STUDIES FROM THE NORTHERN CANADIAN CORDILLERA ............................................................................................................... 355 6.1 INTRODUCTION ................................................................................................................................... 356 6.1.1 Issues Inherent to Magmatic Fluid Compositional Studies ....................................................... 357 6.2 GEOLOGICAL BACKGROUND .......................................................................................................... 360 6.2.1 Previous Mineral Chemistry Investigations of Intrusion-related Systems ................................ 360 6.2.2 The Northern Cordillera ............................................................................................................ 361 6.2.3 Justification for This Study ....................................................................................................... 363 6.2.4 Description of the Study Plutons ............................................................................................... 364 6.3 ANALYTICAL METHODS ................................................................................................................... 365 6.3.1 Apatite.............. ......................................................................................................................... 365 6.3.2 Biotite............... ......................................................................................................................... 366 6.4 RESULTS ................................................................................................................................................ 366 6.4.1 Apatite Chemistry ..................................................................................................................... 366 6.4.2 Biotite Chemistry ...................................................................................................................... 369 xii 6.5 DISCUSSION .......................................................................................................................................... 370 6.5.1 Magmatic Differentiation from Apatite and Biotite Compositions ........................................... 370 6.5.1.1 Unusually Iron-Rich Biotite .................................................................................... 371 6.5.2 Approximations of Magmatic Fluid Compositions ................................................................... 371 6.5.2.1 The Reciprocal Solid Solution Model ..................................................................... 372 6.5.2.2 Pressure–Temperature Estimates ............................................................................. 373 6.5.2.3 The Application of F:Cl:OH Variations and Activity Ratios .................................. 374 6.5.2.4 Unmineralized Case Studies .................................................................................... 376 6.5.2.5 Mineralized Case Studies ........................................................................................ 379 6.6 IMPLICATIONS ..................................................................................................................................... 381 6.6.1 Magmatic Fluid Compositions of Tungsten-Skarn Systems in the Northern Cordillera ........... 381 6.6.1.1 HCl/HF and H2O/HF Activity Ratios as Tungsten Mineralization Vectors ............ 382 6.6.2 Apatite- versus Biotite-Based Estimates of Magmatic Fluid Composition ............................... 382 6.7 CONCLUSIONS ..................................................................................................................................... 383 6.7.1 Magmatic Petrogenesis.............................................................................................................. 383 6.7.2 Magmatic Fluid Composition .................................................................................................... 384 6.7.2.1 A Model for Magmatic Tungsten Skarn Mineralization.......................................... 384 6.7.2.2 Final Remarks .......................................................................................................... 385 7 CHAPTER: PETROGENETIC, THERMAL, AND FLUID EVOLUTION OF MAGMATISM IN THE EASTERN SELWYN BASIN: INTEGRATING GEOCHRONOLOGICAL AND MINERAL COMPOSITIONAL DATASETS .................................................................................................................................................................. 403 7.1 INTRODUCTION ................................................................................................................................... 403 7.1.1 Thermal Evolution of Magmatic–(Hydrothermal) Systems ...................................................... 403 7.1.1.1 Limits Associated with Dating Magmatic Events ................................................... 406 7.1.2 Magmatic Fluid Compositions from Mineral Chemistry .......................................................... 407 7.1.2.1 Limits Associated with Studies of Magmatic Volatile and Fluid Composition ....... 408 7.1.3 The Importance of Magmatic Petrogenesis and Differentiation ................................................ 410 7.2 GEOLOGICAL BACKGROUND .......................................................................................................... 411 7.2.1 Thermal and Magmatic Fluid Conditions for Intrusion-related Tungsten Skarns ..................... 413 7.2.2 Thermal and Magmatic Fluid Conditions for Intrusion-related Gold Mineralization ............... 414 7.3 ANALYTICAL METHODS ................................................................................................................... 416 7.3.1 40Ar/39Ar Geochronology .......................................................................................................... 417 7.3.2 Mineral Analysis ....................................................................................................................... 418 7.3.2.1 Apatite ..................................................................................................................... 418 7.3.2.2 Biotite ...................................................................................................................... 419 7.4 DISCUSSION .......................................................................................................................................... 419 7.4.1 Magmatic Thermal Timescales: emplacement to sub-solidus conditions ................................. 419 7.4.1.1 Magmatic Thermal Evolution of the Intrusions ....................................................... 420 xiii 7.4.1.2 Reset 40Ar/39Ar Ages ............................................................................................... 421 7.4.2 Mineral Compositions as Monitors of Magmatic Petrogenesis and Metallogenesis ................. 423 7.4.2.1 Apatite Compositions .............................................................................................. 423 7.4.2.2 Biotite Compositions ............................................................................................... 425 7.4.2.3 Variations in Fluorine and Chlorine ± SO3 and Associated Metallogeny ............... 427 7.4.3 Regional Characterization of Magmatic Fluid Compositions ................................................... 428 7.4.3.1 Small Variations in Apatite- and Biotite-based Estimates of Fluid Composition ... 430 7.4.3.2 Significant Variations in Apatite- and Biotite-based Estimates of Fluid Composition .............................................................................................................................. 432 7.4.4 Integrated Magmatic Thermal and Fluid Compositional Datasets ............................................ 433 7.5 CONCLUSIONS ..................................................................................................................................... 434 8 CHAPTER: SUMMARY OF CONCLUSIONS AND A GEODYNAMIC MODEL FOR SYN- TO POST-ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERA ........................................ 461 8.1 AGE, COMPOSITION, MAGMA SOURCES, MECHANISMS OF PARTIAL MELTING, AND METALLOGENY .............................................................................................................................................. 461 8.1.1 Cassiar Terrane .......................................................................................................................... 461 8.1.2 Selwyn Basin.... ......................................................................................................................... 462 8.1.3 Summary of Metallogenic Associations .................................................................................... 465 8.2 MAGMATIC FLUID COMPOSITIONS AND THERMAL HISTORIES IN THE SELWYN BASIN.......... ...................................................................................................................................................... 465 8.3 PLATE INTERACTIONS, DEFORMATION, AND MAGMATISM: A PROPOSED TECTONO– MAGMATIC MODEL FOR SYN- TO POST-ACCRETIONARY IGNEOUS ROCKS IN THE NORTHERN CORDILLERAN MIOGEOCLINE ................................................................................................................... 467 8.3.1 Early Cretaceous (~115–100 Ma) ............................................................................................. 467 8.3.2 ‘Mid’-Cretaceous (~103–94 Ma and ~99–90 Ma) .................................................................... 468 8.3.3 Late Cretaceous (~85–67 Ma) ................................................................................................... 469 8.3.4 Latest Cretaceous to Eocene (~67–47 Ma) ................................................................................ 469 8.4 SUGGESTIONS FOR FUTURE WORK ................................................................................................ 470 8.4.1 The Petrogenesis of Syn- to Post-Accretionary Magmatism ..................................................... 470 8.4.2 Age and Composition of Miogeoclinal Basement ..................................................................... 471 8.4.3 A Final Comment on ‘Mid’-Cretaceous Magmatism in the Selwyn Basin ............................... 472 REFERENCES ........................................................................................................................................................ 480 APPENDICES ......................................................................................................................................................... 517 APPENDIX A - SAMPLE DESCRIPTIONS ................................................................................................... 518 A.1 Geochronological samples ........................................................................................................ 519 xiv A.2 Geochemical samples ................................................................................................................ 533 A.3 Images of geochronological and geochemical samples ............................................................. 536 APPENDIX B – SUPPORTING MATERIALS FOR REPORTED U–PB AGES ........................................... 545 B.1 U–Pb geochronological methods. .............................................................................................. 546 B.2 Zircon separation details for samples analyzed by LA–ICP–MS methods. .............................. 552 B.3 Summary of observed zircon textures from polished grain mounts for samples analyzed by LA–ICP–MS methods. .............................................................................................................. 565 B.4 Representative zircons images. ................................................................................................. 570 B.5 LA–CP–MS age determinations. ............................................................................................... 587 B.6 LA–CP–MS geochronological data. .......................................................................................... 673 B.7 ID-TIMS age determinations. .................................................................................................... 674 B.8 ID–TIMS geochronological data. .............................................................................................. 683 B.9 Summary of Zircon Inheritance. ................................................................................................ 684 APPENDIX C – SUPPORTING MATERIALS FOR REPORTED 40AR/39AR AGES ................................... 691 C.1 40Ar/39Ar plateau plots for plutonic and mafic dyke samples from the eastern Selwyn Basin. . 692 C.2 40Ar/39Ar analysis conditions and individual analyses used for final plateau age calculations. 724 APPENDIX D – GEOCHEMICAL DATASETS ............................................................................................. 725 D.1 ALS Chemex major and trace element geochemical data and relevant calculated geochemical parameters........ ......................................................................................................................... 726 D.2 Accuracy analysis for 2006–2007 samples analyzed at ALS Chemex. ..................................... 738 D.3 Comparison between trace elements analyzed by inductively coupled plasma – mass spectrometry (ICP–MS) methods at ALS Chemex and trace element data determined by high resolution ICP–MS methods at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) for 2005 samples. ....................................................................................................... 739 D.4 Sliding normalization to the high-K calc-alkaline Yenchichi–Telabit Series (after Liégeois et al., 1998).......... .......................................................................................................................... 740 APPENDIX E – SUPPORTING MATERIALS FOR ZIRCON COMPOSITIONS ......................................... 741 E.1 Zircon images and analysis spots. ............................................................................................. 742 E.2 Complete zircon trace element analyses. ................................................................................... 773 E.3 Complete zircon Lu–Hf isotope analyses. ................................................................................. 774 E.4 Details for zircon Lu–Hf isotope calculations. .......................................................................... 775 APPENDIX F – SUPPORTING MATERIALS FOR REPORTED MAGMATIC FLUID COMPOSITIONS .............................................................................................................................................. 779 F.1 Complete compositions, molar calculations, and activity calculations for all apatites used in interpretations and figures in Chapter 6. ................................................................................... 780 F.2 Complete compositions, molar calculations, and activity calculations for all biotites used in interpretations and figures in Chapter 6. ................................................................................... 781 F.3 Complete compositions, molar calculations, and activity calculations for all apatites used in interpretations and figures in Chapter 7. ................................................................................... 782 xv F.4 Complete compositions, molar calculations, and activity calculations for all biotites used in interpretations and figures in Chapter 7. ................................................................................... 783 APPENDIX G – DATA COMPILATIONS ...................................................................................................... 784 G.1 Geochronological data compilation for Cretaceous to Eocene magmatic rocks. ...................... 785 G.2 Geochemical data compilation for Cretaceous to Recent magmatic rocks. ............................... 786 G.3 Isotopic data compilation for Cretaceous to Recent magmatic rocks. ....................................... 787 G.4 Isotopic data compilation for northwest Laurentian basement rocks, and accreted pericratonic and oceanic terrane rocks. ......................................................................................................... 788 xvi LIST OF TABLES Table 1.1. Basic characteristics of syn- to post-accretionary, Early Cretaceous to Eocene magmatic suites discussed in the text. .......................................................................................................................... 36 Table 2.1. Summary of basic characteristics of for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. .............................................................................................................135 Table 2.2. Final U–Pb ages of crystallization for felsic plutonic and dykes phases, and 40Ar/39Ar age determinations for mafic dykes in the study area. .....................................................................................137 Table 2.3. Detailed summary of patterns of magmatism, macroscopic and microscopic features, and references for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. ...........141 Table 2.4. Details of geochemical discrimination plots used in Figure 2.4-2.13. .....................................145 Table 2.5. Summary of geochemical observations in Figures 2.4-2.14 for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. ..........................................................................147 Table 3.1. Basic sample descriptions, crystallization ages, trace element ratios, and geochemical parameters for all samples that underwent radiogenic and (or) stable isotopic analysis. ..........................251 Table 3.2. Radiogenic and stable isotopic data generated by this study for magmatism in the Cassiar Terrane.......................................................................................................................................................252 Table 4.1. Basic sample descriptions and crystallization ages for all samples that underwent radiogenic and (or) stable isotopic analysis. ..............................................................................................296 Table 4.2. Radiogenic and stable isotopic data generated by this study for magmatism in the Selwyn Basin. .........................................................................................................................................................297 Table 4.3. Stable isotopic data generated by this study for samples from the eastern Selwyn Basin from Heffernan (2004). .............................................................................................................................301 Table 5.1. Summary of relevant features of representative samples chosen for zircon compositional analysis. .....................................................................................................................................................344 Table 5.2. Results of zircon geochemical analysis. ...................................................................................345 Table 5.3. Results of zircon hafnium isotopic analysis. ............................................................................353 Table 6.1. Abbreviated descriptions for samples analyzed by this study. .................................................395 Table 6.2. Representative apatite compositions and calculated magmatic fluid activity ratios. ...............396 Table 6.3. Representative biotite compositions and calculated Fe#, ∑ Al, and magmatic fluid activity ratios. .........................................................................................................................................................399 Table 6.4. Apatite saturation temperatures for samples with available whole rock chemistry. ................402 Table 7.1. Sample information and final U-Pb crystallization ages for plutons in the eastern Selwyn Basin. .........................................................................................................................................................451 Table 7.2. 40 Ar/ 39 Ar ages and calculated cooling periods. .........................................................................452 Table 7.3. Average apatite compositions for each sample. .......................................................................454 Table 7.4. Average biotite compositions for each sample. ........................................................................457 xvii LIST OF FIGURES Figure 1.1. Basic elements of northern Cordilleran geology. ...................................................................... 29 Figure 1.2. Approximate surface and sub-surface distribution of upper crustal rocks and Precambrian Laurentian basement in the northern Cordillera after the restoration of dextral Cenozoic displacement along major transpressional structures. ....................................................................................................... 31 Figure 1.3. A schematic cross-section of facing lower plate and upper plate rifted margins. ..................... 33 Figure 1.4. Syn- to post-accretionary magmatism and important mineral occurrences and deposits in the northern Cordillera. ............................................................................................................................... 34 Figure 2.1. Basic elements of northern Cordilleran geology, and the distribution of Early Cretaceous to Eocene syn- to post-accretionary igneous rocks in allochthonous and autochthonous Laurentian rocks. ........................................................................................................................................................... 92 Figure 2.2. Syn- to post-accretionary, Early Cretaceous to Eocene magmatism in the northern Cordillera. .................................................................................................................................................... 94 Figure 2.3. The distribution of syn- to post-accretionary plutons and associated volcanic rocks in the study area. .................................................................................................................................................... 96 Figure 2.4. Geochemical variation diagrams for the Cassiar plutonic suite. ............................................... 99 Figure 2.5. Geochemical variation diagrams for the Anvil and Hyland plutonic suites............................ 102 Figure 2.6. Geochemical variation diagrams for the Seagull plutonic suite. ............................................. 105 Figure 2.7. Geochemical variation diagrams for the Tay River plutonic and South Fork volcanic suites in the western Selwyn Basin...................................................................................................................... 108 Figure 2.8. Geochemical variation diagrams for the Tay River plutonic suite and the Coal River sub- suite in the eastern Selwyn Basin. ............................................................................................................. 111 Figure 2.9. Geochemical variation diagrams for the Tungsten plutonic suite. .......................................... 114 Figure 2.10. Geochemical variation diagrams for the Mayo plutonic suite. ............................................. 117 Figure 2.11. Geochemical variation diagrams for the Tombstone plutonic suite. ..................................... 120 Figure 2.12. Geochemical variation diagrams for the Rancheria plutonic suite. ...................................... 123 Figure 2.13. Geochemical variation diagrams for selected Paleocene to Eocene, plutonic and volcanic magmatism, including the Black River suite. ............................................................................................ 126 Figure 2.14. Comparison of the effects of differentiation on SNY and SNX values for different plutonic suites. ........................................................................................................................................... 129 Figure 2.15. Discrimination plots for niobium-richer continental (crustal or sub-crustal) and niobium- poorer asthenospheric (OIB-like) sources for the A-type plutonic suites. ................................................ 130 Figure 2.16. Distribution of intrusion-related and probable intrusion-related mineral occurrences (showings, prospects, drilled prospects, and producers) in the study area. ............................................... 131 Figure 2.17. Basement to the western margin of North America after restoration of dextral strike-slip displacements along major structures, with the approximate area of magmatic suites and occurrences located. ...................................................................................................................................................... 133 xviii Figure 3.1. Basic elements of northern Cordilleran geology. .................................................................... 219 Figure 3.2. Approximate surface and sub-surface distribution of upper crustal rocks and Precambrian Laurentian basement in the northern Cordillera after the restoration of dextral Cenozoic displacement along major transpressional structures. ..................................................................................................... 221 Figure 3.3. Distribution of Early Cretaceous to Eocene, syn- to post-accretionary, magmatism in the northern Cordillera. ................................................................................................................................... 223 Figure 3.4. Distribution of syn- to post-accretionary plutonic suites in the Cassiar Terrane study area. .. 225 Figure 3.5. Summary of isotopic data for each plutonic suite in the Cassiar Terrane. .............................. 226 Figure 3.6. Neodymium and strontium isotopic compositions of plutonic suites in the Cassiar Terrane. 228 Figure 3.7. Lead isotopic compositions of plutonic suites in the Cassiar Terrane. ................................... 230 Figure 3.8. Stable isotopic compositions of plutonic suites in the Cassiar Terrane. ................................. 232 Figure 3.9. Stable and radiogenic isotopic comparisons for plutonic suites in the Cassiar Terrane. ........ 234 Figure 3.10. Expanded 87 Sr/ 86 Sr100 versus δ 18O and δ34S with available basement compositions and ranges for plutonic suites in the Cassiar Terrane. ...................................................................................... 236 Figure 3.11. Comparison of lead isotopic comparisons for plutonic suites in the Cassiar Terrane and similar coeval magmatism in the northern and southern Cordillera. ......................................................... 238 Figure 3.12. Comparison of neodymium and strontium isotopic comparisons for plutonic suites in the Cassiar Terrane and similar coeval magmatism in the northern and southern Cordillera. ........................ 240 Figure 3.13. Whole rock immobile element ratios and isotopic comparisons for plutonic suites in the Cassiar Terrane. ......................................................................................................................................... 242 Figure 3.14. Whole rock geochemical parameters and isotopic comparisons for plutonic suites in the Cassiar Terrane. ......................................................................................................................................... 245 Figure 3.15. Summary of εNd and 87Sr/86Sr for plutonic suites, proposed magma sources in the Cassiar Terrane, and northwest Laurentian basement rocks. .................................................................... 246 Figure 3.16. Comparison of TDM versus εNd100 for northwest Laurentian basement and syn- to post- accretionary magmatism in the northern Cordillera. ................................................................................. 248 Figure 4.1. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous plutonic suites in the northern Cordillera. ................................................................................................................................... 277 Figure 4.2. Distribution of syn- to post-accretionary plutonic suites in the Selwyn Basin study area. ..... 278 Figure 4.3. Summary of isotopic data for each plutonic suite in the Selwyn Basin. ................................. 279 Figure 4.4. Neodymium and strontium isotopic compositions of the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin at 100 Ma. ......................................................................................... 281 Figure 4.5. Lead isotopic compositions of the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. ............................................................................................................................................ 283 Figure 4.6. Stable isotopic compositions of the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. ............................................................................................................................................ 285 Figure 4.7. Strontium and stable isotopic comparisons and available basement compositions and ranges for the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. .................................. 286 xix Figure 4.8. Comparison of TDM versus εNd100 for northwest Laurentian basement and the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. .................................................................... 288 Figure 4.9. Neodymium and strontium isotopic compositions of the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin at 100 Ma. ......................................................................................... 289 Figure 4.10. Lead isotopic compositions of the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. ............................................................................................................................................ 291 Figure 4.11. Stable isotopic compositions of the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. ...................................................................................................................................... 293 Figure 4.12. Strontium and stable isotopic comparisons and available basement compositions and ranges for the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. ............................. 294 Figure 4.13. Comparison of TDM versus εNd100 for northwest Laurentian basement and the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. ..................................................................... 295 Figure 5.1. Distribution of Early to ‘mid’-Cretaceous, syn- to post-accretionary plutonic suites in the northern Cordillera. ................................................................................................................................... 325 Figure 5.2. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous magmatism in the eastern Selwyn Basin study area. .............................................................................................................. 326 Figure 5.3. The link between zircon morphology, composition, and temperature of crystallization. ....... 327 Figure 5.4. Classification of internal zircon morphologies. ...................................................................... 328 Figure 5.5. Chondrite-normalized multi-rare earth element plots of zircon compositions. ....................... 331 Figure 5.6. Trace elemental variation related to magmatic evolution. ...................................................... 333 Figure 5.7. Trace element ratios related to magmatic evolution. .............................................................. 335 Figure 5.8. Magnitude of chondrite-normalized europium and cerium anomalies. .................................. 337 Figure 5.9. Comparison of trace element ratios in zircon rims and cores. ................................................ 338 Figure 5.10. Hafnium isotopic composition and model ages in zircon. .................................................... 340 Figure 5.11. Radiogenic and stable isotopic compositions versus hafnium isotopic compositions in zircon. ........................................................................................................................................................ 341 Fig. 5.12. Comparison of depleted mantle model ages estimated from whole rock neodymium isotopic compositions and zircon hafnium isotopic compositions .......................................................................... 342 Figure 5.13. Comparison of zircon crystallization temperatures estimated from zircon morphology, whole rock chemistry (zircon saturation temperature), magmatic zircon composition (Ti–in–zircon). ... 343 Figure 6.1. Distribution of Early to ‘mid’-Cretaceous, syn- to post-accretionary plutonic suites in the northern Cordillera. ................................................................................................................................... 386 Figure 6.2. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous magmatism, and tungsten deposits and occurrences in NTS map sheets 105I, 105H, 095L, and 095E. .............................. 387 Figure 6.3. Apatite chemistry. ................................................................................................................... 388 Figure 6.4. Biotite chemistry. .................................................................................................................... 390 xx Figure 6.5. Linear relationship between log[aHOH/aHF] values from apatite and (uncorrected) biotite datasets. ..................................................................................................................................................... 391 Figure 6.6. Log(aHCl/aHF) and log(aHF/aHOH) for apatite and biotite. ......................................................... 392 Figure 7.1. Distribution of Early to ‘mid’-Cretaceous, syn- to post-accretionary plutonic suites in the northern Cordillera. ................................................................................................................................... 436 Figure 7.2. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous magmatism in the eastern Selwyn Basin study area. .............................................................................................................. 437 Figure 7.3. Logarithmic molar concentrations and ratios of fluorine, chlorine, and hydroxyl in the hydroxyl site of apatite and biotite, versus differentiation index. ............................................................. 439 Figure 7.4. Apatite chemistry. ................................................................................................................... 440 Figure 7.5. Biotite chemistry. .................................................................................................................... 442 Figure 7.6. Summary of cooling periods. .................................................................................................. 444 Figure 7.7. Fluorine, chlorine, (and SO3) concentrations versus differentiation index. ............................ 445 Figure 7.8. Log[aHCl/aHF] and log[aHOH/aHF] estimates for the magmatic fluid composition determined from apatite and biotite (red symbols) compositions, versus differentiation index................................... 448 Figure 7.9. Corrected aHOH...........................................................................................................450 Figure 8.1. Tectonic setting of syn- to post-accretionary magmatism in the northern Cordillera......................................................................................................................................474 xxi ACKNOWLEDGEMENTS So many people and organizations went into making this project a success. First, funding for the sample collection and analyses for Selwyn Basin study area was provided by Parks Canada, the Northwest Territories Geoscience Office (NTGO), and a National Sciences and Engineering Research Council (NSERC) of Canada Discovery Grant to J. Mortensen. Funding for sample collection and analyses for the Cassiar Terrane study area was provided by the Yukon Geological Survey (YGS), the Society of Economic Geologists (SEG) Canada Foundation, and a NSERC Discovery Grant to J. Mortensen. Sample collection in the Selwyn Basin was also kindly facilitated by North American Tungsten Corporation Ltd. and Selwyn Resources Ltd. Annual travel to regional and international conferences to present the results of my research were mainly funded by a NSERC Discovery Grant to J. Mortensen; however, funding from the NTGO also paid for several conference-related trips to Yellowknife, NWT. A 3-year NSERC PGS Scholarship, a Hugh E. McKinstry Student Research Award (SEG) and an Egil H. Lorntzen Scholarship (UBC), and 4 years of UBC PhD Tuition Fee Awards enabled me to focus on my research and not worry about paying the Hydro bill. People at UBC who helped with my sample preparation and data acquisition include: J. Gabites, B. Kieffer, T. Ullrich, R. Friedman, H. Lin, and M. Raudsepp. M. Raudsepp, in particular, spent many days helping me collect data, while regaling me with his thoughts the sub-standard quality of today’s graduate students and the utility (or lack thereof) of my apatite analyses. D. Weis is also thanked for always finding time to assist with interpreting and trouble-shooting my radiogenic isotopic datasets, and I. Coulson, J. Lang, G. Dipple, and L. Groat were all instrumental in advising on and interpreting the magmatic fluid compositional data. A. Toma from the Mineral Deposits Research Unit (MDRU) has been a constant source of help in his role as the Resource Centre Coordinator. G. Arehart, S. Poulson, and C. Sladek at the Nevada Stable Isotope Laboratory were a great source of help during my time in Reno and made me feel welcome in their lab. M. Tubrett and R. Lam (CREAIT Network; Memorial University) quickly analyzed my samples despite the last-minute submission, and patiently answered my questions regarding the resulting data. S. Heffernan kindly allowed me access to his sample material so that I could expand my stable isotope dataset. C. Leslie also donated 8 samples, his field notes, and geochemical data. From the YGS, D. Murphy, M. Colpron, G. Abbott, and L. Pigage were an invaluable source of knowledge on northern Cordilleran geology. S. Gordey and C. Roots from the Geological Survey of Canada (GSC) were also a fantastic source of information. M. Colpron, L. Pigage, and S. Gordey also kindly shared their unpublished data with me and made themselves available for my questions. My peers, R. Staples (Simon Fraser University) and W. Cielowitch (UBC), provided many interesting geological discussions and patiently answered my queries regarding the direction and results of their research. Those who read portions of this thesis prior to its submission and provided feedback in various forms include: D. Kontak (Laurentian University), I. Coulson (University of Regina), S. Rowins (British Columbia Geological Survey), T. Liverton, G. Arehart (NSIL), M. Colpron (YGS), J. Kerswill (GSC), J. Scoates (UBC), S. Cairns (NTGO), H. Falck (NTGO). L. Ootes (NTGO), J. Walker, and S. McClenaghan also provided feedback on preliminary datasets in their role as editors of the Gangue Newsletter, Mineral Deposits Division, Geological Association of Canada. My supervisory committee, G. Dipple and C. Hart, were a source of much help and discussion over the years. Rebecca Trainor, the Director of Student Academic Services (Faculty of Graduate Studies, UBC), has reserved judgement and instead provided thoughtful and encouraging support through the last two years of my program. My friend and mentor, Hendrik Falck, was an excellent field ‘assistant’ over the course of two summers, including one previous to the start of this project. In addition to introducing me to field work and sample collection, Hendrik taught me everything I know about operating safely in the field and mine environments, particularly how to work in xxii the mountains with helicopter support. He also had a great playlist that kept us awake and humming, even singing, along on many long drives. My supervisor, Jim, always maintained an upbeat attitude and was a constant source of stimulating conversation, encouragement, and help. Although some may argue putting two data-hungry geologists together on any project is a bad idea, I still cannot believe how lucky I have been to have such an understanding and knowledgeable supervisor that allowed me to work independently and with such few limitations. Jim not only allowed, but also encouraged me to work in the summers to gain real-world experience and to add to my knowledge of Cordilleran geology. This has obvious benefits for the writing of this dissertation, but it has also helped prepare me for the next step in my geological career, which what a great supervisor should do. Many of my peers at UBC and several close friends made even the hardest days at school tolerable. My former office-mates and co-hosts of the legendary Office 059 wine tastings, Dianne, Amber, and Victoria, have become some of my closest friends. The ‘original’ MDRU crowd—namely Karie, Nick and Kylie, Emma, Lyle, Al, Rummy, Janina, Scheel, and Stefan—have provided me with a wealth of happy memories. Similarly, Luke, Chris, Reza, and Tyler from Jim’s research group have been a great source friendship and information. Elspeth, Inês, and Tashia have helped me in innumerable ways. I can’t thank those of you (including Ralf and Tera, and Laurel and Jeff) that also shared their homes with me enough, making it possible for me to create a life up north over the last few years and still visit Vancouver to work on this dissertation. Lesley, Stacia, and Tamara: thank you for being good friends despite neglect on my part. Finally, I also owe a great deal to Lizzie, who endeavoured to read the vast majority of this dissertation before it was submitted for external examination. My brother Will, my closest friend Amanda, and my partner Larry have been a huge source of emotional and physical support throughout this process, and have never doubted my abilities to complete this dissertation. My parents, Roxanna and Knud, have provided encouragement from the start of my academic career and have done everything in their power to help me succeed, especially in the last couple years when morale and funding were both low. I thank them all for their unwavering support. 1 CHAPTER: INTRODUCTION Syn- to post-accretionary magmatism forms the most widespread and voluminous, intermediate to felsic magmatic event in the northern Cordillera. Most of this magmatism occurred during the Early to ‘mid’-Cretaceous1, and in the Late Cretaceous (~84–67 Ma), although scattered occurrences of latest Cretaceous to Eocene (~67–47 Ma) are also present (e.g., Nelson and Colpron, 2007; Chapter 2). Many localized studies exist on Cretaceous and younger intrusions, or suites of intrusive and volcanic rocks throughout the northern Cordillera (e.g., Anderson, 1983; Gordey and Anderson, 1993; Driver et al., 2000; Lang, 2001; Heffernan, 2004; Mair et al., 2006b, 2011; Dilworth et al., 2007; Morris and Creaser, 2008). Only a few, however, have attempted to compartmentalize these igneous rocks from a regional perspective into distinct temporally, geographically, and compositionally defined events (Anderson, 1988; Mortensen et al., 2000; Hart et al., 2004a). Despite this work to categorize the age, composition, and petrogenesis of Cretaceous (and younger) intrusions and volcanic rocks, interpretations have been limited by inadequate geochronological controls in certain key areas (e.g., the Dawson Range, Cassiar Platform, and Selwyn Basin). The absence of associated mafic phases with much of the syn- to post-accretionary magmatism has also limited the utility of geochemical interpretations (e.g., tectonic discrimination diagrams; Pearce et al., 1984), and large isotopic datasets have generally been lacking. Where regional compositional datasets do exist, the timing of magmatic events is commonly poorly constrained (e.g., Driver et al., 2000; Morris and Creaser, 2008). The restrictions imposed by the rarity of robust regional age constraints and isotopic datasets, and the relatively felsic magma compositions have hindered previous petrogenetic interpretations of the syn- to post-accretionary magmatic events. Despite the economic significance attributed to syn- to post-accretionary magmatism in the northern Cordillera (e.g., Dawson and Dick, 1978; Dick, 1979; Hart et al., 2004a), many aspects concerning its origin remain unresolved. This study contributes new information on the age, and the chemical and isotopic compositions of syn- to post- accretionary magmatism. This study focuses on igneous rocks emplaced in the Cassiar Platform and Selwyn Basin regions, from south–central Yukon to the southwestern Northwest Territories. The geochronological results of this study, and compiled U–Pb geochronological datasets (some of which have only become available in recent years; e.g., Mortensen and Hart, 2010, unpub. data; Pigage et al., 2011; Gordey, in press), have led to a better definition of the timing of syn- to post-accretionary magmatic events across the northern Cordillera. Furthermore, in the realm of ‘granite’ research in general, a better understanding of intermediate to felsic magmatic provinces has started to change previously held views on how, and from where, these magmas are derived (e.g., Clemens and Stevens, 2012). Thus, the integration of extensive geochemical and isotopic datasets (as well as a compilation of pre-existing data) with the better age constraints have produced new ideas (and questions!) on the origin magma sources and mechanisms of magma generation for much of the Cretaceous and younger igneous rocks examined here. The timing, composition, potential melt sources, mechanisms of magma generation, and tectonic setting of syn- to post- accretionary in the northern Cordillera are, thus, reconsidered from new perspectives throughout this thesis. This introductory chapter has been sub-divided into four primary sections. Section 1.1 presents an overview of 1 The term ‘mid’-Cretaceous refers to the earliest Late Cretaceous (~100–90 Ma). 1 northern Cordilleran geology—with emphasis on aspects that are relevant to syn- to post-accretionary magmatism— to provide an introduction to the geological background to subsequent chapters. Previous work on Cretaceous and selected younger magmatic episodes in the northern Cordilleran miogeocline are then reviewed in Section 1.2, and pros and cons of previously proposed petrogenetic models are discussed. In Section 1.3, the main objectives of the dissertation are outlined, and the various methods that have been employed to address these questions are introduced. Finally, in Section 1.4 a brief overview of the dissertation is presented with a summation of the specific research directions taken by each chapter. 1.1 NORTHERN CORDILLERAN GEOLOGY 1.1.1 Nomenclature The term “northern Cordillera” is used throughout this dissertation to refer to the region west of the eastern extent of Cordilleran deformation in northwest Canada. For the purposes of this thesis, this region is considered to comprise most of Yukon, together with the southwest Northwest Territories, northern British Columbia, and geologically similar portions of east–central Alaska (Fig. 1.1a). The northern Cordillera is an amalgamation of several main entities (Fig. 1.1a). From east to west, these are: • North American or Laurentian basement comprising predominantly of Paleoproterozoic to Paleozoic, rift- related to miogeoclinal supracrustal rocks that were deposited along the western margin of northwest Laurentia; • pericratonic terranes with Laurentian affinities (i.e., Laurentian basement), displaced from their original position along the western margin of Laurentia by rifting and accretion, and (or) by strike-slip faulting; and • the accreted Insular, Arctic, and Pacific terranes of exotic (non-Laurentian) origin. Fragments of terranes representing the former oceanic basins that separated Laurentia, pericratonic, and exotic terranes are also preserved throughout the Cordillera (Fig. 1.1a). Additional terms referring to the tectonostratigraphic sub-division of the Laurentian and peri-Laurentian terranes that host the syn- to post-accretionary magmatic rocks are also used throughout this dissertation. The northern Cordilleran miogeocline encompasses all exposed sedimentary rocks of North American affinity (Fig. 1.1b; e.g., Dietz and Holden, 1966), although the term “miogeocline” is not an appropriate descriptor for late Neoproterozoic and Devono–Mississippian rift-related packages, as well as Permo–Jurassic rocks deposited during accretionary events. It is used here, however, to describe the basement to magmatic events that is hosted mainly in rocks of autochthonous Laurentian affinity as opposed to accreted pericratonic, oceanic, and exotic terranes. Although several tectonostratigraphic elements comprise the northern Cordilleran miogeocline, the two main miogeoclinal or Laurentian hosts of syn- to post-accretionary magmatism are: • the Selwyn Basin, located northeast of the Tintina Fault in Yukon and the westernmost Northwest Territories (Fig. 1.1a); and • the Cassiar Platform, located southwest of the Tintina Fault in Yukon and British Columbia (Fig. 1.1a). 2 The “Cassiar Terrane” is a term traditionally used to refer to the Cassiar Platform, which was translated north during Cenozoic dextral strike-slip displacement along the Tintina Fault (e.g., Gabrielse et al., 2006). Here, in addition to the displaced rocks of the Cassiar Platform, this designation also includes three tectonostratigraphic entities located between the Teslin Tectonic Zone and the Tintina fault that also host syn- to post-accretionary magmatic rocks. These are the: • eastern portions of the Yukon–Tanana terrane, mainly underlain by the Snowcap Assemblage (of Laurentian affinity; Piercey and Colpron, 2009), and located west of the Cassiar Platform (Fig. 1.1a), • the Yukon–Tanana terrane inlier that, although it contains mainly arc-related volcano-sedimentary rocks (e.g., Murphy et al., 2001, 2002), forms only a thin crustal slice obducted onto miogeoclinal rocks (e.g., Cook et al., 2004), and • the Klinkit Assemblage, part of a pericratonic Late Mississippian to Early Permian arc developed on rifted Laurentian crust (Simard et al., 2003), now positioned between the Cassiar Platform and Yukon–Tanana terrane in south–central Yukon (not distinguished from the Yukon–Tanana terrane in Fig. 1.1a). These three components that form the Cassiar Terrane are pericratonic entities that have been displaced as one cohesive block from their original location along the Laurentian margin. The western Cassiar Terrane is bound by the mylonitic Teslin Tectonic Zone (Fig. 1.1a): a high pressure, east–west-trending ductile shear zone that experienced lower pressure, dextral movement in the late Early Cretaceous (Gabrielse, 1991; Gabrielse et al., 2006; Colpron et al., 2011). The Cassiar Terrane and Selwyn Basin are separated by the Tintina Fault, along which about 430 km of latest Cretaceous to early Tertiary dextral strike-slip offset movement occurred, causing the current juxtaposition of the Cassiar Terrane with the western Selwyn Basin and an inlier of eastern Yukon–Tanana terrane (Fig. 2.1a; Gabrielse et al., 2006). In Yukon, the main pericratonic entity consists primarily of the Yukon–Tanana terrane, although Quesnellia and Stikinia become more important basement components in British Columbia (e.g., Nelson and Colpron, 2007). The boundaries between these pericratonic terranes have not been fully resolved, largely because of their shared allochthonous Laurentian basement and similar overlap and island arc assemblages (e.g., Gabrielse and Yorath, 1991). For simplicity, therefore, these three pericratonic terranes have not been distinguished in the figures and will be collectively referred to as the Yukon–Tanana terrane throughout this dissertation. The accreted pericratonic terranes and oceanic rocks of the Slide Mountain and Cache Creek terranes collectively form the Intermontane Superterrane (e.g., Gabrielse and Yorath, 1991; Nelson and Colpron, 2007; Fig. 1.1). The Intermontane Superterrane is bordered on the west by the Insular Superterrane, as well as by several Arctic and Pacific terranes. These exotic accreted terranes do not host any of the syn- to post-accretionary igneous rocks examined here and are not discussed further (e.g., Fig. 1.1). 1.1.2 The Amalgamation and Composition of the Northern Cordillera The tectonic evolution of the entire Cordillera was complicated by multiple rifting events, the accretion of several terranes, and relatively late but large displacements along dextral strike-slip faults (Fig. 1.1a; e.g., Gabrielse et al., 3 2006). Crystalline basement rocks are not exposed, and seismic studies have not been able to conclusively distinguish between crystalline basement, mantle, and overlying strata. Only the extent of Neoproterozoic and younger sedimentary rocks within the first upper 5–7 km of the crust have been well-constrained by surface mapping and stratigraphic studies (Fig. 1.1b). As a result, many aspects concerning the evolution of the western margin of North America are not well understood. However, some background on the age and composition of Cordilleran basement is justified in order to better understand the possible age and composition of magma sources for the syn- to post-accretionary magmatic events, and the mechanisms needed to induce anatexis in this crust (and (or) the underlying mantle). Despite several outstanding issues, the following synthesis of several reviews of northern Cordilleran geology (Griffith, 1991; Gordey and Anderson, 1993; Plafker and Berg, 1994; Thorkelson et al., 2005; Colpron et al., 2006; Nelson and Colpron, 2007) is generally accepted. 1.1.2.1 Crystalline Basement Little is known about the age and composition of crystalline basement in the northern Cordillera, including where it may have been (or still is) present. Despite its lack of exposure, crystalline basement is inferred to underlie much of the northern Cordillera. Southwest of the Tintina Fault, interpretations of seismic studies suggest that Paleoproterozoic or older crystalline basement tapers to the west and terminates halfway between the Teslin Fault and the western margin of the Cassiar Platform in Yukon (Fig. 1.1a; Cook et al., 2004). In northern British Columbia, this basement is thought to end near the Kutcho Fault (Fig. 1.1a; Cook et al., 2004), but may extend as far west as the Stikine River (Evenchick et al., 2005). The location of the inferred boundary between crystalline basement and overlying metasedimentary rocks is also poorly constrained. Crustal scale geophysical cross-sections across the northern Cordillera have been interpreted to indicate the presence of both thin (~5–10 km; Cook et al., 2004) and thick (>20 km; Evenchick et al., 2005) crystalline basement northeast of the Tintina Fault. It is also possible that the transition between crystalline basement and the overlying rocks has been obscured by overprinting high-grade metamorphism (Snyder et al., 2009) and now lies below the MOHO (i.e., at depths greater than 33–36 km; Clowes et al., 2005). Heat flow and aeromagnetic studies suggest that the inferred crystalline basement, or Nahanni terrane, has low gravity (Cook et al., 1991, 1999) and magnetic (Pilkington et al., 2000; Aspler et al., 2003) geophysical character. Gneissic rocks exhumed in metamorphic core complexes in southern British Columbia (e.g., the Malton Gneiss Complex; Chamberlain, 1983) and undeformed ca. 1.85 Ga granitic rocks intercepted in drill core from the Fort Simpson magnetic high (Fig. 1.2b; Villeneuve et al., 1991; Ross et al., 2000) provide the closest temporal and compositional proxies for crystalline basement below the northern Cordillera. This basement, commonly referred to as the “Nahanni terrane” (Fig. 1.2b; Hoffman, 1989), is thought to comprise a unique accreted terrane similar to other Paleoproterozoic domains in northwest Laurentia (Fig. 1.2b; e.g., Hoffman, 1989)—and possibly similar to exhumed basement in southeast British Columbia. It could also comprise attenuated crust of the Fort Simpson magnetic high (Cook et al., 1999), which itself is interpreted to be a Paleoproterozoic arc built on the Nahanni terrane (Hoffman, 1989; Villeneuve et al, 1991). If this interpretation is correct, then depleted mantle model ages for granitic rocks in the Fort Simpson magnetic high would suggest an age between 1.85 Ga and 2.45 Ga for the 4 Nahanni terrane (Villeneuve et al, 1991). Alternatively, it has been proposed that crystalline basement in the northern Cordillera may have comprised oceanic crust, or a transitional continental–oceanic crust (Snyder et al., 2002, 2009), although there is no geophysical evidence for such crust below most of the northern Cordillera today (e.g., Clowes et al., 2005). Finally, there is also the possibility that the Nahanni terrane is as young as ~1.2 Ga (i.e., Grenvillian; e.g., Aspler et al., 2003), as there are multiple lines of evidence for unexposed crust or lithosphere of this age in the northern Cordillera (e.g., Jefferson and Parrish, 1989; Gehrels et al., 1991; Parrish and Reichenback, 1991; Rainbird et al., 1997; Gehrels and Ross, 1998; Mortensen and Colpron, 1998; Peslier et al., 2000; Ross et al., 2005; Lemieux et al., 2007; Leslie, 2009). Whether this Grenvillian signature represents a unique accreted terrane, or reflects tectonic activity of this age that has been superimposed on pre-existing Cordilleran basement, is uncertain (see Milidragovic et al., 2011). 1.1.2.2 Proterozoic Metasedimentary Assemblages Overlying the inferred crystalline basement, a westward tapering wedge of geophysically ‘slow’ and resistive material makes up most of the middle to upper northern Cordilleran crust (e.g., Clowes et al., 2005). These rocks are interpreted to comprise metasedimentary rocks of mainly Mesoproterozoic age (e.g., Cook et al., 2004). Thick successions of siliciclastic and calcareous supracrustal rocks were deposited during major extensional and (or) rifting events, occurring from late Paleoproterozoic to late Neoproterozoic time along the northwest margin of Laurentia (Fig. 1.2b; Aitken and McMechan, 1991; Gabrielse and Campbell, 1991; Souther, 1991; Woodsworth et al., 1991). The development of more than one epicratonic basin and rifted margin led to the deposition of several thick supracrustal packages (Figs. 1.1b and 1.2a; e.g., Delaney, 1981; Aitken and McMechan, 1991; Gabrielse and Campbell, 1991; Rainbird et al., 1997; Ross et al., 2001; Cook et al., 2004; Evenchick et al., 2005; Thorkelson et al., 2005; Leslie, 2009; Furlanetto et al., 209; Medig et al., 2012). These include the: • <1.8 Ga Mesoproterozoic Muskwa Assemblage; • 1.72–1.2 Ga Wernecke Supergroup (~13 km thick); • <1.4-1.0 Ga Pinguicula Group (~3.5 km thick); • 1.2(0.88?)–0.78 Ga Mackenzie Mountain Supergroup (>5–7 km thick); and • 0.78–0.54 Ga Windermere Supergroup, including the Hyland Group (>3.7 km) in the Selwyn Basin and the Ingenika Group (>3–4.5 km) in the Cassiar Platform. In addition to the extensional events that resulted in the deposition of the above-listed sedimentary packages, these rocks have been affected by at least two Proterozoic compressional events (the ~1.7 Ga Racklan, and ~0.9 Ga Corn Creek orogenies). A major hydrothermal event at circa 1.6 Ga in central Yukon hosted within the Wernecke Supergroup resulted in the formation of the Wernecke Breccia, widespread metasomatism of supracrustal basement, and localized iron–copper–gold mineralization (Thorkelson et al., 2001, 2005). Several episodes of alkaline to tholeiitic, mantle-derived magmas have intruded or erupted on these supracrustal packages. These magmatic events include: the 1.7 Ga Bonnet Plume River intrusions; the 1.4 Ga Hart River sills; the 1.27 Ga Bear River dykes; the 0.78 Ga Tzetzotene sills, Little Dal basalts, and Hottah dykes; the ~0.72 Ga Mount Harper Volcanic Complex; and the localized ~0.65 Ga Pool Creek syenite (Fig. 1.1b; e.g., Abbott, 1997; Dudas and Lustwerk, 1997; Pigage and Mortensen, 2004; Schwab et al., 2004; Thorkelson et al., 2001, 2005; MacDonald et al., 2010). 5 Although the suggestion that Proterozoic rocks form a metasedimentary ‘wedge’ in the lower-middle crust is consistent with low magnetic and gravity geophysical character (e.g., Cook et al., 2004; Clowes et al., 2005), other data are also consistent with alternate interpretations (e.g., Ledo et al., 2004; Evenchick et al., 2005; Jones et al., 2005). Evenchick et al. (2005) have interpreted the wedge-like feature to comprise crystalline basement that is interleaved with lesser Proterozoic metasedimentary rocks. References to the proportion of supracrustal versus infracrustal rocks in middle to lower crust of the northern Cordilleran, based on geophysical information, remain open to interpretation (R. Clowes, pers. comm.). This has implications with respect to the type of crust from which the syn-to post-accretionary magmas may have been derived. 1.1.2.3 Paleozoic Sedimentation, Rifting, and Magmatism Paleozoic passive margin sedimentation and intermittent rifting led to the formation of two extensive and long-lived sedimentary facies, which together comprise the northern Cordilleran miogeocline (Fritz et al., 1991; Gabrielse and Yorath, 1991). An eastern platformal or “shelf” facies consists of the shallow water Mackenzie and MacDonald platforms (Fig. 1.1a), which include carbonate and clastic rocks (Gordey and Anderson, 1993). A western basinal facies deposited on the rapidly subsiding rifted margin comprises shale, chert, basinal limestone, and turbiditic sediments of the Selwyn Basin and Kechika Trough (Fig. 1.1a; Gordey and Anderson, 1993). At least one extensive offshelf platform was also formed (e.g., the Cassiar–McEvoy Platform; Fig. 1.1a). The transition from platformal to basinal settings was controlled by structures that were formed during Proterozoic rifting events (see Section 1.1.3.1; e.g., Gabrielse and Yorath, 1991). A Late Devonian to Early Mississippian influx of turbiditic and cherty clastic sediments interrupted Lower Paleozoic miogeoclinal sedimentation in the northern Cordillera. Back-arc rifting in the northern Cordilleran miogeocline at this time led to the creation of several pericratonic entities that were separated from the western margin of Laurentian by the opening of the Slide Mountain ocean basin (Creaser et al., 1999; Erdmer et al., 2005; Murphy et al., 2006; Dusel–Bacon et al., 2006). Relative tectonic quiescence followed Devono–Mississippian rifting and sedimentation, allowing the re-establishment of basinal (east-derived) and platformal sedimentation on the newly formed western margin of North America by the Carboniferous. This miogeoclinal setting persisted until a reversal in subduction direction led to the closure of the Slide Mountain ocean at the end of the Permian, although sedimentation continued locally until the Triassic (e.g., Dusel–Bacon et al., 2006; Murphy et al., 2006; Nelson and Colpron, 2007; Beranek and Mortensen, 2011). Seismic reflection and refraction survey interpretations suggest that the (mainly Paleozoic) miogeoclinal strata in the northern Cordilleran miogeocline presently comprise only a thin, uppermost portion of the crustal section within ~6 km of the current erosional surface (e.g., Clowes et al., 2005). Paleozoic igneous rocks related to asthenospheric upwelling and decompression during basin subsidence and extension are concentrated near basin–platform transitions in the northern Cordilleran miogeocline (Figs. 1.1b and 1.2a). These include basaltic flows and feeder dykes with an alkaline oceanic-island basalt (OIB) geochemical affinity (Souther, 1991; Woodsworth et al., 1991; Pigage and Mortensen, 2004; Pigage, 2004, 2009; MacDonald et al., 2011). This mainly mafic magmatism was commonly concentrated along basin–platform transitions where normal and transform faults are localized (Cecile, 1982; Gabrielse and Yorath, 1991; Goodfellow et al., 1995; Mair 6 et al., 2006a; Leslie, 2009). Three main magmatic episodes comprising OIB-like igneous rocks include a broadly Cambrian event, an extensive mid- to Upper Ordovician event, and a Devonian event (Figs. 1.1b and 1.2a; Cecile, 1982; Tempelman–Kluit, 1981; Goodfellow et al., 1995; Pigage, 2004, 2009). 1.1.2.4 Mesozoic to Cenozoic Terrane Accretion It is uncertain exactly when the Slide Mountain basin began to close, but the accretion of the first pericratonic entity west of what is now Yukon appears to have occurred in the latest Permian to early Triassic (Beranek and Mortensen, 2011). At this time, the Slide Mountain and Yukon–Tanana terranes were obducted as thin klippen and upper crustal slices, respectively, onto miogeoclinal basement (e.g., Cook et al., 2004). This was the first of a series of closely spaced accretionary events that also involved the obduction and (or) tectonic wedging of the Cache Creek oceanic terrane, and the accretion of the outermost pericratonic island arc terrane as a complete (or nearly complete) lithospheric block (Stikinia terrane; Abraham et al., 2001; Cook et al., 2004; Beranek and Mortensen, 2006; Nelson and Colpron, 2007; Beranek, 2009). Allochthonous rocks were tectonically transported over 115 km to the east onto the northern Cordilleran miogeocline. Subduction of oceanic crust related to these accretionary events produced continental arc magmatism in the Yukon–Tanana terrane during the Late Triassic to Early Jurassic (e.g., Nelson and Colpron, 2007). By the mid- to Late Jurassic, significant Cordilleran uplift related to increased crustal and lithospheric thickening is thought to have occurred. One potential cause of this thickening is the closure of the Gravina oceanic basin separating the Insular Superterrane from North America sometime between the Late Jurassic and the Late Cretaceous (e.g., Trop and Ridgway, 2007). Locally preserved, Cretaceous clastic sedimentary rocks (circled; Fig. 1.1b) were deposited with coal seams in terrigenous basinal, fluvial, and (or) deltaic settings are also indicative of widespread Cordilleran uplift at this time (Blusson, 1971; Yorath, 1991; Gordey and Makepeace, 2001; Long and Gordey, 2007; Colpron, 2011). In the Mackenzie Plain, Cretaceous sedimentary rocks interspersed with bentonite layers were deposited in a progressively deepening foreland basin known as the ‘mid’-Cretaceous inland sea (e.g., Yorath, 1991). Prior to ~100 Ma, this detritus was derived from eastern cratonic sources and northern rift margins related to the opening of the Canada Basin in the Arctic (Yorath, 1991; Hadlari et al., 2009). A sea level drop at circa 100 Ma is indicated by a significant unconformity and has been attributed to uplift in the foreland to the west and an influx of westerly-derived detritus to the Cretaceous inland sea (Yorath, 1991; Hadlari et al., 2009; Thomson et al., 2011). During the earliest Late Cretaceous, the foreland basin depocentre shifted eastward until about 93 Ma, then returned to a more westerly position around 89 Ma (Fig. 1.1b; Thomson et al., 2011). These changes to the Cretaceous depositional regime coincided with changes in the locus and composition of syn- to post-accretionary magmatism across the Selwyn Basin. Through the Late Cretaceous and into the Eocene until ~45 Ma, subduction and associated arc magmatism continued intermittently along the western margin of North America, although the angle and rate of convergence varied with time (e.g., Engebretson et al., 1985; Kelley, 1993; Madsen et al., 2006). Orogenic uplift and a central Cordilleran source of detritus and volcanic ash in the foreland basin continued until at least the Paleogene (Yorath, 1991). 7 Clastic sedimentary sequences, together with mainly Eocene bimodal volcanic rocks and lamprophyre dykes were also deposited in fault-controlled basins. These deposits record localized extension in small pull-apart basins associated with regional strike-slip structures that were active at this time (Gabrielse, 1985; Jackson et al., 1986; Yorath, 1991; Gabrielse et al., 2006). 1.1.3 Deformation in the Northern Cordilleran Miogeocline Several events involving deformation of the ancestral western margin of North America have occurred between the Paleoproterozoic and the present day (e.g., Thorkelson et al., 2005). Three main deformational styles are the most relevant to the current architecture of the northern Cordilleran miogeocline. Each of these deformational regimes appears to have been important with respect to the geographic extent of syn- to post-accretionary magmatism in the region, particularly the widespread Early to ‘mid’-Cretaceous magmatic event. 1.1.3.1 Neoproterozoic Rifting and Basin Formation The geometry of the basin–platform transition, which is still apparent today (Fig. 1.1a), was imparted by rifting associated with the Neoproterozoic breakup of Rodinia (Cecile et al., 1997). This produced a lower plate rifted margin and a large eastward embayment in the Selwyn Basin (Gabrielse and Yorath, 1991; Hansen et al., 1993; Cecile et al., 1997). Lower plate rifted margins are characterized by structurally thinned basement, highly rotated normal faults, and thick packages of clastic or basinal rocks sedimentary rocks (Fig. 1.3; e.g., Lister et al., 1991). Several smaller embayments along the margin of the Selwyn Basin also formed as discrete rift systems (Fig. 1.1), although the these sub-basins may be more closely associated with later incipient rifting in the Paleozoic (Cecile et al., 1997). Features of an upper plate rifted margin are more typical of north–central British Columbia, as well as the dextrally displaced Laurentian rocks southwest of the Tintina Fault that make up the Cassiar Terrane (Fig. 1.3; Hansen et al., 1993; Cecile et al., 1997). Upper plate rifted margins are characterized by weakly rotational normal faults—which limit crustal thinning—and produced a narrow sub-basin (the Kechika Trough) and an extensive marginal platform (the Cassiar-McEvoy Platform) in the northern Cordillera (Figs. 1.1 and 1.3). Normal faults trending north–northwest to south–southeast throughout the northern Cordillera are interpreted to be transfer or tear faults between upper and lower plate segments of the rifted margin (Aitken and McMechan, 1991; Gabrielse and Campbell, 1991). There are also several northeasterly trending features identified on the basis of geophysical surveys, geological mapping and stratigraphic studies that are also interpreted to represent transverse offset faults related to Neoproterozoic rifting, including the Liard Line (near the transition between the lower and upper plate rifted margins; Cecile et al., 1997) and several others labeled in Figure 1.1a (Aitken and Pugh, 1984; Cecile et al., 1997; Morrow and Miles, 2000; Turner and Long, 2008). These basement structures were formed in the Proterozoic, but were likely reactivated during Paleozoic extension and, again, during Mesozoic compression (e.g., Norris, 1985; Gabrielse, 1991; Gordey and Anderson, 1993). 8 1.1.3.2 Mesozoic Thin-Skinned Deformation Compressional deformation throughout the Mesozoic, especially in the mid- to late Jurassic, caused thickening of the Laurentian crust, which was important for generating the temperatures and pressures required for extensive syn- to post-accretionary crustal anatexis. During this period of time, miogeoclinal rocks were affected by northeastward- migrating, thin-skinned deformation (e.g., Gordey and Anderson, 1993; Gordey, in press). Most of the contractional structures are east-verging; however, west-verging structures also formed in the western Selwyn Basin and parts of the Cassiar Terrane, probably as a result of the proximity to obducted and (or) tectonically wedged terranes (Price, 1986; Gabrielse et al., 1991). Reactivation of older extensional structures during contractional deformation may be responsible for the greater breadth of deformation across the Selwyn Basin – Mackenzie Platform (where highly rotated normal faults typical of a lower plate rifted margin were likely present in the basement) when compared to the Cassiar Terrane – Kechika Trough – MacDonald Platform (e.g., Fig. 1.1). Overall, the exposed Paleozoic miogeoclinal rocks experienced between 30 to 50% shortening related to this deformational regime, mainly by imbrication on several broadly folded thrust sheets with a relatively shallow detachments (Gabrielse, 1991; Gabrielse and Yorath, 1991; Gordey and Anderson, 1993; Gordey, in press). The relatively low topographic relief throughout the Selwyn Basin suggests that the underlying Neoproterozoic and older strata were not significantly shortened (Gordey, in press). Bedrock mapping studies across the northern Cordilleran miogeocline also indicate that thrust stacking generally does not typically involve Proterozoic stratigraphy (e.g., Gabrielse, 1991; Gordey and Anderson, 1993; Gordey, in press). Extensive Paleozoic–Mesozoic rocks are not thought to be tectonically interleaved with, or in the footwall to, exposed Proterozoic metasedimentary rocks in the northern Cordillera (e.g., G. Abbott, 2009, pers. comm.), although this interpretation is not universally accepted (e.g., Evenchick et al., 2005). Estimated emplacement depths (Anderson, 1982, 1985; Gordey and Anderson, 1993; Murphy, 1997; Baker and Lang, 2001; G.R. Thompson, unpub. data, 2000; Stephens et al., 2004; Moynihan, 2013) for the exposed plutons suggest that between 8 to 15 km of overlying rocks must have eroded from the northern Cordilleran miogeocline since the Early to ‘mid’-Cretaceous. Regional chlorite-zone greenschist facies (Proterozoic strata) to sub-greenschist facies (Paleozoic and younger strata) metamorphic grades of exposed rock units support these shallow burial depths for northern Cordilleran miogeoclinal rocks (Greenwood et al., 1991; Gordey and Anderson, 1993; Gordey, in press). Furthermore, the western Selwyn Basin and Cassiar Platform may have been structurally overlain by up to 10 km of the Yukon–Tanana and Slide Mountain terranes (Gabrielse, 1991; Gordey, in press). These obducted rocks may account for some (or all) of the burial metamorphism that appears to have affected the exposed miogeoclinal rocks. East-verging fold and thrust-style deformation in the northern Cordilleran miogeocline pre-dated syn- to post- accretionary igneous rocks (e.g., Gabrielse, 1991; Gabrielse and Yorath, 1991; Gordey and Anderson, 1993; Gordey, in press). Only small, north- to northeast-trending normal faults interpreted to be post-magmatic relaxation structures are observed crosscutting the Cretaceous intrusions (e.g., Gordey and Anderson, 1993). The deformation 9 front migrated east and north until at least the Late Cretaceous-Paleocene (Blusson, 1971; Aitken and Cook, 1974; Aitken et al, 1982; Gabrielse, 1991; Gabrielse and Yorath, 1991; Yorath, 1991; Gordey and Anderson, 1993; Gordey, in press). Therefore, it is possible that intrusive bodies emplaced into the upper crust were separated from their lithospheric roots and transported northeast along basal detachment faults. 1.1.3.3 Cretaceous–Cenozoic Dextral Strike-slip Deformation Significant dextral displacements along ‘margin-parallel’ strike-slip faults beginning in the Early Cretaceous (~140 Ma) and continuing through to the Cenozoic disrupted the western margin of North America (e.g., Gabrielse, 1991; Mihalynuk et al., 1999; Wyld and Wright, 2001; Gabrielse et al., 2006; Mair et al., 2006b). These faults were important in localizing, and perhaps even generating, Cretaceous (and younger) intrusive and volcanic rocks (e.g., Gabrielse et al., 2006; Colpron et al., 2011). Dextral transpressional deformation in the northern Cordillera was initiated during oblique convergence along the western margin of North America in the Early Cretaceous. Strike-slip faulting occurred well inboard of the continental margin (Gabrielse et al., 2006), where the Kechika and Cassiar faults (and several anastomosing faults) were active in offsetting portions of the Cassiar Platform. Coeval activity on the Teslin Tectonic Zone (and fault) and the Big Salmon fault formed the western boundary of the Cassiar Terrane. In the Selwyn Basin, a large fault system—extending from the Hyland fault in the southeast, to the Hess–MacMillan fault system in the northeast, to the Robert Service and Tombstone thrust faults in the northwest—was also initiated in response to the transpressional regime (Figs. 1.1a and 1.2; Gabrielse, 1991; Gordey and Anderson, 1993; Murphy, 1997; Gabrielse et al., 2006; Hart and Lewis, 2006). After restoration of (mainly) Cenozoic dextral strike-slip offset along the Tintina Fault, it is apparent that this lengthy fault system fringing the Selwyn Basin was contiguous with the Big Salmon and Teslin faults west of the Cassiar Terrane (Fig. 1.2; Gabrielse et al., 2006; Colpron et al., 2011). Many Early Cretaceous strike-slip structures remained active through to the ‘mid’-Cretaceous (and into the Cenozoic in some cases; Gabrielse, 1985; Gabrielse et al., 2006; Colpron et al., 2011). Beginning in the Late Cretaceous, however, the deformational regime across the northern Cordillera began to shift from transpressional to transtensional (Gabrielse et al., 2006; Nelson and Colpron, 2007) as plate vectors changed along the northwest margin of North America (Engebretson et al., 1985; Kelley, 1993; Miller et al., 2002; Haeussler et al., 2003; Madsen et al., 2003). Coupling between an increasingly oblique collisional margin, where oceanic crust was subducting below the western margin of North America, led to large-scale margin-parallel dextral transtensional shearing in the Cenozoic (Dewey, 1980; Beck, 1983; Gabrielse and Yorath, 1991; Kelley, 1993). Displacements totaling about 400 km each began in the Late Cretaceous along the Tintina Fault, and were gradually transferred to further west along the Denali Fault with time (Fig. 1.1a; Gabrielse, 1991; Hart et al., 2004a). Localized extension related to this deformational regime was manifested through the development of fault-bounded clastic and volcanic basins (Yorath, 1991), bimodal volcanism (e.g., Pride, 1988) and north-trending dyke swarms (Gabrielse, 1985, 1991; Sevigny and Thériault, 2003; Adams et al., 2005). 10 1.2 AN OVERVIEW OF SYN- TO POST-ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERA East-dipping subduction along the western margin of North America led to widespread, multi-stage syn- to post- accretionary igneous rocks in the northern Cordillera beginning at ~123 Ma, but becoming especially prevalent at ~115 Ma (Fig. 1.4a; e.g., Mortensen et al., 2000; Hart et al., 2004a). This magmatism is divided into three main events: an Early to ‘mid’-Cretaceous event (~123–90 Ma), and Late Cretaceous event (~84–67 Ma), and a latest Cretaceous to Eocene event (~67–47 Ma). The spatial and temporal distributions, basic compositional features, metallogenic associations, and suggested petrogenetic models for these three syn- to post-accretionary magmatic events across the northern Cordillera are summarized in Table 1.1. Early to ‘mid’-Cretaceous magmatism in the northern Cordillera includes several broad belts located in different portions of the Mesozoic orogen (Fig. 1.4a; Table 1.1; Anderson, 1988; Woodsworth et al., 1991). (1) An Early Cretaceous, mainly oxidized, metaluminous, intermediate to felsic, calc-alkaline magmatic episode that includes abundant associated mafic phases and volcanism (red belts in Fig. 1.4a). This belt was emplaced in a mainly outboard position within accreted pericratonic terranes of the western margin of North America. These igneous rocks are interpreted to represent continental arc magmatism associated with east–northeast dipping, oblique subduction. (2) An Early Cretaceous, mainly reduced, peraluminous, intermediate to felsic, calc-alkaline magmatic episode that lacks associated mafic phases and volcanism (purple belts in Fig. 1.4a). This belt was emplaced in a mainly inboard position within the northern Cordilleran miogeocline and easternmost portions of accreted pericratonic terranes. These igneous rocks are interpreted to represent back-arc magmatism related to transpression- and (or) decompression-assisted partial melting of orogen-thickened continental crust. (3) A ‘mid’-Cretaceous, reduced to oxidized, metaluminous to peraluminous to peralkaline, intermediate to felsic, calc-alkaline to alkaline magmatic episode that mostly lacks associated volcanism (the ex8ceptions being two or three clusters of calderas in the western Selwyn Basin and in east–central Alaska; Fig. 1.4a), although the youngest intrusions are associated with abundant mafic phases (orange belts in Fig. 1.4a). This belt was emplaced in an inboard position within the northern Cordilleran miogeocline and easternmost portions of accreted pericratonic terranes. These igneous rocks are interpreted to represent back-arc magmatism related to transpression- and (or) decompression-assisted partial melting of orogen-thickened continental crust, with localized extension associated with the youngest plutons. The Late Cretaceous, and latest Cretaceous to Eocene, magmatic events (Fig. 1.4b; Table 1.1) are broadly associated with transitions from the Early Cretaceous transpressional regime to the Cenozoic transtensional regime (e.g., Nelson and Colpron, 2007). During these two intervals, extensive calc-alkaline continental arc-related igneous rocks likely persisted in the west (Fig. 1.4b; Table 1.1). In a more inboard position farther east, however, peralkaline to peraluminous igneous rocks with an A-type affinity were sporadically emplaced as narrow belts and isolated occurrences of rare- and trace element-enriched felsic plutons (± associated volcanism), as well as bimodal sub- volcanic to volcanic centres (Fig. 1.4b; Table 1.1; Anderson, 1988; Woodsworth et al., 1991). All of these examples 11 of inboard A-type magmatism are associated with major structures (although it may not be immediately obvious in Figure 1.4b). 1.2.1 The Early to ‘Mid’-Cretaceous Magmatic Event Syn- to post-accretionary, Early to ‘mid’-Cretaceous (123–90 Ma) magmatism in the northern Cordillera comprises widespread, primarily intermediate–felsic intrusive rocks together with minor extrusive equivalents in Yukon and adjacent parts of the southwestern Northwest Territories, north–central British Columbia, and east–central Alaska (Fig. 1.3). Previous petrogenetic studies have sub-divided the intrusions and associated volcanic into several magmatic belts and (or) suites that are distinct in terms of age, lithology, composition (e.g., oxidation state, aluminicity), and geographic distribution (Fig. 1.4a; e.g., Anderson, 1988; Woodsworth et al., 1991; Souther, 1991; Aleinikoff et al., 2000; Mortensen et al., 2000; Hart et al., 2004, 2005; Heffernan, 2004). Some of the plutonic suites are also partly distinguished by features including the size of the exposed bodies, mineralogical and textural similarities, and metallogenic associations. Volcanic rocks are linked to some of the plutonic suites on the basis of age, spatial association, and composition. 1.2.1.1 Early Cretaceous Arc Magmatism The mainly oxidized and metaluminous magmatic episode comprising the westernmost syn- to post-accretionary magmatism of Early Cretaceous-age in the northern Cordillera was emplaced mainly into continentally-derived sedimentary rocks within the pericratonic terranes. The extent of these igneous rocks is defined by the red polygons in Figure 1.4a, and includes the Dawson Range, the Whitehorse, and (possibly) the Teslin plutonic suites (e.g., Mortensen et al., 2000; Hart et al., 2004a). These calc-alkaline igneous rocks have been previously classified as I- type (Anderson, 1988; Woodsworth et al., 1991) and oxidized magnetite-series (Hart et al., 2004a) granitoids. Andesitic to dacitic Early Cretaceous volcanic rocks were also deposited in this region and are spatially and temporally associated with calc-alkaline plutons and batholiths of continental magmatic arc affinity. These volcanic rocks are known as the Mount Nansen Group in western Yukon, where they are closely correlated with the Whitehorse plutonic suite (e.g., Souther, 1991; Mortensen et al., 2000; Colpron, 2011). The composition and mineralogy of this calc-alkaline belt are consistent their interpreted petrogenesis as a continental magmatic arc resulting from the east–northeast dipping subduction of oceanic crust along the western margin of ancestral North America during the Early Cretaceous (e.g., Mortensen et al., 2000; Hart et al., 2004a; Crawford et al., 2005; Mortensen and Hart, 2010). While existing isotopic ages for these intrusions range nearly continuously from 123 Ma to 90 Ma (e.g., Hart et al., 2004a), it should be noted that many of these age determinations employed the K–Ar and Rb–Sr methods, which are unreliable and inaccurate. U–Pb ages of monazite or zircon provide the most robust and accurate ages, and with few exceptions the U–Pb ages in the Early Cretaceous continental magmatic arc are all older than ~99 Ma (Fig. 1.4a)2. After restoration of ~400 km of Cenozoic offset 2 The Tok–Tetlin plutonic suite—for which minimal information is available, but is thought to represent fore-arc partial melting of underplated flysch—may be younger (94–92 Ma; 40Ar/39Ar in hornblende; Richter et al., 1975a; 12 along the Tintina Fault, these continental arc rocks would be located in present-day southwestern Yukon and northern British Columbia. 1.2.1.2 Early Cretaceous Back-Arc Magmatism An Early Cretaceous (ca. 115–100 Ma) belt of relatively aluminous, ilmenite > magnetite-series plutons that is broadly coeval to the continental magmatic arcs was emplaced in a more inboard position (Fig. 1.4a). The extent of these rocks is defined by the purple polygons in Figure 1.4a, and they include the Yukon–Tanana Uplands, Cassiar, Anvil, and Hyland plutonic suites. These plutonic suites broadly correspond to previous classifications as S-type (Anderson, 1988; Woodsworth et al., 1991) and reduced ilmenite-series to weakly oxidized–series (Hart et al., 2004a) granitoids. This magmatic belt is characterized by abundant batholiths, an absence of widespread associated volcanic rocks, and a rarity of associated mafic phases. In the Cassiar Terrane, back-arc magmatism was collectively referred to as the Cassiar plutonic suite, which comprises peraluminous to weakly metaluminous, biotite ± muscovite or hornblende monzogranite to granodiorite batholiths with lesser quartz monzonite to quartz monzodioritic phases (Table 1.1; Fig. 1.4a). In the Selwyn Basin this Early Cretaceous belt is represented by the Anvil and Hyland plutonic suites, and comprises batholiths of mainly peraluminous, biotite ± muscovite, granodiorite to monzogranite ± minor hornblende-bearing quartz monzodiorite composition (Table 1.1; Fig. 1.4a). The compositionally similar and coeval Yukon–Tanana Uplands plutonic suite, located in east–central Alaska (Fig. 1.4a), is thought to represent the northwest continuation of the Anvil and Hyland suites prior to Cenozoic offsets along the Tintina Fault (Aleinikoff et al., 2000; Werdon et al., 2004; Dilworth et al., 2007). Local rhyolitic calderas associated with Yukon–Tanana Uplands plutonic suite (Fig. 1.4a; Bacon et al., 1990; Mortensen, 2008) are probably the source of ~107 Ma ignimbrites in Early Albian sedimentary rocks several 100’s of km farther east in the foreland basin (Thomson et al., 2011). 1.2.1.3 ‘Mid’-Cretaceous Post-Arc Magmatism A ‘mid’-Cretaceous (99–90 Ma) belt is widespread in the eastern and northern Selwyn Basin where it is found inboard of the Early Cretaceous back-arc belt, and is narrower in the Cassiar Terrane where it overlaps with the Early Cretaceous back-arc belt (Fig. 1.4a). Based on available U–Pb geochronological data, the ‘mid’-Cretaceous belt comprises several post-arc plutonic suites emplaced after continental arc magmatism to the west–southwest had terminated (Fig. 1.4b). ‘Mid’-Cretaceous igneous rocks comprise metaluminous to peraluminous (to peralkaline), ilmenite to magnetite-dominant plutons that are younger than the continental magmatic arc rocks located farther to the west–southwest. Of the ‘mid’-Cretaceous plutonic suites in the Selwyn Basin, the Tay River plutonic is the most proximal to the Early Cretaceous Anvil and Hyland plutonic suites (Fig. 1.4b). This suite comprises reduced to oxidized, slightly Hart et al., 2004a). This suite comprises a very narrow belt of plutons immediately northeast of the Denali Fault in east–central Alaska that are not part of any of the magmatic belts shown in Figure 1.4a. 13 peraluminous, quartz monzodiorite to granite intrusions, which contain biotite ± hornblende or muscovite. Associated dacitic volcanic rocks in the ~97 Ma South Fork calderas are preserved in down-dropped blocks in the western Selwyn Basin (e.g., Souther, 1991; Gordey, in press). The other three ‘mid’-Cretaceous plutonic suites in the Selwyn Basin are the Tungsten, Mayo, and Tombstone suites, which were mainly emplaced inboard from the Tay River suite (Fig. 1.4b). Collectively, these three plutonic suites are referred to as the Tombstone–Tungsten Belt (TTB; e.g., Lang, 2000; Mortensen et al., 2000; Hart et al., 2005). The Tungsten plutonic suite comprises reduced, peraluminous, biotite ± muscovite ± garnet monzogranite and leucogranite. The Mayo plutonic suite comprises mainly reduced, metaluminous, biotite ± hornblende ± clinopyroxene quartz monzonite to quartz monzodiorite to monzogranite. Late lamprophyre dykes are commonly associated with the Mayo plutonic suite. The Tombstone plutonic suite comprises reduced to oxidized, metaluminous to peraluminous to peralkaline, monzonite to granodiorite to monzogranite, but is also locally associated with gabbroic and silica-undersaturated intrusions. The Mayo (± Tay River) suite is compositionally analogous to the Fairbanks–Salcha suite in east–central Alaska, whereas the Tombstone plutonic suite is compositionally analogous to the Livengood suite in east–central Alaska (Fig. 1.4a; e.g., Hart et al., 2004a). Based on compositional similarities and anomalously young K–Ar and Rb–Sr ages, the Tombstone suite was previously grouped with “Late Cretaceous A-type” plutons (Anderson, 1988; Woodsworth et al., 1991). Although the A-type affinity is correct for at least some phases of the Tombstone plutonic suite, this suite is unlike younger examples of A-type granitoids elsewhere in the northern Cordillera (e.g., Fig. 1.4b; Chapter 2). In the Cassiar Terrane, ‘mid’-Cretaceous igneous rocks are present as the peraluminous (to peralkaline) Seagull– Thirtymile “granites” (Fig. 1.4b; Liverton and Alderon, 1994; Mortensen et al., 2007). Several scattered, small bodies of coeval granite extend this belt to the north-northwest (Fig. 1.4b; Gallagher, 1999; Colpron et al., 2003; Piercey and Colpron, 2009; M. Colpron, 2011, pers. comm.). 1.2.1.4 Previously Proposed Petrogenetic Models for Early to ‘Mid’-Cretaceous Magmatism Most workers generally accept that the western belt of calc-alkaline igneous rocks in the northern Cordillera reflects continental arc magmatism related to subduction and dehydration of an oceanic plate (e.g., Anderson, 1985; Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004a). In continental magmatic arcs, the initial melts are produced by hydration and anatexis of the lithosphere above down going oceanic lithosphere (Tatsumi et al., 1983). These mantle-derived melts are thought to pond at the base of the crust, where partial melting of the lower crust and the generation of hybrid magmas ranging from gabbroic to granitic compositions is thought to take place (e.g., Hildreth and Moorbath, 1988; Annen et al., 2006). Away from the subduction zone in the back-arc region, however, magmatic processes are less well understood. A source of heat and magma in the mantle below an extensional back-arc basin is easily attributed to the result of decompression in the mantle. Specific mechanisms that are capable of generating enough heat in a compressional back-arc setting to induce extensive crustal anatexis, however, are somewhat more complicated and contentious. More than one possible scenario has been proposed to explain intermediate to felsic, back- and post-arc igneous 14 rocks in the northern Cordillera, which—in the Early to ‘mid’-Cretaceous—was characterized by a transpressive (i.e., compressive) tectonic regime. This magmatism is located well to the northeast of the inferred continental magmatic arc and its associated subduction zone. However, all of the Early to ‘mid’-Cretaceous plutonic suites and magmatic belts (and even elongate batholiths) are oriented sub-parallel to the western margin of the Cordillera, which implies a role for continental margin processes in the genesis of thesinboard igneous rocks (Fig. 1.4a; e.g., Gabrielse and Yorath, 1991). Similarly, spatially associated (and in many cases, apparently contemporaneous) regional strike-slip faults may also have been important in focusing―and even generating―much of this back- and post-arc magmatism (Gabrielse et al., 2006; Colpron et al., 2011). Changes in the composition and distribution of the back-arc and post-arc magmatism may even be directly tied plate reorganizations in the west (Gabrielse and Yorath, 1991). The pros and cons of four petrogenetic models previously proposed for magmatism in this region are discussed as follows. It is possible (even likely), however, that the petrogenesis of the Early to ‘mid’-Cretaceous igneous rocks in the northern Cordilleran miogeocline may have combined elements of more than one of the following scenarios. (a) Anatexis of over-thickened crust: One interpretation for widespread Cretaceous igneous rocks in the northern Cordilleran miogeocline is that they are post-tectonic, compositionally restricted, S-type granitic plutons (Anderson, 1988; Gordey and Anderson, 1993; Driver et al., 2000). These authors propose Cretaceous partial melting of continental crust that was over-thickened by between 50 to 100% in the mid- to late Jurassic, possibly as a result of an accretionary event (e.g., the Insular Superterrane; Monger et al., 1982; Souther, 1991; Woodsworth et al., 1991). This model does not require that active subduction (and continental arc magmatism) was occurring at the same time, and can be applied to both the Early Cretaceous back-arc and ‘mid’-Cretaceous post-arc magmatic belts. The most recent version of this model is based on extensive isotopic and geochemical data for the Cassiar batholith (of the Cassiar plutonic suite) in north–central British Columbia and south–central Yukon. Driver et al. (2000) suggested that the intermediate to felsic magmas comprising this composite intrusion were sourced primarily from the Meso- to Neoproterozoic metasedimentary rocks(± basement gneisses) and interlayered mafic to intermediate amphibolite rocks. This model generally accounts for the range of isotopic and compositional variation exhibited by the intrusions of the Cassiar plutonic suite (e.g., Morris and Creaser, 2008). A similar petrogenetic scenario has also been extended to similar plutonic suites emplaced across the Cassiar Terrane and the Selwyn Basin (e.g., Heffernan, 2004; Morris and Creaser, 2008). Most experimental studies and modeling of fluid-absent partial melting in the crust indicate that heat, and commonly magma, input from the mantle are required for crustal anatexis to proceed (e.g., Clemens and Wall, 1981; England and Thompson, 1984; Clemens and Vielzeuf, 1987; Roberts and Clemens, 1993; Brandon et al., 1994; Barbarin, 1996; Soesoo and Nicholls, 1999; Clemens, 2003; Bonin, 2004; Bea, 2011). Extensive anatexis of the crust without the addition of external heat, however, is thought to occur only in situations where the crust was substantially thickened (i.e., to total thicknesses of 50–70 km) about 30–40 m.y. prior to the main period of granitic magmatism (e.g., Patiño-Douce et al., 1990; Bea, 2011). In this scenario, the crust cannot have undergone gravitational collapse or thinning of any type prior to the main magmatic event. Furthermore, the melt source must have been enriched by 15 50% (or more) in radiogenic heat-producing elements (e.g., K, Th, U) compared to average continental crust (Patiño-Douce et al., 1990; Bea, 2011). Thus, the burial of radiogenic sedimentary rocks to lower-middle crustal depths is a viable option for the generation of extensive crustally derived magmatism through the fluid-absent dehydration of micas, if the above conditions are met. The partial melting of less radiogenic mafic–intermediate lithologies to produce the hornblende- (or pyroxene-) bearing magmas, however, requires temperatures higher than those that can be attained by this model (i.e., >850°C; e.g., Thompson, 2001). Mantle-derived heat, or mantle- derived heat and magma, must have been involved the generation of less siliceous (hornblende-bearing quartz monzodiorite to granodiorite) intrusive phases of the above-described Early to ‘mid’-Cretaceous plutonism. A role for mantle-derived heat (and possibly magma) in the genesis of some of these intrusions is compatible with coeval “lamprophyre” dykes reported in the vicinity of the Cassiar batholith (Cooke and Godwin, 1984; Ash, 2001). There is no evidence for Early Cretaceous crustal thickening greater than about 50 km in the northern Cordilleran miogeocline. The depth to the MOHO below the northern Cordilleran miogeocline today is ~33–36 km (Clowes et al., 2005), and regional (sub)greenschist metamorphic grades with local andalusite at the present surface (e.g., Greenwood et al., 1991; Gordey and Anderson, 1993). The exposed rocks in the northern Cordilleran miogeocline cannot have been buried more than 15 km depth, and in many examples may have experienced even less burial (i.e., 8–10 km; see Section 2.2.2.1, Chapter 2; M. Colpron, 2012, pers. comm.; Moynihan, 2013). It seems that the continental crust cannot have been more than 50 km thick following the deposition of the exposed upper crustal rocks, making the model of crustal anatexis problematic. Nor can this model account for the generation of the intermediate, hornblende- or clinopyroxene bearing intrusive phases associated with nearly all of the Early to ‘mid’- Cretaceous plutonic suites detailed above. Therefore, a simple model of crustal over-thickening and anatexis, even in a relatively radiogenic sedimentary protolith, may not provide a sufficient mechanism for the voluminous melting responsible for syn- to post-accretionary igneous rocks across much of the northern Cordilleran miogeocline. This suggests that the generation of both back-arc and post-arc suites was not as simple as a model of over-thickened continental crust would imply. (b) Back-arc collapse, exhumation, and extension: More recently, both Early to ‘mid’-Cretaceous igneous rocks in the northern Cordillera have been interpreted to be the result of back-arc decompression and extension-related partial melting of the crust (and possibly mantle) during and following the collapse of orogenically thickened crust (Hart et al., 2004a, 2005; Mair et al., 2006b; Dilworth et al., 2007). This model has been based mainly on observations from the northwest Selwyn Basin in central Yukon and east–central Alaska. It accounts for two distinct styles of magmatism in the northern Cordilleran miogeocline that correspond to the Early Cretaceous belt and the ‘mid’-Cretaceous belt. The Early Cretaceous Anvil and Yukon–Tanana Uplands plutonic suites are syn- to post-kinematic plutons that have been emplaced into greenschist to amphibolite facies metamorphic basement (e.g., Murphy et al., 2001, 2002, 2006; Dusel–Bacon et al., 2006; Mair et al., 2006b; Dilworth et al., 2007). In east–central Alaska, widespread Early Cretaceous exhumation of upper amphibolite facies basement is evidenced by pre- to syn-magmatic metamorphic 16 and cooling ages, as well as regional shallow-dipping faults ( Dusel–Bacon et al., 2002; Pavlis et al., 1993; Berman et al., 2007). The Yukon-Tanana Uplands plutonic suite was emplaced into this region during exhumation (Fig. 1.4a). Similar cooling ages in the host rocks to the Anvil plutonic suite in the Yukon-Tanana inlier (Fig. 1.4a) indicates that these rocks were also exhumed in conjunction with Early Cretaceous magmatism (Murphy et al., 2001). The schistose to gneissic basement of east–central Alaska and the Yukon–Tanana terrane inlier in the western Selwyn Basin may have been in close proximity when they were exhumed in the Early Cretaceous prior to dextral offset along the Tintina Fault (e.g., Gabrielse et al., 2006). Early Cretaceous igneous rocks in east–central Alaska are interpreted to have been generated during an elevated crustal geothermal gradient, and decompression associated with regional uplift (Mair et al., 2006b). A model invoking the convective removal of over-thickened lower lithosphere following the attainment of maximum crustal thicknesses (~60 km in the Jurassic) was proposed to explain the timing of regional metamorphism, and extensive exhumation and magmatism several 10’s of m.y. following the delamination event (Mair et al., 2006b). Thus, the Yukon–Tanana Uplands suite (and by proxy the Anvil suite) is interpreted to represent large-volume, near-in situ, decompression-related crustal melts derived from their metasedimentary hosts (e.g., Hart et al., 2004a). The Hyland, Tay River, and Cassiar plutonic suites have also been interpreted to represent near-in magmas, also possibly derived from their (mainly Neoproterozoic) host rocks (Hart et al., 2004a). In the Cassiar Terrane and Selwyn Basin (Yukon and the Northwest Territories), however, high- grade metamorphic rocks, Early Cretaceous exhumation, and shallow-dipping syn-magmatic extensional faults are absent. Instead, plutonic suites in the Cassiar Terrane and Selwyn Basin are more closely associated with dextral transpressional structures (Gabrielse et al., 2006; Mair et al., 2006b; Colpron et al., 2011). The ‘mid’-Cretaceous TTB plutonic suites comprise smaller, generally rounded (post-deformational) plutons thought to have been derived from a mainly sedimentary protolith, although magma from an underlying enriched lithospheric mantle was apparently involved in the genesis of (at the very least) the Tombstone plutonic suite (Hart et al., 2004a). In this model of back-arc collapse, exhumation, and extension, the ‘mid’-Cretaceous magmatic event in the northwest Selwyn Basin would have been located on the far flank of the main region of crustal exhumation occurring in east–central Alaska (Mair et al., 2006a). This area did not experience as of a high heat flux from the mantle, and less pronounced uplift and extension produced smaller magma volumes (Hart et al., 2004a, 2005; Mair et al., 2006a). Late-stage, lithospheric extension is envisioned to have caused decompression of the underlying mantle and the intrusion of mantle-derived magmas into the crust in association with the youngest TTB magmas in the most outboard position relative to the continental margin (Fig. 1.4a; Hart et al., 2004a). In this model, upper crustal and lithospheric extension are thought to be directly linked to ongoing continental arc magmatism in the west (Hart et al., 2004a). However, robust isotopic age constraints (e.g., U–Pb, 40Ar/39Ar, and Re–Os ages) for rocks of the interpreted magmatic arc lying to the west indicate that voluminous arc-related magmatism ended by about 99 Ma west of the Selwyn Basin (and 107 Ma west of the Cassiar Terrane; Table 1.1; Fig. 1.4a; e.g., Hart, 1995; Gordey et al., 1998; Mortensen et al., 2000; Mortensen and Hart, 2010, unpub. data). Subduction along the northwest margin of North America, therefore, cannot be the driving force responsible for the processes that led to ‘mid’-Cretaceous magmatism in the northern Cordilleran miogeocline. Despite this issue, the above mechanisms for partial melting in the northern Cordilleran miogeocline are consistent with field observations and this model provides a reasonable interpretation for how intermediate–felsic granitoids may be generated in compressive back- 17 arc regimes. (c) Dextral transpressional tectonism: Several of the batholiths in Figure 1.4b are elongated sub-parallel to the regional structural grain (i.e., to the north- northeast). An alternate explanation for the generation of Early to ‘mid’-Cretaceous magmatism, thus, is syn- to late- kinematic decompression-related melting along major dextral transpressional structures (Murphy et al., 1995; Murphy, 1997; Gabrielse et al., 2006). This model is partly supported by the local development of strongly foliated to mylonitic, syn- to late-emplacement fabrics observed in several 96 to 110 Ma plutons and batholiths in the Cassiar Platform and Selwyn Basin that are juxtaposed against large strike-slip faults (e.g., Poole, 1956; Gabrielse, 1969; Murphy et al., 1995; Harvey et al., 1997; Gallagher, 1999; Johnston, 1999; Gabrielse, 2003; Colpron et al., 2011). Furthermore, the Hyland plutonic suite the Selwyn Basin west of, or in the hanging wall of the Hyland fault (Figs. 1.1a and 1.4a) is localized in a region underlain by the oldest and some of the highest grade metamorphic rocks (staurolite-grade regionally, with staurolite-garnet in contact aureoles; e.g., Roots et al., 1966) in the northern Cordilleran miogeocline. This association between the Hyland plutonic suite and amphibolite grade host rocks suggests that exhumation related to movement along the Hyland Fault was important in the genesis of intermediate– felsic batholiths in its hanging wall (Murphy et al., 1995; Gabrielse et al., 2006). Conversely, reverse faults in the Selwyn Basin may be more closely associated with smaller-volume plutons of the TTB (e.g., the Tombstone and March thrust faults; Figs. 1.1a and 1.4a). The timing of northwest-directed thrusting along the Tombstone fault, however, is coeval to the onset of the Hyland magmatic event, linking these two distinctly different structures located at opposite ends of the Selwyn Basin (Figs. 1.1a and 1.4a; Murphy et al., 1995; Murphy, 1997; Gabrielse et al., 2006; Colpron et al., 2011). A model invoking syn- to late-kinematic decompression-related magmatism near strike-slip faults may be applicable throughout much of the Cassiar Terrane and Selwyn Basin. It may not, however, fully explain plutonism that is located more distally to regional strike-slip structures (e.g., in the central Selwyn Basin, and east–central Alaska; Figs. 1.1a and 1.4a) or undeformed and rounded intrusions (e.g., several Tombstone suite plutons). Regardless, extension related to dextral transpressional structures may have been particularly important in localizing, if not generating, Early to ‘mid’-Cretaceous magmas across the Selwyn Basin and Cassiar Terrane. (d) West-dipping subduction and the accretion of orogenic float: The most recently proposed model for syn- to post-accretionary Cretaceous magmatism in the northern Cordilleran miogeocline is drastically different from previously proposed models in that it suggests these igneous rocks are pre- collisional (Johnston, 2008). The model postulates that a cryptic suture zone of Eocene age is present below, or at, the eastern margin of the Selwyn Basin. In this model, the Selwyn Basin is described as an accretionary prism located between an accreted ribbon-continent (comprising the Cassiar Terrane and previously accreted pericratonic terranes) and North American rocks (comprising platformal rocks east of the Selwyn Basin and Kechika Trough; Fig. 1.1a). The accretion of “orogenic float” (cf. Oldow et al., 1990) to the ribbon-continent, and the accretion of the 18 ribbon-continent to North American platformal rocks would have occurred as a consequence of the west-dipping subduction of oceanic crust below the ribbon-continent. The progressive addition of orogenic float to the upper plate (ribbon-continent) led to the northeastward progression of magmatism with time, providing an explanation for emplacement of younger plutonic suites farthest from the current western margin of North America (the orange belt; Fig. 1.4a). In this model, compositional variations among the plutonic suites are attributed to changes in basement composition during the northeast-ward migration of a continental magmatic arc from the ribbon-continent into argillaceous orogenic float during the accretionary process, leading to younger and more radiogenic S-type plutons further to the northeast. Similarly, the relatively radiogenic compositions of I-type plutons farther to the west in the accreted ribbon-continent (the red and purple belts; Fig. 1.4a) could be due to the subduction of sediment on the down-going plate. The main problem with this model is the stratigraphic continuity of numerous Cambrian to Triassic sedimentary units from the Mackenzie Platform across the Selwyn Basin, and in particular across the locus of the proposed suture zone. Although a concentration of Paleozoic volcanic rocks and thrust faults occur near the basin–platform boundary in the northern Cordillera (Figs. 1.1b and 1.2a), these features do not require that this boundary was a tectonic suture. Instead, the transitional stratigraphy, and the localization of intrusive rocks and Mesozoic structures at this boundary, were probably controlled by normal faults related to earlier rifting (e.g., Cecile, 1982; Gordey and Anderson, 1993; Goodfellow et al., 1995; Mair et al., 2006a; Leslie, 2009). Furthermore, the basin–platform transition is not characterized by the high-grade metamorphic rocks and penetrative deformation that would indicate the presence of a nearby suture zone (although some thrust panels are intensely deformed in the northwest Selwyn Basin; e.g., Murphy, 1997; Mair et al., 2006a; Gordey, in press). Lastly, compositional variations exhibited by the plutonic suites do not support a model of arc magmatism in the northeast Selwyn Basin, even if the arc magmas were highly crustally contaminated. The vast majority of these intrusions are intermediate to felsic, and lack abundant mafic and volcanic equivalents (although large felsic calderas are locally present in the central Selwyn Basin). This is contrary to most continental magmatic arcs, where voluminous granodioritic plutonic rocks are associated with abundant mafic phases and effusive andesitic volcanism, regardless of the thickness and composition of the crustal column (e.g., Pitcher et al., 1985). 1.2.2 The Late Cretaceous Magmatic Event In addition to the Early to ‘mid’-Cretaceous plutonic suites in the northern Cordilleran miogeocline, Late Cretaceous episodes of magmatism are also present in the northern Cordillera (Table 1.1; Fig. 1.4b; e.g., Anderson, 1988; Woodsworth et al., 1991). The majority of this later magmatism, however, has not been formally defined in Yukon, and most geochronological constraints rely largely on less reliable isotopic dating methods (K–Ar or Rb–Sr). The Late Cretaceous magmatic episode may be simplified into two main, ‘outboard’ suites (Anderson, 1988; Souther, 1991; Woodsworth et al., 1991; Mihalynuk, 1999; Selby et al., 1999): (1) the 78–67 Ma Carmacks Group of plutonic and volcanic rocks in western Yukon and northwest British Columbia (Fig. 1.4b), and (2) the 84–78 Ma Surprise Lake plutonic suite, in north–central to northwest British Columbia (Fig. 1.4b). 19 The widespread Carmacks Group (Fig. 1.4b) comprises relatively potassic, mainly andesitic plutonic and volcanic rocks (and associated volcaniclastic sediments), and is found mainly within the Yukon–Tanana terrane (e.g., Souther, 1991; Mihalynuk, 1999; Mortensen and Hart, 2010). These rocks are currently the subject of a detailed petrogenetic study (e.g., Ciolkiewicz et al., 2012) and an extensive geochronological study (Mortensen and Hart, 2010, unpub. data) that will undoubtedly change current ideas regarding their petrogenesis. For now, however, the older Carmacks Group rocks (78–72 Ma; Mortensen and Hart, 2010, unpub. data) are primarily plutonic and are likely correlative with the Windy Table volcanic suite in northern British Columbia (e.g., Mihalynuk, 1999). This older episode is thought to represent a narrow calc-alkaline continental magmatic arc that was emplaced during a dextrally transpressional regime related to oblique subduction along the western margin of North America (Souther, 1991; Selby et al., 1999; Mortensen and Hart, 2010). Younger Carmacks Group rocks (72–67 Ma; Mortensen and Hart, 2010, unpub. data) are more oxidized, alkaline, and possibly more metallogenically important suite (Table 1.1; Selby and Creaser, 2001; Ciolkiewicz et al., 2012). This younger episode is dominated by effusive volcanic rocks across much of western Yukon, and has been postulated to have resulted from lithospheric delamination following the cessation of subduction along the western margin of North America (Gabrielse and Yorath, 1991; Mortensen and Hart, 2010). The Surprise Lake plutonic suite comprises a west–northwest trending belt of slightly peralkaline to strongly peraluminous plutons with A-type (or within-plate) geochemical affinities (Anderson, 1988; Woodsworth et al., 1991; Ray et al., 2000). This plutonic suite has correlated to the calc-alkaline Carmacks Group on the basis of its similar age (e.g., Woodsworth et al., 1991; Mihalynuk, 1999). It is compositionally distinct, however, from the calc- alkaline Carmacks Group. Instead, the Surprise Lake plutonic suite is more akin to the ‘mid’-Cretaceous Seagull batholith in south–central Yukon with respect to mineralogy, composition, texture, and metallogenic signature (Anderson, 1988; Woodsworth et al., 1991; Liverton and Botelho, 2001). Nelson and Colpron (2007) have grouped the Surprise Lake plutonic suite with other similar-age intrusions (see the following paragraph), and attributed it to crustal anatexis following orogenic thickening. However, A-type plutonic suites are petrogenetically complex, and thought to be generated during unusually high temperature partial melting of either anhydrous crust and (or) undepleted to enriched mantle (Whalen et al., 1987; Sylvester, 1989; Eby, 1990, 1992; Turner et al., 1992; Bonin et al., 1998; Liégeois et al., 1998; Bonin, 2004; Kemp et al., 2005; Frost and Frost, 1997, 2011). Farther east, in an inboard position within in the northern Cordilleran miogeocline, evidence for additional Late Cretaceous magmatism is found only locally in the Cassiar Terrane. Late Cretaceous ages from K–Ar and Rb–Sr (and U–Pb at one locality; Gregory, 2007) datasets for plutonic, metamorphic (hornfels), and hydrothermal micas and whole rocks form a short belt near the north end of the Kechika fault in the Cassiar Terrane (Figs. 1.1a and 1.4b). Additionally, 64–74 Ma (U–Pb) intrusive rocks occur in the western Cassiar Terrane near the Teslin Tectonic zone (Figs. 1.1a and 1.4b; e.g., Gordey et al., 1998). There is no Late Cretaceous magmatism of this age documented in the Selwyn Basin, except for the latest Cretaceous McQuesten plutonic suite, which will be discussed in the following section. 20 1.2.3 The Latest Cretaceous to Eocene Magmatic Event Latest Cretaceous to Eocene magmatism (Table 1.1; Fig. 1.4b) has not been studied extensively, nor are most of these occurrences formally defined in the northern Cordillera (with the exception of the McQuesten plutonic suite; e.g., Murphy, 1997). Geochronological constraints for much of this magmatism also largely rely on unreliable and inaccurate dating methods. As this is the last widespread magmatic event to affect the northern Cordillera, however, K–Ar and Rb–Sr ages for latest Cretaceous to Eocene rocks are less likely to have been re-set, although some ‘apparently’ Paleogene–Eocene intrusions may actually be older (e.g., Woodsworth et al., 1991; Mihalynuk, 1999). As with the previous section, this magmatic event is divided into rocks located in an outboard position, and rocks located in an inboard position. The circa 55 Ma volcanic rocks of the Sloko and Skukum suites (in northwest British Columbia and southwest Yukon, respectively) are found mainly in the Yukon–Tanana terrane and the adjoining oceanic and Insular terranes (Fig. 1.4b). Both of these calc-alkaline volcanic suites are interpreted to be derivatives of continental arc magmatism above a shallow-dipping subduction zone (Souther, 1991; Mihalynuk, 1999; Morris and Creaser, 2003; Crawford et al., 2005; Simmons et al., 2006, 2007). These extrusive rocks are likely related to the Coast Plutonic Complex, which encompasses the Ruby Range continental magmatic arc in southwest Yukon (Souther, 1971; Lambert, 1974; Smith, 1983; Gehrels et al., 1991; Souther, 1991; Currie, 1994; Mihalynuk et al., 1995; Mihalnyuk, 1999; Simmons et al., 2006, 2007; S. Israel, 2012, pers. comm.). At the northern tip of the Coast Plutonic Complex, the Nisling Range plutonic suite (Fig. 1.4b) comprises high-level alaskite plutons, dyke swarms, and felsic pyroclastic rocks ranging from 52 to 62 Ma (K–Ar and Rb–Sr ages; e.g., LeCouteur and Tempelman–Kluit, 1976) with features that are indicative of an A-type affinity (e.g., Anderson, 1988; Souther, 1991) A possible explanation for the composition of the Nisling Range plutonic suite is that it may be related to the subduction of one or more slab windows (e.g., Madsen et al., 2006) rather than the continental magmatic arc represented by the Coast Plutonic Complex (Fig. 1.4b). Hints of Paleocene-Eocene magmatism in the eastern Cassiar Terrane include K–Ar and Rb–Sr ages for several intrusions (Fig. 1.4b; Christopher et al., 1972; Canadian Geochronological Knowledgebase, 2013). Some Late Cretaceous to Eocene plutonic and volcanic suites have also been briefly described in the western Selwyn Basin (the ca. 65 Ma McQuesten plutonic suite and the ~55 Ma bimodal Ross Group of volcanic and sub-volcanic rocks; Fig. 1.4b; Murphy, 1997; Pigage, 2004). Other isolated or clustered examples of igneous rocks of this age are scattered across the inboard regions of the northern Cordillera (Fig. 1.4b). The majority of this latest Cretaceous to Eocene magmatic event in the northern Cordillera has A-type or within plate geochemical affinities. This is consistent with the dextral transtensional tectonic regime, and the extension of over-thickened continental crust along regional structures (e.g., Souther, 1991; Yorath, 1991; Pigage and Mortensen, 2004; Nelson and Colpron, 2007). 21 1.3 RESEARCH OBJECTIVE AND METHODOLOGIES This dissertation aims to provide a comprehensive discussion of the petrogenesis of the syn- to post-accretionary magmatism that was emplaced inboard of the continental margin, in the northern Cordilleran miogeocline. To do so, two regions of the northern Cordilleran were chosen for extensive sampling. Two hundred and twenty-seven samples of plutonic rocks, dykes, veins, and country rocks were collected by K. Rasmussen in the eastern Selwyn Basin over the course of one field season in 2005. An additional 10 samples from the eastern Selwyn Basin (collected by J. Mortensen in 2003), and 8 samples from the northeast Selwyn Basin (collected by C. Leslie in 2006) were also added to this sample suite. Samples from twenty-eight plutonic rocks, and one volcanic rock, were collected by J. Mortensen and K. Rasmussen in the Cassiar Terrane and western Selwyn Basin in 2007. An additional 6 samples from the Yukon–Tanana inlier (collected or compiled by J. Mortensen over the course of several years) contribute U–Pb ages to this sample set. Whole rock powders and light mineral Wilfley table fractions for plutons in the eastern Selwyn Basin from the work of Heffernan (2004) were also available for stable isotope analysis. Representative plutonic and dyke samples from the eastern and northeastern Selwyn Basin, and all plutonic samples from the Cassiar Terrane and western Selwyn Basin, were analyzed by a number of methods to generate several distinct datasets. U–Pb ages determined from zircons and monazites are used to constrain the range of crystallization ages for igneous rocks in the two study regions. Whole rock geochemical (major and trace element) and isotopic (whole rock Sm–Nd, Rb–Sr, and S; Pb–Pb in feldspar; and O in quartz) data used to distinguish suites on the basis of composition, and to provide information on potential magma sources in the crust and mantle. More detailed work in the magmatically complex eastern Selwyn Basin (Fig. 1.4a) employed zircon composition (reconnaissance-scale trace element and Lu–Hf isotopes). Also in the eastern Selwyn Basin, 40Ar/39Ar cooling ages in plutonic biotite (and several mafic dykes), and magmatic fluid compositions (with respect to F, Cl, and OH) to better differentiate between plutonic suites and their petrogenetic histories, and to discuss their potential for associated mineralization. The collection of these data were made possible through collaboration with several research groups and laboratories3, and with the assistance of several sources of funding4. Extensive compilations of geochronological, geochemical, and isotopic data for syn- to post-accretionary igneous rocks in the northern Cordillera were assembled and integrated with the data produced by this study, extending the implications of this study to a larger region. With the new analytical datasets generated during the course of this study, and an extensive compilation of all known relevant data, an updated or modified interpretation of the petrogenesis of syn- to post-accretionary magmatism may be considered. More detailed work on potential magma sources also has broad implications for the age and composition of unexposed northern Cordilleran crust. Ultimately, however, this dissertation aims to 3 The Mineral Deposits Research Unit and the Pacific Centre for Isotopic and Geochemical Research at the University of British Columbia, the Nevada Stable Isotope Laboratory at the University of Nevada-Reno, and the INCO Innovation Centre Microanalytical Facility at the Memorial University of Newfoundland. 4 The Yukon Geological Survey, the Northwest Territories Geoscience Office, Parks Canada, the National Science and Engineering Research Council (through grants to J. Mortensen and K. Rasmussen), and the Society of Economic Geologists (through Canada Foundation Grants to K. Rasmussen). 22 contribute directly to our knowledge of the age, composition, and distribution of this prolific period of magmatism in the northern Cordillera. More generally, this synthesis of syn- to post-accretionary igneous rocks in the northern Cordilleran miogeocline will add to our understanding of intermediate–felsic, syn- to post-accretionary granitoid provinces world-wide. The following sections provide greater detail on the methods and number of samples used to accomplish the above- stated objectives. 1.3.1.1 U–Pb and 40Ar/39Ar Geochronology Geochronological datasets providing accurate control on the age of igneous rocks are invaluable in terms of differentiating between different magmatic events, and for defining the temporal limits of specific magmatic (i.e., plutonic) suites. The foundation of any petrogenetic study of igneous rocks must be precise and reliable age determinations for the rocks studied. Furthermore, any study utilizing the Sm–Nd and Rb–Sr radiogenic isotopic systems requires good geochronological control on the age of the samples analyzed to allow for the correction of isotopic decay following crystallization of the intrusions. Thus, a major focus of this study was to contribute a large set of U–Pb geochronological data in regions of the Selwyn Basin and Cassiar Terrane for which limited robust geochronological datasets existed and (or) that previously had proved difficult to date using Isotope Dilution – Thermal Ionization Mass Spectrometry (ID–TIMS) methods because of inheritance and (or) lead-loss problems in zircon. In such instances, single spot or line analysis targeting pristine magmatic zircon—revealed by back-scattered electron (BSE) or cathodoluminescent (CL) images—is more likely to produce concordant ages. The Sensitive High Resolution Ion Microprobe (SHRIMP), and Laser Ablation – Inductively Coupled Plasma – Mass Spectrometry (LA–ICP–MS) are both capable of the localized analysis of zircon. Seventy-seven new U–Pb ages for intermediate to felsic plutonic rocks (and two felsic dykes) across the northern Cordilleran miogeocline are presented in this dissertation. These data were predominantly collected using LA–ICP–MS methods (62 samples), although samples collected prior to this study by J. Mortensen were analyzed by ID–TIMS methods. Twenty-six 40Ar/39Ar plutonic ages were also determined for biotites from plutonic rocks in the eastern Selwyn Basin, and five 40Ar/39Ar ages were determined for biotites from mafic dykes in the eastern Selwyn Basin. 1.3.1.2 Whole Rock Geochemistry Detailed major and trace element geochemistry are standard for studies involving igneous rocks. Following the sub- division of plutonic suites on the basis of field relationships and crystallization ages, major and trace element geochemical characterizations are typically used to describe the basic compositional range exhibited by the various plutonic suites. Geochemical compositions are also used to make inferences regarding the tectonic environment, source(s) of magma, and the effects of differentiation (assimilation, fractional crystallization, and contamination) processes for individual plutons, and for plutonic suites as a whole (e.g., Christiansen and Keith, 1996). Major and trace elemental variations provide useful information on the main controls on magma compositions, whereas trace element variations are more commonly used for petrogenetic interpretations. In particular, trace elements normalized 23 to a generalized frame of reference using one of Earth’s many inferred reservoirs (e.g., chondritic or bulk Earth, primitive mantle, mid-ocean ridge basalt, upper continental crust, or even an igneous liquid-line of descent) reduces the effects of higher concentrations of some elements and allows us to compare compositions between plutonic suites and to identify geochemical anomalies. Furthermore, geochemical and tectonic discrimination plots are useful for classifying plutonic rocks on the basis of major or trace element concentrations compared to existing global datasets, allowing for generalized inferences to be made regarding basic melt source composition(s) and the tectonic setting in which magmatism may have been generated. This study contributes whole rock major element oxide and trace element data from 93 samples of plutonic and dykes rocks in the Cassiar Terrane and Selwyn Basin. 1.3.1.3 Radiogenic Isotopes Radiogenic isotopic datasets (e.g., Rb–Sr, Sm–Nd, and Pb–Pb) are increasingly considered essential information for petrogenetic studies of igneous rocks. Radiogenic isotopic systems are based on the predictable and measurable decay of a radioactive isotope to its radiogenic daughter product. They have been widely applied in igneous studies as tracers for specific source regions or reservoirs in Earth’s crust and processes. Due to the similar ionic behaviour (related to ionic radii and the valency) of both the parent and daughter elements, radiogenic isotopic compositions are not strongly fractionated during partial melting (unlike trace elements, which are also used as petrogenetic indicators). The small degree of isotopic fractionation that does occur, however, imparts variable parent:daughter elemental ratios (e.g., Rb/Sr, Sm/Nd, and U/Pb) that are used to distinguish isotopic reservoirs in the mantle and crust. This ratio, and the amount of time that has passed since the extraction of the melt from the mantle, are the two primary controls on radiogenic isotopic compositions in unaltered igneous rocks. As both factors controlling isotopic composition quantitatively increase, the daughter isotope will become more enriched and the isotopic composition of the igneous rocks will become more radiogenic with respect to that isotopic system. These isotopic decay systems provide valuable information on the origin of igneous rocks, including insight into the general age and composition of the magma source(s). Used together, radiogenic isotopic systems are a powerful petrogenetic tool in the study of igneous rocks. This study contributes radiogenic isotopic data for 45 to 50 (depending on the isotopic system) plutonic and mafic dyke samples. 1.3.1.4 Stable Isotopes Stable isotopic compositions of igneous rocks, although less frequently employed than radiogenic isotopes, also provide information on crustal sources of magma and (or) contamination (e.g., Taylor and Sheppard, 1986). Unlike the radiogenic isotopes, stable isotopes are only are minimally affected by high temperature partial melting and magmatic processes, but are strongly fractionated by low temperature processes at the Earth’s surface. Due to the high abundance of oxygen in the mantle and crust, the oxygen isotopic composition of magmas typically reflects the composition of the melt source(s). ‘Heavy’, or positive, oxygen compositions in igneous rocks are suggestive of a derivation by partial melting of rocks comprising materials that were weathered at the Earth’s surface (i.e., supracrustal rocks), or rocks that were affected by low temperature hydrothermal alteration near the Earth’s surface (i.e., altered volcanic rocks, including oceanic ‘basalts’). Conversely, sulfur concentrations are low in all magmas 24 (especially those derived from the mantle). Thus, minimal amounts of assimilation or contamination of a magma by sulfur-rich lithologies in the crust (e.g., evaporitic or pyritic sediments) can substantially alter the sulfur isotopic composition of a magma with very little effect on the other isotopic systems (Poulson et al., 1991). Thus, the sulfur isotopic system is more sensitive than the oxygen and radiogenic isotopic systems to crustal contamination. Stable isotopic compositions of igneous rocks clearly provide complimentary, and in some cases more specific information than radiogenic isotopic compositions. This study contributes stable isotopic data for 100 (O) and 105 (S) igneous samples.5 1.3.1.5 Zircon Compositions Zircon compositions are a relatively new tool used for deciphering magmatic processes and the petrogenesis in igneous rocks (e.g., Woodhead et al., 2001; Schmitz et al., 2004; Kemp et al., 2005). This is partly because in situ analytical techniques (e.g., SHRIMP, LA–ICP–MS) are still relatively new (e.g., Kinney and Mass, 2003). Zircon is a particularly useful mineral for isotopic dating and tracer studies because it contains detectable quantities of uranium, thorium, and lead, as well as lutetium, hafnium, samarium, neodymium, and oxygen (e.g., Corfu et al., 2003; Hoskin and Schaltegger, 2003; Kinney and Maas, 2003). This combination of trace element contents and isotopic compositions (including the ability to accurately determine the age of the mineral) make zircon analysis a potentially important petrogenetic tool for igneous rocks. When combined with isotopic compositions (e.g., Lu–Hf or O), variations in zircon geochemical composition become an even more powerful petrogenetic tracer. Zircon morphologies also provide information on the magma composition and temperature (e.g., Pupin, 1980), and titanium concentrations in zircon provide a promising thermometer (Watson et al., 2006; Ferry and Watson, 2007). Finally, additional information on the composition, provenance, and (or) age of the magma protolith(s) may also be obtained by systematically analyzing inherited zircon cores (e.g., Pupin, 2000). This study presents a reconnaissance-scale study of zircon trace element and Lu–Hf isotopic compositions for magmatic and inherited sectors in zircon separated from nine representative samples of the ‘mid’-Cretaceous igneous rocks in the eastern Selwyn Basin. 1.3.1.6 Magmatic Fluid Compositions The study of volatile species in crystallizing magmas (e.g., F-, Cl-, HS-, OH-) is of broad interest to the field of economic geology. In the magmatic–hydrothermal environment, these ionic species may act as ligands that facilitate the transport of metals in exsolved aqueous or vapourous solutions (e.g., Seward and Barnes, 1979; Crerar et al. 1985; Piccoli and Candela, 1994). The ability of a magmatic fluid to transport metals is strongly dependent on the activity or thermodynamic concentration of specific ligand(s) (e.g., Seward and Barnes, 1979; Crerar et al. 1985). Here, the activities of chlorine and fluorine are considered. Chlorine has a strong preference for partitioning into the magmatic fluid (e.g., Webster, 1997a,b), whereas fluorine is partitioned preferentially into minerals, then melt, and least preferentially into fluid (e.g., Zhu and Sverjensky, 1991). The early exsolution of a Cl/F-rich fluid, therefore, should lower the concentration of chlorine relative to fluorine in the magma body, which is then thought to also be 5 The sulfur isotope data produced here comprise one of the largest existing datasets of whole rock sulfur isotopic composition for igneous rocks. 25 reflected in the composition hydrous minerals crystallizing from the magma. The late exsolution of a Cl/F-rich magmatic fluid—which is typical of tungsten skarn mineralization (e.g., Newberry and Swanson, 1986)—instead, should increase the concentration of chlorine relative to fluorine in the magma body and hydrous minerals. These concepts may be tested by analyzing magmatic apatites and biotites using the electron probe microanalyzer (EPM) and using an appropriate thermodynamic calculation scheme that accounts for halogen partitioning not related to the magmatic fluid composition. This allows the relative activities of XF, XCl, and (by stoichiometry) XOH in the magmatic fluid to be estimated. Here, regional magmatic fluid compositions for 25 plutons in the eastern Selwyn Basin document the variation exhibited between the plutonic suites and relative to the degree of differentiation experienced by the sample. Additionally, magmatic fluid compositions for multiple intrusive phases from 5 composite plutons, three of which are associated with tungsten skarn mineralization in the eastern Selwyn Basin, provides information on the evolution of magmatic volatiles within several mineralizing and non-mineralizing intrusions. 1.4 OUTLINE OF THE DISSERTATION This dissertation is composed of eight chapters in total, including this introductory chapter (Chapter 1). A summation of the main topics addressed by Chapters 2–7 is given below. Chapter 8 concludes the dissertation with an overview of the main conclusions reached during the course of this study, and some highlighted questions that remain for future researchers. Chapter 2 – Geochronological and Geochemical Characterization of Syn- to Post-Accretionary Magmatic Suites Emplaced into the Northern Cordilleran Miogeocline Chapter 2 provides the basis for the sub-division of igneous rocks of Cretaceous to Eocene age across the northern Cordilleran miogeocline. This chapter presents the U–Pb geochronological data and whole rock geochemical data generated by this study. These data are integrated with compilations of existing data, and—along with mineralogical, textural, and field characteristics—are used to characterize syn- to post- accretionary magmatism throughout the Cassiar Terrane and Selwyn Basin. These plutonic suites are sub- divided into high-K calc-alkaline to medium-K calc-alkaline, I- and S-type series, and alkaline to alkali- calcic A-type series. General magma source rock compositions and metallogenic signatures are suggested for these magmatic series on the basis of the observed compositional ranges. The spatial and temporal patterns of igneous rocks in the Cassiar Terrane and Selwyn Basin are discussed, separately, relative to continental arc magmatism to the west–southwest. Chapter 3 – Isotopic Characterization of Syn- to Post-Accretionary, Early Cretaceous to Eocene Magmatism in the Cassiar Terrane Chapter 4 – Isotopic Characterization of Syn- to Post-Accretionary, Early to ‘mid’-Cretaceous Magmatism in the Selwyn Basin These two chapters present the radiogenic and stable isotopic compositions of Early Cretaceous to Eocene plutonic suites in the Cassiar Terrane, and for Early Cretaceous to ‘mid’-Cretaceous magmatism in the 26 eastern Selwyn Basin. Extensive compiled isotopic data are also integrated with the results of this study. These data are discussed in more detail in Chapter 3, which also provides a basis for interpretations regarding the age and composition of specific magma sources—both in the crust and in the mantle—for both chapters. Broader implications regarding the age and composition of basement rocks below the Cassiar Terrane and Selwyn Basin are also considered. The petrogenesis of the igneous rocks in these two regions is discussed in light of the conclusions reached by Chapter 2, and incorporates elements of previous models (Section 1.2). A brief description of the potential tectonic setting(s) that may be responsible for the syn- to post-accretionary magmatic events are also suggested for both study areas. Chapter 5 – A Reconnaissance-Scale Study of Zircon Trace Element and Hafnium Isotopic Compositions in the Eastern Selwyn Basin This chapter presents zircon compositions from representative ‘mid’-Cretaceous plutons in the eastern Selwyn Basin where several plutonic suites of similar composition and (or) overlapping age also converge geographically (Fig. 1.4a). Zircon morphologies, and trace element concentrations and ratios in magmatic and inherited sectors provide information on magmatic temperatures and compositions. The hafnium isotopic compositions are also integrated with the isotopic data for the same samples (from whole rock powders and quartz separates) presented in Chapter 4. Chapter 6 – Approximating Magmatic Fluid Compositions of Barren and Tungsten Skarn-Associated Plutons with Apatite and Biotite: case studies from the northern Canadian Cordillera The chapter looks at multiple intrusive phases belonging to several plutons in the eastern Selwyn Basin, including two plutons that do not appear to be associated with mineralization, and three plutons that are associated with tungsten skarn deposits of varying size and economic importance. This detailed study on the evolution of magmatic fluid compositions relative to fluorine, chlorine, and water in these intrusions demonstrates how magmatic fluid compositions evolve with respect to magmatic differentiation. From this, specific intrusive phases that have anomalous magmatic fluid compositions are identified and hypothesized to be related to the main tungsten mineralizing event (based on the current understanding of magmatic– hydrothermal systems and how intermediate–felsic magmas generate large deposits of tungsten; e.g., Newberry and Swanson, 1986). Chapter 7 – Petrogenetic, Thermal, and Fluid Evolution of Magmatism in the Eastern Selwyn Basin: integrating geochronological and mineral compositional datasets Chapter 7 attempts to assess intrusions in the eastern Selwyn Basin on the basis of their potential to be associated with mineralization. It is generally thought that intrusions with lengthier thermal histories have experienced multiple pulses of magma of longer periods of time (e.g., Hart, 2004), which may increase their prospectivity for associated mineralization. Thus, the length of thermal histories (taken to be the difference between the U–Pb crystallization age in zircon, and the 40Ar/39Ar “cooling” age in biotite) are estimated and compared to calculated magmatic fluid compositions. Differences existing between the various plutonic suites in this region that can be related to petrogenesis or associated metallogeny, as well 27 as the utility of thermal histories and magmatic fluid compositions in predicting mineralization, are discussed. 28 Figure 1.1. (next page) Basic elements of northern Cordilleran geology. a) Tectonostratigraphic map; the upper left inset shows the location of the map area relative to Canada. Basinal and platformal rocks of the northern Cordilleran miogeocline have been shaded pale yellow and pale blue, respectively. Shades of grey represent terranes accreted to the western margin of Laurentia in the latest Paleozoic to the Mesozoic. Bolded tectonostratigraphic abbreviations are: SB = Selwyn Basin, YTT = Yukon–Tanana terrane (not distinguished from the Stikinia and Quesnellia terranes), CC = Cache Creek terrane, CP = Cassiar Platform, CT = Cassiar Terrane, MeP = McEvoy Platform, MCE = Misty Creek Embayment, MkP = Mackenzie Platform, MdP = McDonald Platform, and KT = Kechika Trough. Regions of the Selwyn Basin commonly referred to in the text are indicated with the transparent NW, NE, W, and SE labels (YTU = Yukon–Tanana uplands). Geographic abbreviations are: YK = Yukon, AK = Alaska, NWT = Northwest Territories, and BC = British Columbia. Modified from Nelson and Colpron (2007), with contributions from Aitken and Pugh (1984), Cecile et al. (1997), Gordey and Anderson (1993), Gordey and Makepeace (1999), Morrow and Miles (2000), Gabrielse et al. (2006), and Turner and Long (2008). b) Exposed Proterozoic metasedimentary rocks, pre- to syn-magmatic Cretaceous sedimentary rocks, and Neoproterozoic, Paleozoic (Goodfellow et al., 1995), and Quaternary–Recent (Abraham et al., 2001) alkaline igneous rocks. The pale purple shading defines the known sub-surface of the Muskwa Assemblage detected by drilling (e.g., Snyder et al., 2002, 2009). The two red circles highlight areas where Cretaceous sedimentary rocks are preserved in the northern Cordillera. The two dashed red lines illustrate the approximate boundaries of the western edge of the foreland basin during the Albian (112– 99.5 Ma), Cennomanian (99.6–93.5 Ma), and Turonian (93.5–89.3 Ma; Thomson et al., 2011). Neoproterozoic magmatic rocks are: MH = Mount Harper Volcanic Complex, BP = Bonnet Plume intrusions, BR = Bear River dyke, LD = Little Dal ‘basalts’, and PC = Pool Creek syenite. Map lines are as in (a) and undivided Insular, Arctic, and Pacific terranes are indicated by the same shade of dark grey as in (a). 29 30 Figure 1.2. (next page) Approximate surface and sub-surface distribution of upper crustal rocks and Precambrian Laurentian basement in the northern Cordillera after the restoration of dextral Cenozoic displacement along major transpressional structures (modified from Nelson and Colpron, 2007). The western continental margin is represented by the heavy dashed line; it is not known what lay west of this line in the Early Cretaceous. Mainly Cenozoic oroclinal bending in Alaska may have begun by the Early Cretaceous (e.g., Gabrielse and Yorath, 1991), but has not been restored here. Map lines and tectonostratigraphic elements are as in Figure 1.1. The Tintina Fault, along which 430 km of strike-slip movement has been restored, is shown with a bolded solid black line. a) 1.8 Ga to present: latest Paleoproterozoic to Neoproterozoic metasedimentary packages (SG = Supergroup, G = Group, and Assem = Assemblage), and Proterozoic, Paleozoic, and Recent alkaline igneous rocks. The sub-surface extent of the Muskwa Assemblage indicated by drill holes is indicated by pale purple shading. b) 3.8–1.8 Ga: Archean to Late Paleoproterozoic basement domains (modified from Ross, 1991, and Aspler et al., 2003). Basement domains are labeled: Na = Nahanni terrane, FN = Fort Nelson “magnetic high” (of unknown origin; shaded light grey), FS = Fort Simpson magmatic arc or “magnetic high”, H = Hottah Terrane, GB = Great Bear magmatic arc, S = Slave craton, N = Nova domain, K = Ksituan domain, C = Chinchaga domain, BH = Buffalo Head domain, T = Taltson domain, and W = Wabamun (a cross-cutting mylonitic domain associated with the Snowbird Tectonic Zone). The approximate locations of the two drill holes that sampled the Fort Simpson magmatic arc are indicated by the circled crosses. The ‘known’ western sub-surface extent of the Nahanni Terrane is indicated by the red dashed line; the Nahanni terrane is inferred to continue further west to the edge of the Laurentian miogeocline based on geophysical characteristics (e.g., Aspler et al., 2006). 31 32 Figure 1.3. A schematic cross-section of facing lower plate and upper plate rifted margins (after Lister et al., 1991). The main tectonostratigraphic sub-divisions from Figure 1.1a are indicated by red labels (Cecile et al., 1997). Note that this diagram not meant to suggest that rifting occurred between the Selwyn Basin and the Cassiar Platform. 33 Figure 1.4. (next page) Syn- to post-accretionary magmatism and important mineral occurrences and deposits in the northern Cordillera. The white stars locate notable mineral deposits or occurrences that are mentioned in Table 1.1. The study area is outlined in black. Intrusive and volcanic polygons are primarily from Gordey and Makepeace (1999) and Massey et al. (2005), and modified based on the regional maps of Dusel–Bacon et al. (1993), Newberry et al. (1998), Mortensen et al. (2000), Baker and Lang (2001), Murphy et al. (2001), Gordey and Ryan (2005), Hart et al. (2004a), Pigage (2004, 2009), Pigage et al. (2011), and Gordey (in press). Map lines and tectonostratigraphic elements are as in Figure 1.1a. a) Early to ‘mid’-Cretaceous magmatic suites in the northern Cordillera based on the geochronological results of this study, compiled robust age determinations, and the work of Mortensen et al. (2000) and Hart et al. (2004a). The red belts represent continental arc magmatism; the purple belts represent back-arc magmatism; and the orange belts represent post-arc magmatism. The approximate location of plutonic sub-divisions described in the text and summarized in Table 1.1 are labeled: A = Anvil, C = Cassiar, DR = Dawson Range, F–S = Fairbanks–Salcha, H = Hyland, M = Mayo, S = Seagull–Thirtymile, Te = Teslin, To = Tombstone, TR = Tay River, Tu = Tungsten, W = Whitehorse, and YTT = Yukon–Tanana Uplands. b) Late Cretaceous to Eocene magmatic belts and (or) occurrences in the northern Cordillera based on the geochronological results of this study, compiled robust age determinations, and the work of Anderson (1988), Murphy (1997), Mortensen and Hart (2010, unpub. data). The approximate location of plutono– volcanic sub-divisions described in the text and summarized in Table 1.1 are labeled: Ca = Carmacks Group (volcanic > plutonic), RR = Ruby Range (plutonic), NR = Nisling Range (plutonic), CPC = Coast Plutonic Complex (plutonic > volcanic), Sk = Skukum (volcanic), Sl = Sloko (volcanic), WT = Windy Table (volcanic), SL = Surprise Lake (plutonic), and MQ = McQueston (plutonic). 34 35 Table 1.1. Basic characteristics of syn- to post-accretionary, Early Cretaceous to Eocene magmatic suites depicted in Figure 1.4 and discussed in the text. Note that the Livengood plutonic suite is not shown in Figure 1.4. Ages are based on U–Pb data (compiled and this study), unless italicized and denoted with K– Ar, Rb–Sr, or Ar–Ar. Tectonic Setting Suite Age (Ma) Locationa Oxidization, Aluminicity Metallogeny (example) Representative References Early to 'mid'-Cretaceous continental magmatic arc Teslin 123–115 south–central Yukon oxidized, metaluminous typically non–mineralized Hart (1995, 1997); Gordey et al. (1998); Mortensen et al. (2000). continental magmatic arc Whitehorse (associated with the Mount Nansen volcanic suite) 115–104 south–central Yukon oxidized, metaluminous Cu ± Au, Ag, Mo porphyry, skarn (Whitehorse Copper) Souther (1991); Woodsworth et al. (1991); Hart (1995, 1997); Mihalnyuk (1999); Mortensen et al. (2000); Hart et al.(2004a). continental magmatic arc Dawson Range/Casino, (associated with the Mount Nansen volcanic suite) 112–99 west–central Yukon oxidized, metaluminous Cu ± Au, Ag, Mo porphyry, skarn Souther (1991); Selby et al. (1999); Selby and Creaser (2001); Hart et al. (2004a); Mortensen and Hart (2010, unpub. data). back-arc Yukon-Tanana Uplands (offset from the Anvil suite by the Tintina Fault) 109–102 (as young as 94?) east–central Alaska reduced– oxidized, peraluminous– metaluminous Au vein (Pogo); small W or Mo porphyry, skarn Aleinikoff et al. (2000); Werdon et al. (2004); Dilworth et al. (2007); Hart et al. (2004a); Mortensen and Hart (2010, unpub. data). back-arc Cassiar 112–104 south–central Yukon weakly oxidized, peraluminous– metaluminous abundant W–(Mo) or Sn skarn, porphyry, vein (Northern Dancer, Risby, Stormy); distal Ag–Pb–Zn vein; Au manto (Ketza River) Woodsworth et al. (1991); Mortensen and Hansen (1992); Mortensen et al. (2000); Hart et al. (2004a); Mortensen et al. (2007). back-arc Anvil 109–100 south–central Yukon reduced, peraluminous W–(Mo) or Sn–(Ag) skarn, vein; Au vein; distal Ag–Pb– Zn vein Pigage and Anderson (1985); Anderson (1988); Woodsworth et al. (1991); Mortensen et al. (2000); Murphy et al., 2002; Pigage (2004); Hart et al. (2004a). back-arc Hyland (or Hyland River) 106–98 southeast Yukon weakly oxidized, peraluminous abundant W–(Mo) or Sn skarn, porphyry, vein; distal Ag–Pb–Zn vein Mortensen et al. (2000); Hart et al. (2004a); Heffernan (2004); Hart and Lewis (2006). post-arc Seagull–Thirtymile sub-suite of the Cassiar suite 103–96 western CT (Harper Ranch sub-terrane) reduced, peralkaline– peraluminous abundant Sn skarn, vein (JC) Liverton and Alderton (1994); Liverton and Bohtelo (2001); Mortensen et al. (2006b, 2007). post-arc Tay River (associated with the South Fork volcanic suite) 99–96 south–central to south–east Yukon reduced– oxidized, peraluminous typically non-metallogenic; Pb–Zn–Ag skarn, manto, vein; W skarn Wood and Armstrong (1982); Pigage and Anderson (1985); Anderson (1988); Woodsworth et al. (1991); Souther (1991); Mortensen et al. (2000); Heffernan (2004); Pigage (2004). post-arc Tungsten 98–94 southwestern NT reduced, peraluminous W–Cu ± Zn–Bi–Au (Mactung, Cantung) and distal Pb–Zn–Ag veins Anderson (1988); Woodsworth et al. (1991); Mortensen et al.(2000); Hart et al. (2004a,b). post-arc Mayo (or South Lansing) 96–93 east-central Yukon reduced > oxidized, metaluminous Cu–, Au–, or W skarn (Ray Gulch); Sn, Sb–As or U–Th– REE occurrences; Au vein (Scheelite Dome, Dublin Gulch) Anderson (1988); Woodsworth et al. (1991); Mortensen et al.(2000); Hart et al. (2004a,b). post-arc Fairbanks– Salcha (offset from Mayo suite by the Tintina Fault) 94–90 (Ar– Ar, Rb–Sr) east-central Alaska reduced, peraluminous Au ± Ag vein (Fort Knox); small W skarn Mortensen et al. (2000); Hart et al. (2004a,b). post-arc Tombstone 93–90 east-central Yukon, southwestern NT reduced– oxidized (zoned), mainly metaluminous (peralkaline– peraluminous) Cu–, Au–, or W skarn (Marn, Horn); Sn, Sb–As or U–Th– REE occurrences; Au vein (Brewery Creek) Anderson (1988); Woodsworth et al. (1991); Mortensen et al.(2000); Hart et al. (2004a,b). 36 Tectonic Setting Suite Age (Ma) Locationa Oxidization, Aluminicity Metallogeny (example) Representative References post-arc Livengood (offset from Tombstone suite by the Tintina Ft) 92–90 (Ar– Ar) east-central Alaska reduced– oxidized (variably zoned), mainly metaluminous (peralkaline– peraluminous) Sb–precious metal vein; U– Th–REE occurrences Hart et al. (2004a,b). Late Cretaceous pre- (Late K) arc Surprise Lake 84–78 northwest to north–central British Columbia reduced, peralkaline– peraluminous Mo–W porphyry, Sn–(U–W– F) skarn/vein (Ruby Creek) Anderson (1988); Souther (1991); Woodsworth et al. (1991); Mihalnyuk (1999); Ray et al. (2000); Morris and Creaser (2008); Smith (2009). continental magmatic arc ± lithospheric delamination Carmacks (and Windy Table volcanic rocks, and Prospector Mountain plutonic suite) 78–67 (up to 86 Ma in BC) northwest British Columbia, western Yukon, and east–central Alaska oxidized, metaluminous Cu–Mo ± Au porphyry/epithermal (Casino) Woodsworth et al. (1991); Souther (1991); Mihalnyuk (1999); Selby and Creaser, (2001); Simmons et al. (2006, 2007); Mortensen and Hart (2010, unpub. data); Ciolkiewicz et al. (2012). post-arc Allen Stock (at north tip of ca. 69- 82 Ma hydrothermal- metamorphic belt) 76–75 eastern CT oxidized, peraluminous U–REE-porphyry; Pb–Zn– Ag manto (Midway) Gregory (2007); Canadian Geochronological Knowledgebase (2013). post-arc Cassiar Stock and portions of the Cassiar batholith 74–60 (K– Ar, Rb–Sr) eastern CT in north–central British Columbia oxidized, peraluminous Mo ± W porphyry and skarn (Storie, McDame) Cooke and Godwin(1984); Canadian Geochronological Knowledgebase (2013). post-arc Deadman batholith and nearby magmatism 74–64 western CT nd typically non-metallogenic Gordey et al. (1998); Gareau and Mortensen (1992, unpub. data); Gordey and Makepeace (1999); Canadian Geochronological Knowledgebase (2013). Latest Cretaceous to Eocene slab window? Nisling Range 62–52 (K– Ar; Rb–Sr) southwest Yukon oxidized, peraluminous small Mo vein Ridgeway(1973); LeCouteur and Tempelman–Kluit (1976); Templeman-Kluit (1977a,b, 1979); Lynch and Pride (1984); Anderson (1988); Woodsworth et al. (1991); Souther (1991). continental magmatic arc Skukum and Sloko volcanic rocks (and Coast Plutonic Complex and Ruby Range plutonic suite) 64–51 southwest Yukon and northwest British Columbia oxidized, metaluminous– peraluminous Au vein (Mount Skukum) Souther (1971); Lambert (1974); Smith (1983); Souther (1991); Currie (1994); Mihalynuk et al. (1995); Mihalynuk(1999); Simmons et al. (2006, 2007); S. Israel (2012, pers. comm.). post- (Late Cretaceous) arc McQuesten 67–64 northwest SB reduced, peraluminous Sn–Ag–(Au–Bi) vein/breccia Emond (1986, 1992); Emond and Lynch (1990); Murphy (1997). post- (Late Cretaceous) arc Cream and Ivo plugs, and associated Jox, Royall, and Salivo plugs 69–60 (K– Ar) southeast SB reduced, peraluminous W skarn Main and Cathro (1982). back-arc Ross Group (bimodal volcanic and sub-volcanic rocks) 56–54 western SB and YTT inlier reduced, peraluminous base and polymetallic vein (Grew Creek; Sa Dena Hes) Gabrielse (1985); Jackson et al.(1986); Pride (1988); Dilworth (2003); Neufeld (2004); Pigage (2004); Gordey and Ryan (2005); Colpron (2011); Gordey (in press); Geochronological Knowledgebase (2013). post- (Paleocene– Eocene) arc Black River pluton and Tintina plug 49–47 eastern CT adjacent to TF reduced, peraluminous Ag ± Zn, Cu, and (or) Pb vein/replacement Canadian Geochronological Knowledgebase (2013) 37 Tectonic Setting Suite Age (Ma) Locationa Oxidization, Aluminicity Metallogeny (example) Representative References back-arc and post- (Paleocene– Eocene) arc? Mt. Reed, Mt. Haskin, Balourdet, and portions of the Cassiar batholith 58–42 (K– Ar) eastern CT in north–central British Columbia reduced, peraluminous Mo ± W porphyry (Mt. Reed, Mt. Haskin) Christopher et al. (1972); Evenchick (1988); Canadian Geochronological Knowledgebase (2013). back-arc Ting 52 southeast SB oxidized, alkaline typically non-metallogenic Fallas et al. (2004); Pigage and Mortensen (2004); Pigage (2009). Notes: a CT = Cassiar Terrane, SB = Selwyn Basin. 38 2 CHAPTER: GEOCHRONOLOGICAL AND GEOCHEMICAL CHARACTERIZATION OF SYN- TO POST-ACCRETIONARY MAGMATIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE 2.1 INTRODUCTION Voluminous, syn- to post-accretionary, Early Cretaceous and younger plutonic rocks of mainly intermediate to felsic composition have been emplaced into upper crustal Laurentian rocks of Neoproterozoic to Triassic age, which herein will be referred to as the “northern Cordilleran miogeocline” (Fig. 2.1). Exposed rocks of probable North American, or Laurentian, miogeoclinal affinity in the northern Cordillera mainly comprise the miogeoclinal Selwyn Basin and the “Cassiar Terrane” (Fig. 2.1a). Allochthonous and autochthonous tectonostratigraphic entities thought to be underlain by Laurentian basement hosts the majority of the syn- to post-accretionary magmatism discussed here (Fig. 2.1b). These igneous rocks have been broadly subdivided into three main episodes (e.g., Nelson and Colpron, 2007): an Early to ‘mid’-Cretaceous event (ca. 123–90 Ma; Fig. 2.2a), a Late Cretaceous event (ca. 84–67 Ma; Fig. 2.2b), and a latest Cretaceous-Eocene event (ca. 67–47 Ma; Fig. 2.2b) 6. The timing of magmatic episodes is tied to changes in the deformational regime of the northern Cordillera, which in turn coincide with the timing of changes in plate vectors (orientation and speed) along the western margin of North America (e.g., Anderson, 1988; Gabrielse and Yorath, 1991; Woodsworth et al., 1991; Kelley, 1993; Nelson and Colpron, 2007). In the Early to ‘mid’-Cretaceous, the northern Cordillera was experiencing primarily compressional, or transpressional, deformation (e.g., Nelson and Colpron, 2007). Coeval igneous rocks were emplaced in response to oblique convergence along the western margin of North America at this time. During the Late Cretaceous, and especially by the Eocene, magmatism associated with transtensional tectonism became more important (e.g., Nelson and Colpron, 2007). This change in the geodynamic regime also affected the composition of magmatism occurring through the Late Cretaceous to Eocene, particularly in the more inboard regions (i.e., distal to the continental margin). The three syn- to post-accretionary magmatic episodes have been further subdivided into several “suites” (Fig. 2.2). Much of these subdivisions rely on extensive K–Ar and Rb–Sr datasets produced in the mid-1970’s to early 1980’s (e.g., Le Couteur and Tempelman–Kluit, 1976; Anderson, 1988; Souther, 1991; Woodsworth et al.,1991; Gordey and Anderson, 1993; Sinclair, 1995; Canadian Geochronological Knowledgebase, 2013; Yukon Igneous Database, 2013). These data suggested that most of the syn- to post-accretionary igneous rocks in the northern Cordillera, particularly in the northern Cordilleran miogeocline, were broadly Cretaceous in age. Some isolated examples of Paleocene–Eocene igneous rocks and (or) Cenozoic re-setting of the K–Ar and Rb–Sr systems in older rocks were also apparent (e.g., Woodsworth et al., 1991). Many of the K–Ar and Rb–Sr ages, however, are problematic because of the ease with which these two isotopic systems can be partially to fully re-set during post-magmatic hydrothermal or thermal events. Intrusions with high Rb/Sr and (or) plutons that experienced open system behaviour during 6 Reported age ranges are based on U–Pb geochronological data, including the ages resulting from this study. 39 hydrothermal, metamorphic, and (or) later igneous events are particularly difficult to date with the Rb–Sr and K–Ar methods. Recent geochronological work has focused on more robust isotopic systems—such as U–Pb, 40Ar/39Ar, and Re– Os—and, where possible, has employed more refractory mineral phases that are less likely to have been disrupted by sub-solidus processes (e.g., zircon, monazite, titanite). The resulting ages are more accurate and have smaller external errors. This recent geochronological work has been partly focused on constraining metallogenic events. As such, districts with known metallogenic potential have been targeted, including the Pogo and Mount Nansen areas (circled; Fig. 2.1b; e.g., Werdon et al., 2004; Carlson, 1987, Selby et al., 2002; Mortensen et al., 2003); Selby et al., 2002, 2003; Dilworth et al., 2007). Other U–Pb studies have focused on regions that were generally lacking widespread geochronological control to advance our knowledge of Cordilleran geology overall (Fig. 2.1b; Hart, 1995, 1997; Gordey et al., 1998; Mortensen et al., 2000; Heffernan, 2004; Pigage, 2004; Pigage et al., 2011). There are also several, yet unpublished U–Pb ± 40Ar/39Ar datasets coming out of recent projects (Fig. 2.1b; Mortensen and Hart, 2010, unpub. data; Gordey, in press). The regional characterization of syn- to post-accretionary magmatism has gradually improved with the increase in the amount of high quality geochronological, geochemical, and isotopic data available (e.g., Anderson, 1982, 1983; Souther, 1991; Woodsworth et al., 1991; Gordey and Anderson, 1993; Mortensen et al., 2000; Hart et al., 2004, 2005). The data presented here are from two study areas underlain by the northern Cordilleran miogeocline that were chosen for their lack of robust geochronological and regional compositional datasets when this project was initiated (Fig. 2.1a): 1) the Cassiar Platform, Klinkit Assemblage (not distinguished in Fig. 2.1a), and adjacent portions of the Yukon–Tanana terrane—collectively referred to as the Cassiar Terrane—as well as the western Selwyn Basin and the Yukon–Tanana inlier, all in south–central Yukon; and 2) the eastern Selwyn Basin, in southeast Yukon and the southwest Northwest Territories. U–Pb dating was carried out on 77 samples of mainly plutonic bodies in these two study areas (Figs. 2.1b and 2.3a). Most of these were thought to be of Early to ‘mid-‘Cretaceous age (Gordey and Makepeace, 2001)―although there were K–Ar and Rb–Sr (and, in one case U–Pb; Gregory, 2007) indications of the Late Cretaceous to Eocene age magmatic events among the group of intrusions sampled in both the Selwyn Basin (e.g., Royall–Jox–Cream, and Ivo–Salivo; Fig. 2.2b) and in the Cassiar Terrane (e.g., the Black River and Tintina plutons, and portions of the Allen Stock; Fig. 2.2b). The integration of the ages resulting from this study with other U–Pb, 40Ar/39Ar, and Re–Os ages across the northern Cordillera provides an extensive, robust dataset for the classification syn- to post- accretionary magmatism (albeit one that still has regional gaps; Fig. 2.1b). Field observations, macroscopic to petrographic sample descriptions, and geochemical compositions from this study (as well as previously published studies) were also compiled to characterize the syn- to post-accretionary plutonic suites emplaced in the northern Cordilleran miogeocline. The distribution of magmatic belts and suites illustrated in Figure 2.2 and detailed in Table 2.1 are based on the most up-to-date data, including the results of this study. The newly expanded dataset has resulted modifications to some of the previous sub-divisions, and has allowed the 40 introduction of several new syn- to post-accretionary plutonic suites or sub-suites based on timing, geographic distribution, and (or) composition. The results of the new data have broad implications with respect to the petrogenesis and metallogenesis of magmatism across the northern Cordilleran miogeocline, and better illustrate patterns of magmatism in the Selwyn Basin and Cassiar Terrane, allowing the tectono–magmatic regime in each region to be compared. These topics are briefly addressed here, but remain general themes that are further elaborated on with the addition other isotopic, zircon, and magmatic fluid composition datasets throughout this thesis. 2.2 GEOLOGICAL BACKGROUND The northern Cordilleran miogeocline comprises mainly Paleozoic rocks of Laurentian (or North American) affinity that are characterized by easterly sourced, basinal and platformal assemblages deposited along the ancestral western margin of North America (Hansen et al., 1993; Cecile et al., 1997). The two main tectonostratigraphic sub-divisions of the northern Cordilleran miogeocline, shown in Figure 2.1a, are the: 1) the Cassiar Platform, an off-shore carbonate platform deposited as a marginal plateau along an “upper plate rifted margin”, located in south–central Yukon and north–central British Columbia; and 2) the Selwyn Basin, an offshelf basinal package deposited on a highly attenuated “lower plate rifted margin”, located in southeast Yukon and southwest Northwest Territories. The term “miogeoclinal” implies that deposition in these regions occurred along a passive continental margin, although it is unlikely that such a tectonic setting existed in the northern Cordillera through most of the Paleozoic (e.g., Nelson and Colpron, 2007). Despite the implied quiescent sedimentation at the subsiding Paleozoic Laurentian margin, at least three episodes of extension or rifting the produced mafic volcanic rocks of OIB-like composition during this era (Fig 2.1b; e.g., Goodfellow et al., 1995). One major episode of clastic sedimentation interrupted miogeoclinal sedimentation in the mid-Paleozoic, and was marked by an influx of detritus from the north and locally where Devono–Mississippian rifting was occurring in the basement (e.g., Gordey and Anderson, 1993). West of the Cassiar Platform is an accreted, penetratively deformed pericratonic entity comprising three Jurassic oceanic island arc assemblages built on Devono–Mississippian (or older) metasedimentary rocks with Laurentian affinities (the Snowcap Assemblage; Gabrielse, 1991; Piercey and Colpron, 2009). The pericratonic Yukon–Tanana terrane, and adjacent portions of Stikinia and Quesnellia (not distinguished; Fig. 2.1a) are thought to comprise the continental crust that was rifted from the western Laurentian margin during the Early Mississippian opening of the Slide Mountain Ocean (e.g., Stevens et al., 1996; Murphy et al., 2002; Nelson and Colpron, 2007). Between the Cassiar Platform and Yukon–Tanana terrane the Klinkit Assemblage, part of a pericratonic Late Mississippian to Early Permian arc developed on rifted Laurentian crust (Simard et al., 2003). Together, the Cassiar Platform, the Klinkit Assemblage, and portions of the accreted pericratonic terranes located east of the Teslin Tectonic Zone (and west of the Tintina Fault; Fig. 2.1a) comprise what is referred to here as the Cassiar Terrane. The eventual closure of the Slide Mountain Ocean through westerly dipping subduction also led to the obduction of thin slices of Slide Mountain terrane onto the Cassiar Platform, Selwyn Basin, and easternmost Yukon–Tanana terrane (Fig. 2.1a; e.g., Murphy et al., 2002; Nelson and Colpron, 2007; Cook et al., 2004). 41 The Paleozoic Selwyn Basin and Cassiar Platform cover rocks (and obducted pericratonic portions of the Cassiar Terrane west of the Cassiar Platform) are interpreted to thinly overlie North American basement, which comprises mainly late Paleoproterozoic to Neoproterozoic metasedimentary rocks (e.g., Murphy et al., 2001, 2002; Cook et al., 2004; Clowes et al., 2005; Evenchick et al., 2005; Snyder et al., 2009). The Proterozoic metasedimentary packages were deposited during several basin-forming events in northwest Laurentia and are associated with minor mafic alkaline volcanic and intrusive rocks (Fig. 2.1b; e.g., Thorkelson et al., 2005). Older crystalline basement is not exposed in the northern Cordillera, although its presence is inferred on the basis of geophysical surveys and exposures in central to southern British Columbia (Evenchick et al., 1985; Cook et al., 2004; Clowes et al., 2005; Evenchick et al., 2005; Jones et al., 2005). The thickness, extent, and composition of crystalline basement below the northern Cordilleran miogeocline, if it exists, remains open to interpretation (R. Clowes, 2011, pers. comm.). 2.2.1 Neoproterozoic to Cenozoic Structural Evolution of the Northern Cordilleran Miogeocline The broad structural and stratigraphic differences between the Selwyn Basin and Cassiar Platform observed in outcrop (and interpreted to be present at greater depths on the basis of balanced cross-sections and geophysical imaging) were largely inherited from the rifted margin geometry that was a result of the ~760 Ma break-up of the Rodinia super-continent. Rocks of the Selwyn Basin were deposited a lower plate rifted margin (in the sense of Hansen et al., 1993, and Cecile et al.,1997). This lower plate rifted basement to the Selwyn Basin is thought to be characterized by structurally thinned crystalline basement (presumed to be between 1.8 and 2.1 Ga in age) that is overlain by thick sequences of Proterozoic metasedimentary rocks (Cook et al., 2004; Snyder et al., 2009). Conversely, the Cassiar Platform is interpreted to be a dextrally offset, marginal plateau that formed basin-ward of an upper plate rifted margin in northeast British Columbia. The basement to the Cassiar Platform is thought to be disrupted by steeply dipping normal faults, which limited structural thinning during rifting in this region (Hansen et al., 1993; Cecile et al., 1997; Evenchick et al., 2005). Many of the extensional and transform structures formed during the Neoproterozoic rifting event (or possibly were re-activated from earlier extension) appear to have been important controls on later stratigraphic and structural breaks in the northern Cordilleran miogeocline (e.g., Norris, 1985; Gabrielse, 1991; Gordey and Anderson, 1993; Morrow and Miles, 2000; Hart and Lewis, 2006; Mair et al., 2006b). The basin–platform sedimentological transitions related to the geometry of the rifted margin, for example, also appear to have localized Paleozoic volcanism and major faults (e.g., the Dawson thrust fault and the Tintina Fault; Fig. 2.1). Paleozoic tectonic ‘quiescence’ ended in the latest Permian to Early Triassic with the first of several accretionary events (Nelson and Colpron, 2007; Beranek and Mortensen, 2011). Ongoing northeast-directed convergence along the western margin of Laurentia, and the accretion and obduction of pericratonic, terranes, and exotic terranes continued throughout most of the Mesozoic, terminating sometime between the Late Jurassic and the Late Cretaceous with the accretion of the Insular Superterrane (Trop and Ridgway, 2007). This compressive regime produced a northeast-verging and migrating, thin-skinned fold and thrust belt (Gabrielse and Yorath, 1991; Gordey and Anderson, 1993; Nelson and Colpron, 2007). Paleozoic rocks in both the Selwyn Basin and Cassiar Platform 42 experienced up to 50% shortening (Gabrielse, 1991; Gabrielse and Yorath, 1991; Gordey and Anderson, 1993; Gordey, in press). However, the lack of high relief in much of the northern Cordilleran miogeocline is compatible with an upper crustal detachment, and suggests that the Neoproterozoic metasedimentary rocks were not as affected by the thin-skinned deformation (e.g., Gordey, in press). In contrast to penetratively deformed western portions of the Cassiar Terrane (i.e., the Yukon–Tanana terrane), the miogeoclinal rocks have retained their primary compositional layering and only display evidence for one compressive phase of deformation (Greenwood et al., 1991; Murphy et al., 2002). The low regional metamorphic grades—which range from chlorite-zone greenschist (Proterozoic strata) to sub-greenschist (Paleozoic and younger strata) facies (Greenwood et al., 1991; Gordey and Anderson, 1993)—indicate that the exposed crustal rocks were not buried to depths greater than ~15 km. The syn- to post-accretionary granitoids post-date thin-skinned deformation in their host rocks (Pigage and Anderson, 1988; Gordey and Anderson, 1993; Murphy, 1997; Gordey, in press), although this structural thickening most likely contributed to the generation of these intrusive rocks. Dextral strike-slip movement along continental margin–parallel structures beginning at about 140 Ma also appears to have been important in terms of focusing, and perhaps generating, magmas in both the Early Cretaceous transpressional to Eocene transtensional tectonic regimes (Figs. 2.1 and 2.2; Gabrielse, 1991; Mihalynuk et al., 1999; Wyld and Wright, 2001; Gabrielse et al., 2006; Mair et al., 2006b). Post-magmatic dextral displacements along regional structures, some of which appear to have penetrate the modern-day crust (e.g., the Tintina Fault; Clowes et al., 2005), may have displaced many of the syn- to post- accretionary plutons from their underlying lithosphere. Similarly, movement along basal detachment faults during northeast-directed deformation in the northeast Cordillera (that continues to this day; Hyndman et al., 2005a) has probably also contributed to the displacement of igneous rocks from their roots (e.g., Gordey and Anderson, 1993). 2.2.2 Syn- to Post-accretionary Magmatism in the Northern Cordilleran Miogeocline The three main magmatic events in the northern Cordillera have been sub-divided on the basis of age into Early to ‘mid’-Cretaceous, Late Cretaceous, and latest Cretaceous to Eocene events. The widespread Early to ‘mid’- Cretaceous episode comprises most of the plutonic activity in the northern Cordilleran miogeocline and generally ranges within 115–90 Ma in the Cassiar Terrane and Selwyn Basin (Figs. 2.2a and 2.3b), although in the Cassiar Terrane intrusive rocks ranging from 82–75 and 49–47 Ma were also identified in the course of this study (Figs. 2.2b and 2.3b). These intrusions were emplaced in different tectonic regimes related to plate motions and convergence along the western margin of North America. Early to ‘mid’-Cretaceous magmatism has mainly been attributed to partial melting of over-thickened continental crust forming the basement to the Cassiar Terrane, due to one, or a combination of, the following processes: • over-thickening of the compression-driven orogen (Woodsworth et al., 1991; Driver et al., 2000), • decompression melting during post-orogenic collapse and associated exhumation (Hart et al., 2004a; Mair et al., 2006a; Dilworth et al., 2007), or • localized decompression melting along regional dextral transpressional structures (Gabrielse et al., 2006; Colpron et al., 2011). 43 Associated coeval continental arc magmatism was also active through much of this time (e.g., Hart, 1995, 1997; Mortensen et al., 2000; Hart et al., 2004a), suggesting that that these mainly intrusive rocks in the northern Cordilleran miogeocline were emplaced into a back-arc setting. Late Cretaceous igneous rocks are generally thought to comprise crustal melts from orogenically thickened crust undergoing dextral transtension related to highly oblique subduction along the western margin of North America (e.g., Engebretson et al., 1985; Hart, 1995, 1997; Murphy, 1997; Ray et al., 2000; Nelson and Colpron, 2007; Mortensen and Hart, 2010; Gordey, in press). Voluminous continental arc magmatism in western Yukon and northwest British Columbia may also have been associated with more inboard plutonic rocks in the northern Cordilleran miogeocline (Morris and Creaser, 2003; Simmons et al., 2006, 2007). Like the Late Cretaceous magmatism, examples of latest Cretaceous to Early Tertiary igneous rocks are also interpreted to represent crustal melts related to dextral transtensional deformation of previously thickened crust (e.g., Murphy, 1997; Gabrielse et al., 2006). The bimodal volcanic rocks in particular appear to have been erupted into pull-apart volcano–sedimentary basins related to movement along dextral strike-slip faults (Jackson et al., 1986; Gabrielse, 1991; Gordey, in press). 2.2.2.1 Estimated Depths of Emplacement for Syn- to Post-accretionary Magmatism Most of the large syn- to post-accretionary intrusions in the northern Cordilleran miogeocline were emplaced into predominantly quartzo–feldspathic metasedimentary rocks of the Neoproterozoic Windermere Supergroup (e.g., the Hyland and Ingenika groups in the Selwyn Basin and Cassiar Terrane, respectively; Fig. 2.3a). These batholiths— which include the Glenlyon, Big Salmon, Nisutlin, Cassiar, Anvil, and Mount Billings intrusions (Fig. 2.3a)— typically have wide, amphibolite-grade (garnet to staurolite to andalusite to sillimanite ± cordierite) contact aureoles with schistose, gneissic, and (or) migmatitic textures (Roots et al., 1966; Smith and Erdmer, 1990; Greenwood et al., 1991; Mortensen and Hansen, 1992; Gordey and Anderson, 1993). Although these higher-grade contact aureoles may be partly related to a greater depth of magma emplacement, they could instead be the result of the emplacement of a large, slow-cooling, body of hot magma into country rocks metamorphosed to regional greenschist facies. Regional chlorite-zone greenschist facies metamorphic grades suggest that batholithic intrusions in Neoproterozoic rocks were emplaced at depths between 12 to 15 km (Greenwood et al., 1991; Moynihan, 2013). Smaller plutonic bodies are more commonly hosted in, or near miogeoclinal Paleozoic rocks (Figs. 2.1b and 2.3a). These smaller plutons may represent higher-level portions of batholiths contained within the Neoproterozoic rocks that shallowly underlie the Paleozoic miogeocline (e.g., Gordey and Anderson, 1993; Cook et al., 2004; Gordey, in press). Contact aureoles containing biotite with andalusite, sillimanite, or cordierite up to several 100’s of m thick enclose most of the smaller syn- to post-accretionary plutons (e.g., Gordey and Anderson, 1993; Baker and Lang, 2001). These mineral assemblages, together with the presence of miarolitic cavities in some intrusions and the results of fluid inclusion studies (Baker and Lang, 2001; G.R. Thompson, unpub. data, 2000; Stephens et al., 2004), suggest that emplacement of most of the smaller intrusions in sub-greenschist-grade Paleozoic host rocks occurred depths of 8 km or less (Anderson, 1982, 1985; Gordey and Anderson, 1993; Murphy, 1997). 44 2.2.2.2 Previously Defined Plutonic Suites in the Cassiar Terrane and Selwyn Basin All Mesozoic (and younger) plutonic rocks in the Cassiar Terrane were historically classified as the Cassiar plutonic suite (e.g., Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004a), although evidence for less voluminous Late Cretaceous and Eocene plutonism in the same region also existed (see references in Table 2.1). Much of the previous work on magmatism in the western Selwyn Basin initially grouped these intrusions into the Anvil plutonic suite (Pigage and Anderson, 1985), although the later recognition of a younger, less siliceous episode was subsequently distinguished as the Tay River plutonic suite (Mortensen et al., 2000; Pigage, 2004). Prior to this study, controls on the age of plutonic rocks in the eastern Selwyn Basin was limited mainly to several published and unpublished 96–77 Ma K–Ar ages compiled by Anderson (1982, 1983), mainly from the northeast Selwyn Basin. Magmatism in this region was initially grouped into the all-encompassing ‘mid’-Cretaceous Selwyn plutonic suite (Anderson, 1983; Gordey and Anderson, 1993), which was noted to be compositionally and temporally similar to rocks of the Anvil and Tay River plutonic suites in the western Selwyn Basin (Anderson, 1988; Gordey and Anderson, 1993; Pigage and Anderson, 1988). The Selwyn plutonic suite was subdivided on a mineralogical basis into three groupings: two-mica (muscovite ± biotite) or hornblende-free plutons (further subdivided into composite and homogenous types); transitional plutons (biotite > hornblende); and hornblende- bearing (± biotite) plutons. The sub-divisions were described on the basis of textural, mineralogical, geochemical, and strontium isotopic characteristics from four representative bodies (Anderson, 1982, 1983; Gordey and Anderson, 1993). However, just as magmatism in the western Selwyn Basin was subdivided into two distinct plutonic suites (Mortensen et al., 2000; Pigage, 2004), more recent U–Pb, geochemical, and isotopic work on ‘mid’- Cretaceous intrusions across the northwest and north–central Selwyn Basin in Yukon led to the introduction of several new plutonic suite names in this region (e.g., Fig. 2.2a). These newer plutonic suites continue into the eastern Selwyn Basin (Mortensen et al., 2000; Hart et al., 2004a, b; Heffernan, 2004), where correlations have been made by extrapolation of the plutonic suites to the east and south. These cross-border correlations were aided by limited geochemistry, and K–Ar and Rb–Sr geochronological data (that are anomalously younger or older than newer U–Pb crystallization ages) for the northeast Selwyn Basin (Anderson, 1982, 1983; Gordey and Anderson, 1993), and by U–Pb ages, geochemistry, and radiogenic isotopic compositions reported for the southeast Selwyn Basin (Heffernan, 2004). This study provides a substantially improved characterization of the plutonic suites in the eastern Selwyn Basin and Cassiar Terrane. 2.3 METHODS AND RESULTS 2.3.1 Geochronology U–Pb crystallization ages have been determined from zircon or monazite extracted from 74 plutonic, one volcanic, and two felsic dyke samples of known or suspected Cretaceous age across the study area (Fig. 2.3a). Additionally, 45 ages for five mafic dykes ranging in composition from quartz dioritic to kersantitic (shoshonitic) lamprophyres in the eastern Selwyn Basin (Fig. 2.3a), have also been determined by 40Ar/39Ar dating of biotite. Sample locations, and macro- and microscopic details for each sample are tabulated in Appendix A.1. All sample preparation and analyses were done at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia (UBC). A heavy mineral concentrate was separated from samples using conventional crushing, grinding, and wet shaking table techniques. Heavy liquid and magnetic separation was then done in order to further concentrate zircon, monazite, and titanite for U–Pb analysis7. U–Pb geochronology was done primarily by Laser Ablation – Inductively Coupled Plasma – Mass Spectrometry (LA–ICP–MS), although some samples were analyzed by Isotope Dilution – Thermal Ionization Mass Spectrometry (ID–TIMS) methods. Additional details concerning sample preparation and these analytical methods are provided in Appendix B.1, and detailed descriptions of the zircon separates can be found in Appendix B.2. A summary of zircon textures observed in back-scattered electron (BSE) and cathodoluminescent (CL) images obtained from the scanning electron microprobe (SEM) for all samples analyzed by LA–ICP–MS methods, as well as images of representative zircon grains for each sample, can be found in Appendices B.3 and B.4, respectively. All isotopic ages and concentration data for the individual analyses used for final age calculations (as well as conditions of analysis), and weighted mean plots (LA–ICP–MS) and concordia (ID–TIMS) plots for calculating crystallization ages, are presented in Appendices B.5 through to B.8. Biotite fractions for 40Ar/39Ar analysis were hand-picked from a light mineral concentrate taken off the wet shaking table. Isotopic ratios, ages, associated uncertainties for individual steps, and plateau and concordia plots for the 40Ar/39Ar plateau ages of five mafic dykes are presented in Appendices C.1 and C.2. Final interpreted U–Pb crystallization ages for plutons and felsic dykes (as well as those of any detected older magmatic or inherited cores) and final 40Ar/39Ar ages for mafic dykes are reported in Table 2.2. 2.3.1.1 Summary of Geochronological Results The geochronological results of this study (Table 2.2), together with an extensive dataset of compiled robust geochronological data (Appendix G.1), have led to several modifications to the previous sub-divisions for igneous 7 Heavy accessory mineral phases were also hand-picked from the magnetic fractions (initially separated for LA– ICP–MS analysis), mounted in epoxy, and examined on the SEM, with reconnaissance mineral identifications tabulated in Appendix A.1. The identification of accessory minerals by SEM and petrographic methods have allowed us to characterize the accessory mineral assemblages of each of the plutonic suites. These are in many cases related to some the geochemical variations exhibited by the plutonic suites. Variable quantities of zircon and apatite are ubiquitous in every sample, and iron–titanium oxide minerals (e.g., magnetite, ilmenite, and titanite) were detected in nearly all of the samples in varying proportions related to the redox state of the intrusion. Additionally, cerium-rich allanite and monazite are present in most samples, and (in approximate order of decreasing abundance) iron- or calcium-rich garnet, rutile (frequently Nb–Ta-rich), xenotime (frequently Ta-rich), tourmaline, pyroxene, thorite, uraninite, niobium–iron–fluorine-rich titanite, epidote, scheelite, wolframite, at least one unidentified REE- bearing oxide, betafite, topaz, and corundum were also detected. The accessory mineral phases are not detailed in this chapter, except when necessary for petrogenetic interpretations, but they have been listed in Table 2.3 for reference purposes. 46 rocks in the northern Cordilleran miogeocline. Our current knowledge of the age and distribution of igneous rocks in the northern Cordilleran miogeocline is summarized below and in Table 2.1 (where a more extensive list of relevant references is given). Some plutonic suite age ranges have been slightly modified from previous interpretations (Mortensen et al., 2000; Hart et al., 2004, 2005; Heffernan, 2004) and several new plutonic suites in the Cassiar Terrane are introduced. The 115–104 Ma Cassiar plutonic suite has been found to be much older than initially reported (110-98 Ma; mainly based on K–Ar ages; see Table 2.1 for references), and several intrusions previously included within this suite are either Late Cretaceous or Eocene in age (Table 2.2). The older age range determined for the Cassiar suite plutons also led to the temporal distinction of several “ultra-fractionated”, highly siliceous, and strongly reduced granitoids in the southwest Cassiar Terrane (Liverton and Alderton, 1994). Previously published U–Pb ages for what was known as the Seagull–Thirtymile “granite” sub-suite ranged from 103 to 96 Ma (see Table 2.1 for references). Here, the Seagull–Thirtymile “granites” are considered to be temporally, spatially, and compositionally distinguishable from the Cassiar plutonic suite, and are referred to herein as the 103–96 Ma Seagull plutonic suite (Fig. 2.2a and 2.3b). Although only two samples of this suite were successfully dated by this study, other researchers have produced comparable ages for compositionally similar plutons forming a narrow, north-trending belt in the western Cassiar Terrane (e.g., Gallagher, 1999; Mortensen et al., 2007; Piercey and Colpron, 2009; M. Colpron, 2011, pers. comm.). Together, these coeval and compositionally similar intrusions comprise the Seagull plutonic suite (Figs. 2.2a and 2.3b). Northeast of the Tintina Fault, several analyses of the Anvil plutonic suite plutons in the Yukon–Tanana inlier ranged from ca. 114–111 Ma (Fig. 2.3b). These are the first published U–Pb ages for Cretaceous plutons in the Yukon–Tanana inlier, although they had been previously compiled by the Canadian Geochronological Knowledgebase, 2013 as unpublished ages. This range of ages for the Anvil plutonic suite in the Yukon–Tanana inlier demonstrate that these plutons are actually older than the 109–104 Ma Anvil plutonic suite in the western Selwyn Basin (Pigage, 2004; Gordey, in press). The 106–100 Ma Hyland plutonic suite (Heffernan, 2004) in the eastern Selwyn Basin (Figs. 2.2a and 2.3b) was not sampled by this study, but will be discussed regardless. One analysis each of a Tay River suite pluton and nearby coeval South Fork volcanism in the western Selwyn Basin (Fig. 2.3b) returned ages that are consistent with the U–Pb age range of 99–96 Ma for the Tay River plutonic suite rocks and 97–(98) Ma for the South Fork volcanic rocks (see Table 2.1 for references)—and well within the 100–96 Ma range of 40Ar/39Ar ages for biotites from these rocks (Mortensen and Ballantyne, 1992; Mortensen et al., 2000; Pigage, 2004; Gordey, in press). The Anvil and Tay River magmatic events in the western Selwyn Basin, therefore, were approximately coeval to the Cassiar and Seagull plutonic suites, respectively. Intrusions in the eastern Selwyn Basin that had been previously classified as the Tay River plutonic suite (Heffernan, 2004) also returned coeval ages of 99–96 Ma in this study (Fig. 2.3b). The Tay River suite in the eastern Selwyn Basin includes examples of coeval, but compositionally distinct plugs of mainly quartz monzodiorite composition (herein referred to as the Coal River sub-suite) and quartz monzodiorite dykes, and it lacks associated volcanic rocks (although this may be a factor of the relative level of exhumation). Geographically and temporally 47 overlapping with northeast expressions of the Tay River suite is a group of small to moderate-sized, biotite ± muscovite monzogranitic plutons of the 98 to 95 Ma Tungsten plutonic suite (Fig. 2.3b). The northernmost examples of the Tungsten plutonic suite geographically (and temporally) overlap with the 98 to 95 Ma Mayo plutonic suite (or “South Lansing” suite; Mortensen et al., 2000; Fig. 2.2a), which is distinguished on a first order basis from the Tungsten suite by its biotite ± hornblende ± clinopyroxene bearing mineralogy and less siliceous compositions. Late-stage shoshonitic lamprophyre dykes have been reported in association with Mayo suite plutons across the Selwyn Basin (Mortensen et al., 2000; Hart et al., 2004, 2005; Mair et al., 2011). In the eastern Selwyn Basin, however, kersantite (shoshonitic) lamprophyres are spatially and temporally associated with both the Tungsten and Mayo plutonic suites (Fig. 2.3b). The northernmost and easternmost intrusions in the Selwyn Basin belong to the Tombstone plutonic suite (Fig. 2.2a), which ranges from 94–91 Ma (but clusters at ~93 Ma; Fig. 2.3b) in the eastern Selwyn Basin. The Tombstone suite plutons are typically rounded, and range from sub-alkalic to moderately alkalic monzonitic to monzogranitic phases in the eastern Selwyn Basin (Fig. 2.2a). This suite is unique, however, because it is associated with less abundant syenite, monzodiorite, gabbro, and silica-undersaturated intrusive phases outside of the study area in the northern Selwyn Basin. Collectively, the Tungsten, Mayo, and Tombstone suites are known as the Tombstone–Tungsten Belt (TTB; e.g., Hart et al., 2005). Two younger episodes of magmatism in the Cassiar Terrane were also confirmed by the new U–Pb ages. Late Cretaceous intrusive rocks in the northern Cordilleran miogeocline are markedly less voluminous than Early to ‘mid’-Cretaceous intrusive rocks (Fig. 2.2), and are mainly confined to the eastern Cassiar Terrane. Late Cretaceous magmatism forms a 82–76 Ma belt that includes the Gravel Creek and Meister Lake plutons (KR-07-23 and -15; Fig. 2.3), as well as a 75–77 Ma (U-Pb; Gregory, 2007, and this study) phase of the composite ‘mid’-Cretaceous Cabin Creek pluton (or Allen Stock; Figs. 2.2b and 2.3). A series of Late Cretaceous plutonic, hornfels, and hydrothermal mica and whole rock ages (ranging from 69 to 82 Ma; K–Ar and Rb–Sr ages; Canadian Geochronological Knowledgebase, 2013) has been obtained from a discrete belt that extends from the Gravel Creek pluton to at least the Midway lead-zinc deposit in north–central British Columbia, and possibly as far south as the Cassiar Stock (Fig. 2.2b; Cook and Godwin, 1984; Canadian Geochronological Knowledgebase, 2013). Late Cretaceous plutonism, and associated ~69–82 Ma hydrothermal and thermal metamorphism in the Cassiar Terrane, is assigned here to the Rancheria plutonic suite (Fig. 2.3b). Similar K–Ar and U–Pb ages have been obtained for biotite-bearing granite phases of the Deadman Creek batholith and nearby intrusions in the western Cassiar Terrane (Fig. 2.2b; Gordey et al., 1998; Canadian Geochronological Knowledgebase, 2013), but these intrusions appear to be compositionally distinct from the Rancheria suite (based on geochemical data from Morris and Creaser, 2008) and are excluded from this designation. There are no Late Cretaceous intrusions of similar age in the Selwyn Basin (Fig. 2.2b). Latest Cretaceous to Eocene magmatism in the northern Cordilleran miogeocline is also volumetrically small, but is more widely scattered. One sample each of the Tintina and Black River plutons immediately southwest of the Tintina Fault produced Eocene ages (49–47 Ma), which are similar to previous K–Ar ages for the same intrusions (Canadian Geochronological Knowledgebase, 2013). This localized Eocene episode of plutonism is referred to here as the Black River plutonic suite (Fig. 2.3b), but may be correlative with other scattered Eocene plutons in northern 48 British Columbia, including the ~50 Ma Mount Reed and Mount Haskin plutons (Fig. 2.2b; Christopher et al., 1972; Canadian Geochronological Knowledgebase, 2013) and the ~42 Ma Balourdet pluton (Evenchick, 1988). Older examples of latest Cretaceous to Eocene intrusions are also present in the western Selwyn Basin (Fig. 2.2b), specifically the 67–63 Ma intermediate to felsic McQuesten plutonic suite (Murphy, 2004) and the Rau or Rackla pluton (Thiessen et al., 2012, unpub. data), and scattered Paleocene to earliest Eocene, sub-volcanic quartz–feldspar porphyries and associated bimodal volcanic rocks and feeder dykes of the Ross Group (Fig. 2.2b)8. Suggestions of Late Cretaceous felsic magmatic activity in the eastern Selwyn Basin include two Latest Cretaceous to Earliest Paleocene K–Ar ages (60.3 ± 1.0 Ma and 69.4 ± 1.2 Ma; D. Archibald, 1981, unpub. data) for biotites from the Ivo and Cream leucogranites (Fig. 2.2b). Despite reports of these latest Cretaceous to Paleocene K–Ar ages in the eastern Selwyn Basin, no U–Pb evidence for magmatism younger than ~98 Ma exists in these samples (this study; J. Mortensen, 2009, unpub. data; Figs. 2.2b and 2.3b). These micas must have been re-set by a subsequent thermal event, such as the emplacement of younger, but buried or undiscovered, magmas (see Section 7.4.1.2, Chapter 7). Younger igneous rocks farther to the southeast may provide an example of the types of magmas that could be responsible for this thermal re-setting. These are the ~52 Ma biotite-bearing syenite plugs and breccias of the Ting plutonic suite (Fig. 2.2b; Pigage and Mortensen, 2004). 2.3.2 Major and Trace Element Chemistry A total of 93 representative multi-kilogram samples from plutons and dykes throughout the Selwyn Basin and Cassiar Terrane study areas were also submitted for geochemical analysis at ALS Chemex, Vancouver. Sample locations, and macro- and microscopic details for each sample are tabulated in Appendix A.2. The crushed, split, and finely pulverized (75 μm) sample material underwent analysis for major element oxide concentrations by lithium borate fusion and X-ray fluorescence (XRF) spectroscopy, for ferrous iron content by HF–HCl acid digestion and titrimetric finish, and for rare earth and trace element concentrations by lithium borate fusion and ICP–MS. All geochemical data generated by this study and obtained from ALS Chemex are reported in Appendix D.1. Rare earth and trace element concentrations were found to be remarkably consistent with a subset of thirty samples from the eastern Selwyn Basin that were also analyzed by high resolution ICP–MS methods at the PCIGR at UBC (this study, Appendix D.3). An accuracy analysis for the chemical data presented here is provided in Appendix D.2, and compiled geochemical data are tabulated in Appendix G.2. 2.3.2.1 Basis of Geochemical Composition and Classification of the Plutonic Suites9 In addition to well-defined age ranges and geographic distributions, the plutonic suites are also characterized by fixed range of geochemical compositions. Major element abundances are used to broadly classify the igneous rocks, whereas trace element abundances provide information on the source of the parent melt(s) of magmatic suites and 8 Similar-age volcanic and sub-volcanic rocks have also been reported in the Yukon Tanana Uplands region of east– central Alaska (e.g., Dilworth et al., 2005) and localized along dextral transpressional faults in the western Cassiar Terrane and in northern British Columbia (Fig. 2.2b). 9 The term “granite” is used senso stricto (interchangeably with monzogranite or syenogranite), whereas the terms “granitoid” or “granitic” include all silica-saturated, coarse-grained intrusive igneous rocks. 49 the tectonic environment in which they were emplaced. All geochemical interpretations presented and discussed here are based on plutonic samples and any associated mafic dykes. Data from felsic dykes were excluded from plots in order to minimize anomalous compositional variations exhibited by late-stage and highly fractionated phases. The same geochemical plots were prepared for each plutonic suite to better compare and contrast inter-suite differences. Additionally, due to the width of the Selwyn Basin, several plutonic suites emplaced in this region were subdivided into geographic groupings to identify any significant intra-suite mineralogical, textural, and geochemical (and isotopic, as discussed in Chapters 3 and 4) variations. Some of the observed compositional variations were sufficiently distinctive to warrant the proposed “sub-suites” described below. No significant within-suite variations were observed in the Cassiar Terrane. The compiled geochemical data include multiple datasets produced by various analytical techniques (although the vast majority of trace element data was acquired using ICP–MS methods) in different geochemical laboratories. In any study incorporating compiled datasets, there is the potential for anomalous data (particularly for trace elements). Perhaps the most obvious trace elemental variation observed in several of the compiled geochemical datasets here is exhibited by zirconium. Zirconium is commonly employed in petrogenetic interpretations (e.g., high vs. low temperature I-type magmas; Chappell et al., 1998, 2000), magma type discriminatory plots (e.g., Zr vs. 10000*Ga/Al for the discrimination of A-type magmas; Whalen et al., 1987), and calculations (e.g., zircon saturation temperature; Watson and Harrison, 1983). Due to anomalously elevated zirconium concentrations in several of the compiled ICP–MS geochemical datasets, however, this element was only used for petrogenetic interpretations if XRF datasets were available instead. The zirconium versus silica Harker variation diagram shown in the figures that follow distinguishes zirconium concentrations that were determined by XRF analysis from those that were determined by less reliable methods of analysis. No other plots using zirconium are used here; however, a corollary of problematic zirconium concentrations is that hafnium (which forms a solid solution with Zr in zircon; e.g., Hoskin and Schaltegger, 2003) should also be anomalously enriched in the samples analyzed by ICP–MS. The discrimination plot of Harris et al. (1986) for collisional granitoids is used here, but employs hafnium as one of three chemical components that are plotted. Unless these data were derived from XRF analyses, geochemical plots relying on either of these two elements should be interpreted with caution. It may be that the overall trends are reliable, even if the zirconium and hafnium concentrations are not. Some datasets were also anomalous with respect to other trace elements. The most apparent example of these are the elevated uranium, tantalum, vanadium, and cobalt concentrations reported for more siliceous samples in the data from Heffernan (2004; e.g., “Co” in Fig. 2.8a). Basic geochemical trends and characteristics of the plutonic suites are first addressed here by discussing general plutonic suite compositions evident in Harker variation diagrams, several fractionation indices, and multi-element profiles normalized to chondritic and primitive mantle references (Sun and McDonough, 1989). The plutonic suites are then classified using a series of standard compositional and tectonic classification plots that were chosen in order to demonstrate basic compositional parameters. These classifications allow inferences regarding the general composition of magma types or their melt sources and the tectonic regimes into which they were emplaced. Table 2.4 provides details on each of the plots used in the chemical classification of magmatism in the northern Cordilleran miogeocline. As Clemens (2003) has stated, however, “it is the geology of a region that should tell us the particular 50 setting of the magmatism, not the types of granites we find or the geochemistry of some basalts.” The fact that upwards of 20 different classification schemes on the basis of the nomenclature, petrography, enclaves, mineralogy and mineral composition, zircons morphology, opaque oxide mineralogy, major and trace element geochemistry, associated mineralization, and tectonic environments of granitoid rocks have previously been proposed (see Barbarin, 1999) indicates that a satisfactory method of classifying granitic magmas has yet to be found. Although all of the geochemical diagrams in Table 2.4 have some validity with respect to granitoid classification, they were selected mainly for their widespread use in the literature and (or) for their specific utility in discriminating the plutonic suites examined here. Here, three main classification schemes employing different terminologies are used to describe the granitoids emplaced into the northern Cordilleran miogeocline. Although the relevance of these classifications cannot be fully evaluated without incorporating information from isotopic datasets (Chapters 3 and 4), this chapter concludes with a basic overview of the granitoid types exemplified by the syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. The ‘alphabetical’, or I-type (including the “high” and “low” temperature distinctions) versus S-type versus A-type, classification is used here. This scheme was originally developed in the Lachlan Fold Belt of southeast Australia for a 3600-km long and 750-km wide region of extensive, mainly plutonic, magmatism (e.g., Chappell and White, 1974, 1992; Loiselle and Wones, 1979)—not unlike the northern Cordillera. The Lachlan Fold Belt granitoids are interpreted by most researchers to unrelated to subduction processes, but remain similar in composition (and apparent petrogenesis) to syn- to post-accretionary igneous rocks in the northern Cordilleran miogeocline. This terminology is probably also the most widely used classification of granitoid rocks in the literature (a simple database search returned over 10,000 published results using the I-, S-, or A-type terminology), and it is a relatively simple classification that most readers will find familiar. Furthermore, the I- and S-type classifications generally (although imperfectly) correspond to the magnetite versus ilmenite-series of Ishihara (1981), another well known classification scheme (e.g., Sato et al., 2004). The largest drawbacks associated with the use of this classification, however, are the genetic connotations imparted by the I- vs. S-type terminologies (e.g., igneous vs. sedimentary protoliths). A second issue with this scheme is a correlative of the first: the difficulty in interpreting suites that have indeterminate or mixed compositions with respect to the criteria (e.g., mixed melt sources). The use of the shoshonitic and high-, medium-, and low-K calc-alkaline terminology of Peccerillo and Taylor (1976) here stems from their popularity as descriptors of overall magma series composition with researchers working on igneous rocks associated with the Variscan orogeny in southern Europe, the Pan–African orogeny in northwest Africa, and elsewhere. Post-collisional examples of magmatism from these regions are also similar in composition to syn- to post-accretionary magmatism in the northern Cordilleran miogeocline. This classification, therefore, provides a link between the descriptions of plutonic suites in the northern Cordillera to descriptions of similar episodes of magmatism where the ‘alphabetic’ terminology is not as widely used. Generally, the I- and S-type granites range from medium- to high-K calc-alkaline in composition (although they are typically high-K calc- alkaline rocks, at least in the Lachlan Fold Belt; Roberts and Clemens, 1993). Unlike the alphabetical classification, 51 the use of the “shoshonitic” and “high-, medium-, and low-K calc-alkaline” terms does not have any major genetic connotations, beyond the role for enriched mantle-derived melts in the shoshonitic suites (and inferred for some high-K calc-alkaline suites; e.g., Liégeois et al., 1998). This classification scheme may be better at identifying mixed granitic melt sources or magmas, as granitic series (or plutonic suites) do not normally cross the boundaries that define the low-, medium-, high-K and shoshonitic evolutionary paths (Roberts and Clemens, 1993; Clemens, 2003). Finally, the incorporation of tectonic descriptors, such as syn-collisional and late- or post-collisional, have been used to provide some indication of the tectonic settings into which the plutonic suites in the northern Canadian Cordillera were emplaced. This assumes that magmatism during continent–continent collision comprises simple stages that can be easily compartmentalized, which is not likely to be entirely the case. Collisional events may be complex, and the boundaries between pre-, syn-, and post-collisional processes and related magmatism are often transitional or diffuse (e.g., Harris et al., 1986). This is also problematic for intrusive rocks in the northern Cordillera as the timing of the accretion of the Insular Superterrane in particular is questionable, occurring sometime between the Late Jurassic and the Late Cretaceous (e.g., Trop and Ridgway, 2007). However, all post-Triassic magmas in the northern Cordillera were technically emplaced into a post-collisional or post-accretionary setting (e.g., Hart et al., 2004a); however, it is not certain whether any of these suites should instead be considered “syn-collisional” given the poor constraints on the timing of accretionary events. 2.4 CHARACTERIZATION OF PLUTONIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE A “suite” of magmatic rocks is a group of spatially and temporally associated rocks (i.e., they belong to the same “rock association”; Bowen, 1928) that share common modal, mineralogical, petrographic, geochemical, and petrogenetic histories (e.g., Chappell et al., 1987; Chappell and White, 1992). One of the primary considerations for assigning intrusions to a plutonic suite is the crystallization age of the intrusions. Even if intrusions share all of the above characteristics, if they are of significantly different age they are not grouped into a single plutonic suite. The age ranges given in the following characterization of plutonic suites in the northern Cordilleran miogeocline are based solely on robust U–Pb dating of zircon and (or) monazite, which are generally taken to most closely represent the true age of crystallization or magma emplacement (reported 40Ar/39Ar ages in the literature are commonly slightly older or younger than the U–Pb ages). Within the suites described below, there are also common variations in composition, texture, and mineralogy that may be related to small differences in the melt source composition across the northern Cordillera. The range of chemical variation exhibited by a suite, however, is a result of several processes experienced by each magma during partial melting, extraction from the restite, ascent, and emplacement. These processes include: anatectic disequilibrium between magma and restite, mixing with other magma(s), assimilation of or contamination by crustal material, restite-unmixing, and (or) mineral fractionation. Of these, anatectic processes are now thought to be the most important in the generation of intermediate to felsic granitoids in the crust (Clemens et al., 2010; Clemens and Stevens, 2012). Magma mixing, assimilation, contamination, restite- unmixing, and crystal fractionation are rarely capable of explaining the entire range of chemical (and isotopic) variation exhibited by suites of magmas (Clemens et al., 2010), although there are certainly exceptions (Clemens et 52 al., 2009). For relatively minor differences in the range of chemical variation, the magmatic suites below are described based on their geographic location. When these minor variations are extensive enough to identify distinct groups of plutonic and volcanic rocks, however, the plutonic suites have been divided into “sub-suites”. Although many of the ‘mid’-Cretaceous magmatic suites described in the northern Cordilleran miogeocline have been partially to fully described by other studies (see below), new information produced by this study justified a re- characterization of some of the previously described plutonic and volcanic suites (from a regional context), the sub- division of some of the plutonic suites into sub-suites, and new descriptions for a handful of newly proposed plutonic suites in the Cassiar Terrane and (mainly) the eastern Selwyn Basin. Geochronological, macroscopic, microscopic, and geochemical data and observations documented by this study (Tables 2.2, 2.3; Appendices A.1– A.3) are combined with information available from other studies throughout the region. The following characterizations (listed in order of decreasing age) summarize only the most relevant age, macroscopic, microscopic, and geochemical10 details of Cretaceous to Eocene suites and sub-suites in the northern Cordilleran miogeocline. This information, as well as additional observational and geochemical data not discussed below, is summarized in detail in Tables 2.3 and 2.5. (References for compiled geochemical datasets are also detailed for each plutonic suite in Table 2.5.) 2.4.1 Cassiar Plutonic Suite The Cassiar plutonic suite is the most widespread magmatic event in the Cassiar Terrane (Fig. 2.3b). Although previous workers have already described this suite well (e.g., Noble et al., 1984; Lowenstern and Sinclair, 1996; Driver et al., 2000; Black et al., 2003; Mortensen et al., 2007; Colpron et al., 2011; Gordey, in press), much of the following description also relies on the large number of samples collected by this study. Samples have been sub- divided into muscovite ± biotite-bearing (“muscovite-bearing”), hornblende ± biotite-bearing (“hornblende- bearing”), and “biotite-bearing” lithologies wherever possible, although several samples lacking mineralogical descriptions were included in the biotite-bearing category by default. Plutons and batholiths comprising the 115–104 Ma Cassiar plutonic suite display a range of textures and grain sizes (fine-to coarse-grained, equigranular to K-feldspar megacrystic to porphyry). Samples from the intrusions in Yukon comprise both: reduced, ilmenite-bearing, biotite ± muscovite monzogranite and leucogranite; and, oxidized, magnetite-rich, biotite ± hornblende ± titanite-bearing monzogranite to granodiorite (and minor quartz monzodiorite). This suite was classified as weakly oxidized by Hart et al. (2004a). Except for a greater abundance of reduced, muscovite-bearing intrusions in the northern portions of the Cassiar batholith (Driver et al., 2000), there are no obvious spatial trends with respect to the distribution of reduced versus oxidized phases. There are, however, temporal trends that can be tied to the mineralogical sub-divisions and the oxidation state of the magma. New and compiled U-Pb ages indicate that, on average, the muscovite-bearing phases are older (113.0 ± 2.0 Ma; 8 samples) than the biotite-bearing phases (109.1 ± 2.3 Ma; 19 samples), which in turn are older than the hornblende-bearing 10 The geochemical data presented in this chapter have not been re-calculated to anhydrous (i.e., excluding loss-on- ignition, of LOI) values. 53 phases (104.2 ± 1.3 Ma; 2 samples; Fig. 2.3b; Table 2.2). Earlier felsic magmas were also documented in the Glenlyon batholith (Fig. 2.3a), where a moderately foliated “two mica” granite is intruded by an undeformed but co- magmatic biotite–hornblende-bearing granodiorite (although the latter is cut by minor muscovite-bearing granite and pegmatite dykes; Black et al., 2003). Associated quartz monzodioritic to monzonitic phases are rare (Driver et al., 2000), but evidence for synchronous mafic magmatism in the region includes rare coeval shoshonitic lamprophyre (kersantitic) dykes that cross-cut miogeoclinal and obducted oceanic rocks east of the Cassiar batholith (110–112 Ma; Cooke and Godwin, 1984; Ash, 2001). Rafts of schist, amphibolite, hornfels, and quartzo–feldspathic gneiss observed in the Cassiar batholith could be from contact aureoles or roof pendants, as well as entrained mid-crustal xenoliths or source rocks (Driver et al., 2000). The Cassiar plutonic suite encompasses a wide range of geochemical compositions, with SiO2 contents ranging from 56 to 77 wt.%, and relatively heterogeneous trace element contents (Fig. 2.4a,c). Of particular interest are several oxidized hornblende-bearing samples, including a low-SiO2 outlier sampled from the Cassiar batholith by Driver et al. (2000). The Teslin plutonic suite (an inferred continental magmatic arc comprising mainly oxidized hornblende- bearing intrusions) may overlap geographically with the Cassiar plutonic suite and is also plotted (Fig. 2.4). Their geochemical similarity, and the lack of age constraints on several of the Cassiar and Teslin suite samples from the studies of Morris and Creaser (2008) and Driver et al. (2000), makes conclusively distinguishing between these two suites challenging. All of the hornblende-bearing samples that are thought to belong to the Cassiar plutonic suite, and possibly several highly siliceous biotite- and muscovite-bearing samples, however, delineate an alkaline–alkali- calcic–calc-alkaline evolutionary trend in Figure 2.4g. This trend is distinct from the calc-alkaline trend exhibited by most of the biotite- and muscovite-bearing phases comprising the majority of the Cassiar plutonic suite, as well as hornblende-bearing phases belonging to the Teslin plutonic suite. The least siliceous of the alkaline–alkali-calcic Cassiar suite samples are also enriched in LILEs and HFSEs (K, U, Th, Ba, Sr, Nb, Ta; Fig. 2.4d,h,n). The alkaline trend becomes calc-alkaline at the highest SiO2 contents; therefore, mixing between two source rocks or geochemical reservoirs of distinctly different composition appears to have been important in the genesis of the Cassiar plutonic suite. Overall, the compositional variability exhibited by the Cassiar plutonic suite in most of the tectonic discrimination diagrams is typical of post-collisional, high-K calc-alkaline granitoids that may have had multiple and (or) heterogeneous magma sources (e.g., Pearce et al., 1984). The Cassiar suite is also relatively geochemically (e.g., LILE) enriched with respect to typical pre-collisional or subduction-related I- and S-type granitoids (Driver et al., 2000), and has trace element contents that are compatible with highly fractionated I- or S-type granites. The geochemical enrichments are apparent on plot of SNX (average Yenchichi–Telabit-normalized Y–Ce–Sm–Yb) versus SNY (average Yenchichi–Telabit-normalized Rb–Th–U–Nb; Fig. 2.4n), although this plot is not influenced by magmatic differentiation11. Some of the alkaline–alkali-calcic hornblende-bearing samples may be classified as 11 To demonstrate that elevated SNX and SNY values or trends are not the result of fractionation, an example using the Cassiar plutonic suite was made by plotting the Rb/Sr fractionation index (see Fig. 2.14a). The lack of correlation between SNX or SNY and an increasing degree of fractionation demonstrates that the sliding normalization of Liégeois et al (1998) is effective at removing fractionation-induced heterogeneities for all but the most fractionated granitoids (e.g., Seagull plutonic suite; Fig. 2.14b; see Section 2.4.4). 54 aluminous–alkaline post-collisional granitoids, which may be more akin to shoshonitic post-collisional magmatism (Bonin et al., 1998; Liégeois et al., 1998). Thus, at least two compositionally distinct melt sources are required to generate the range of geochemical compositions exhibited by the Cassiar plutonic suite: a sub-ordinate alkaline and oxidized lithology or magma, and a dominant calc-alkaline and reduced lithology or magma. The latter may have been derived from partial melting of compositionally heterogeneous siliciclastic rocks of overall greywacke composition (± less important calcareous lithologies; Driver et al., 2000). The former compositional signature may include magmas derived from the mantle, or from a mafic crustal melt source (e.g., Driver et al., 2000). The relationship between these two melt sources and their role in the petrogenesis of the Cassiar plutonic suite cannot be fully evaluated without an isotopic discussion (Chapter 3), as well as better geochronological control on the alkaline–alkali-calcic samples. 2.4.2 Anvil Plutonic Suite The Anvil plutonic suite is present into two areas: 1) the Yukon–Tanana inlier immediately northeast of the Tintina Fault, where several small, highly evolved plutons ranging from 114–111 Ma occur; and 2) the western Selwyn Basin north of the Yukon–Tanana inlier, where the 109–104 Ma Anvil batholith and a small number of scattered plutons trending easterly occur (Fig. 2.3b). Although the intrusions in these two regions are grouped together as one plutonic suite, the Anvil plutonic suite in the western Selwyn Basin may better correlate with the Hyland plutonic suite emplaced southeast of the Yukon– Tanana inlier in terms of timing and possibly depth of emplacement. However, the observational and compositional data available for the Anvil and Hyland plutonic suites is too limited to justify any changes to the groupings at this time. Data from the two Anvil suite localities, for example are limited to four samples from the Anvil batholith that were published as one averaged composition (the original dataset has since been lost), a handful of samples from the highly fractionated Tsa Da Glisza pluton, and one sample from this study (Pigage and Anderson, 1988; Neufeld, 2004; Neufeld et al., 2005; Pigage, 2004). Some additional observations for the Yukon–Tanana inlier plutons have been merged with prior descriptions from Pigage (2004) and Mortensen et al. (2000). In the western Selwyn Basin, the Anvil batholith is described as two, reduced, commonly K-feldspar megacrystic, biotite ± muscovite monzogranitic laccolithic bodies separated by a layer of metasedimentary rock. A shallowly dipping foliation is present within both laccoliths. Muscovite forms either large euhedral primary grains, ragged sub- solidus grains, or clots with localized fibrolite that are thought to be pseudomorphs of earlier aluminous minerals (e.g., garnet, cordierite). Late-stage fine-grained hornblende–(plagioclase)-phyric diorite dykes, although probably of Tay River suite age, crosscut the Anvil batholith. In the Yukon–Tanana inlier, the Tsa Da Glisza pluton has been similarly described as a weakly foliated to massive biotite–muscovite monzogranite to muscovite leucogranite, with associated late-stage massive to highly strained (boudinaged) aplitic dykes (with tourmaline, and beryl-bearing pegmatitic and quartz-rich segregations) and quartz–tourmaline–(beryl) veins. Other plutons in the Yukon–Tanana inlier are described as fine- to medium-grained, equigranular, biotite–muscovite bearing granites (Murphy et al., 2001). One intrusion in the Yukon–Tanana inlier that was sampled for this study is a strongly reduced, quartz- 55 phyric, biotite–muscovite–tourmaline monzogranite to leucogranite. Garnet and staurolite with local fibrolite and cordierite in contact aureole of the Anvil batholith suggest a relatively deep level of emplacement (but <13 km depth, based on the absence of kyanite; Pigage and Anderson, 1988). The Anvil suite samples are similar to siliceous examples of the Hyland plutonic suite (which has a wider compositional range and degree of scatter or heterogeneity on the Harker variation diagrams; Fig. 2.5a). A limited range of SiO2 concentrations are present in the available samples of the Anvil plutonic suite, largely because all but two of the data points are from the highly fractionated granite associated with the Tsa Da Glisza emerald showing. The highly siliceous composition (71–74.5 wt.% SiO2; Fig. 2.5a) and only slight to moderate peraluminicity (Fig. 2.5f) imply that the Anvil suite samples were derived by near-minimum partial melting of crustal rocks (e.g., Chappell et al., 2000). However, an association with emerald mineralization at Tsa Da Glisza suggests that at least some of these plutons are also highly differentiated, which also produces haplogranitic compositions. High Rb/Sr (Sr <200 ppm) and low Eu/Eu* (<0.5), Th, Nb, Ti, Yb, and Y compared to other the high-K calc-alkaline suites examined in this study characterize the Anvil suite as “low-Sr leucogranitic”, low-temperature, minimum melt magmas lacking entrained restite, as described by Kemp and Hawkesworth (2003). Despite anomalously elevated Ga/Al (Fig. 2.5k), the lack of other mineralogical and chemical features indicative of A-type magmatism implies that the high Ga/Al might have been inherited from a pelitic protolith (London, 1995). Using the geochemical discriminators of Sylvester (1998) for strongly peraluminous granites, a pelitic protolith for the Anvil suite samples is also indicated on the basis of low CaO/Na2O and FeT+MgO+TiO2, and high Al2O3/TiO2 (not shown). This supracrustal origin is also consistent with higher normative corundum in the least siliceous samples (Fig. 2.5j). In general, these intrusions appear to comprise post-collisional, high-K calc-alkaline, S-type granitoids derived mainly by mica dehydration reactions, and extracted from a pelitic protolith. 2.4.3 Hyland Plutonic Suite The Hyland plutonic suite comprises several batholithic bodies emplaced into the eastern Selwyn Basin (Fig. 2.3b). This suite was not sampled by this study, and most of the following description is summarized from observations reported by Heffernan (2004) for several 106–100 Ma samples collected in the study area, as well as previously published work (Hart et al., 2004a; Hart and Lewis, 2006). The circa 106–100 Hyland plutonic suite (or “Hyland River” suite; Hart et al., 2004a) in the eastern Selwyn Basin comprises several batholiths that were emplaced into “migmatitized” metasedimentary rocks of the Neoproterozoic Hyland Group (Fig. 2.2a; Roots et al., 1966). This plutonic suite temporally overlaps with the Anvil batholith in the western Selwyn Basin, and for this reason was correlated with the Anvil plutonic suite by Heffernan (2004). Different oxidation states, however, led Hart et al. (2004a) to distinguish the Hyland and Anvil plutonic suites, which is also compatible with their differing mineralogical and chemical compositions (Pigage and Anderson, 1988; Heffernan, 2004). The Hyland plutonic suite consists of reduced to oxidized (or “weakly oxidized”; Hart et al., 2004a), typically massive (although occasionally weakly foliated) biotite granodiorite and monzogranite, with subordinate granodioritic to dioritic plagioclase–biotite porphyry phases, dykes, and enclaves. Most descriptions of the Hyland 56 suite imply that the intrusions are solely biotite-bearing, lacking hornblende or muscovite. There are, however, reports of hornblende-bearing granodiorite phases associated with lesser quartz monzonite and quartz monzodiorite, abundant mafic xenoliths, and dioritic dykes at some tungsten skarn occurrences (e.g., TAI and WOAH, see Fig. 2.16; Brock, 1978; Durgin, 1979). Similarly, muscovite has been observed in small-volume leucocratic phases spatially associated with tungsten skarn occurrences (e.g., CALI, see Fig. 2.16; Kuehnbaum, 1977), although the textures of this muscovite have not been described. Like the Anvil batholith, the Hyland plutonic suite was emplaced as thick laccolithic bodies with mainly concordant contacts and staurolite- to garnet-grade contact aureoles displaying schistose, gneissic, and migmatitic textures (Roots et al., 1966; Hart and Lewis, 2006). The geochemical classification of the Hyland suite is limited by the small number of samples analyzed, and by contamination or analytical issues. Despite the limited data, this suite appears to have a relatively variable composition that is similar to the heterogeneity displayed by the Cassiar plutonic suite (Fig. 2.5). Peraluminicity in the Hyland suite (Fig. 2.5f) broadly decreases with increasing degree of fractionation (e.g., SiO2, Rb/Sr, Eu/Eu*; not shown), and has a rough positive correlation with general maficity (e.g., FeT; Fig. 2.5j). This is suggestive of a supracrustal melt source, perhaps one that is similar in composition to the fine-grained biotite–plagioclase-bearing enclave, one of the least siliceous and most peraluminous (A/CNK = 1.30 and crnnorm = 2.7; Fig. 2.5f,j) samples from this suite. The Hyland suite is similar to the strongly peraluminous, biotite-rich (and locally cordierite-bearing) S-type granite classification of Barbarin (1996). That only one sample has slightly elevated Rb/Sr (i.e., >2) suggests that the phases sampled are not highly fractionated, and (or) that these intrusions were formed by large-fraction partial melting involving plagioclase as a reactant and (or) entrained mineral. This interpretation is also compatible with an origin of near-in situ anatexis that is suggested by the wide migmatitic contact aureoles (Roots et al., 1966), the high Eu/Eu* in the mafic enclave (Fig. 2.5b; assuming the enclave represents entrained restite; Chappell et al., 1987), and the relatively high strontium and Na/K (Fig. 2.5c,d,i ; e.g., Kemp and Hawkesworth, 2003). Associated quartz monzodioritic porphyry phases and dykes are slightly metaluminous to moderately peraluminous, but lack the normative corundum that would be suggestive of a supracrustal protolith. Whether these less siliceous phases are cogenetic with the Hyland plutonic suite cannot be determined without added geochronological work. Although generally of high-K calc-alkaline composition, some of the less siliceous plutonic samples are of medium-K calc- alkaline composition (Fig. 2.5h). This could be due to either source heterogeneity (as is suggested by the variable trace and REE concentrations), or to mixing of compositionally distinct materials. If the Hyland suite is in part derived from a mixed melt source, and the quartz monzodiorite dykes are coeval to the main phase of plutonism, then these dykes may represent magma derived from an infracrustal component of mafic–intermediate composition. Here, the Hyland plutonic suite is classified as a post-collisional, high-K calc-alkaline, intermediate–felsic granitoid suite that may have been derived from partial melting of supracrustal rocks ± lesser interlayered infracrustal rocks (e.g., Hart et al., 2004a). 2.4.4 Seagull Plutonic Suite Detailed work on the Seagull, Hake, and Thirtymile intrusions by Liverton (1992) is the main source of information on the Seagull plutonic suite (e.g., Liverton, 1990; Liverton and Alderton, 1994; Liverton and Botelho, 2001; 57 Liverton et al., 2005; Liverton, 2012, pers. comm.), although Colpron et al. (2003) and unpublished data from M. Colpron provide much of the data available on the northern portion of this plutonic suite. Additional information on the Seagull plutonic suite also comes from Sinclair and Richardson (1992) and Colpron et al. (2003). Intrusions comprising the 103–96 Ma Seagull plutonic suite are typically small bodies, although two batholithic examples also exist (the Seagull and Hake batholiths; Fig. 2.3). This plutonic suite typically comprises composite, “one mica” (biotite) bodies displaying a range of heterogeneously zoned textures (e.g., Liverton and Alderton, 1994). It was emplaced as a narrow, north-northwest-trending belt, extending from the Seagull batholith near the Yukon–British Columbia border to the Glenlyon map sheet (Fig. 2.3b). Only a few phases of this plutonic suite were sampled during the course of this study. These comprise reduced, ilmenite-bearing, fine- to medium-grained equigranular biotite (with rare muscovite) leucogranite to monzogranite. Fine-to coarse-grained, equigranular to K- feldspar porphyritic to megacrystic, to porphyry textures have all been observed in the Seagull, Hake, and Thirty- mile intrusions, and partially developed rapakivi textures were also observed in a border facies to the Hake batholith (Liverton and Alderton, 1994; T. Liverton, 2012, pers. comm.). Volumetrically minor stocks of lepidolite– (zinnwaldite)–albite–topaz leucogranite with fluorite are unique to this suite, and occur as massive high-level dykes and sills adjacent to larger batholiths (Liverton and Alderton, 1994). HFSE-rich (i.e., fergusonite, xenotime) and LREE-rich (i.e., monazite, allanite) minerals, as well as locally abundant fluorite, tourmaline, and topaz are observed in variable abundances within the most evolved members of this suite (T. Liverton, 2012, pers. comm.). In the eastern Seagull batholith, miarolitic cavities contain tourmaline and fluorite ± beryl and topaz, pegmatitic pods, and (or) quartz–tourmaline orbicules with aplitic margins, and are compatible with relatively shallower levels of emplacement (when compared to the Cassiar plutonic suite, which lacks miarolitic cavities). Minor magma mingling textures in the least geochemically evolved phase of the Thirtymile Stock (a porphyry phase) occur as disaggregated cm-scale hornblende–biotite-phyric dioritic enclaves, and enclaves within enclaves (T. Liverton, 2012, pers. comm.). Cross-cutting hornblende-bearing lamprophyric dykes have also been emplaced near the margins of the Thirtymile Stock, but it is not known whether these dykes are cogenetic to the Seagull plutonic suite (T. Liverton, 2012, pers. comm.). Samples of the Seagull plutonic suite are strongly siliceous; SiO2 increases from porphyry to megacrystic granite to equigranular granitic phases (Liverton and Alderton, 1994), although some of the most magmatically evolved samples have slightly lower SiO2 (T. Liverton, 2012, pers. comm.). Silica loss is particularly obvious on the Rb/Sr plot (Fig. 2.6b), but is also apparent on the TiO2, FeT, Na2O, K2O, zirconium, and thorium plots in Figure 2.6a. The REEs are enriched in this plutonic suite (Fig. 2.6c), and the elevated HREEs in particular are a feature that distinguishes A-type and highly fractionated S-type magmas (e.g., Chappell et al., 2000). Anomalously enriched HFSEs (e.g., Zr, Ce, Y, and Ga—as well as Nb, which is not shown; Fig. 2.6a,k;) for many of the biotite-bearing phases, combined with relatively low CaO, MgO, and Al2O5, are also indicative of an A-type magmatic affinity (Fig. 2.6; Collins et al., 1982) and (or) unusually high degrees of magmatic fractionation (which may also characterize some A-type suites; Turner et al., 1992; Chappell et al., 2000; Liverton and Botelho, 2001; Vander Auwera et al., 2003; Namur et al., 2011). Extreme magmatic fractionation is evidenced by high Rb/Sr ratios, most of which are greater than 2 (this is anomalous in comparison to most of the other plutonic suites discussed) but ranging 58 as high as ~9500 for a lepidolite topaz leucogranite (Fig. 2.6b). Other elements also exhibit extreme enrichments (e.g., Th) and depletions (e.g., Eu, Ba, P, Ti) that are features of “ultra-fractionated” late-orogenic A-type intrusions that have an association with tin mineralization (Liverton and Alderton, 1994; Chappell et al., 2000; Liverton and Botelho, 2001; Liverton et al., 2005). Ultra-fractionation is envisioned to occur during latest-stage liquid–vapour fractionation, in which HFSEs (and Sn) are scavenged from residual melts (Liverton and Botelho, 2001; Liverton et al., 2005). It is notable that the most highly fractionated phases (lepidolite-bearing) are not as strongly enriched in HFSEs, which implies that either this enrichment is a primary feature of the parental magma, or that liquid-vapour fractionation removed much of the HFSEs from these residual magmas before they solidified. The Seagull plutonic suite comprises peralkaline to peraluminous, relatively sodic, A-type granitoids of post-collisional affinity (e.g., Liverton et al., 2005). These intrusions share compositional features with other alkaline–peralkaline (HFSE enriched), post-orogenic A-type magma series (Fig. 2.6n; Bonin et al., 1998). The LILE enrichment in lepidolite– (zinnwaldite)–albite–topaz leucogranite phases is similar to the “rapakivi-topaz” A-type series (Fig. 2.6n; Bonin et al., 1998). Overall, the Seagull plutonic suite is not unlike post-orogenic A-type magmatism emplaced near shear zones in the Tuareg Shield of the African craton (e.g., the Taourirt suite; Liégeois et al., 1998; Bonin et al., 1998)12. 2.4.5 Tay River Plutonic Suite Although the 99–96 Ma Tay River plutonic suite has been extended across the Selwyn Basin (Fig. 2.2a), minor textural, mineralogical, and geochemical differences are present between plutons from the western and eastern Selwyn Basin. Although isotopic data for the Tay River and South Forks suites in the western Selwyn Basin are extremely limited, there are also differences across the Selwyn Basin that suggest some variation in the age and composition of the magma protolith(s) (Chapter 4). The Tay River suite intrusions in each region are discussed separately. 12 As an aside, the trend towards the origin of the SNY:SNX plot is a fractionation trend that should correspond to increasing SiO2 in the samples, despite the use of a reference series-based normalization to remove or reduce the effects of fractionation on elemental comparisons (Liégeois et al., 1998). Samples with the lowest SNX (i.e., <1) are from the northwest Seagull batholith where several tin occurrences are clustered, two samples from the Thirtymile pluton, and one sample is from an apophysy near the Quiet Lake batholith. The middle cluster of data is largely composed of what are interpreted to be the least fractionated intrusive members of the Seagull plutonic suite (T. Liverton, 2011, pers. comm.), with some highly fractionated exceptions (Fig. 2.14b). However, the cluster of analyses at high SNX (i.e., >3; average normalized SNX) is composed of leucogranite samples from the east Seagull batholith and leucogranite (similar to the lepidolite topaz leucogranite from the Thirtymile stock) from the northwest Seagull batholith that have some (but not exclusively) of the highest SiO2 concentrations in the Seagull suite (i.e., 76–78 wt. %) and are arguably the most fractionated phases sampled. Therefore, SNX is weakly but positively correlated with Rb/Sr, a fractionation index (Fig. 2.14b). Additionally, when SNY is less than 2.25 then Rb/Sr is weakly negatively correlated to SNY, but when SNY is greater than 2.25 then Rb/Sr is positively correlated with SNY (Fig. 2.14b). Therefore, the apparent sub-vertical trends indicative of SNY-enrichment (average normalized Rb-Th-U-Nb) exhibited by the individual clusters in Figure 2.6n are also attributable to fractionation towards alkaline rapakivi topaz granitic compositions of Bonin et al. (1998), as exemplified by the single lepidolite topaz leucogranite sample. These fractionation patterns are typical of post-orogenic trends for alkaline magmatism (Bonin et al., 1998), where the samples furthest from the origin are the most fractionated, and is contrary to trends typifying calc-alkaline suites (Liégeois et al, 1998). 59 2.4.5.1 Western Selwyn Basin Only one sample from the Tay River and South Fork suites in the western Selwyn Basin each were examined by this study, and much of the following description of these suites relies on the observations made by previous workers (Wood and Armstrong, 1982; Anderson, 1988; Pigage and Anderson, 1988; Mortensen et al., 2000; Pigage, 2004; Gordey, in press). The 99–96 Ma Tay River suite in the western Selwyn Basin comprises reduced (ilmenite > magnetite), monzogranite and granodiorite batholiths and plutons that are associated with quartz–feldspar porphyry intrusions (typically biotite–hornblende bearing) and abundant dykes and sills. Texturally, the plutons range from massive to strongly foliated, fine- to coarse-grained, and equigranular to porphyritic. Biotite ± hornblende (± Ca-rich clinopyroxene) are the main ferromagnesian minerals, and K-feldspar phenocrysts and megacrysts are rare. The Tay River intrusions cross-cut the Anvil plutonic suite (Fig. 2.3b). The one sample of the Tay River suite from the western Selwyn Basin that was examined by this study is an altered, sub-volcanic plagioclase–hornblende–quartz porphyry of monzogranitic–granodioritic composition. Spatially, temporally, and compositionally associated volcanic rocks are the 96.8–97.7 Ma (Gordey, in press) extrusive rocks in the South Fork calderas (Fig. 2.2a, 2.3b; Pigage and Anderson, 1988; Mortensen et al., 2000). The South Fork volcanic suite comprises a cluster of eight calderas filled with thick sequences (up to 950 m thick in the Connolly Caldera) of predominantly dacitic composition, with lesser basaltic to andesitic flows. Dacitic ash fall or ash flow deposits are inferred to have been much more extensive, but have been removed by erosion everywhere except in the calderas (Yorath, 1991; Gordey, in press). The South Fork calderas are down-dropped, syn-magmatic blocks that developed in response to magma chamber evacuation. The South Fork volcanic sample examined here is a welded, reduced, dark grey plagioclase– quartz–biotite–(hornblende) lapilli tuff that contains zircon, apatite, ilmenite, cerium-rich allanite, and rare corundum. Normative mineralogies based on geochemical composition suggest that magnetite is more abundant than ilmenite in the volcanic sample. The South Fork volcanic rocks and the Tay River plutonic suite in the western Selwyn Basin exhibit relatively little compositional variability, although the South Fork rocks have slightly lower SiO2 and Al2O3 (Fig. 2.7a). This lack of inter- and intra-suite variation in geochemical composition suggests a remarkably homogeneous melt source. In general, the Tay River and South Fork suites in the western Selwyn Basin comprise relatively magnesium-, calcium- , and potassium-rich, calk-alkaline melts (e.g., Wood and Armstrong, 1982) that have mainly high-K calc-alkaline I- type characteristics, which are consistent with trace element-based tectonic discrimination plots suggesting a volcanic arc affinity for this plutonic suite (Fig. 2.7l,m). Compositionally, the Tay River and South Fork suites are not unlike voluminous high-K calc-alkaline magmatism emplaced in a post-collisional setting in the Tuareg Shield (Liégeois et al., 1998), although like the other high-K calc-alkaline northern Cordilleran suites presented here (e.g., Cassiar plutonic suite), they are characterized by a well-defined negative titanium and (to a lesser degree Ca) anomaly (Fig. 2.7d). 60 2.4.5.2 Eastern Selwyn Basin This study extensively sampled the Tay River plutonic suite in the eastern Selwyn Basin. The following observations are based largely on this work, although the summary below benefits substantially from the work of others (Anderson, 1983; Mortensen et al. , 2000; Heffernan, 2004; Bezzola, 2009; Pigage et al., 2011). Intrusions of the 99–96 Ma Tay River suite in the eastern Selwyn Basin are mineralogically and texturally similar to coeval plutonism in the western Selwyn Basin. They are, however, almost always more oxidized in the eastern Selwyn Basin (i.e., titanite > magnetite > ilmenite), where there is no associated volcanism and dykes range from fine-grained plagioclase–(quartz)-phyric quartz monzodiorite, to plagioclase–biotite-phyric microgranite, to aplite and rare pegmatite. Several examples of Tay River suite in the eastern Selwyn Basin are noticeably coarser grained, and comprise K-feldspar porphyritic to megacrystic, biotite monzogranite to leucogranite that also contain appreciable amounts of monazite and magmatic muscovite ± tourmaline. These Tay River suite intrusions are referred to as ‘evolved’ (Fig. 2.3b) on the basis of their more evolved geochemical (and isotopic; Chapter 4) compositions. Although they are otherwise similar to the Tungsten plutonic suite (and were previously classified as a transitional “Tay River–Tungsten” before isotopic data was available; Rasmussen et al., 2007), the ‘evolved’ Tay River suite intrusions are isotopically more akin to the Tay River plutonic suite than they are to the Tungsten plutonic suite (Chapter 4). The ‘evolved’ plutons may have small to large tourmaline-infilled miarolitic cavities with aplitic margins, as well as myrmekite ± spherulitic quartz–feldspar intergrowths suggesting relatively shallow levels of emplacement. Igneous-textured, fine-grained to K-feldspar porphyritic, biotite ± hornblende-bearing enclaves with acicular apatite (suggestive of magma mingling and quenching) have been observed in several of these ‘evolved’ intrusions. Several coeval, but small, isolated plugs in the southeast Selwyn Basin have been sub-divided into the Coal River sub-suite (Fig. 2.3b). These plugs may be described as equigranular or fine-grained plagioclase ± hornblende ± biotite ± quartz porphyritic, granodiorite to quartz monzodiorite bodies that typically display massive textures, although are occasionally foliated. Compared to coeval magmatism in the western Selwyn Basin, there is a wider compositional range in the Tay River suite in the eastern Selwyn Basin—particularly if the Coal River sub-suite and three quartz monzodiorite dykes are also considered (Fig. 2.8a). Major element concentrations and trends are similar to those in the western Selwyn Basin with equivalent SiO2 concentrations, apart from slightly lower total iron and CaO, but trace elements in the western Tay River suite samples are more variable. As implied above, the ‘evolved’ Tay River intrusions are some of the most siliceous and fractionated members of this suite. Conversely, in all major and trace element plots the less siliceous Coal River samples appear to form an uninterrupted liquid line of descent with the rest of the Tay River suite samples, except in the Na2O Harker variation diagram. Na2O is enriched in the less siliceous Coal River sub- suite, and is negatively correlated with, converging with the main Tay River suite trend at circa 65 wt.% SiO2. Diverging Na2O concentrations with lower SiO2 suggest that the Coal River sub-suite had a different melt source than the bulk of the Tay River plutonic suite. There are numerous other changes in trends on the trace element Harker variation diagrams (e.g., Ce, Zr, Th, and Cr) that correspond roughly with the change from Coal River to Tay River samples, although some of the more siliceous Coal River samples do occupy the same compositional ranges as 61 many of the Tay River suite samples. Despite its less geochemically evolved compositions, the Coal River suite comprises the most peraluminous examples of the Tay River plutonic suite (Fig. 2.8f). This A/CNK pattern may be related to the low K2O in most of the Coal River sub-suite intrusions. However, it is also notable that some of the plutonic samples of the Tay River suite with the lowest peraluminicity are typically the most siliceous and geochemically (and isotopically; Chapter 4) ‘evolved’ (Fig. 2.8f). This association between decreasing peraluminosity with decreasing maficity would imply that the protolith to these rocks was relatively aluminous (i.e., supracrustal; Chappell et al., 1987, 2000). Conversely, the overall trace element homogeneity for all members of the Tay River plutonic suite—the Coal River sub-suite included—combined with a trend of normative corundum decreasing with increased total iron (Fig. 2.8j) is more consistent with an infracrustal melt source for the Tay River suite. Trace element concentrations range from those typical of volcanic arc to post-collisional granites, although some of the ‘evolved’ samples plot within the syn-collisional granite fields due to an enrichment in rubidium (Fig. 8l-n). As with the Tay River and South Fork suites in the western Selwyn Basin, the Tay River plutonic suite in the eastern Selwyn Basin is similar to highly fractionated I- or S-type granites (in this case I > S) or high-K calc-alkaline post-collisional suites potentially derived from mixed sources (Fig. 2.8n). An association with quartz monzodiorite dykes that may have a shoshonitic affinity (Fig. 2.8n) also implies some role for heat and (or) melt input derived from the lithospheric mantle during the genesis of the Tay River plutonic suite in the eastern Selwyn Basin. 2.4.6 Tungsten Plutonic Suite The Tungsten suite has only been identified in the eastern Selwyn Basin near the Yukon–Northwest Territories border (Fig. 2.3b). Previous classifications have generally lumped all of the muscovite ± biotite bearing plutons in the eastern Selwyn Basin (i.e., the “two mica” granitic series for this region; Anderson, 1983; Gordey and Anderson, 1993) into the Tungsten plutonic suite (Lang, 2000; Mortensen et al., 2000; Hart et al., 2004, 2005). Mortensen et al. (2000) suggested on the basis of geochemical and geochronological evidence that the Tungsten suite may represent more geochemically evolved and (or) crustally contaminated phases of the roughly coeval Tay River and Mayo plutonic suites, rather than a distinct plutonic suite. However, even the large isotopic dataset produced by this study (Chapter 4) does not allow for the easy classification of these plutons as a distinct plutonic suite, or as differentiates of coeval and spatially associated plutonic suites. The following characterization of the Tungsten plutonic suite is based largely on samples collected for this study, but also incorporates the work of others (e.g., Anderson, 1982, 1983, 1985; van Middelaar, 1988; Gordey and Anderson, 1993; Mortensen et al., 2000; Hart et al., 2004, 2005; Rasmussen, 2004; Rasmussen et al., 2007, 2011 in press). The 98–95 Ma Tungsten plutonic suite comprises fine-grained and equigranular, medium- to coarse-grained and K- feldspar porphyritic or megacrystic, ilmenite- and monazite-bearing biotite monzogranite and leucogranite. Intrusions are small round plugs (possibly cupolas) and plutons. Plutons are commonly massive-textured, although biotite ± K-feldspar megacrysts may be locally aligned along contacts, and several samples near the Hyland Fault have strongly foliated augen to mylonitic textures (KR-05-207 and 208; Fig. 2.3a; Appendix A.1). This plutonic suite is typically dominated by a K-feldspar-phyric or megacrystic phase; however, abundant, finer-grained highly evolved intrusive phases (equigranular leucocratic muscovite ± biotite, garnet, and (or) tourmaline-bearing phases) 62 are commonly localized along plutonic margins and contacts between intrusive phases are usually gradational. Abundant cross-cutting and satellitic, locally miarolitic, aplite ± pegmatite ± quartz–feldspar–(biotite) porphyry dykes are also associated with all of the Tungsten suite plutons. Although some Tungsten suite plutons (e.g., Lened; Fig. 2.3a) and smaller-volume marginal phases and dykes may contain some primary magmatic muscovite, this mineral is more typically found as a minor sub-solidus phase in the groundmass (e.g., Cantung; Fig. 2.3a). “Dark” xenoliths in the apical Mine Stock at Cantung and rare micro- to porphyritic diorite (biotitic) enclaves have been reported in some Tungsten suite plutons (e.g., Lened, Cac, Rudi; Fig. 2.3a). Late-stage mafic dykes are rare, but kersantitic (shoshonitic) lamprophyre dykes are exposed in the Cantung area, as well underground in the Cantung mine where they cross-cut the tungsten skarn mineralization (Rasmussen et al., 2011). A kersantite dyke was also observed in association with the Lened pluton (Gordey and Anderson, 1993). Sheeted quartz ± tourmaline veins and associated greisenized selvages, or gossanous fractures/joints cross-cutting the upper portions of several Tungsten suite plutons, may be mineralized with scheelite or auriferous pyrite ± arsenopyrite. The Tungsten suite samples in the eastern Selwyn Basin have been subdivided into two regional groupings in the eastern Selwyn Basin (northeast and southeast) due to small variations in their geochemical composition (Fig. 2.9). These differences may be significant with respect to their origin, although most field, macroscopic, and petrographic observations are essentially the same in both regions. This northeast versus southeast subdivision is also consistent with the inferred geographic extent of the Mayo and Tay River suite intrusions, respectively, both which are potentially cogenetic to the Tungsten suite (Fig. 2.3b). Arguably, in the southeast Selwyn Basin the intrusions have more of an I-type affinity (i.e., slightly elevated Al and Na) as well as a greater proportion of inherited zircon cores. Conversely, in the northeast Selwyn Basin the Tungsten suite has slightly elevated incompatible elements (K, Rb, U, Th, Ta, Nb, Sm, Y, and Yb) that, when combined with a positive correlation between normative corundum and total iron, are more suggestive of an S-type13 affinity. On a cautionary note, the positive correlation between corundum and iron relies largely on one outlier sample—a fine-grained leucocratic phase from the Cirque Stock at MacPass— that might be closer in age to the Tombstone plutonic suite (Selby et al., 2003; Gebru and Lentz, 2008). The Tungsten suite as a whole is not particularly highly fractionated, even when associated with tungsten mineralization (e.g., only some of the most siliceous samples have Rb/Sr ranging over 2). Geochemical trends are also similar to those exhibited by the Tay River and Mayo suites, although the compositional range with respect mainly to SiO2 is smaller for the Tungsten suite. The limited range of SiO2 contents is probably also responsible for the apparent geochemical homogeneity exhibited by the multi-element plots in Figure 2.9c. Associated kersantitic and monzonitic porphyry dykes (with 53 wt.% SiO2) have similar normalized trace element compositions, but significantly smaller anomalies and large enrichments in LILEs (Fig. 2.9c,d). Overall, the Tungsten plutonic suite has compositions similar to those of other post-collisional, high-K calc-alkaline magmas (Fig. 2.9n), which are also associated with shoshonitic lamprophyre dykes (Liégeois et al., 1998). 13 A large component of argillaceous supracrustal material in the melt source of the Lened pluton has also been inferred on the basis of nearly pervasive coarse-grained igneous muscovite (Chapter 6), magmatic apatite and biotite compositions characterized by a combination of depleted Ca and elevated MnO, FeT, and Al2O3 (markers of an S- type petrogenesis; Sha and Chappell, 1999; Belousova et al., 2001), and unusually low differentiation index (i.e., a lower proportion of non-normative minerals―feldspars and quartz―relative to micas; Rasmussen et al., 2012; Chapter 7) given the apatite and biotite compositions. 63 2.4.7 Mayo Plutonic Suite The Mayo plutonic suite has received much attention, due largely to its association with gold mineralization (e.g., Scheelite Dome, Dublin Gulch; Fig. 2.2a; Hart et al., 2000). Observations and data from this previous work (e.g., Anderson, 1982, 1983, 1985; Gordey and Anderson, 1993; Lang, 2000; Mortensen et al., 2000; Hart et al., 2004, 2005; Mair et al., 2006a, 2011; Rasmussen, 2011) are combined with the results of this study in the following characterization. The 98–93 Ma Mayo plutonic suite is composed primarily of small, massive to foliated (mainly aligned biotite along pluton contacts), reduced to slightly oxidized (titanite ± magnetite bearing), homogenous to composite plutons (Fig 2.3b). Several well-known examples of Mayo suite plutons in the western Selwyn Basin are the Scheelite Dome, Dublin Gulch, Clear Creek, and Roop Lakes intrusive complexes (Fig. 2.2a). This plutonic suite designation now also applies to several plutons in the northeast Selwyn Basin (Fig. 2.3). In the northeast Selwyn Basin, the Mayo suite ranges widely in texture and composition. Fine- to coarse-grained, equigranular to weakly K-feldspar porphyritic to coarsely megacrystic textures have all been observed. Some samples are crowded with K-feldspar crystals, displaying a texture that is commonly exhibited by Tombstone suite intrusions (Appendices A.1–A.3). Hornblende > biotite (± trace clinopyroxene) quartz monzonite or monzodiorite to biotite > hornblende (± trace augite) monzogranite to granodiorite compositions are typical. Intrusions may be cut by aplite–pegmatite dykes, as well as graphic quartz–tourmaline veins and tourmaline-infilled vugs, base metal mineralized veins and sheeted fractures. Enclaves or xenoliths of mafic–intermediate igneous material (e.g., fine- grained biotite ± hornblende diorite; some inclusions contain smaller inclusions) may make up to 5% of some plutonic phases belonging to this suite in the northeast Selwyn Basin (e.g., the Pelly River pluton; Gordey and Anderson, 1993). In the northwest Selwyn Basin, late-stage intrusions associated with most of the Mayo suite plutons range from monzonitic to dioritic compositions, or are sub-alkaline kersantite or spessartite to alkalic minette lamprophyres (e.g., Mortensen et al., 2000; Hart et al., 2004, 2005; Mair et al., 2011). Geochemical plots for the Mayo plutonic suite display well-defined and mostly continuous trends over a wide range of SiO2 concentrations (Fig. 2.10). The Mayo suite has also been subdivided into regional groupings due to variations in geochemical (and isotopic; Chapter 4) compositions across the Selwyn Basin. Some exceptions to simple, nearly linear trends on most of the major element Harker variation plots are also evident and help to identify K-feldspar cumulate phases (e.g., anomalously K2O-rich for SiO2 concentrations in the northeast Selwyn Basin) and (or) hypersolvus phases in which plagioclase did not saturate until the latest-stages of magmatic crystallization, if at all (e.g., particularly for samples mainly in the northwest Selwyn Basin that exhibit a well-defined trend of increasing Eu/Eu* with SiO2, and that have elevated BaO; Bonin, 1990). This suite typically displays a high-K calc- alkaline trend with a range of 58–75 wt.% SiO2 (Fig. 2.10h). Several samples, however, plot within the shoshonitic field of Pecerillo and Taylor (1976; Fig. 2.10h) and along the alkali-calcic trajectory of Frost and Frost (2001; Fig. 2.10g). Most of these potassium-rich plutonic rocks are siliceous and have K-feldspar-crowded textures (particularly in the northeast Selwyn Basin) that may account for the potassium enrichment. The late-stage shoshonitic lamprophyres that appear to form one compositional end-member in Figures 2.10g,h do not appear to represent a parental magma to this plutonic suite. Furthermore, textural evidence for extensive magma mixing, has yet to be 64 reported. Mixing between enriched mantle-derived magma and a crustally derived magma, or the fractionation of mantle-derived melts to form the more siliceous magmas, therefore, cannot be responsible for the chemical variation exhibited by the high-K calc-alkaline phases that form the bulk of this plutonic suite. The abundance of late-stage mafic dykes associated with this suite—particularly in the northwest Selwyn Basin—is consistent with the influx of enriched mantle-derived magmas into the crust during and (or) immediately following high-K calc-alkaline Mayo suite magmas. These enriched mantle-derived magmas provide a possible heat source for extensive crustal melting, but do not appear to have mixed extensively with the crustally derived Mayo suite magmas. 2.4.8 Tombstone Plutonic Suite The Tombstone plutonic suite has probably been the focus of the majority of previous geochemical studies of ‘mid’- Cretaceous magmatism, also largely because of its well-known association with gold mineralization (e.g., Brewery Creek; Fig. 2.2a; Hart et al., 2000). The following summary includes observations and data from this and several previous studies (Anderson, 1983, 1985; Abercrombie, 1990; Gordey and Anderson, 1993; Murphy, 1997; Duncan et al., 1998; Duncan, 1999; Lang, 2000; Mortensen et al., 2000; Coulson et al, 2002; Hart et al., 2004, 2005; Rasmussen, 2011). Several of the more well known examples Tombstone suite plutons include the Tombstone and Deadman plutons in the northwest Selwyn Basin, and the O’Grady (Fig. 2.3) and Emerald Lake intrusions in the northeast Selwyn Basin (Figs. 2.2a and 2.3a). This 94–89 Ma plutonic suite comprises scattered occurrences or clusters of rounded, commonly concentrically zoned, composite plutons. For example, the alkalic and silica-undersaturated Horn Peak syenite is the oldest phase of the Emerald Lake pluton, and a biotite monzogranite core is the youngest phase (Fig. 2.2a; Duncan, 1999). Despite gradational contacts between intrusive phases, however, at least three of the four main intrusive phases at Emerald Lake do not appear to be related by internal magmatic differentiation (Duncan, 1999). Instead, this composite intrusion was formed by the sequential emplacement of several geochemically distinct batches of magma (Duncan, 1999). Biotite monzogranite and highly fractionated, peraluminous leucogranitic plutonic phases and dykes with magmatic muscovite and tourmaline are commonly observed in central portions of composite intrusions. Tombstone suite plutons range from slightly to moderately oxidized with magnetite > titanite in the northwest Selwyn Basin, but are typically reduced and ilmenite or titanite-bearing in the eastern Selwyn Basin. Massive, medium- to coarse-grained, syenites dominated by K-feldspar (commonly crowded textures), quartz, or aegerine– augite are the most voluminous intrusive phases observed in the northwest Selwyn Basin. Conversely, most of the Tombstone suite plutons in the eastern Selwyn Basin are dominated by sub-alkaline, fine- to coarse-grained, hornblende > biotite or hornblende > clinopyroxene, quartz monzonite and monzogranite. Some K-feldspar crowded phases are also present in the northeast Selwyn Basin: up to 40% perthite K-feldspar with subordinate anhedral plagioclase is a typical texture of the compositionally zoned O’Grady batholith (Fig. 2.3a; Appendix A.3). Less abundant but coeval sub-alkaline to alkaline, granite, monzonite, granodiorite, hornblende diorite, and pyroxenite, (syeno)gabbro, and localized silica-undersaturated phases locally also form small satellitic bodies and intrusive phases associated with this suite. Fine-grained, round to elongate enclaves of biotite ± hornblende diorite composition form up to 3% of some intrusive phases (although hornblende–plagioclase porphyry, enclaves within 65 enclaves, and other mineralogical and textural varieties of uncertain protolith are also observed locally). Finally, several intrusive phases contain significant amounts of tourmaline in the form of orbicules (e.g., Lost Horses and O’Grady batholiths) that represent infilled or proto-miarolitic cavities. Late aplite and pegmatite dykes, locally abundant miarolitic cavities, and sheeted mineralized quartz veins are also locally present (e.g., the Emerald Lake pluton and the O’Grady batholith (Fig. 2.2a). As with the Mayo plutonic suite, the Tombstone suite samples have been subdivided into regional groupings due to variations in geochemical (and isotopic; Chapter 4) composition. A moderate to high degree of compositional variation is event in the geochemical plots, and in general the wide range of SiO2 can be correlated to alkalic (43–62 wt.%) versus sub-alkalic (62–72 wt.% SiO2), highly fractionated and “felsite” compositions (72–84 wt.% SiO2; Fig. 2.11; Lang, 2000). Compositional scatter is particularly pronounced for the least siliceous samples, although at least FeT and CaO Harker variations plots display two low-SiO2 trends (Fig. 2.11a): one population descends from high cafemic concentrations, and one population ascends from low cafemic concentrations until they converge at ~65 wt.% SiO2. These trends are indicative of at least two mafic parent magma compositions: a subordinate population of ferroan alkaline composition with the moderately siliceous tinguaite end-member; and, a dominant population of shoshonitic composition with the lamprophyre dykes and mafic plutonic phases end-member. Unlike most of the other plutonic suites, K2O is generally enriched at low SiO2 contents, and particularly in syenitic samples following the alkaline trend towards tinguaite compositions (Fig. 2.11a). As with the Mayo suite, high barium (and Pb and Sr; not shown) for many of the least siliceous samples is consistent with K-feldspar-crowded textures (e.g., the O’Grady batholith). The limited sample suite from the southeast Selwyn Basin is more sodic than the majority of the samples from the northwest and northeast Selwyn Basin (Fig. 2.11i), which is also a characteristic of several Tay River and Tungsten suite intrusions in the same region. A poorly defined upward inflection at ~76 wt.% SiO2 suggests that many of the trace elements and REEs become incompatible in the latest-stages of magmatic differentiation (as is indicated by the slight increase in Rb/Sr at >75 wt.% SiO2) or are enriched during sub-solidus alteration by alkaline fluids. Chromium may be anomalously enriched in high SiO2 samples as it increases with SiO2 in at least the Emerald Lake pluton (Coulson et al., 2002). Geochemical compositions for the Tombstone plutonic suite are largely indicative of an A-type magmatic affinity (e.g., relatively elevated Zn, Ce, Ba, Pb, and Nb; high Ga/Al; ferroan compositions for alkaline samples; shoshonitic > high-K calc-alkaline magma series), although its trace element chemistry is atypical of within-plate or post-orogenic environments. These intrusions are more akin to shoshonitic to high-K to medium-K calc-alkaline post-collisional magmatism, and they parallel the A-type “rapakivi topaz” granite series trend of Bonin et al. (1998; Fig. 2.11n). Many composite Tombstone suite plutons overlap temporally with the slightly older Mayo plutonic suite, and most Tombstone suite intrusions have similar sub-alkaline compositions to those exhibited by the Mayo suite (Figs. 2.10 and 2.11). In addition to its slightly younger age, however, the characteristic association with mafic and (or) alkalic plutonic phases, evidence for both shoshonitic and high-K calc- alkaline liquid lines of descent, and geochemical scatter between the alkaline and sub-alkaline magma series demonstrate that the petrogenesis of the Tombstone plutonic suite is much more complex than that of the Mayo plutonic suite. 66 2.4.9 Rancheria Plutonic Suite Three small intrusions comprise the 82–76 Ma Rancheria plutonic suite, although a series of magmatic, metamorphic, and hydrothermal ages in intrusive, hornfels, and mineralized rocks (Nelson and Colpron, 2007; Wojdak and Febbo, 2008; Canadian Geochronological Knowledgebase, 2013) extend from the Gravel Creek pluton to at least the Midway lead-zinc deposit in northern British Columbia (Fig. 2.2b). One coeval phase of the Cassiar batholith (Canadian Geochronological Knowledgebase, 2013), and possibly the Cassiar Stock (Christopher et al., 1972; Cooke and Godwin, 1984) might represent equivalent magmatic–thermal activity farther south (although the Cassiar Stock has also been correlated to the 84–78 Ma Surprise Lake plutonic suite; Anderson, 1988). Two of the three Rancheria suite plutons sampled in Yukon are highly reduced, medium- to coarse-grained, equigranular biotite–muscovite leucogranite and monzogranite. The third pluton, the Allen Stock (Figs. 2.2b and 2.3b), is composed of both an oxidized, a medium-grained, equigranular biotite monzogranite (containing trace muscovite), as well as a highly oxidized, coarse-grained biotite monzogranite with K-feldspar megacrysts typically 1 x 5 cm in size (Appendix A.3). Gregory (2007) described coarse-grained megacrystic pods that are enriched in uranium (up to 2.1% U3O8) and REEs within the K-feldspar megacrystic phase of this pluton. Small aplite dykes with granophyric intergrowths, quartz veins, and rare narrow (<5 cm) biotite-bearing and quartz-absent “mafic” dykes cut the megacrystic granite and aplite dyke phases (Gregory, 2007). It is unclear whether or not the late-stage mafic dykes are cogenetic to the Rancheria plutonic suite. Rancheria suite plutons display a moderate amount of compositional variability and do not form well defined linear trends on Harker variation diagrams, despite exhibiting a range of SiO2 contents (Fig. 2.12). Some trace elements— such as niobium (not shown), cerium, and chromium—are anomalously enriched (Fig. 2.12a), but the moderately high Rb/Sr (up to 4 in more siliceous samples; Fig. 2.12b) indicates that magmatic fractionation in this suite is not extreme. Normalization to the high-K calc-alkaline reference series produces a pattern similar to that exhibited by late-shear alkaline–peralkaline magmatism of the Tuareg Shield, which has light and heavy element enrichments and massively depleted middle elements (e.g., Ba, Sr, Ca; Fig. 2.12d). Unlike similar igneous rocks in the Tuareg Shield, however, the lightest elements (Rb to Nb; Fig. 2.12d) are more enriched in the Rancheria plutonic suite. The exception to this are several samples from the mineralized Allen Stock that display an enrichment in heavier elements (e.g., Zr to Yb; Fig. 2.12d). It could also be argued that if the handful of samples displaying the most anomalous enrichments and depletions are not considered, then the Rancheria plutonic suite would have a pattern similar to alkaline shoshonitic magma series, like the Tombstone plutonic suite (i.e., a greater light element enrichment and lesser heavy element enrichment in comparison to the alkaline–peralkaline series; Liégeois et al., 1998). Also like the Tombstone (and Mayo) suite(s), the Rancheria suite trends towards shoshonitic compositions typical of post-collisional settings (Fig. 2.12n). This plutonic suite is, thus, best classified as a predominantly sub- alkaline, alkali-calcic (shoshonitic) to high-K calc-alkalic, A-type series. 2.4.10 Black River Plutonic Suite The Black River plutonic suite is defined by two reduced 49–47 Ma intrusions of sub-porphyry to medium-grained, 67 K-feldspar-phyric, biotite (± trace muscovite) monzogranite and leucogranite. The Tintina and Black River plutons were the only intrusions sampled in the study area that have Eocene U–Pb ages (Fig. 2.3). For comparative purposes, somewhat older (ca. 66–50 Ma) scattered plutonic–volcanic rocks located in the Selwyn Basin and Cassiar Terrane are discussed together with the Black River plutonic suite. The Paleocene to Eocene igneous rocks include: • the ~65 Ma McQuesten plutonic suite in the northwest Selwyn Basin; • ~55 Ma bimodal magmatism comprising basaltic flows and rhyolitic plugs of the Ross Group scattered across the western Selwyn Basin and Yukon–Tanana inlier; • the ~52 Ma Ting pluton in the southeast Selwyn Basin; and • the ~50 Ma Mount Reed and Mount Haskin plutons located in the Cassiar Terrane in north–central British Columbia (Fig. 2.2b). Most of these samples have similar felsic compositions (except for the basaltic intrusions associated with the Ross Group volcanic rocks, and intermediate phases of the McQuesten plutonic suite) and geochemical features suggestive of an A-type affinity. Due to the wide range in SiO2 compositions and the general geochemical similarity between all of these magmatic events, Harker variation diagrams display scattered but reasonably consistent major element trends, although minor and trace element plots are less consistent (Fig. 2.13). The decreased influence of SiO2 content on normalized multi- element REE and trace element plots makes differences and similarities between the Black River suite plutons and Paleocene–Eocene igneous rocks become more apparent (Fig. 2.13c,d). The Mount Reed sample, rhyolitic members of the Ross Group, and felsic samples from the McQuesten suite plutons are the most compositionally similar to the Black River and Tintina intrusions. The Black River suite is an A-type suite of post-collisional aluminous–alkaline shoshonitic affinity, and it follows a peralkaline–alkaline trend similar to that of the Seagull plutonic suite (Fig. 2.13n; Bonin et al., 1998). 2.5 PETROGENESIS OF SYN- TO POST-ACCRETIONARY PLUTONIC SUITES IN THE NORTHERN CORDILLERAN MIOGEOCLINE The above-described plutonic suites have been classified as members of either the high-K calc-alkaline magmatic series or the alkali-calcic (to alkaline) magmatic series. A review of the occurrence and petrogenesis of these two types of magma series precedes the petrogenetic discussions for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. 2.5.1 A Review of the High-K Calk-alkaline Magma Series High-K calc-alkaline rocks are abundant in post-Archean orogenic belts and represent a large proportion of the intermediate to felsic granitoids described in the literature (e.g., Roberts and Clemens, 1993; Black et al., 1994; Liégeois et al., 1998; Kemp and Hawkesworth, 2003). Previous workers have proposed several possible petrogenetic scenarios for high-K calc-alkaline rocks, ranging from the anatexis of mafic to intermediate crust to mixtures of mantle and crustal components (e.g., Clemens, 2006). Most isotopic datasets, as well as the incompatible element 68 enrichment (typically K, U, Th, Rb, and light REEs) that characterizes high-K calc-alkaline rocks, are consistent with a protolith of mainly or entirely crustal origin (Roberts and Clemens, 1993; Clemens and Stevens, 2012). Additionally, the large volumes and general geochemical similarities exhibited by high-K calc-alkaline suites across many different orogens imply that the source rocks are not only abundant in the crust, but are of relatively consistent compositions (Roberts and Clemens, 1993). Most post-collisional high-K calc-alkaline suites commonly plot as “volcanic arc” granites on tectonic discrimination diagrams—which is thought to be a reflection of the composition of their source rocks (e.g., Roberts and Clemens, 1993)—although some trace element enrichments may cause these igneous rocks to plot in other fields (Pearce et al., 1984). Widespread crustal protoliths to these suites, therefore, may be magmatic arc-related rocks of average andesitic and medium-K to high-K calc-alkaline composition (Clemens et al., 2011). The enrichment in incompatible elements in high-K calc-alkaline magmas is partly inherited from similar (if smaller) enrichments in their medium-K to high-K calc-alkaline crustal protolith (Roberts and Clemens, 1993), and is partly the result of partial melting processes. If enriched mantle-derived magmas are also involved in the genesis of high-K calc-alkaline suites, they may also impart some incompatible element enrichment. Lithospheric mantle-derived magmas do appear to have been intimately involved in the genesis some examples of high-K calc-alkaline magmatism (e.g., as described in the West African and Variscan orogenies; Liégeois et al., 1998; Barbarin, 1999; Janousek et al., 2000; Ferré and Leake, 2001; Bonin, 2004). Mantle-derived melts, however, do not necessarily interact directly with high-K calc-alkaline magmas, although they are usually viewed as an important source of heat for anatectic processes in the crust (e.g., Roberts and Clemens, 1993; Clemens, 2003). A common association between high-K calc-alkaline suites and subordinate shoshonitic lamprophyre dykes attests to advective transfer of heat to the crust at the time of high-K calc-alkaline magmatic events (e.g., Liégeois et al., 1998) as well as high heat flow in the upper mantle (a source of conductive heat to the crust). The main source of chemical variation in most intermediate to felsic, high-K calc-alkaline (i.e., non-arc) magmatic suites, however, is not usually attributed to mixing mantle and crustal magmas (e.g., Clemens, 2003; Clemens et al., 2010; Clemens and Stevens, 2011, 2012). This is due in large part to the chemical and mechanical difficulty of mixing two compositionally distinct magmas of inherently different viscosity, density, and temperature (Watson, 1982; Sparks and Marshall, 1986; van der Laan and Wyllie, 1993; Kay et al., 1992). The assimilation of crustal rocks by basaltic magmas (e.g., Patiño Douce, 1995) is equally unlikely to be capable of altering the bulk composition of a large volume of basaltic magma (except at the basalt–host rock interface; Glazner, 2007, 2011; Clemens and Stevens, 2012). Likewise, it is thermodynamically impossible for a small volume of basaltic magma to assimilate a large quantity of crustal rocks to modify its composition to that of a granitic magma (e.g., Glazner, 2011). Finally, contamination is, by definition, a localized process that is incapable of altering the bulk composition of large volumes of magma. The most efficient method of chemical (and isotopic) ‘mixing’ between compositionally distinct components, therefore, is mechanical mixing in the source rocks prior to partial melting. In this way, the chemical variation in intermediate–felsic continental granitoid provinces primarily reflects the variation already inherent to a single crustal melt source (e.g., Kemp and Hawkesworth, 2003; Clemens et al., 2009; Clemens and Stevens, 2012). Anatectic conditions controlling the amount of peritectic minerals and entrained accessory minerals in the extracted magma are an additional control on the chemical variation in the resulting granitoids (Ayres and Harris, 1997; Zeng et al., 2005; Stevens et al., 2007; Clemens et al., 2009, 2011; Villaros et al., 2009; Farina and Stevens, 2011; 69 Clemens and Stevens, 2012). 2.5.1.1 I-Type and S-type Variants of High-K Calc-alkaline Magmas Based on partial melting experiments involving potassic minerals and the highly incompatible behaviour of potassium in general, I-type high-K calc-alkaline suites are best interpreted as derived from andesitic (to basaltic– andesitic) metaigneous rocks (Roberts and Clemens, 1993). The most voluminous high-K calc-alkaline, hornblende ± biotite bearing, intermediate–felsic magmas are generated by biotite and amphibole dehydration under fluid-absent conditions, which can produce between 30 to 60% melt (Clemens and Vielzeuf, 1987; Roberts and Clemens, 1993; Castro et al., 1999; Petford and Gallagher, 2001). This melt is hydrous, peraluminous, and leucogranitic to leucogranodioritic in composition, and in equilibrium with an anhydrous, metaluminous peritectic mineral assemblage consisting of clinopyroxene ± plagioclase, ilmenite, orthopyroxene, and possibly garnet (Patiño Douce and Beard, 1995; Clemens and Stevens, 2011, 2012; Stevens et al., 2011). The resulting magmas, thus, comprise the fluid-rich melt, entrained peritectic minerals (up to 30 mol%) (Chappell et al., 1987; Clemens and Stevens, 2011). The greater the entrainment of peritectic minerals, the less siliceous the magma becomes, which is why these I-type magmas range from dioritic to granodioritic to monzogranitic compositions (e.g., Chappell et al., 1987; Clemens et al., 2011; Chappell et al., 2012; Clemens and Stevens, 2012). If peritectic clinopyroxene or plagioclase are entrained, the magma will tend towards metaluminous compositions, but if large amounts of peritectic garnet are entrained, the magma will instead tend towards peraluminous compositions (Clemens et al., 2011). Refractory accessory minerals formerly included in the hydrous minerals (Villaros et al., 2009) and restitic components (Chappell et al., 1987) may also be entrained in the magma if they are capable of being transported by the ascending magma through narrow crustal conduits (e.g., 2–20 m thick; Clemens and Mawer, 1992; Petford et al., 1993, 2000, 2004; Clemens and Stevens, 2012). Mainly infracrustal melt sources have been proposed for the high-K calc-alkaline magmatic series; however, magmas produced by mixed infracrustal–supracrustal or solely supracrustal melt sources are not excluded from this designation (Clemens and Vielneuf, 1987; Roberts and Clemens, 1993; Clemens et al., 2011). Supracrustal rocks typically invoked for the generation of slightly to strongly peraluminous intermediate to felsic granitoids range from mature muscovite-rich metapelites to immature plagioclase and biotite-rich quartzo–feldspathic greywackes or psammites (e.g., Clemens and Wall, 1981; Harris and Inger, 1992; Sylvester, 1999). Most workers identify biotite- bearing quartzo–feldspathic protoliths to be the source of the largest volumes of S-type magmas (e.g., Miller, 1985; Patiño Douce and Johnson, 1991; London, 1995; Clemens, 2003; Clemens and Stevens, 2011, 2012). The decomposition of biotite in the source rock produces a peritectic mineral assemblage consisting of garnet, ilmenite, and orthopyroxene ± plagioclase. Differential entrainment of the peritectic assemblage and accessory mineral phases that were encased in the decomposing hydrous minerals is thought to be responsible for the compositional variation exhibited by S-type suites (Stevens et al., 2007; Villaros et al., 2009; Clemens and Stevens, 2011; Stevens et al., 2011). Magmas generated from biotite ± muscovite dehydration follow a granodioritic–monzogranitic evolutionary trend, and are usually peraluminous and potassic (e.g., Clemens, 2006). Magmas with a greater proportion of less mature quartzo–feldspathic melt sources will typically be more sodic and mafic (e.g., London, 1995), which could 70 make them difficult to chemically distinguish from I-type magmas in some cases. Infracrustal and supracrustal magmas are closely associated in time and space in many collisional and post- collisional orogenies (e.g., Chappell and White, 1974, 2000; Bonin, 2004; Anderson and Rowley, 1981; Castro et al., 1999; Pankhurst et al., 2000; Bonin, 2004; Fernandez-Suarez et al., 2000; Clemens et al., 2011). Although interactions between these two magma types may occur, and (or) magmas may be derived from ‘mixed’ melt sources, transitional magmas do not appear to be particularly common (e.g., Clemens et al., 2011). However, they do occur (e.g., Healy et al., 2004; Clemens and Stevens, 2011), and this may be one reason why some of the plutonic suites in the northern Cordilleran miogeocline share features of both I- and S-type granitoids (e.g., Table 2.5). The most likely source of transitional I- and S-type characteristics in intermediate to felsic granitoids is pre-anatectic mixing of two or more chemically distinct components. This is envisioned to occur by several mechanisms, including: • the deposition of mixed detritus eroded from source regions comprising mixed rock types (e.g., Erdmer and Creaser, 1997); • the anatexis of interlayered clastic (pelite, greywacke, and arenite) ± chemical (e.g., carbonate or evaporite) metasedimentary rocks (e.g., Kemp and Hawkesworth, 2003); • the anatexis of immature packages of layered volcano–sedimentary successions (Clemens and Stevens, 2011); or • metasedimentary rocks intruded by felsic to mafic igneous rocks (e.g., sills, dykes; e.g., Driver et al., 2000). Any, or several, of these potential avenues of ‘source mixing’ may be responsible for the range of chemical compositions and mineralogies exhibited by the igneous rocks in the northern Cordilleran miogeocline. 2.5.1.2 Temperatures and Mechanisms of Crustal Anatexis Muscovite dehydration reactions beginning at ~650°C in the presence of fluid produce volumetrically small, hydrous magmas that crystallize to form strongly peraluminous muscovite-bearing leucogranitic plutons (e.g., Barbarin, 1996, 1999). These intrusions are usually associated with compressional regional structures that are capable of focusing crustal fluids (e.g., Barbarin, 1996). The fluid-present conditions, however, produce H2O-rich magmas that cannot ascend to higher structural levels and instead are essentially ‘frozen’ in (or near) their protolith rocks (e.g., Clemens and Wall, 1981; England and Thompson, 1986; Clemens and Vielneuf, 1987; Zen, 1988; Barbarin, 1996, 1999). Under fluid-absent conditions, however, higher temperatures are required for the anatexis of a muscovite-bearing melt source (~750oC), but the resulting melt volumes remain small (<15%) and, in some cases, may not be extracted from their source (forming migmatitic complexes; Patiño Douce et al., 1990; Thompson and Connelly, 1995; Clemens and Vielneuf, 1987; Nabelek and Bartlett, 2000; Rushmer, 2001; Thompson, 2001). Large volume, peraluminous granites dominated by biotite—with lesser local muscovite, cordierite, and (or) Al2O5 polymorph minerals—requiring higher temperatures of 850oC or more at fluid-absent crustal conditions are generated from hydrous quartzo–feldspathic metasedimentary rocks (Patiño Douce et al., 1990; Barbarin, 1996, 1999; Kemp and Hawkesworth, 2003). A biotite-bearing quartzo–feldspathic protolith produces melt volumes of up to about 40% (Clemens and Vielneuf, 1987; London, 1995; Thompson and Connelly, 1995; Patiño Douce and 71 Beard, 1995; Nabelek and Bartlett, 2000; Rushmer, 2001). These magmas are likely derived in the middle to lower crust (>25 km depth; e.g., Sawyer et al., 2011), but are capable of ascending to middle to upper crustal levels because of their low initial H2O contents (Clemens and Vielneuf, 1987; Patiño Douce and Beard, 1995; Barbarin, 1999). Hornblende-bearing, tonalitic to granitic magmas are also generated by fluid-absent dehydration, although typically from amphibolitic protoliths containing amphibole ± biotite. These rocks require temperatures between 850–950oC to produce melt fractions of 30–60% (Clemens and Vielzeuf, 1987; Roberts and Clemens, 1993; Castro et al., 1999; Petford and Gallagher, 2001). The temperatures required for biotite, and especially amphibole, dehydration are rarely attained in the lower to middle crust without heat input from the mantle. Exceptions are situations in which metasedimentary rocks enriched in radioactive elements (U, Th, and K) were deeply buried, intensely over-thickened, and maintained their anomalous thickness for several 10’s of m.y. (e.g., England and Thompson, 1984; Patiño Douce et al., 1990; Nabelek and Liu, 2004; Bea, 2011). Such a scenario may be applicable to magmas sourced mainly from micaceous supracrustal rocks, which concentrate radioactive elements (e.g., K). Partial melting of hornblende-bearing amphibolitic rocks in the lower to middle crust, however, requires externally sourced heat that can only come conductively or advectively from the underlying mantle (England and Thompson, 1984; Patiño Douce et al., 1990; Brandon and Lambert, 1994; Dorais and Paige, 2000; Clemens, 2003, 2006). Upwelling of asthenosphere below an extended or delaminated lithosphere and (or) the injection of hot basaltic magmas into the crust are two important mechanisms of heat transfer to the crust (e.g., Clemens and Wall, 1981; England and Thompson, 1984; Clemens and Vielneuf, 1987; Roberts and Clemens, 1993; Barbarin, 1996; Thompson, 2000; Petford and Gallagher, 2001; Clemens, 2003; Kemp and Hawkesworth, 2003; Bonin, 2004; Bea, 2011). Of the two, basaltic under- and inter- plating is probably the most effective way of raising crustal geotherms (Roberts and Clemens, 1993; Petford and Gallagher, 2001; Bea, 2011), particularly as lithospheric thinning alone is unlikely to produce voluminous crustal melts except in crust that was already anomalously hot (Bea, 2011). 2.5.2 High-K Calc-Alkaline Plutonic Suites in the Northern Cordilleran Miogeocline The most voluminous magmas in the northern Cordilleran miogeocline are the high-K calc-alkaline plutons of Early Cretaceous age in the Cassiar Terrane, and of Early to ‘mid’-Cretaceous age in the Selwyn Basin. Early Cretaceous magmatism in both regions appears to be of mainly to entirely supracrustal origin, whereas ‘mid’-Cretaceous magmatism in the Selwyn Basin appears to be mainly (if not entirely for some plutonic suites) of infracrustal origin. Different petrogenetic interpretations for the Cassiar, the Anvil and Hyland, the Tay River, and the Mayo and Tungsten plutonic suites are discussed as follows. 2.5.2.1 Cassiar Plutonic Suite: high-K calc-alkaline (to alkaline), S > I-type Generally, the observed range of rock types, mineralogies, and major and trace element geochemical compositions exhibited by the Cassiar plutonic suite would be typical high-K calc-alkaline rocks of “low temperature I-type” affinity (Table 2.3; e.g., Chappell et al., 1998). However, normalized multi-element plots and tectonic classification 72 diagrams display a relatively wide spread of trace element data, and magma mixing and assimilation processes are unlikely to have operated extensively in this intermediate–felsic plutonic suite (e.g., Chappell et al., 2000; Driver et al., 2000; Clemens and Stevens, 2012). The chemical, mineralogical, and lithological heterogeneity exhibited by this suite implies that the protolith to these magmas was also a chemically, mineralogically, and lithologically heterogeneous crustal package (e.g., Kemp and Hawkesworth, 2003). Driver et al. (2000) proposed a model invoking the partial melting of a dominantly supracrustal package of rocks (± felsic orthogneisses) that had been intruded by intermediate to mafic Proterozoic sills for the genesis of the Cassiar batholith, the largest intrusion belonging to this plutonic suite. This model, with some modifications, may also be applicable to the remainder of the Cassiar plutonic suite. An immature, feldspathic (clay-poor) metagreywacke protolith (or a mixture of siliciclastic rocks with an average greywacke composition) could account for many of the “I-type” geochemical features exhibited by the reduced, calc-alkaline biotite ± muscovite-bearing examples of the Cassiar suite. These features include relatively high sodium and calcium, and relatively low potassium and aluminicity (A/CNK < 1.2), both of which are compatible with a plagioclase-rich melt source. Carbonate or evaporite layers, or interlayered mafic to intermediate sills and flows could also be responsible for these geochemical features (e.g., Kemp and Hawkesworth, 2003). Another possible source of elevated sodium and calcium in the supracrustal rocks may be lithic detritus that was eroded from mafic rocks. There is no shortage of possible protoliths for this plutonic suite in the northern Cordilleran miogeocline. Thick packages of Proterozoic to Paleozoic, quartzo–feldspathic sedimentary rocks are interlayered with less calcareous, pelitic, and tholeiitic to OIB-like amphibolitic lithologies (Delaney, 1981; Chamberlain, 1983; Armstrong et al, 1991; Evenchick et al., 1985; Aitchen and McMechan, 1991; Cook et al., 1991; Gabrielse and Campbell, 1991; Gordey and Anderson, 1993; Cook et al., 1999; Driver et al., 2000; Ross et al., 2001; Evenchick et al., 2005). These supracrustal rocks are inferred to comprise much of the middle, and possibly lower, crust below the Cassiar Terrane (Cook et al., 2004; Snyder et al., 2009). The association with incompatible element enriched (e.g., U, Th, Ta, Nb), oxidized, alkaline to alkali-calcic (to calc- alkaline) hornblende-bearing phases belonging to the Cassiar plutonic suite requires a similarly oxidized, alkaline, and incompatible element enriched component to the hornblende-bearing magmas. An HFSE enrichment (relative to other high-K calc-alkaline suites; Fig. 2.4d) argues for a relatively deep mantle source, whereas the LILE enrichment may be more compatible with a subduction-enriched lithospheric mantle (e.g., Liégeois et al., 1998). These intermediate–mafic magmas could be derived from large melt fractions of mafic crustal rocks, or as smaller melt fractions from the underlying mantle; both possibilities will be evaluated from an isotopic standpoint in Chapter 3. Partial melting conditions in the geochemically diverse source rocks may also partly explain the compositional variability in the Cassiar plutonic suite. Rates of melt extraction may have exceeded those of diffusive homogenization between melts or magmas generated from different lithologies (Patiño Douce, 1995). Furthermore, hornblende-bearing lithologies may not have undergone anatexis until higher crustal temperatures, which may have occurred after the genesis of most of the reduced, calc-alkaline magmas. The average ages of muscovite-bearing and 73 biotite-bearing phases are significantly older than the available ages for hornblende-bearing phases (see Table 2.2; Section 3.5.2.3, Chapter 3; Appendix G.1). This implies that the earliest Cassiar suite magmas may have been derived at lower temperatures during the muscovite ± biotite dehydration of mainly supracrustal rocks, whereas the latest Cassiar suite magmas may have been derived at higher temperatures during hornblende ± biotite dehydration of supracrustal and infracrustal rocks. This, combined with the variable entrainment of peritectic minerals and (or) restite, would produce a series of geochemically (and isotopically) variable magmas. The chemical variation in these magmas were heavily dependent on the composition and proportion of the melt fractions derived from each compositional end-member that experienced anatexis. 2.5.2.2 Anvil and Hyland Plutonic Suites: (medium-K to) high-K calc-alkaline, S > I-type As with the Cassiar plutonic suite, the intrusions that comprise the high-K calc-alkaline Anvil and medium-K to high-K calc-alkaline Hyland plutonic suites exhibit relatively heterogeneous compositions (particularly the Hyland plutonic suite) and appear to be predominantly supracrustally derived. However, the scope of the petrogenetic interpretations for the Anvil and Hyland plutonic suites presented in this section is limited by the number of samples available for these suites, making the following interpretations potentially subject to later revisions. Additional work on these two plutonic suites (including detailed petrography, geochemistry, and isotope chemistry) is needed in order to better constrain their composition and potential melt sources. The high-K calc-alkaline Anvil suite comprises slightly to moderately peraluminous, highly siliceous, near- minimum temperature, leucocratic monzogranite magmas. The inferred low magma temperatures are compatible with zircon textures (grains from the one sample examined here have an abundance of large resorbed inherited cores with thin magmatic rims; Appendices B.3 and B.4) and with the presence of older inheritance detected in one of the samples dated by ID–TIMS methods (Table 2.2). The Anvil suite is similar to the strontium-poor leucogranites in the Himalayas described by Kemp and Hawkesworth (2003), and like the Himalayan examples, is compositionally and mineralogically compatible with a derivation from metapelitic rocks. Near-minimum temperature leucogranites are thought to have formed by the partial melting of pelitic lithologies via mica dehydration at moderate to low magmatic temperatures (ca. 650–850oC; Clemens and Vielzeuf, 1987; Harris and Inger, 1992; Barbarin, 1996; Kemp and Hawkesworth, 2003). Magmas resulting primarily from the dehydration of mica in such rocks are also moderately to strongly hydrous (e.g., ca. 4–8 wt.% H2O, depending on the amount of biotite and muscovite involved, respectively; Wyllie, 1977; Miller, 1985). These fluid-rich magmas become saturated in H2O with small decreases in pressure and (or) temperature and typically crystallize in an in situ, or near-in situ, position as migmatite complexes or segregated small leucocratic plutons, both hosted in amphibolite-grade metapelitic rocks (Patiño Douce et al., 1990; Kemp and Hawkesworth, 2003). The biotite–muscovite bearing Anvil plutonic suite appears to have ascended above its source rocks, but is still emplaced into relatively high-grade host rocks (e.g., Smith and Erdmer, 1990). The Hyland plutonic suite is a medium- to high-K calc-alkaline suite that displays more heterogeneous compositions that are suggestive of a derivation from lithologically heterogeneous supracrustal rocks. Pre-anatectic mixing of 74 interlayered lithologies or depositionally mixed detritus, is the most likely explanation for the observed heterogeneity. All mineralogical and geochemical evidence presented here suggest that the supracrustal melt source to this suite was dominantly of quartzo–feldspathic composition. Quartzo–feldspathic rocks will produce mainly biotite-bearing (± cordierite) magmas, although slightly more or less peraluminous muscovite ± biotite-bearing leucogranites (e.g., Barbarin, 1996) phases are possible with smaller melt fractions or greater entrainment of peritectic garnet (e.g., Stevens et al., 2007). Cordierite and Al2O5 polymorph minerals have not been reported for the Hyland plutonic suite (e.g., Zen, 1988; Barbarin, 1996, 1999), although this may be due to the limited sampling as these minerals are not necessarily ubiquitous to all intrusive phases. Like the Cassiar plutonic suite, the associated hornblende-bearing intrusive phases require some input of melt from intermediate to mafic igneous rocks— however, a genetic link between the bulk of the biotite-bearing Hyland plutonic suite and locally observed hornblende-bearing phases has not been tested chemically or chronologically. Migmatitic contact aureoles and concordant laccolithic geometries characterizing the Hyland suite batholiths (e.g., Roots et al., 1966; Hart and Lewis, 2006) may be compatible with a near-in situ derivation from their host Neoproterozoic Hyland Group rocks. Certainly the Hyland Group—which comprises immature quartzo–feldspathic “grits” interlayered with calcareous, pelitic, and amphibolitic rocks (e.g., Gabrielse and Blusson 1969; Gabrielse et al., 1973a,b; Gordey and Anderson, 1993)—is a compositionally appropriate melt source for the Hyland plutonic suite. However, biotite-bearing magmas emplaced into migmatitic complexes are not necessarily derived in situ; instead, they are thought to originate at greater depths and higher crustal temperatures (Patiño Douce et al., 1990). Metamorphic minerals in the host rocks suggest burial depths of ~15 km (Moynihan, 2013), which is not deep enough to generate voluminous biotite-bearing magmas. 2.5.2.3 Tay River Plutonic Suite: high-K calc-alkaline, I-type The Tay River plutonic suite in both the eastern and western Selwyn Basin, and associated South Fork volcanic rocks in the western Selwyn Basin, are also classified as high-K calc-alkaline magma series. These igneous rocks have a geochemical homogeneity that—aside from a few small regional variations (e.g., a depletion in HFSEs and a lower range of SiO2 in the western Selwyn Basin)—is not apparent in the S-type plutonic suites. The similar geochemical compositions exhibited by this suite across the Selwyn Basin may be indicative of either a widespread infracrustal melt source (Chappell et al., 2000; Clemens, 2003) or widespread supracrustal melt source that is more homogeneous than that inferred for the Cassiar and Hyland plutonic suites. The Tay River intrusions also appear to have evolved from a relatively sodium-poor parental magma or melt sources in comparison to the S-type Cassiar and Hyland plutonic suites. A homogeneous sodium-poor protolith could be a mature pelitic sedimentary rock (e.g., Chappell and White, 1974); however, it could also be an infracrustal rock that did not contain abundant sodic plagioclase. The abundance of hornblende-bearing intrusive phases of intermediate composition, and the lack of chemical heterogeneity in this plutonic suite, are more compatible with a mainly to entirely infracrustal protolith. There are several possible types of infracrustal rocks that could be present in the basement to the Selwyn Basin. Mafic rocks of tholeiitic affinity (i.e., relatively siliceous and Na-poor; Sun and McDonough, 1989) may have been underplated to the crust during numerous Proterozoic episodes of mafic magmatism in the crust (Thorkelson et al., 75 2005). Alternatively, a mixture of magmas derived from enriched lithospheric mantle and juvenile mafic crust has also been invoked in the generation of some high-K calc-alkaline rocks elsewhere (Liégeois et al., 1998). Finally, a derivation from mafic to intermediate-composition rocks similar to those generated above subduction zones has been suggested for high-K calc-alkaline magmatism with I-type affinities (Roberts and Clemens, 1993). Of these, the latter possibility seems to be the most appropriate given the slight LILE and HFSE (Rb, Th, U, Nb, and Ta) enrichments and cafemic (e.g., Fe, Ca, Ti) depletions compared to the high-K calc-alkaline suites examined by Liégeois et al. (1998). These geochemical features imply a protolith of intermediate composition for the Tay River plutonic suite (e.g., containing ~60 wt.% SiO2 in the western Selwyn Basin, and ~65 wt.% SiO2 in the eastern Selwyn Basin). Experimental studies suggest that voluminous magmas with a range of compositions similar to those of the Tay River suite are generated by high-temperature (850–950oC), vapour-absent partial melting of rocks of bulk intermediate composition in the lower crust (e.g., Clemens and Vielzeuf, 1987). The associated high-K calc- alkaline (quartz) monzodioritic dykes in the eastern Selwyn Basin are strongly enriched in uranium and thorium, suggesting that they were sourced in an LILE enriched (although not quite shoshonitic) mantle (e.g., Sun and McDonough, 1989; Liégeois et al., 1998). These mafic intrusions do provide evidence of advective heat transfer to the crust in the ‘mid’-Cretaceous, which would have been important in hornblende ± biotite dehydration reactions. Thus, an infracrustal protolith of slightly LILE enriched, intermediate composition may best explain the mineralogical and geochemical variation (or lack thereof) within this plutonic suite. One implication of this interpretation is that crystalline basement of intermediate, possibly magmatic arc, composition could underlie much of the Selwyn Basin. Compared to the bulk of the plutonic Tay River (and volcanic South Forks) samples across the Selwyn Basin, the Coal River sub-suite in the eastern Selwyn Basin displays even less geochemical variability. This sub-suite also appears to have been derived from either a more sodic melt source, or from relatively sodic parental magmas. The higher strontium and calcium (as well as lower Rb, Ti, Fe, Y, and REEs—even when normalization accounts for differing SiO2 content; Fig. 2.8d)—are compatible with a more mafic melt source (Petford and Atherton, 1996). The partial melting of mafic rocks containing hornblende and plagioclase typically produces a sodic, peraluminous melt and a metaluminous peritectic assemblage containing anorthitic plagioclase and clinopyroxene, although garnet can also be common (Ellis and Thompson, 1986; Rutter and Wyllie, 1988; Patiño Douce et al., 1990; Patiño Douce and Beard, 1995; Jung and Pfander, 2007; Clemens et al., 2011). The small size of the Coal River intrusions, and their increasing peraluminicity (related to increasing Na2O and increasing Al2O3) with decreasing SiO2, suggests that these intrusions were formed as smaller melt fractions that may have been dominated by peritectic garnet rather than metaluminous peritectic minerals. It is also conceivable that the Coal River sub-suite instead formed from partial melting of an immature but homogeneous, lithic- and sodic plagioclase-rich metasedimentary rock, which would also produce similar trends with respect to peraluminicity. This type of supracrustal melt source, however, would have to have been in close proximity its source of detritus, which itself would have to consist of mafic rocks containing hornblende and plagioclase. 76 2.5.2.4 Mayo and Tungsten Plutonic Suites: high-K calc-alkaline to alkali-calcic (with late shoshonitic), I- (± S-) type Both the Mayo and Tungsten suites comprise high-K calc-alkaline intrusions that are associated with late-stage shoshonitic lamprophyre dykes, including kersantites (eastern Selwyn Basin), and minnettes and spessartites (western Selwyn Basin; Mair et la., 2011). The Mayo plutonic suite has two main compositional trends that are defined by bimodal K2O concentrations at low SiO2 contents. These include: a) a dominant calcic to (high-K) calc-alkalic trend (Fig. 2.10g) with a quartz diorite–quartz monzodiorite– granodiorite–monzogranite rock association (Fig. 2.10e), which represents the evolution of a typical calc- alkaline magma series and is similar to the compositional range of the Tay River plutonic suite; and b) a subordinate alkali-calcic (shoshonitic) to calc-alkalic trend (Fig. 2.10g) with a monzonite–quartz monzonite–monzogranite rock association (Fig. 2.10e), that includes samples with K-feldspar-crowded textures. Associated mafic dykes are mostly monzonitic and form the least siliceous end-member to the latter rock association (although one mafic dyke is granodioritic and considered to be a member of the calc-alkaline series). The mainly high-K calc-alkaline composition of the Mayo plutonic suite is consistent with a derivation from mainly intermediate (>60 wt.% SiO2 and Na-poor) infracrustal rocks, similar to the protolith inferred for the Tay River plutonic suite. This protolith would have been responsible for geochemical and mineralogical variations in the Mayo plutonic suite that correspond to the dominant calc-alkalic–calcic trend towards lower SiO2 concentrations. Overall, the petrogenesis of the Mayo plutonic suite appears to be similar to that inferred above from the Tay River plutonic suite. In the Mayo plutonic suite, late-stage shoshonitic lamprophyres provide a heat source for high temperature partial melting of intermediate infracrustal rocks, but they may also have interacted to a small degree with the small population of crustally derived magmas that trend towards alkali-calcic compositions. The Tungsten plutonic suite is monzogranitic–(granodioritic) in the northeast Selwyn Basin, and monzogranitic in composition in the southeast Selwyn Basin. Associated mafic dykes (at least in the southeast Selwyn Basin) are kersantitic lamprophyres with silica-undersaturated monzonitic compositions. Despite similar field, macroscopic, and petrographic observations for all of the Tungsten suite plutons, these intrusions could potentially be classified as either S- or I-type granites. Higher Na/K and an increase in normative corundum with decreasing total iron for the Tungsten suite in the southeast Selwyn Basin are indicative of low-temperature I-type granites. In the northeast Selwyn Basin, however, this suite is characterized by lower peraluminicity (typical of minimum temperature melts of either I- or S-type affinity) and Na/K, a slight relative enrichment in LILEs, HFSEs, and REEs, and a possible (although questionable) positive correlation between normative corundum and total iron (Fig. 2.09f,i,j). These chemical features suggest that Tungsten suite plutons in the northeast Selwyn Basin may be near-minimum melts of supracrustal origin, but may also have been imparted by a combination of a small-fraction partial melting, greater refractory mineral entrainment (Villaros et al., 2009), and (or) magmatic fractionation or differentiation (which includes contamination by sedimentary host rocks). 77 Chemical similarities, and temporal and geographic overlap between the Tungsten suite and the Mayo and Tay River suites could imply that the former are highly fractionated equivalents of the latter two infracrustally derived suites (e.g., Mortensen et al., 2000). Alternatively, given more recent studies of the source of chemical variation exhibited by granitic rocks (e.g., Clemens and Stevens, 2012), the Tungsten suite could comprise more siliceous and incompatible element enriched, magmas derived by small-fraction partial melting of the same infracrustal melt source as the Mayo and Tay River plutonic suites (as is also consistent with their slightly elevated Rb/Sr and Eu/Eu*, and lower Na contents). An association with late shoshonitic dykes for all Tungsten suite intrusions may indicate a cogenetic relationship to the Mayo plutonic suite, which is also associated with late-stage lamprophyre dykes. The Tungsten (and Tombstone; Figs. 2.2a and 2.3b) suite intrusions in the southeast Selwyn Basin, therefore, may be the only representatives of the TTB in this region. 2.5.3 A Review of the Alkali-Calcic to Alkaline Magma Series There is a great variety of definitions and interpretations for intrusions of “alkaline” composition in the literature. In the broadest sense, the basic definition of alkaline magmatism is based on the abundance of total alkalis (K2O + Na2O) relative to SiO2 concentrations from whole rock chemistry (Irvine and Barager, 1976). The alkaline versus calc-alkaline classification has evolved, however, to distinguish differing proportions of alkaline and calcic components within a magma series, using the “MALI” index (K2O + Na2O – CaO; Frost et al., 2001, after Peacock, 1931). Alkaline plutonic suites are considered to be a variation of the “A-type” granitoids (although A-type granitoids include many magmatic series that would not be classified as “alkaline” based on the above definition). The A-type granitoids are generally subdivided into two main categories (e.g., Whalen et al., 1987; Poitrasson et al., 1995; King et al., 1997; Bonin et al., 1998): • dominantly alkali-calcic (aluminous) granitoids that range from metaluminous to peraluminous compositions, and • dominantly alkaline (sub-aluminous) granitoids that range from peralkaline to peraluminous compositions. A-type granites, however, are rarely limited to the above end-members and various hybrid descriptions thereof are abundant in the literature. These ‘hybrid’ compositions have been attributed to either mixing of magmas from distinctly different melt sources, or extensive crustal contamination of magmas sourced in the mantle (e.g., Eby, 1990, 1992; Turner et al., 1992; Frost and Frost, 2011). Magmas of A-type affinity include rocks that have a wide range of SiO2 concentrations and are as diverse as bimodal (rhyolite–basalt) ring complexes and calderas, peralkaline to metaluminous to strongly peraluminous granitoids, topaz leucogranites, rapakivi granites, syenites, gabbros, and anorthosites (e.g., Bonin et al., 1998; Frost and Frost, 2011). All of the alkali-calcic to alkaline plutonic suites discussed above (e.g., Tombstone, Seagull, Rancheria, and Black River) may be generally classified as “alkali-calcic” (aluminous) A-type granitoids after Poitrasson et al. (1995) and Bonin et al. (1998), although the Tombstone plutonic suite also has a significant alkaline component. The A-type plutonic suites in the northern Cordilleran miogeocline are of post-collisional or post-orogenic, rather than anorogenic, origin. Truly anorogenic settings are regions of within-plate magmatism where deep mantle 78 upwelling at either a hot spot or the locus of continental rifting produces mantle and (or) crustal melts. Post- collisional or post-orogenic A-type plutono–volcanic rocks are emplaced in association with localized, late- to post- orogenic extension along lithosphere-scale structures, or in response to post-collisional asthenospheric upwelling triggered by the delamination of over-thickened or subducted lithosphere (e.g., Whalen et al., 1987; Eby, 1992; von Blankenburg and Davies, 1995; Davies and von Blankenburg, 1995; Frost and Frost, 1997; Liégeois et al., 1998; Kemp et al., 2005). Although post-collisional alkali-calcic and alkaline granitoids share compositional affinities with within-plate or anorogenic granitoids, they commonly have trace element signatures that are suggestive of volcanic arc or syn-collisional origins (e.g., Pearce et al., 1984)—as exemplified by the Tombstone plutonic suite (see Fig. 2.11m). These apparent pre- to syn-collisional affinities the greater input of crustal or subduction-enriched sub- crustal melts in post-collisional to post-orogenic A-type magmas. 2.5.3.1 Mantle and Crustal Melt Sources of Alkali-Calcic to Alkaline Magmas A transition from calc-alkaline I- and S-type (or hybrids thereof) magmas to alkaline A-type magmas has been documented in many locations, including the Lachlan Fold Belt (e.g., King et al., 1997; Soesoo and Nicholls, 1999), the Variscan orogeny in southwest Europe (e.g., Bonin et al., 1998; Paquette et al., 2003; Bonin, 2004), the Tuareg Shield of the West African craton (e.g., Liégeois et al., 1998), and the Damara orogen in Namibia (Jung et al., 1998). This transition from high-K calc-alkaline to alkali-calcic and (or) alkaline magmatism typically occurs in response to a changing geodynamic regime, commonly marked by the end of a collisional orogeny and the initiation of post- collisional processes to anorogenic processes (e.g., Rogers and Greenberg, 1990; Liégeois et al., 1998; Bonin, 2004). One result of the wide range of compositional variations permitted under the classification of “A-type granites” is that there are as many suggested melt sources as there are compositional variations of these magmas (Collins et al., 1982; Clemens et al., 1986; Whalen et al., 1987; Eby, 1990, 1992; Creaser et al., 1991; Turner et al., 1992; Bonin et al., 1998; Liégeois et al., 1998; Liverton and Bohtelho, 2001; Bonin, 2004; Frost and Frost, 2011). A-type magmas are, however, linked by several common characteristics, including high alkali (Na or K) components relative to aluminum, high temperatures, low viscosities and H2O contents, and distinct fluorine–iron–HFSE–REE enrichments. Suggested protoliths cover both crust and mantle end-members, but also include a variety of crust– mantle hybrids (Eby, 1990, 1992, 2011; Bonin et al., 1998; Kemp and Hawkesworth, 2003; Bonin, 2004; Frost and Frost, 2011). Proposed crustal sources vary from anhydrous mafic to felsic lower crust (e.g., dehydrated paragneiss, granulite/restite, anhydrous mafic mineral cumulates, “charnockites”, and underplated tholeiitic basalt), to hydrous or non-restitic intermediate to felsic infracrustal rocks (e.g., quartz diorite, tonalite, and granodiorite). Even if these magmas are sourced entirely from the crust, the addition of mantle-derived heat is needed to generate partial melts in at least the anhydrous protoliths, as well as the observed lack of zircon inheritance and high zircon saturation temperatures in these suites. Given the high magma temperatures, it is conceivable that the contamination of crustally derived magmas by, and possible hybridization with, magmas derived from the sub-continental mantle would be a common process (e.g., Bonin et al., 1998; Barbarin, 1999; Faure, 2001; Kemp and Hawkesworth, 2003; Eby, 2011; Frost and Frost, 2011), barring physiochemical restrictions on magma mixing. Possible mantle sources of magma include small melt fractions derived from enriched portions of the lithospheric mantle, extensively 79 fractionated tholeiitic ‘basalts’ (i.e., large melt fractions of depleted lithosphere; Turner et al., 1992; Vander Auwera et al., 2003; Namur et al., 2011), or smaller-fraction partial melting in undepleted upwelling asthenospheric mantle. Although the involvement of the mantle is usually obvious in A-type magmas generated in continental rift environments and (or) above hot spots, it is also apparent in many examples of post-collisional to post-orogenic A- type granites (Frost and Frost, 1997; Liégeois et al., 1998; Faure, 2001; Bonin, 2004). It may be argued that the characteristic trace element enrichments of A-type magmas are determined by the composition of the protolith rocks. Additional controls on the extreme geochemical enrichments or depletions that typify A-type granitoids is the derivation of these magmas by either unusually small partial melt fractions (e.g., <5% melt; Liégeois et al., 1998; Faure, 2001) and (or) extreme degrees of magmatic fractionation (Turner et al., 1992; Faure, 2001; Kemp and Hawkesworth, 2003; Vander Auwera et al., 2003; Namur et al., 2011). Some specialized processes have also been proposed for the generation of incompatible element enrichments and alkaline or peralkaline compositions (e.g., volatile-induced partial melting of the lower crust, assimilation of alkali-rich rocks such as evaporite, and alkali–halogen metasomatism during and after emplacement, among others), but may not be as important. Thus, two main A-type magma series derived from distinct source regions have been geochemically discriminated on the basis of HFSEs (Eby, 1992). The alkali-calcic (aluminous) A-type granitoid series is typically thought to have been emplaced in post-orogenic, continental settings where either the crust or sub-continental mantle are important sources of magma (Eby, 1992). The alkaline (sub-aluminous) A-type granitoid series is typically thought to have been emplaced in continental rifting or within-plate settings where an anorogenic asthenospheric (OIB-like) magma source appears to be dominant (Bonin, 1990; Eby, 1992) 14. The A-type plutonic suites in the northern Cordilleran miogeocline are plotted on an HFSE ternary plot from Eby (1992) in Figure 2.15, which is used to distinguish these two main magma sources. The Rancheria suite clearly corresponds to a continental (crustal or mantle) magma source on the basis of elevated niobium and cerium contents relative to yttrium. The Tombstone, Seagull, and the Eocene suites, however, have attributes of both continental and asthenospheric magma sources (Fig. 2.15). While this discrimination is in no way definitive, it does provide a starting point for discussing potential melt sources to these plutonic suites. 2.5.4 Alkali-Calcic to Alkaline Plutonic Suites in the Northern Cordilleran Miogeocline Smaller volume granitoids of A-type affinity emplaced into post-collisional settings typically post-date the main volume of high-K calc-alkaline magmatism, although geographic overlap between these distinct magma series is common (e.g., Bonin et al., 1998; Liégeois et al., 1998; Bonin, 2004). The northern Cordilleran miogeocline is no exception. In the Cassiar Terrane, the Seagull plutonic suite immediately post-dates the cessation of back-arc magmatism and was emplaced as a narrow belt that overlaps with the westernmost Cassiar plutonic suite. The younger Rancheria and Black River plutonic suites in the Cassiar Terrane are also localized in the vicinity of major structures, but generally east of the Cassiar plutonic suite. In the Selwyn Basin, the Tombstone plutonic suite comprises the youngest of the ‘mid’-Cretaceous plutons. Although the Tombstone plutonic suite is the most 14 The removal of aluminum by anorthite fractionation in transitional to alkaline basaltic magmas from the deeper mantle may be responsible for the common peralkaline composition of these magmas (Frost and Frost, 2011). 80 ‘inboard’ magmatism in the northern Selwyn Basin, but it overlaps geographically with earlier high-K calc-alkaline magmas of the Mayo and Tay River plutonic suite, particularly in the eastern Selwyn Basin. 2.5.4.1 Seagull, Rancheria, and Black River Plutonic Suites: aluminous alkali-calcic (high-K calc- alkaline) ± alkaline, A-type The Seagull plutonic suite comprises extremely fractionated monzogranites and leucogranites of reduced, aluminous alkali-calcic (but nearly peralkalic) composition. These compositions are typical of A-type suites that are transitional from high-K calc-alkaline to alkaline post-orogenic settings (Barbarin, 1990; Liverton and Botelho, 2001; Liverton et al., 2005). This plutonic suite plots within the post-collisional to within-plate fields on tectonic discrimination diagrams, although the lepidolite topaz leucogranites from the Thirtymile pluton trend towards syn-collisional compositions as a result of their greater degree of fractionation (i.e., elevated Rb; Fig. 2.6; e.g., Liverton and Botelho, 2001; T. Liverton, 2012, pers. comm.). Mineral fractionation and late-stage “ultra-fractionation” (Liverton and Botelho, 2001) may have partly controlled the most extreme trace element enrichments (e.g., Rb, Th, Nb, Ta, F, HREE) and depletions (e.g., Sr, Ba, P, Ti, Eu) exhibited by the Seagull suite. Other trace elements (Ce, Ga, Y, and Yb), however, do not appear to have been enriched during fractionation (i.e., they do not correlate well with Rb/Sr and Eu/Eu*), and attest to a role for either the derivation of this suite from a protolith enriched in these trace elements, or a derivation by particularly small-fraction partial melting. The peralkaline–peraluminous trends exhibited by the Seagull plutonic suite in the Shand’s Index plot is typical of the alkaline A-type magma series. The greater ‘compatible’ (“SNY”, or Y, Ce, Sm, and Yb) to ‘incompatible’ (“SNX”, or Rb, Th, U, and Nb) element enrichment also characterizing this plutonic suite is compatible with an alkaline, lower lithospheric to asthenospheric mantle origin (e.g., Eby, 1992; Liégeois et al., 1998). Alkaline granitic magmas may be modified to more aluminous alkali-calcic, or even calc-alkaline compositions, if extensive interactions with the crust have occurred (e.g., Frost and Frost, 1997, 2011). A few of the less geochemically enriched (but highly siliceous), peraluminous Seagull suite magmas may have been ‘diluted’ by the incorporation of material from an intermediate quartzo–feldspathic infracrustal protolith15 (at depth; Frost and Frost, 2011). In this way, the three SNX populations of variably enriched ‘compatible’ elements in Figure 2.6n could have been controlled by differing degrees of ‘mixing’ between siliceous alkaline and high-K calc-alkaline magmas. The samples with the lowest SNX components (i.e.., <1) may have been diluted by the incorporation of high-K calc- alkaline infracrustal magma. These samples include a porphyry phase of the Thirtymile pluton that locally contains granodioritic enclaves, which is suggestive of a mixed A- and I-type petrogenesis (intermingled textural and chemical zoning in the Seagull batholith has also been observed and may provide additional evidence for magma mingling; ; T. Liverton, 2012, pers. comm.). Overall, given the extreme trace element enrichments and depletions in the majority of the Seagull suite plutons, a highly fractionated A-type magma derived from an undepleted, upper 15It is notable that the sample of the Quiet Lake batholith (KR-07-03) displays the smallest geochemical enrichments and depletions, indicating that it is neither highly fractionated nor derived entirely by small-fraction partial melting. This sample was included in the Seagull plutonic suite initially on the basis of age rather than chemical composition; it may represent the composition of crustally derived high-K calc-alkaline magma that interacted with alkaline mantled-derived magma. 81 asthenospheric parental magma appears to dominate this plutonic suite. Almost all samples of this plutonic suite, however, may have been modified to some degree through interactions with cogenetic peraluminous calc-alkaline magmas sourced from high-K calc-alkaline magmas that may have been derived from a tonalitic to granodioritic infracrustal protolith (e.g., Frost and Frost, 2011). The Rancheria plutonic suite is mainly alkali-calcic and monzogranitic in composition, although two samples approach monzonitic compositions (Fig. 2.12e). This suite is peraluminous and has trace element enrichments (e.g., Ga, Ta, Yb, and Nb) that are more typical of A-type magmas thought to have been derived from the crust or underlying lithospheric mantle (e.g., Eby, 1992; Bonin et al., 1998). This shoshonitic to high-K calc-alkaline association has been attributed to partial melting of enriched layers in the lithospheric mantle (Liégeois et al., 1998), although a crustal magma source cannot be entirely ruled out (particularly as inherited zircon cores were observed in some of these intrusions (Appendices B.3 and B.4). The most geochemically enriched Rancheria suite samples are also the least siliceous and the least fractionated (relative to Rb/Sr), which is compatible with the suggested incompatible element enriched lithospheric magma source. The Black River plutons are less definitive in terms of their petrogenesis, in part due to the limited range of compositions exhibited by these two highly siliceous samples. These samples do, however, appear to have compositions that are compatible with an A-type magmatic affinity. Like the Seagull plutonic suite, the Black River plutonic suite appears to have been more enriched in ‘compatible’ elements (SNX) than ‘incompatible’ elements (SNY), as would be consistent with the alkaline–peralkaline A-type series (Bonin et al., 1998)—even though these two intrusions are not alkaline. This plutonic suite may, therefore, have been derived from highly fractionated basaltic magmas in the deep mantle (e.g., Liégeois et al., 1998), a hypothesis that is also supported by the absence of textural evidence for zircon inheritance in these two plutons (Table 2.2). Scattered examples of latest Cretaceous to Eocene igneous rocks, many of which are similarly associated with strike-slip faults, throughout the northern Cordilleran miogeocline (Fig. 2.2b) are geochemically similar to the Black River intrusions. The Black River plutonic suite, therefore, is likely one of many episodic, small-volume, and highly localized A-type magmatic events in the northern Cordilleran miogeocline (Fig. 2.2b; e.g., Nelson and Colpron, 2007). 2.5.4.2 Tombstone Plutonic Suite: aluminous alkali-calcic (high-K calc-alkaline) > peralkaline, mixed A-/I-type Generally, the Tombstone plutonic suite is dominated by reduced to slightly oxidized, alkali-calcic monzogranites to alkalic monzonites. The macroscopic, mineralogical, microscopic, and compositional characteristics of these intrusions collectively suggest that this suite was derived from a mixture of two parental magmas. An alkaline (shoshonitic) and LILE-rich magma produced the plutons with the largest range of SiO2 contents. Several tinguaite and mafic phases are characteristic of this shoshonitic (K-rich) end-member (although the least siliceous mafic intrusions border on high-K calc-alkaline to shoshonitic compositions; Fig. 2.11h). Enrichment in LILEs and LREEs in the Tombstone suite is correlated with total alkali content (specifically K) and is the most elevated in the tinguaite phases (Fig. 2.11c,d). These enrichments are clearly inherent to this alkaline end-member, and may be the product of 82 its derivation in an enriched lithospheric or upper mantle (e.g., Hart et al., 2004a). A high-K calc-alkaline magma with SiO2 contents greater than ~65 wt.% forms the second end-member composition to the Tombstone plutonic suite, although a small number scattered low to medium-K calc-alkaline samples (some with lower SiO2) are also observed (Fig. 2.11h). This high-K calc-alkaline end-member was probably derived from the partial melting of infracrustal rocks underlying the Selwyn Basin that are similar to those suggested for the other ‘mid’-Cretaceous plutonic suites (Tay River, Mayo, and Tungsten) in this region. It is clear from the geochemical data plotted in Figure 2.11, however, that the crustally derived high-K calc-alkaline magmas must have variably interacted with the alkaline mantle-derived magmas. Not only do these two compositionally distinct trends merge at moderately high SiO2 contents (>67 wt.% SiO2), but there is much scatter between the two end-member compositions at lower SiO2 contents (Fig. 2.11h). The appropriateness of widespread magma mixing as an explanation for the geochemical and isotopic variation within plutonic suites is questionable in many instances (e.g., Glazner, 2007, 2011; Clemens and Stevens, 2012), particularly for intermediate–felsic granitoid provinces (Clemens, 2003; Clemens et al., 2010; Clemens and Stevens, 2011). An origin by mixing of such disparate magmas, however, is consistent with the presence of cogenetic mafic intrusive phases, and previous petrogenetic interpretations made for this plutonic suite (e.g., Mortensen et al., 2000; Hart et al., 2004, 2005). Mixing is also typically responsible for branched, but otherwise continuous linear arrays on Harker variation plots (including samples crossing the fields of low- to high-K calc-alkaline and shoshonitic rocks; Fig. 2.11a,h) that cannot be attributed to fractional crystallization or restite removal (e.g., Roberts and Clemens, 1993; Ferré and Leake, 2001). Other evidence for the mixing of magmas in the Tombstone suite includes complex feldspar zoning (Table 2.3; Appendix A.1), a major resorption event observed in zircons from at least the eastern Selwyn Basin (e.g., 96–95 Ma magmatic zircon cores with ca. 93 Ma rims; Tables 2.2, 2.3), and the presence of mafic through to felsic intrusive phases in zoned intrusive complexes. Mixing or mingling textures are also observed at the level of emplacement in some of the Tombstone suite intrusions in the northeast Selwyn Basin. Disaggregated dm-scale fine-grained mafic xenoliths or dykes and abundant cm-scale rounded fine-grained mafic-rich (biotite ± hornblende) ‘clots’ (Appendix A.1) might provide some evidence of magma mingling processes. 2.6 METALLOGENIC ASSOCIATIONS FOR MAGMATISM IN THE NORTHERN CORDILLERAN MIOGEOCLINE Much of the ability of an intrusion to produce mineralization depends on the extent to which magmatic processes operate during the ascent, emplacement, and crystallization of a magma (e.g., Christiansen and Keith, 1996). In reduced tungsten skarn systems, for example, greater potential for related mineralization exists if the associated intrusion is relatively deeply emplaced and has time to cool slowly and evolve a tungsten enriched residual fluid phase (Newberry, 1998; Newberry and Swanson, 1986). Conversely, in copper porphyry systems mineralization is associated with relatively shallow levels of emplacement and early decompression-related fluid exsolution that produces the characteristic porphyry texture of the associated intrusions. These processes are broadly related to the tectonic setting of magmatism and the composition of the magma (e.g., initial H2O content). The composition of the magma is also a major control on the availability of specific ligands and metals for forming mineralization. 83 Chlorine-, fluorine-, and sulfur-rich magmas (presumably derived from melt sources that are themselves relatively enriched in these volatile species) will each produce distinctive mineral deposits that are associated with different tectonic settings (e.g., Christiansen and Keith, 1996). Thus, the temporal and compositional characterization of igneous rocks, which is partly controlled by the geodynamic setting and is commonly associated with a certain style and type of mineralization, is a fundamental aspect of regional metallogenic studies. Metallogenic associations for most of the Cretaceous plutonic suites in the northern Cordillera have been summarized previously (e.g., Sinclair, 1995; Anderson, 1988; Lang, 2000; Mortensen et al., 2000; Hart et al., 2004, 2005; Rasmussen et al., 2007). The following discussion is, therefore, restricted to a basic description of the main metal tenor of the plutonic suites outlined above. Figure 2.16 locates all documented mineral occurrences that are known to be, or appear to be, associated with Cretaceous and younger magmatism in the study area (after Declerk, 2003). Important mineral deposits and occurrences outside of the study area that are mentioned below are located in Figure 2.2. 2.6.1 High-K Calk-Alkaline “S-type” Suites The Cassiar plutonic suite is associated with tungsten ± molybdenum porphyry, skarn, and vein occurrences and deposits (e.g., Northern Dancer, Risby, Stormy; Fig. 2.16), as well as proximal argentiferous lead–zinc vein, skarn, and manto-type mineralization (possibly including the Ketza River Au-rich manto deposits; Fig. 2.16) and copper skarn and vein occurrences. These metallogenic signatures are very similar to those of the compositionally similar, Early Cretaceous Hyland plutonic suite in the eastern Selwyn Basin (see below). Rarer mineral occurrence types associated with some Cassiar suite plutons include uranium porphyries, fluorine-rich veins, gold-rich veins, and precious stones (Fig. 2.16). The Anvil plutonic suite in the Yukon–Tanana inlier is most notably associated with the Tsa Da Glisza vein-hosted emerald deposit (Fig. 2.16), although the majority of associated occurrences are proximal tungsten skarns and distal silver–lead veins. Farther north in the western Selwyn Basin, the Anvil plutonic suite is also associated with minor silver-rich veins (e.g., Ag–Sn–base metal veins and breccias), as well as silver–(Pb–Zn ± Cu) skarn or replacement- type occurrences near the margins of the Anvil batholith (Fig. 2.16). In the eastern Selwyn Basin, the Hyland plutonic suite is associated with a series of tungsten skarns and intrusion-hosted molybdenum or copper occurrences along the outer fringes of the intrusions, concentrated in the northern portions, whereas there is a distinct lead–zinc– (Ag) skarn/manto/vein metal tenor in more central portions of this batholithic suite (Fig. 2.16). 2.6.2 High-K Calk-Alkaline “I-type” Suites ‘Mid’-Cretaceous magmatism in the Selwyn Basin is mainly crustally derived and has a metallogenic signature not unlike that of the Cassiar plutonic suite. The I-type South Fork and Tay River suites in the western Selwyn Basin do not appear to be particularly well-mineralized. South Fork volcanic rocks and associated caldera structures are generally unmineralized and unaltered, although this may be the resulted of erosion of the uppermost volcanic rocks 84 (Gordey, in press). One gold vein showing of unknown affinity is documented in the nearby Connolly Caldera (Fig. 2.16). Tay River suite plutons in the western Selwyn Basin are associated with minor base metal, polymetallic, and (or) gold-rich veins, and there are indications of associated porphyry style mineralization (e.g., TAC: a Cu–Mo occurrence comprising altered and mineralized dioritic float; Fig. 2.16). The Tay River plutonic suite is associated primarily with small base metal (e.g., Pb, Zn, Ag, Cu, Mo, W) skarns and mantos across the Selwyn Basin, although tungsten–copper dominated skarns appear to be more commonly found in association with ‘evolved’ phases of this suite in the eastern Selwyn Basin (Fig. 2.16). Distal base (Pb–Zn–Ag–Cu) metal vein and manto occurrences associated with this suite also appear to be more abundant in the eastern Selwyn Basin (Fig. 2.16). In the far northwest Selwyn Basin, the Mayo plutonic suite is typically associated with contact aureole-hosted and intrusion-hosted (i.e., sheeted veins and fractures) gold mineralization (e.g., Scheelite Dome, Dublin Gulch, and Clear Creek; Fig. 2.2a). Bismuth–tellurium–arsenic–antimony enrichments are typical of this gold mineralization, and nearby but uneconomic tungsten skarns and tin occurrences are also common (Fig. 2.16; e.g., Hart et al., 2007). Conversely, most of the Mayo suite plutons from the north–central to northeast Selwyn Basin have a stronger association with copper or tungsten mineralization, although lead–zinc and metasedimentary-hosted gold occurrences are also present (Fig. 2.16). Tungsten suite intrusions are, not surprisingly, particularly noted for abundant proximal tungsten-copper (± Zn–Bi–Au–Sn–Mo) skarns, including the world-class tungsten skarn deposits at Cantung and Mactung (Fig. 2.16). The I-type petrogenesis interpreted for samples of this plutonic suite is compatible with a review of tungsten skarn mineralization in this area (Newberry, 1998), and the suggestion that I- type granitoids are associated with the largest tungsten skarn deposits (Newberry and Swanson, 1986). One emerald occurrence associated with the Lened pluton is also notable (Fig. 2.16), and intrusion-hosted sheeted gossanous fractures and veins with anomalous gold, arsenic, and bismuth are present at Cantung (Rasmussen et al., 2007). Distal base metal veins are also characteristic of the Mayo and Tungsten plutonic suites (Fig. 2.16). 2.6.3 Alkali-Calcic to Alkaline “A-type” Suites Like the high-K calc-alkaline (mainly S-type) Cassiar plutonic suite, the ‘mid’-Cretaceous Seagull plutonic suite in the Cassiar Terrane is associated with proximal tungsten ± molybdenum and argentiferous lead–zinc occurrences. The Seagull suite, however, is particularly notable for a large number of associated tin skarn and vein occurrences (e.g., JC; Fig. 2.16). This is distinctly different than the base metal-dominated tenor of the Cassiar plutonic suite (Liverton and Alderton, 1994). Elevated tin concentrations—and presumably associated mineralization—are common to A-type granitoids (e.g., Chappell et al., 2000). The Late Cretaceous McQuesten plutonic suite in the western Selwyn Basin provides another example of the A-type granite–tin association (Fig. 2.16; e.g., Emond, 1986; Emond and Lynch, 1990). Late Cretaceous mineralization and magmatism in the Cassiar Terrane is diverse, but appears to be characterized by mainly tungsten and (or) molybdenum occurrences. Rancheria suite plutons in southern Yukon are associated primarily with proximal tungsten skarns and variations on distal and proximal vein or manto styles of base metal mineralization (e.g., Pb, Zn, Ag, Cu, and Mo; Fig. 2.16). The intrusion-hosted uranium–REE porphyry occurrence in 85 the Allen Stock (Alle; Fig. 2.16) is the only known example of this type of mineralization in association with Late Cretaceous plutons in the region. Other examples of coeval plutonism and (or) mineralization south of the Rancheria suite plutons include: the Midway silver-rich lead–zinc manto deposit, the McDame tungsten skarn, and the Storie molybdenum porphyry (the latter two are associated with the Late Cretaceous Cassiar Stock; Figs. 2.2b and 2.16). This molybdenum ± tungsten metallogenic signature appears to characterize coeval intrusive rocks to the west and south, including several molybdenum ± tungsten porphyry deposits and occurrences from southern Yukon through to southern British Columbia. Important examples of these in the northern Cordillera include the Ruby Creek (Adanac) deposit associated with the 84 Ma Surprise Lake batholith, and the Swan showing associated with the 73 Ma Glundebery batholith (Fig. 2.2b; Christopher and Pinsent, 1982; Mihalynuk et al., 1992; Nelson and Colpron, 2007)16. The two Eocene Black River plutons in the Cassiar Terrane are associated with nearby silver-rich base metal (e.g., Zn, Cu, and Pb) vein and replacement-type mineral occurrences (Fig. 2.16). Eocene rhyolitic intrusions in the Anvil batholith area of the western Selwyn Basin are also associated with base and polymetallic veins. One example of Eocene intrusion-related mineralization is the Grew Creek gold-bearing epithermal vein deposit hosted in bimodal Tertiary volcanic rocks of the Ross Group (Fig. 2.16). Farther to the south near the British Columbia border, high grade zinc–lead–silver replacement-style mineralization at Sa Dena Hes (Fig. 2.16) is believed to be the product of retrograde thermal metamorphism related to a buried intrusion represented by quartz–albite porphyry dykes (ca. 50 Ma K–Ar; D. Sinclair, unpub.). These deposits are not unlike the argentiferous or auriferous lead–zinc vein and replacement-style mineralization associated with Eocene mantle-derived lamprophyric to gabbroic magmatism in southeast British Columbia (Beaudoin et al., 1992b,c; Sevigny and Thiérault, 2003; Adams et al., 2005). Finally, in northern British Columbia, the Mount Reed and Mount Haskin intrusions (ca. 50 Ma) are associated with molybdenum–tungsten porphyry occurrences (Fig. 2.2b). Like other members of the ‘mid’-Cretaceous TTB (i.e., the Tungsten and Mayo plutonic suites), Tombstone suite intrusions in the eastern Selwyn Basin are also associated with proximal argentiferous lead–zinc, copper, and gold (including placer) occurrences, as well as one intrusion-hosted gem (elbaite; Fig. 2.16) occurrence in miarolitic cavities of pegmatite–aplite horizons in the O’Grady batholith. Several mafic volcanic-hosted copper–(Au) occurrences near Tombstone suite plutons in the southeast Selwyn Basin are of uncertain origin although lead isotopic compositions of this mineralization are compatible with Cretaceous magmatic affinity (Rasmussen et al., 2007). In the northwest Selwyn Basin, the Tombstone suite is predominantly associated with disseminated or vein- type uranium–thorium–REE–fluorine mineralization and intrusion-hosted sheeted gold-rich veins or contact aureole- hosted gold ± copper–bismuth mantos (e.g., Brewery Creek; Fig. 2.2b). 16 Significant coeval mineralization and magmatism is also found farther south in west–central British Columbia at the Endako, Huckleberry, and Shan molybdenum porphyry deposits (e.g., Nelson and Colpron, 2007; Wodjak and Febbo, 2008). 86 2.7 TEMPORAL AND COMPOSITIONAL PATTERNS OF SYN- TO POST- ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERA Three main syn- to post-accretionary magmatic episodes occurred in the northern Cordillera: an Early to ‘mid’- Cretaceous event, a Late Cretaceous event, and a latest Cretaceous to Eocene (Tertiary) event (e.g., Nelson and Colpron, 2007). These syn- to post-accretionary magmatic events record the evolution of the northern Cordillera from a transpressional setting in the Early Cretaceous to the transtensional setting still present today (Engebretson et al., 1985; Gabrielse and Yorath, 1991; Kelley, 1993; Gabrielse et al., 2006). The differing tectonic regimes are also reflected in the compositions and metallogenic associations exhibited by the magmatic suites (e.g., Anderson, 1988), and changes to the tectonic regimes are partly constrained by the timing of the three magmatic flare-ups. Furthermore, the relationship between the timing and composition of igneous rocks in ‘outboard’ (mainly in accreted pericratonic terranes) and ‘inboard’ (mainly in the northern Cordilleran miogeocline) positions is significant with respect to the specific tectonic processes that may have operated during these three intervals. The following summary of the timing and distribution of syn- to post-accretionary magmatism, with an emphasis on the northern Cordilleran miogeocline, relies heavily on the results of this study as well as compiled U–Pb ages. 2.7.1 The Early to ‘mid’-Cretaceous (~115-100 Ma) 2.7.1.1 ‘Outboard’ Magmatism Although this study did not examine magmatic rocks in the more ‘outboard’ or westerly regions of the northern Cordillera, a brief synopsis is provided since its timing and composition is significant with respect to ‘inboard’ or easterly magmatism in the northern Cordilleran miogeocline. Early Cretaceous magmatism in western portions of the northern Cordillera was emplaced as a syn- to post-accretionary event that resulted from dextrally oblique convergence at the northwest continental margin (e.g., Gabrielse et al., 2006; Nelson and Colpron, 2007). ‘Outboard’ expressions of the Early Cretaceous event comprise mainly oxidized > reduced, metaluminous > peraluminous, intermediate to felsic (with abundant mafic phases), medium-K calc-alkaline magmatic suites of I- type affinity. Plutonic expressions of these continental arc-related igneous rocks are subdivided into several suites (Fig. 2.2a), including: the 123–115 Ma Teslin plutonic suite (e.g., Gordey et al., 1998; Mortensen et al., 2000) and the 115–104 Ma Whitehorse plutonic suite in southwest Yukon (Hart, 1995, 1997; Mortensen et al., 2000; Hart et al., 2004a); and, the 112–99 Ma Dawson Range plutonic suite in east–central Alaska and west–central Yukon (Mortensen and Hart, 2010, unpub. data). The broad region of into which these plutonic suites were emplaced would have lain directly west–southwest to the Selwyn Basin and Cassiar Terrane in the Early Cretaceous (Fig. 2.17a). This belt of medium-K calc-alkaline plutonic suites (and and associated volcanic rocks) is interpreted to represent a continental magmatic arc that resulted from east–northeast dipping subduction of oceanic crust below the western margin of North America (e.g., Mortensen et al., 2000; Hart et al., 2004a; Crawford et al., 2005; Nelson and Colpron, 2007; Mortensen and Hart, 2010). This Early Cretaceous magmatic arc suggests that coeval ‘inboard’ magmatism in the northern Cordilleran 87 miogeocline was emplaced in a back-arc setting. 2.7.1.2 ‘Inboard’ Magmatism Beginning ca. 115 Ma, a voluminous Early Cretaceous, mainly S-type, high-K calc-alkaline magmatic event was emplaced in the northern Cordilleran miogeocline. Early Cretaceous, (mainly) plutonic rocks have been sub-divided into several plutonic suites (Fig. 2.2a), including: the 115–104 Ma Cassiar suite in the Cassiar Terrane (Hart et al., 2004a); the 114–104 Ma Anvil suite in the western Selwyn Basin and the Yukon–Tanana inlier (Mortensen et al., 2000; Pigage, 2004); the 106–100 Ma Hyland suite in the eastern Selwyn Basin (Heffernan, 2004); and the 109–102 Ma (± ~94 Ma) Yukon–Tanana Uplands suite in east–central Alaska (Hart et al., 2004a). These intermediate to felsic intrusive rocks are characterized by predominantly reduced to slightly oxidized magmas with peraluminous compositions, although subordinate oxidized and slightly metaluminous phases are present in the Cassiar and Hyland plutonic suites. When syn- to post-magmatic strike-slip deformation is restored along the dextral structures, these plutonic suites form a belt of S-type magmatism several 1000’s of km long (Fig. 2.17a). This belt is coeval to continental arc magmatism farther west, but was derived primarily from the partial melting of miogeoclinal rocks and their basement. For these reasons, the ‘inboard’ granitoids have traditionally been interpreted to represent back- arc magmatism generated by decompression during the orogenic collapse of previously orogen-thickened continental crust (Anderson, 1988; Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004, 2005; Mair et al., 2006b; Nelson and Colpron, 2007). Around ~103 Ma in the Cassiar Terrane and ~99 Ma in the Selwyn Basin, a significant change in the locus and composition of magmatism across the northern Cordilleran miogeocline occurred. In the Cassiar Terrane, immediately following the emplacement of the youngest known Cassiar suite intrusion, the locus of plutonism shifted slightly to the west where the ‘mid’-Cretaceous (103–96 Ma) aluminous A-type Seagull suite intrusions were emplaced as a narrow belt (Fig. 2.17a). A lack of coeval continental arc magmatism to the west indicates that the Seagull plutonic suite was emplaced in a post-collisional or post-orogenic setting, which is consistent with its A-type affinity. Similarly, a significant change in composition and locus of magmatism also occurred immediately following the emplacement of the S-type Anvil and Hyland plutonic suites in the Selwyn Basin. Here, ‘mid’- Cretaceous magmatic rocks shifted east and north, and became infracrustal in origin. Associated ~97 Ma high-K calc-alkaline mafic dykes in the eastern Selwyn Basin may attest to the influx of mantle-derived magma into the crust at this time. This suggests that advective heat was transferred from the mantle to the crust, which would have been important for at least hornblende and possibly biotite dehydration reactions in the crust. Although they are of very different composition and distribution, the high-K calc-alkaline I-type Tay River plutonic suite (99-96 Ma) and associated volcanic rocks in the South Fork (ca. 97–96 Ma) calderas are roughly coeval to the A-type Seagull plutonic suite in the Cassiar Terrane (Fig. 2.17a). Stepping farther inboard in the Selwyn Basin, younger granitoids with I-type affinities were emplaced through the ‘mid’-Cretaceous until ~90 Ma. As with the Tay River plutonic suite, the involvement of lithospheric mantle first appears to have begun at circa ~97 Ma (Table 2.2) with the emplacement of late-stage shoshonitic dykes in 88 association with the 98–95 Ma Tungsten and 98–93 Ma Mayo plutonic suites. These mafic dykes provided or contributed to the heat necessary for partial melting of intermediate-composition infracrustal rocks, there may be geochemical evidence that the shoshonitic magmas interacted with a small proportion of the Mayo suite intrusions. ‘Mid’-Cretaceous magmatism in the Selwyn Basin terminated with the emplacement of the 94–89 Ma Tombstone suite. This youngest plutonic suite is of mixed I- and A-type character, which reflects continued partial melting of metaigneous crust and the influx of magmas sourced from an enriched, sub-continental lithosphere. Interactions between these two disparate magma types appears to have been widespread and are responsible for the variety of compositions exhibited by the Tombstone plutonic suite. As with the S-type suites, the ‘mid’-Cretaceous plutonic suites have been interpreted by previous workers to represent back-arc magmatism (Anderson, 1988; Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004, 2005; Mair et al., 2006a; Nelson and Colpron, 2007). However, these ‘mid’-Cretaceous intrusions post-date essentially all of the more outboard continental arc rocks (Fig. 2.1b; Mortensen and Hart, 2010, unpub. data). These plutonic suites, therefore, must have been emplaced in a post-arc geodynamic setting. The 103–99 Ma change in the locus and composition of post-arc magmatism indicates that whatever the change to the geodynamic regime was at this time, it was expressed differently along the northwest margin of Laurentia. The change in patterns of magmatism may be related to different geometry of the Neoproterozoic rifted margins inferred for these two regions (i.e., the lower plate rifted margin in the Selwyn Basin, and the upper plate rifted margin in the Cassiar Terrane; Cecile et al., 1997). The basement to the Cassiar Terrane could potentially be overlain by a thinner package of supracrustal rocks and may be ‘underlain’ by underplated rift-related ‘basalts’ (e.g., Lister et al., 1991; Cecile et al., 1997; Evenchick et al., 2005) if inferences regarding the geometries of the rifted margin are correct. The basement to the Selwyn Basin should comprise thicker packages of siliciclastic rift-related and miogeoclinal rocks overlying thinner, extended crystalline basement (Lister et al., 1991; Cook et al., 1991; Evenchick et al., 2005). The dominance of S-type intrusive rocks in the Cassiar Terrane, and the extensive I-type intrusive (and local volcanic) rocks across much of the Selwyn Basin do not really support the above interpretation. There may be, however, something intrinsic to the composition or strength of the basement that differs in these two regions dictated or partly controlled the patterns of syn- to post-accretionary magmatism. 2.7.2 The Late Cretaceous (~84–67 Ma) 2.7.2.1 ‘Outboard’ Magmatism Magmatism in the Late Cretaceous was emplaced in the northern Cordillera during the transition from what was primarily a dextral transpressional regime, to what became primarily a dextral transtensional regime (e.g., Nelson and Colpron, 2007). The changing geodynamic setting, controlled by outboard oceanic plate motions to the west, led to a change to patterns of magmatism through this interval (Engebretson et al., 1985; Anderson, 1988; Woodsworth et al., 1991; Gabrielse and Yorath, 1991; Kelley, 1993; Nelson and Colpron, 2007). A calc-alkaline continental magmatic arc is thought to have persisted from east–central Alaska to northwest British Columbia (Fig. 2.2b) as the 78–67 Ma Carmacks Group (and the up to 86 Ma Windy–Table suite) of volcano-plutonic rocks (Fig. 2.2b; 89 Christopher and Pinsent, 1982; Anderson, 1988; Souther, 1991; Woodsworth et al., 1991; Mihalynuk et al., 1992; Mihalynuk, 1999; Selby et al., 1999; Simmons et al., 2006, 2007; Nelson and Colpron, 2007; Mortensen and Hart, 2010, unpub. data). Plutonic expressions of this magmatic arc might extend as far east as the westernmost Cassiar Terrane where a ~74 Ma biotite-bearing leucogranite was sampled from the Deadman Creek batholith (Gordey et al., 1998). Effusive volcanic rocks characterizing the younger Carmacks Group (i.e., most of the dark green shaded region in Figs. 2.2b and 2.17b) is speculated to have resulted from lithospheric delamination after the cessation of northeast dipping subduction of an oceanic plate west of North America at this time (Gabrielse and Yorath, 1991; Mortensen and Hart, 2010). 2.7.2.2 ‘Inboard’ Magmatism After Early Cretaceous continental arc magmatism ended west–southwest of the Cassiar Terrane, several plutonic suites of peralkaline to peraluminous, rare earth and trace element enriched, felsic pluton of A-type affinity were emplaced as narrow belts near terrane boundaries and (or) major structures (Fig. 2.2b). One of these suites was the ‘mid’-Cretaceous Seagull plutonic suite. After a magmatic lull in the Cassiar Terrane following the emplacement of the Seagull plutonic suite, three plutons defining the 82–75 Ma Rancheria plutonic suite were emplaced in the eastern Cassiar Terrane, near the north end of the Kechika fault (Figs. 2.1a and 2.2b). This belt appears to continue into northernmost British Columbia based on K–Ar and Rb–Sr ages from plutonic, metamorphic (hornfels), and hydrothermal mica and whole rock samples east of the Cassiar batholith (Table 2.1; Fig. 2.2b). Intrusions belonging to this Late Cretaceous, plutonic–metamorphic–hydrothermal belt have A-type compositions that are compatible with a crustal or sub-crustal lithospheric melt source. Although the timing of the continental magmatic arc farther to the west (the Carmacks Group) overlaps with the Rancheria plutonic suite, the earlier onset of this A-type plutonism suggests that it pre-dated arc magmatism and was not necessarily related to concurrent subduction. Instead, ‘inboard’ plutonism in the Late Cretaceous may have been controlled by transpressional to transtensional deformation along crustal discontinuities. 2.7.3 The Latest Cretaceous to Eocene (~67–47 Ma) 2.7.3.1 ‘Outboard’ Magmatism A broad belt of Paleocene to Eocene igneous rocks was also emplaced west of the northern Cordilleran miogeocline (Fig. 2.2b). This belt is generally known as the Coast Plutonic Complex, although it has been locally sub-divided into plutonic and volcanic suites (Fig. 2.2b; Anderson, 1988; Souther, 1991; Woodsworth et al., 1991; Mihalynuk, 1999; Morris and Creaser, 2003). These sub-divisions in the northern Cordillera include: the 64–51 Ma Ruby Range plutonic suite (S. Israel, 2012, pers. comm.); the 52–62 Ma (K–Ar and Rb–Sr ages) Nisling Range plutonic suite; and the ca. 55 Ma volcanic rocks of the Sloko and Skukum suites in northwest British Columbia and southwest Yukon, respectively. The latter are preserved as erosional remnants of more voluminous and widespread volcanic rocks and minor associated sub-volcanic intrusions (Mihalynuk, 1999). 90 All of these suites are of calc-alkaline composition, with the possible exception of the Nisling Range plutonic suite (Anderson, 1988). The bulk of this magmatism has, therefore, been interpreted to represent a continental magmatic arc related to the subduction of an oceanic plate during highly oblique (northward) convergence and (or) above a shallowly dipping slab (Souther, 1991; Mihalynuk, 1999; Morris and Creaser, 2003; Crawford et al., 2005; Simmons et al., 2006, 2007; Nelson and Colpron, 2007). The Nisling Range plutonic suite comprises mainly high-level alaskitic plutons, associated dyke swarms, and felsic pyroclastic rocks and has features that are indicative of an A- type affinity (e.g., such as miarolitic textures, perthite K-feldspar with interstitial albite, and accessory fluorite; Anderson, 1988; Souther, 1991). It may be that Nisling Range plutonic suite was not derived directly from the same processes that generated the coeval continental magmatic arc. 2.7.3.2 ‘Inboard’ Magmatism Felsic plutonism appears to have terminated in the Cassiar Terrane with the emplacement of two Eocene (47–49 Ma), A-type plutons of the Black River plutonic suite immediately southwest of the Tintina Fault (Fig. 2.2b). Several scattered occurrences of compositionally similar magmatic (plutonic ± volcanic) suites or occurrences of latest Cretaceous- to Eocene-age have also been compared here to the Black River samples. Although none of these other ‘suites’ or occurrences are identical in either age or composition to the Black River plutonic suite, collectively these represent scattered, small-volume, and (typically) fault-associated A-type magmatism across much of the northern Cordilleran miogeocline (Fig. 2.2b and 2.17b). The majority of latest Late Cretaceous to Eocene igneous rocks also have A-type or within plate geochemical affinities, and have been associated with dextral transtension and (or) extension of thickened continental crust along regional structures (e.g., Souther, 1991; Yorath, 1991; Nelson and Colpron, 2007). The emplacement of the Black River plutonic suite coincided with the apparent end to highly oblique subduction of the Kula plate (following the subduction of the Kula–Farallon spreading centre) and the formation of a transcurrent plate boundary along the western margin of North America (e.g., Engebretson et al., 1985; Gabrielse and Yorath, 1991; Miller et al., 2002). This change in to a north-directed compression and a transform fault plate boundary along the western margin of the northern Cordillera appears to have brought about an end to plutonism across the northern Cordilleran miogeocline. 91 Figure 2.1. (next page) Basic elements of northern Cordilleran geology, and the distribution of Early Cretaceous to Eocene syn- to post-accretionary igneous rocks in allochthonous and autochthonous Laurentian rocks. The study area is outlined in black. a) Tectonostratigraphic elements of the northern Cordillera. The upper left inset shows the location of the map area relative to Canada. Basinal and platformal rocks of the northern Cordilleran miogeocline are shaded pale yellow and pale blue, respectively, and shades of grey represent terranes accreted in the latest Paleozoic to the Mesozoic. Bolded tectonostratigraphic abbreviations are: SB = Selwyn Basin, YTT = Yukon–Tanana Terrane (not differentiated from adjacent Stikinia and Quesnellia terranes), CC = Cache Creek terrane, CP = Cassiar Platform, CT = Cassiar Terrane, MeP = McEvoy Platform, MkP = Mackenzie Platform, MdP = McDonald Platform, and KT = Kechika Trough. Geographic regions of the Selwyn Basin referred to in the text are indicated with transparent NW, NE, W, and SE labels (YTU = Yukon–Tanana uplands). Geographic abbreviations are: YK = Yukon, AK = Alaska, NWT = Northwest Territories, and BC = British Columbia. Tectonostratigraphic elements are modified from Nelson and Colpron (2007), with contributions from Aitken and Pugh (1984), Cecile et al. (1997), Gordey and Anderson (1993), Morrow and Miles (2000), Gordey and Makepeace (1999), Gabrielse et al. (2006), and Turner and Long (2008). b) Sample locations for robust (i.e., U–Pb, 40Ar/39Ar, or Re–Os) ages. Bright green circles represent U–Pb (plutonic and felsic dyke) and 40Ar/39Ar (mafic dyke) ages determined by this study. Yellow circles are U–Pb results in the Dawson Range to Yukon–Tanana Uplands regions (Mortensen and Hart, 2010, unpub. data). Red circles are published and unpublished results from other studies1; areas of larger datasets have been outlined with blue dashed lines and labeled (data in the “Pogo” area are from Werdon et al., 2004, and Dilworth et al., 2007; and data in the “Mount Nansen” area are from Carlson, 1987, Selby et al., 2002, and Mortensen et al., 2003). Plutonic and volcanic polygons are primarily from Gordey and Makepeace (1999) and Massey et al. (2005), and modified based on the maps of Dusel–Bacon et al. (1993), Newberry et al. (1998), Mortensen et al. (2000), Baker and Lang (2001), Murphy et al. (2001), Hart et al. (2004a), Pigage (2004, 2009), Pigage et al. (2011), and Gordey (in press). Map lines and colours are as in (a), however, oceanic and pericratonic terranes are uncoloured. 1 Compiled geochronological data are from: the Canadian Geochronological Knowledgebase (http://gdr.ess.nrcan.gc.ca/geochron/e/explorer.htm), the Yukon Igneous Database (http://servlet.gov.yk.ca/ygsis/index.do), Richter et al. (1975), Carlson (1987), Mortensen and Hansen (1992), Hart (1995, 1997, 2004), McCoy et al. (1997), Murphy (1997), Duncan (1999), De Keijzer et al. (2000), Baker and Lang (2001), Lang (2001), Liverton et al. (2001), Colpron (et al., 2002, 2006), Coulsen et al. (2002), Selby et al. (2002, 2003), Mortensen et al. (2003, 2006b, 2007), Heffernan (2004), Pigage (2004), Werdon et al (2004), Dilworth et al. (2007), Gregory (2007), Gallagher (1999), Colpron and Friedman (2008), Mortensen, (2008, 2009), Bineli Betsi and Lentz (2011), Colpron (2011, and unpub. data), Pigage et al. (2011), Gordey (in press). 92 93 Figure 2.2. (next 2 pages) Distribution of syn- to post-accretionary, Early Cretaceous to Eocene igneous rocks in the northern Cordillera based on the geochronological results of this study, compiled robust age determinations, and the work of Anderson (1988), Murphy (1997), Mortensen et al. (2000), Hart et al. (2004a) and Mortensen and Hart (2010, unpub.). The study area is outlined in black. Intrusive and volcanic polygons are primarily from Gordey and Makepeace (1999) and Massey et al. (2005), and modified based on the regional maps of Dusel–Bacon et al. (1993), Newberry et al. (1998), Mortensen et al. (2000), Baker and Lang (2001), Murphy et al. (2001), Hart et al. (2004a), Pigage (2004), Pigage et al. (2011), and Gordey (in press). There are some unresolved ages for magmatism in east–central Alaska that have led to the same bodies being labeled Early Cretaceous, Late Cretaceous, and (or) Paleocene. Map lines and tectonostratigraphic units are as in Figure 3.1. a) Early to ‘mid’-Cretaceous magmatic suites. The red belts represent continental arc magmatism; the purple belts represent back-arc magmatism; and the blue, orange, and yellow belts represent post-arc magmatism. The plutonic suites are labeled: A = Anvil, C = Cassiar, DR = Dawson Range, FS = Fairbanks–Salcha, H = Hyland, M = Mayo, S = Seagull, Te = Teslin, To = Tombstone, TR = Tay River, Tu = Tungsten, W = Whitehorse, YTT = Yukon–Tanana Uplands, TT = Tok–Tetlin. Plutons associated with significant mineralization in the northwest Selwyn Basin (outside the study area) and mentioned in the text are labeled. b) Late Cretaceous to Eocene magmatic suites and (or) occurrences. The green belts represent Late Cretaceous magmatism, and the pink belts and stars represent Paleocene–Eocene magmatism. The plutono–volcanic suites are labeled: Ca = Carmacks Group (volcanic > plutonic), RR = Ruby Range (plutonic), NR = Nisling Range (plutonic), CPC = Coast Plutonic Complex (plutonic > volcanic), Sk = Skukum (volcanic), Sl = Sloko (volcanic), WT = Windy Table (volcanic), SL = Surprise Lake (plutonic), and MQ = McQueston (plutonic). Plutono–volcano occurrences mentioned in the text are labeled, and the Midway lead–zinc deposit is indicated by the white star in northern British Columbia. 94 95 Figure 2.3. (next 2 pages) The distribution of syn- to post-accretionary plutons and associated volcanic rocks in the study area. The base map is modified from Nelson and Colpron (2007) and Gordey and Makepeace (1999), with contributions from Mortensen et al. (2000), Baker and Lang (2001), Murphy et al. (2001), Hart et al. (2004a), Pigage (2004), Pigage et al. (2011), and Gordey (in press). a) Locations of geochronology samples collected for this study (yellow circles are felsic pluton or dyke samples, and green circles are mafic dyke samples). Plutonic rocks are indicated by grey polygons and volcanic rocks are indicated by black polygons. Plutons mentioned in the text are labeled. The areas where the Neoproterozoic Windermere Supergroup is exposed have a cross-hatched pattern. Broad within-terrane tectonostratigraphic sub-divisions are indicated, and major Cretaceous to Cenozoic dextral strike-slip faults (Ft) and fault systems (FS) are labeled. The inset at the bottom right is for the Cantung area (outlined by the black box in the eastern Selwyn Basin). b) Distribution of plutonic suites in the study area defined by robust magmatic ages (see Fig. 2.3a caption for data sources) and the geochemical classifications in the text. U–Pb ages for felsic plutonic (black text; results from this study are bolded) and felsic dyke (red text; results from this study are bolded) samples, and 40Ar/39Ar ages (green text) for mafic dyke samples (results from this study are bolded) or cooling ages in plutonic rocks (other studies). U–Pb (black) and 40Ar/39Ar (green) ages determined by other studies are given and located by an “x”; and in areas with less geochronological control selected less reliable K–Ar and Rb–Sr from other studies are given as an italicized “x”. Plutonic suite sub-divisions are modified from Gordey and Anderson (1993), Mortensen et al. (2000), Hart et al. (2004a, 2005) and Heffernan (2004) are given for magmatism northeast of the Tintina Fault and magmatism southwest of the Tintina Fault. Faults and boundaries between tectonostratigraphic subdivisions are as in (a), and grey polygons are undivided intrusive rocks. 96 97 98 Figure 2.4. (next 2 pages) Geochemical variation diagrams for the Cassiar plutonic suite, with fields and (or) fractionation trends for I-, S-, and A-type magmas indicated in red on some plots. The Cassiar suite has been sub-divided on the basis of mineralogy: ms = muscovite ± biotite, bt = biotite, and hbl = hornblende ± biotite. The Teslin plutonic suite is plotted as red filled circled or lines for comparative purposes. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Cassiar plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; black lines are hornblende-bearing samples, dark grey lines are biotite-bearing samples, and light grey lines are muscovite-bearing samples. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for muscovite (light grey), biotite-bearing (medium grey), and hornblende (black) samples. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams. 99 100 101 Figure 2.5. (next 2 pages) Geochemical variation diagrams for the Anvil and Hyland plutonic suites. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Anvil and Hyland plutonic suites. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; the Anvil suite is in light grey Anvil, the Hyland suite is in medium grey (including boxes), the green lines and (or) small boxes are Hyland suite mafic dykes, and the blue lines and small boxes are the Hyland suite enclave. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for Anvil (light grey), Hyland (medium grey) samples; the green dotted lines and circles are the average Hyland suite mafic dykes; and, the blue line and boxes are the Hyland suite mafic enclave. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams. 102 103 104 Figure 2.6. (next 2 pages) Geochemical variation diagrams for the Seagull plutonic suite; shoshonitic lamprophyre dykes that may be cogenetic are plotted as green symbols and lines. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Seagull plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; the cumulate phase is indicated by the red line. d) Average sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); the Seagull suite average is dark grey, and the cumulate phase from the Quiet Lake batholith is the red line. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams; see Table 2.4a for details of discrimination plots; the single bolded square symbol in ‘n’ is a lepidolite–(zinnwaldite)–albite–topaz leucogranite sample. 105 106 107 Figure 2.7. (next 2 pages) Geochemical variation diagrams for the Tay River plutonic and South Fork volcanic suites in the western Selwyn Basin. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Tay River and South Fork suites. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; medium grey lines are the South Fork suite, and black lines are the Tay River suite in the western Selwyn Basin. d) Average sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for samples from the South Fork (medium grey) and Tay River (black) samples in the western Selwyn Basin. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams; see Table 2.4a for details of discrimination plots. 108 109 110 Figure 2.8. (next 2 pages) Geochemical variation diagrams for the Tay River plutonic suite and the Coal River sub-suite in the eastern Selwyn Basin; high-K calc-alkaline mafic dykes that may be cogenetic are plotted as green symbols and lines. See Table 2.4 for details of the geochemical plots, and a Table 2.5 for summary of geochemical characteristics for the Tay River plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. Circled data points are potentially contaminated samples (as discussed in text). b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; dark grey lines are the Tay River samples, light grey lines are ‘evolved’ Tay River samples, and black lines are the Coal River samples. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for samples from the Tay River (medium grey), ‘evolved’ Tay River (light grey), and Coal River (black) samples in the eastern Selwyn Basin. Average compositions for the Tay River plutonic suite (heavy red line) and South Fork volcanic suite (thin red line) from the western Selwyn Basin are also illustrated for comparative purposes. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams; see Table 2.4a for details of discrimination plots. 111 112 113 Figure 2.9. (next 2 pages) Geochemical variation diagrams for the Tungsten plutonic suite; shoshonitic mafic dykes that may be cogenetic are plotted as green symbols and lines. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Tungsten plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. Circled data points are potentially contaminated samples (as discussed in text). b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; light grey lines are samples from the northeastern Selwyn Basin, and medium grey lines are samples from the southeastern Selwyn Basin. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for samples from the northeast (light grey) and the southeast (medium grey) Selwyn Basin. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams; see Section Table 2.4a for details of discrimination plots. 114 115 116 Figure 2.10. (next 2 pages) Geochemical variation diagrams for the Mayo plutonic suite; shoshonitic mafic dykes that may be cogenetic are plotted as green symbols and lines. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Mayo plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. Circled data points are potentially contaminated samples (as discussed in text). b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; samples from the northwest Selwyn Basin are in black, and samples from the northeast Selwyn Basin are in light grey. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for samples from the northwest Selwyn Basin (black) and northeast Selwyn Basin (light grey). e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams. 117 118 119 Figure 2.11. (next 2 pages) Geochemical variation diagrams for the Tombstone plutonic suite; shoshonitic mafic plutonic and dyke phases that may be cogenetic are plotted as green symbols and lines, and tinguaite phases are plotted as blue symbols and lines. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Tombstone plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. Circled data points are potentially contaminated samples (as discussed in text). b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; samples from the northwest Selwyn Basin are in light grey, samples from the northeast Selwyn Basin are in medium grey, samples from the southeast Selwyn Basin are in dark grey. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for samples from the northwest Selwyn Basin (black), northeast Selwyn Basin (light grey), and southeast Selwyn Basin (dark grey). e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams. 120 121 122 Figure 2.12. (next 2 pages) Geochemical variation diagrams for the Rancheria plutonic suite; a high-K calc-alkaline mafic dyke that may be cogenetic is plotted as green symbols and lines. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Tombstone plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams. d) Average sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998). e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams; see Table 2.4a for details of discrimination plots. 123 124 125 Figure 2.13. (next 2 pages) Geochemical variation diagrams for selected Paleocene to Eocene, plutonic and volcanic magmatism, including the Black River suite. See Table 2.4 for details of the geochemical plots, and Table 2.5 for a summary of geochemical characteristics for the Black River plutonic suite. a) Harker variation diagrams; major element oxide data are in wt.%, and minor and trace elements are in ppm. b) Bimodal plots of fractionation indices. c) Normalized multi-element diagrams; medium grey fine lines are the ca. 65 Ma McQuesten plutonic suite in the northwestern Selwyn Basin; pale grey lines are the ca. 55 Ma bimodal Ross Group volcanic rocks in the western Selwyn Basin (the heavy dashed line is the porphyry phase at Tsa Da Glisza; basaltic rocks are not shown), the heavy dashed black line is the ca. 52 Ma Ting suite in southeastern Selwyn Basin; and, black lines are the ca. 48 Ma Black River (heavy black line) and ca. 50 Ma Mount Reed (fine black line and white boxes) plutonic suites in the Cassiar Terrane and northeastern British Columbia, respectively. d) Sliding normalization to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series of Liégeois et al. (1998); lines are averages for samples as described in (c) except the porphyry phase at Tsa Da Glisza is not included in the averaged Ross Group rhyolites. e) CIPW-normative IUGS quartz–K-feldspar–plagioclase ternary for silica-saturated granitoid rocks. f-n) Geochemical and tectonic discrimination diagrams; see Table 2.4a for details of discrimination plots. 126 127 128 Figure 2.14. Comparison of the effects of differentiation on SNY (average Yenchichi–Telabit-normalized Y–Ce–Sm–Yb) and SNX (average Yenchichi–Telabit-normalized Rb–Th–U–Nb) values for different plutonic suites. a) The high-K calc-alkaline Cassiar plutonic suite (symbols are as in Fig. 2.4). b) The alkali-calcic–alkaline Seagull plutonic suite; the white bolded box is the only lepidolite- (zinnwaldite) topaz leucogranite plotted. . 129 Figure 2.15. Discrimination plots for niobium-richer continental (crustal or sub-crustal) and niobium- poorer asthenospheric (OIB-like) sources for the A-type plutonic suites; the two groups are divided by the black line (after Eby, 1992). a) Seagull plutonic suite; a shoshonitic mafic dyke is plotted as the green cross. b) Tombstone plutonic suite; the majority of plutonic samples from the northwest and northeast Selwyn Basin plot within the dashed circle and only the mafic dykes that plot within this circle are shown; all samples that fall outside of this circle are individually plotted. c) Rancheria plutonic suite; a shoshonitic mafic dyke is plotted as the green diamond. d) Selected Paleocene to Eocene, plutono-volcanic magmatism. 130 Figure 2.16. (next page) Distribution of intrusion-related and probable intrusion-related mineral occurrences (showings, prospects, drilled prospects, and producers) in the study area (after Deklerk, 2003). Symbols for rare or unusual commodities are white and are labeled with their main elemental association(s). Major mineral deposits in the study area are indicated by large bolded stars, and mineral occurrences mentioned in the text and (or) listed in Table 2.1 are labeled with capitalized letters. Plutonic suites are shaded according to the legend in Figure 2.3b. 131 132 Figure 2.17. (next page) Basement to the western margin of North America after restoration of dextral strike-slip displacements along major structures, with the approximate area of magmatic suites and occurrences located (see legends in Fig. 2.2b,c for plutonic suite sub-divisions). Magma types for the plutono-volcanic rocks are labeled: I = I-type, S = S-type, and A = A-type. Known major faults are indicated by solid (strike-slip or transform) and dashed (thrust or reverse) lines; inferred basement faults are indicated by dotted lines. The Tintina Fault, along which 430 km of strike-slip movement has been restored, is shown with a bolded solid black line. Tectonic elements outlined by fine black lines are as in Fig. 2.1a. The distribution of restored terranes and magmatic suites is modified from Nelson and Colpron (2007) and Hart et al. (2004a). The ‘known’ western sub-surface extent of the Nahanni Terrane is indicated by the red dashed line (although it is geophysically inferred to continue west to the edge of the Laurentian miogeocline; e.g., Aspler et al., 2003), and the distribution of Archean to Late Paleoproterozoic basement domains is modified from Ross (1991) and Aspler et al. (2003). a) Early to ‘mid’-Cretaceous plutonic suites (123–90 Ma). The possible boundary between mainly infracrustal (I) and supracrustal (S) basement, based on the predominant composition of the plutonic suites, is indicated by the very heavy dashed black line. This line could represent a change in the type of basement present below the Selwyn Basin. b) Late Cretaceous to Eocene magmatic suites and occurrences (84–47 Ma). 133 134 Table 2.1. Summary of basic characteristics of for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. All ages are based on U–Pb data (this study, and compiled) unless otherwise indicated. Plutonic Suite or Intrusion Name Age (Ma)a Location a Basic Descriptionb Metallogenic Association Relevant References CASSIAR TERRANE AND WESTERN SELWYN BASIN Early to 'mid'-Cretaceous: Yukon– Tanana Uplands (offset from Anvil and Tay River suites by TF) 109– 102, ~94 Ma east–central Alaska reduced, peraluminous, laccolithic granodiorite to leucogranite + younger reduced, metaluminous- peraluminous, diorite to quartz diorite quartz vein-hosted Au (possibly the Pogo deposit, Alaska); small W or Mo porphyries and skarns Hart et al. (2004a); Dilworth (2003); Dilworth et al. (2007); Mortensen and Hart (2010, unpub. data). Cassiar 115–104 widespread in CT reduced–oxidized, peraluminous– metaluminous, batholithic intrusions of bt ± ms monzogranite, bt ± hbl granodiorite, and hbl ± bt quartz monzodiorite abundant W–(Mo) skarns/porphyries and veins (Northern Dancer, Risby, Stormy); distal Ag– Pb–Zn veins and/or manto (Ketza River) Noble et al. (1984); Layne and Spooner(1991); Liverton and Alderton (1994); Fonseca (1998); Driver et al. (2000); Mortensen et al. (2000, 2006b, 2007); Lang (2001); Hart et al. (2004a); Colpron et al. (2006); Nockleburg et al. (2005); Brand (2008); this study. Seagull 103–94 western CT reduced, peraluminous, monzogranite to Li-mica toz leucogranite abundant Sn–skarns/vein (JC) Liverton (1990, 1992); Layne and Spooner, (1991); Liverton and Botelho (2001); Liverton et al. (2005); Liverton (2011, pers. comm.); this study. Anvil 114–111 (YTT) and 109–104 (SB) western SB and YTT inlier reduced, peraluminous, laccolithic bt-ms leucogranite W–(Mo) or Sn–(Ag) skarns and veins; Ag–(Pb–Zn ± Cu) skarn; quartz vein– hosted Au; distal Ag–Pb– Zn veins; emerald (Tsa Da Glisza) Pigage and Anderson (1985); Mortensen and Ballantyne (1992); Mortensen et al. (2000); Murphy et al. (2001, 2002); Hart et al. (2004a); Neufeld (2004); Pigage (2004); Gordey (in press). Tay River 99–96 western SB reduced, peraluminous, hbl- bt granodiorite to monzogranite (with lesser quartz monzodiorite); coeval to 96.8–97.7 Ma dacitic calderas (South Fork) typically non–metallogenic; Pb–Zn–Ag skarns, mantos, and veins; possibly Cu– porphyry; W skarns Pigage and Anderson (1985); Gordey (1988, in press); Mortensen et al. (2000); Hart et al. (2004a); Pigage (2004); this study Late Cretaceous: Rancheria (includes: Allen Stock, ca. 69–82 Ma hydrothermal/metamorphic belt, Cassiar Stock, portions of Cassiar batholith) 82–75 eastern CT oxidized, peraluminous, leucocratic bt-bearing monzogranite and associated, quartz–feldspar porphyry, aplite, and kersantite dykes U–REE–porphyry (Alle); Pb–Zn–Ag manto (Midway); Mo ± W porphyries and skarns (Storie, McDame) Christopher et al. (1972); Cooke and Godwin (1984); Gregory (2007); Nelson and Colpron (2007); Wojdak and Febbo (2008); Canadian Geochronological Knowledgebase (2013); this study. Deadman batholith and nearby magmatism 73–74 westernmost CT (Teslin Tectonic Zone) bt granite to granodiorite phases of the composite Early to Late Cretaceous Deadman batholith and associated nearby magmatism nd Gordey et al. (1998); Gareau and Mortensen (unpub. data); Nelson and Colpron (2007); Mortensen and Hart (2010, unpub. data); Canadian Geochronological Knowledgebase (2013). Latest Cretaceous to Eocene: McQuesten 65–64 northwestern SB reduced, peraluminous, bt ± ms monzogranite Sn–Ag–(Au–Bi) veins/breccias (with "two mica" granites) Emond (1986, 1992); Emond and Lynch (1990); Murphy (1997). 135 Plutonic Suite or Intrusion Name Age (Ma)a Location a Basic Descriptionb Metallogenic Association Relevant References Ross Group (and regional equivalents) 56–52 western SB and YTT inlier reduced, peraluminous, sub-volcanic rhyolitic quartz–feldspar porphyry plugs and quartz-phyric dykes; associated with basaltic flows and clastic graben-fill base and polymetallic vein/manto (e.g., Grew Creek; Sa Dena Hes) Tempelman-Kluit (1972); Jackson et al. (1986); Pride (1988); Gabrielse (1991); Murphy et al. (2001, 2002); Dilworth (2003); Neufeld (2004); Pigage (2004); Gordey (in press); Canadian Geochronological Knowledgebase (2013). Black River (F1) pluton and Tintina plug 49–47 eastern CT; adjacent to TF bt ± ms leucogranite Ag ± Zn, Cu, and/or Pb veins/ replacement Nelson and Colpron (2007); Canadian Geochronological Knowledgebase (2013); this study. Mt. Reed, Mt. Haskin, Balourdet, and portions of the Cassiar batholith 42–58 (K–Ar) eastern CT (northern BC) quartz monzonite to monzonite Mo ± W porphyries (Mt. Reed, Mt. Haskin) Christopher et al. (1972); Evanchick (1988); Nelson and Colpron (2007); Canadian Geochronological Knowledgebase (2013). EASTERN SELWYN BASIN Early to 'mid'-Cretaceous: Hyland (or "Hyland River"; Hart et al., 2004a) 106–100 southeastern Yukon reduced–oxidized, peraluminous, laccolithic bt monzogranite to granodiorite abundant W–(Mo) or Sn skarns/porphyries and veins; distal Ag–Pb–Zn veins Roots et al. (1966); Hart et al. (2004a); Heffernan (2004); Hart and Lewis (2006). Tay River 99–95 southeastern SB reduced–oxidized, peraluminous, batholithic bt ± hbl or ms granodiorite to monzogranite and quartz monzodiorite dykes; quartz monzodiorite porphyry plugs small W–Cu skarns (Lucky), distal Pb–Zn–Ag veins/mantos (Roy) Pigage and Anderson (1985); Gordey and Anderson (1993; Heffernan (2004); Pigage (2004); this study. Tungsten ("two mica" plutons of the Selwyn suite) 98–95 Ma southwestern NT reduced, peraluminous, bt ± ms ± grt or tur monzogranite to leucogranite; associated kersatite dykes small to world–class W–Cu +/– Zn–Bi–Au–Mo (Mactung, Cantung, Lened, Clea), distal Pb–Zn–Ag veins Anderson (1982, 1983, 1988); Gordey and Anderson (1993); Lang (2001); Hart et al. (2004a, 2005); Rasmussen et al. (2011); this study. Mayo (or "South Lansing"; Mortensen et al., 2000) 98–93 east–central Yukon reduced > oxidized, metaluminous, bt ± hbl Cu–, Au–, or W–dominant skarns; Sn, Sb–As or U– Th–REE occurrences; quartz vein–hosted Au mineralization (Scheelite Dome) Anderson (1982, 1983, 1988); Emond and Lynch (1990); Emond (1986, 1992); Gordey and Anderson (1993); Murphy (1997); Lang (2001); Hart et al. (2004a, 2005); Mair et al., (2006a, 2011); this study. Tombstone 93–89 east–central Yukon, southwestern NT reduced-oxidized, peralkaline–metaluminous– peraluminous Cu–, Au–, or W–dominant skarns; Sn, Sb–As or U– Th–REE occurrences; quartz vein–hosted Au (Dublin Gulch) Anderson (1983); Emond and Lynch (1990); Emond (1986, 1992); Gordey and Anderson (1993); Murphy (1997); Lang (2001); Hart et al. (2004a, 2005); Emond (1986, 1992); this study. Late Cretaceous: Ivo, Salivo, Royall, Jox, Cream ~98 (Ar– Ar: 67– 65) southeastern SB reduced, peraluminous, bt– ms ± sil leucogranite W skarn (Ivo, Royall) Bezzola (2009); this study. Latest Cretaceous to Eocene: Ting 52 southeastern- most SB oxidized, alkaline, bt syenite plugs and associated intrusive breccia nd Pigage and Mortensen (2004); Fallas et al. (2004); Pigage (2009). Notes: a CT = Cassiar Terrane, SB = Selwyn Basin, YTT = Yukon–Tanana terrane; b mineral abbreviations are from Kretz (1983). 136 Table 2.2. Final U–Pb ages of crystallization for felsic plutons and dykes, and 40Ar/39Ar age determinations for mafic dykes, in the study area. Sample Plutona Method (mode)b Mode b Mineral Age (Ma) 2-δ error (Ma)c 2-δ error (%)c nd re Notes U–Pb Inheritance (Ma)f,g n Typeh Plutonic Suitei Cassiar Terrane and western Selwyn Basin: KR-07-01 Dycer Creek LA–ICP–MS mixed zrn 114.8 1.2 1.0 7 0 4 analyses + 3 outliers 118.1 (0.4) 5 A Cassiar KR-07-02 Quiet Lake LA–ICP–MS mixed zrn 114.7 1.1 1.0 9 0 8 analyses + 1 outlier *121.7 (1.5) 361.3 (3.6) 2 4 A X Cassiar KR-07-03 Quiet Lake LA–ICP–MS mixed zrn 97.7 1.0 1.0 13 0 4 analyses + 13 outliers 100.9 (0.6) 233.5 (1.5) 3 1 A X Seagull KR-07-04 Nisutlin LA–ICP–MS mixed zrn nd nd nd nd nd nd nd nd nd Seagull? KR-07-05 Nisutlin LA–ICP–MS mixed zrn 113.2 1.1 1.0 13 1 11 analyses + 1 outlier 2018 (27) 1 I Cassiar KR-07-06 Nisutlin LA–ICP–MS mixed zrn 105.1 0.7 0.7 14 0 13 analyses + 1 outlier not analyzed n/a n/a Cassiar KR-07-07 Young LA–ICP–MS mixed zrn 107.2 0.5 0.5 14 0 13 analyses + 1 outlier not analyzed n/a n/a Cassiar KR-07-08 St Cyr LA–ICP–MS mixed zrn 107.7 0.6 0.5 16 1 15 analyses not present n/a n/a Cassiar KR-07-09 Fox Mountain LA–ICP–MS mixed zrn 107.8 1.1 1.0 13 0 11 analyses + 2 outliers 112.3 (1.8) 2 A Cassiar KR-07-10 Big Salmon LA–ICP–MS mixed zrn 106.8 1.3 1.2 15 0 10 analyses + 5 outliers not analyzed n/a n/a Cassiar KR-07-11 Glenlyon LA–ICP–MS mixed zrn 107.7 0.7 0.7 11 0 11 analyses 113.3 (3.4) 3 A Cassiar KR-07-12 Tintina LA–ICP–MS mixed zrn 49.1 0.4 0.8 15 1 10 analyses + 4 outliers not present n/a n/a Black River KR-07-13 Square Lake LA–ICP–MS mixed zrn 111.7 1.1 1.0 7 0 6 analyses + 1 outlier not analyzed n/a n/a Anvil KR-07-14 Black River LA–ICP–MS mixed zrn 47.1 0.3 0.7 16 1 14 analyses + 1 outlier not present n/a n/a Black River KR-07-15 Meister Lake LA–ICP–MS mixed zrn 79.2 0.4 0.5 13 0 7 analyses + 5 outliers not present n/a n/a Rancheria KR-07-16 Cassiar LA–ICP–MS mixed zrn 111.5 0.5 0.4 13 1 11 analyses + 1 outlier not analyzed n/a n/a Cassiar KR-07-17 Cassiar LA–ICP–MS mixed zrn 111.1 0.6 0.5 13 1 9 analyses + 3 outliers not analyzed n/a n/a Cassiar KR-07-18 Cassiar LA–ICP–MS mixed zrn 114.3 0.9 0.8 14 1 12 analyses + 2 outliers not analyzed n/a n/a Cassiar KR-07-19 Marker Lake LA–ICP–MS mixed zrn 108.4 0.5 0.4 16 1 14 analyses + 1 outlier not analyzed n/a n/a Cassiar KR-07-20 Cabin Creek LA–ICP–MS mixed zrn 109.8 0.5 0.5 15 0 13 analyses + 2 outliers not analyzed n/a n/a Cassiar KR-07-21 Allen Stock LA–ICP–MS mixed zrn 77.2 0.4 0.5 14 0 7 analyses + 7 outliers not present n/a n/a Rancheria KR-07-22 Cassiar LA–ICP–MS mixed zrn 113.8 0.8 0.7 11 1 9 analyses + 1 outlier not analyzed n/a n/a Cassiar KR-07-23 Gravel Creek LA–ICP–MS mixed zrn 81.9 0.8 1.0 10 0 6 analyses + 4 outliers not analyzed n/a n/a Rancheria KR-07-24 Cassiar LA–ICP–MS mixed zrn 113.7 0.6 0.5 16 0 12 analyses + 4 outliers not present n/a n/a Cassiar 07M-150 Carolyn LA–ICP–MS mixed zrn 95.7 1.4 1.5 11 0 10 analyses + 1 outlier 102.0 (2.2) 107.2 (1.8) 2 1 A A Tay River 07M-151 (volcanic) LA–ICP–MS mixed zrn 97.2 0.8 0.8 10 0 10 analyses 102.5 (1.2) 6 A South Fork 07M-152 Buttle Creek LA–ICP–MS mixed zrn 104.0 0.5 0.5 16 0 16 analyses not analyzed n/a n/a Cassiar 07M-153 Glenlyon LA–ICP–MS mixed zrn 108.5 1.1 1.0 14 0 12 analyses + 2 outliers not analyzed n/a n/a Cassiar 07M-154 Little Salmon LA–ICP–MS mixed zrn 97.5 0.7 0.7 20 1 10 analyses + 10 outliers n/a n/a n/a Seagull ARG-33 Tsa Da Glisza ID–TIMS n/a mnz 112.2 0.5 0.4 n/a n/a 2 near-concordant fractions n/a n/a n/a Anvil FL-46 Square Lake ID–TIMS n/a mnz 112.3 0.3 0.3 n/a n/a 1 near-concordant fraction n/a n/a n/a Anvil 137 Sample Plutona Method (mode)b Mode b Mineral Age (Ma) 2-δ error (Ma)c 2-δ error (%)c nd re Notes U–Pb Inheritance (Ma)f,g n Typeh Plutonic Suitei KG-31 Tsa Da Glisza ID–TIMS n/a mnz 114.3 0.3 0.3 n/a n/a 1 near-concordant fraction n/a n/a n/a Anvil KM-07 Mink ID–TIMS n/a mnz 111.7 0.5 0.4 n/a n/a 1 concordant fraction n/a n/a n/a Anvil MC-CL-01 dyke (Clear Lake) ID–TIMS n/a mnz 108.1 0.2 0.2 n/a n/a 1 concordant fraction n/a n/a n/a Cassiar MP-18 Money ID–TIMS n/a mnz 112.4 0.7 0.6 n/a n/a 1 near-concordant fraction 761 (9) 1660 (7) 2366 (7) 1 1 1 nd nd nd Anvil Eastern Selwyn Basin: KR-05-08 Jorgensen LA–ICP–MS mixed zrn 95.9 0.8 0.8 12 0 11 analyses + 1 outlier 339 (2.8) 471 (4.5) 1 1 I I Tay River (CR) KR-05-10 Powers LA–ICP–MS mixed zrn 98.9 0.7 0.7 15 1 15 analyses 449 (4.2) *496 (7.7) 1 2 I I Tay River (CR) KR-05-22 Roy LA–ICP–MS IC + mixed zrn 96.1 0.9 1.0 16 0 16 analyses not detected n/a n/a Tay River KR-05-26 Fish LA–ICP–MS mixed zrn 96.1 0.6 0.6 10 0 10 analyses not present n/a n/a Tay River KR-05-32 Ivo LA–ICP–MS mixed zrn 96.0 1.0 1.0 12 0 12 analyses 103.2 (0.9) 7 A Tay River ('evolved') KR-05-43 Park LA–ICP–MS mixed zrn 96.1 1.2 1.2 16 0 14 analyses + 2 outliers not present n/a n/a Tay River (CR) KR-05-62 Mount Christie LA–ICP–MS IC zrn 95.1 1.2 1.3 12 0 12 analyses not detected n/a n/a Mayo KR-05-68 Mount Wilson LA–ICP–MS mixed zrn 97.7 0.8 0.8 12 0 12 analyses 103.2 (2.0) 1591 (24) 1766 (19) 4 1 1 A X I Mayo (H–M) KR-05-76 Pelly River LA–ICP–MS mixed zrn 98.0 1.1 1.1 15 0 15 analyses 104.2 (1.4) 1792 (39) 6 2 A I Mayo (H–M) KR-05-77 mafic dyke, Pelly River 40Ar–39Ar n/a bt 95.7 0.7 0.7 n/a n/a 91.9% of 39Ar, steps 3–9 n/a n/a n/a Mayo (H–M) KR-05-97b O'Grady LA–ICP–MS mixed zrn 92.8 0.6 0.7 13 1 11 analyses + 1 outlier 97 (15) 2 A Tombstone KR-05-110 Cac LA–ICP–MS mixed zrn 97.4 1.0 1.0 12 0 12 analyses that all plot above concordia not present n/a n/a Tungsten (NE) KR-05-126 (plug SW of O'Grady) LA–ICPMS IC zrn 95.9 1.4 1.5 12 0 12 analyses 101.7 (2.5) 3 A Tungsten (NE) KR-05-130 Central Nahanni LA–ICP–MS mixed zrn 97.1 0.8 0.8 15 0 15 analyses 101.3 (1.5) 106.1 (1.5) 4 1 A A Mayo KR-05-136 North Nahanni LA–ICP–MS mixed zrn 96.9 0.7 0.8 18 0 17 analyses + 1 outlier 101.8 (1.3) 107.1 (1.6) 5 1 A A Mayo KR-05-143 Circular Stock LA–ICP–MS mixed zrn 96.9 0.6 0.6 18 1 16 analyses + 1 outlier 103.0 (1.6) 5 A Tungsten (SE) KR-05-148 Rifle Range LA–ICP–MS mixed zrn 96.3 1.0 1.0 16 0 16 analyses not present n/a n/a Tungsten (SE) KR-05-157 mafic dyke, Roy 40Ar–39Ar n/a bt 96.5 0.7 0.7 n/a n/a 99.12% of 39Ar, steps 2–8 n/a n/a n/a Tay River KR-05-158 mafic dyke, Roy 40Ar–39Ar n/a bt 97.6 0.7 0.7 n/a n/a 94.1% of 39Ar, steps 3–8 n/a n/a n/a Tay River 138 Sample Plutona Method (mode)b Mode b Mineral Age (Ma) 2-δ error (Ma)c 2-δ error (%)c nd re Notes U–Pb Inheritance (Ma)f,g n Typeh Plutonic Suitei KR-05-164 Hole-in-the- Wall LA–ICP–MS mixed zrn 95.3 1.1 1.2 15 0 11 analyses + 4 outliers not present n/a n/a Tay River ('evolved') KR-05-175 Lened LA–ICP–MS mixed zrn 97.6 0.8 0.8 11 0 10 analyses + 1 outlier 1794 (14) 1 I Tungsten (NE) KR-05-191 Marion LA–ICP–MS IC zrn 96.7 1.4 1.4 10 0 10 analyses 103.7 (2.6) 1 A Tay River KR-05-194 Coal River LA–ICP–MS mixed zrn 92.8 0.9 1.0 10 0 9 analyses + 1 outlier 95.8 (0.8) 5 A Tombstone KR-05-196 Coal River LA–ICP–MS mixed zrn 98.4 1.0 1.0 14 0 13 analyses + 1 outlier not present n/a n/a Tay River KR-05-198 Hole-in-the- Wall LA–ICP–MS mixed zrn 91.3 2.1 2.3 5 0 5 analyses 96.7 (1.0) 10 A Tombstone KR-05-204 mafic dyke, Cantung 40Ar–39Ar n/a bt 91.0 0.8 0.8 n/a n/a 63.4% of 39Ar, steps 4–8 (overlap of points results in a '2-pnt' I.I., but this is within error of the I.I. age for all steps; alteration may have affected age) n/a n/a n/a Tungsten (SE) KR-05-208 Little Hyland LA–ICP–MS mixed zrn 96.8 0.6 0.6 17 1 17 analyses 99.9 (1.0) 1 A Tungsten (SE) KR-05-210 Nahanni Range Road LA–ICP–MS mixed zrn 97.2 0.5 0.5 15 0 15 analyses 100.3 (1.2) 1 A Tungsten (SE) KR-05-213 mafic dyke 40Ar–39Ar n/a bt 96.7 0.8 0.8 n/a n/a 99.99% of 39Ar, steps 1–7 n/a n/a n/a Tungsten CL-06-33 Dechen’La LA–ICP–MS mixed zrn 93.7 0.5 0.6 9 1 7 analyses + 1 outlier 96.4 (0.7) 15 A Tombstone CL-06-34 Natla LA–ICP–MS mixed zrn 92.3 0.6 0.6 12 0 8 analyses + 4 outliers 1807 (18) 1901 (21) 1 1 I I Tombstone CL-06-35 Logan LA–ICP–MS mixed zrn 95.6 1.2 1.3 11 0 7 analyses + 4 outliers not analyzed n/a n/a Mayo CL-06-36 Mount Christie LA–ICP–MS mixed zrn 96.1 0.5 0.5 12 0 8 analyses + 4 outliers not detected n/a n/a Mayo CL-06-37 Mount Christie LA–ICP–MS mixed zrn 96.7 0.6 0.7 13 0 9 analyses + 4 outliers not analyzed n/a n/a Mayo CL-06-38 Ross River LA–ICP–MS mixed zrn 93.0 0.9 1.0 7 0 6 analyses + 1 outlier 95.3 (0.5) 125.9 (0.8) 13 1 A A Tombstone CL-06-39 Keele River LA–ICP–MS mixed zrn 93.3 0.4 0.4 15 0 13 analyses + 2 outliers 96.3 (0.9) 5 A Tombstone CL-06-40 Mile 222 LA–ICP–MS mixed zrn 96.6 0.6 0.6 11 0 11 analyses not detected n/a n/a Mayo 03M-203 felsic dyke, Cantung ID–TIMS n/a zrn 97.4 0.3 0.3 n/a n/a 3 concordant/ near-concordant fractions n/a n/a n/a Tungsten (SE) 03M-204 aplite, Cantung ID–TIMS n/a zrn 98.3 0.3 0.3 n/a n/a 1 concordant fraction 1821 (39) 2180 (22) 2438 (16) 1 1 1 nd nd nd Tungsten (SE) 03M-207 Mine Stock- apical, Cantung pluton ID–TIMS n/a zrn 98.2 0.4 0.4 n/a n/a 6 concordant/near-concordant fractions n/a n/a n/a Tungsten (SE) 03M-214 Shelf Lake ID–TIMS n/a mnz 97.2 0.2 0.2 n/a n/a 1 concordant fraction n/a n/a n/a Tay River ('evolved') 03M-215 Mt Appler ID–TIMS n/a mnz 95.6 0.2 0.2 n/a n/a 2 concordant fractions n/a n/a n/a Tay River ('evolved') 03M-216 Faille ID–TIMS n/a zrn 97.0 n/a n/a n/a n/a 1 near-concordant fraction n/a n/a n/a Tay River 03M-217 Mount Sir James McBrien ID–TIMS n/a zrn 96.8 0.5 0.5 n/a n/a 3 concordant fractions n/a n/a n/a Tay River 139 Sample Plutona Method (mode)b Mode b Mineral Age (Ma) 2-δ error (Ma)c 2-δ error (%)c nd re Notes U–Pb Inheritance (Ma)f,g n Typeh Plutonic Suitei 03M-218 Hole-in-the- Wall (north) ID–TIMS n/a zrn 97.5 0.3 0.3 n/a n/a 2 concordant fractions n/a n/a n/a Tay River 03M-219 Pyramid Mtn ID–TIMS n/a mnz 97.4 0.3 0.3 n/a n/a 2 near-concordant fractions n/a n/a n/a Tay River ('evolved') 03M-220 Ragged ID–TIMS n/a mnz 98.2 0.2 0.2 n/a n/a 1 near-concordant fraction n/a n/a n/a Tungsten (SE) Notes: Preliminary U–Pb ages for the Eastern Selwyn Basin were previously reported in Rasmussen et al. (2007). a italicized = names given in this study; b Mode during LA–ICP–MS analysis: IC = ion-counting mode for all isotopes; mixed = IC (206, 207, 208, 235) + analogue (232, 238) modes; c 95% conf. interval unless annoted " * " (reported as internal 2-sigma error); d number of analyses robust enough to calculate age; e number of analyses rejected by Isoplot; f non-systematically measured ages of inheritance (1-δ uncertainty in brackets) with questionable but still relevant data italicized; g not analyzed = inherited cores observed but not analyzed; not detected = inherited cores were not observed but cannot be precluded; not present = inherited cores appear to be absent; h type of inheritance analyzed, as defined in the text: A = antecrystic I = old core; X = xenocrystic; nd = type of inheritance was not determined (applies to ID–TIMS upper intercept ages); i plutonic sub-suites or geographic subdivisions discussed in the text are given in brackets. 140 Table 2.3. Detailed summary of patterns of magmatism, macroscopic and microscopic features, and references for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. Mineral abbreviations are from Kretz (1983), nd = no data, and geographic abbreviations are: CT = Cassiar Terrane, SB = Selwyn Basin, YTT = Yukon–Tanana terrane. Cassiar Patterns of Magmatism batholithic to small plutons that get younger to the east over ~150 km; felsic phases are most abundant in north and intermediate phases are most abundant in the south in the Cassiar batholith (the largest intrusion in this suite) Macroscopic Textures generally massive, but may be moderately foliated along plutonic margins (e.g., Buttle Creek pluton: strain during emplacement) to mylonitized adjacent to bounding or cross-cutting regional structures (e.g., Cassiar batholith; syn-magmatic tectonic deformation); fine- to coarse-grained, equigranular to Kfs porphyritic (to megacrystic) to porphyry Composition apparently equal quantities of (a) reduced, ilm-bearing, bt ± ms monzogranite and leucogranite, and (b) oxidized bt ± hbl ± ttn-bearing monzogranite to granodiorite (and minor qtz monzodiorite) with abundant mag; abundant cross-cutting aplite/pegmatite dykes; rare cross- cutting lamprophyre dykes; contain bt-rich xenoliths and rarer metasedimentary rafts Mineralogy and Microscopic Textures ms is fine- to coarse-grained (magmatic), or very fine-grained and ragged (sub-solidus), or as felted patches (secondary); minor myrmekite; a few samples with moderately foliated mica, finer grain size, and (or) intense mylonitization (including rotated porphyroclasts); minor clay speckling in microcline; minor sericitization ± saussuritization of pl; minor to moderate chloritization of bt Accessory Minerals ubiquitous zrn, ap, ilm or mag ± mag; Ce-aln; Ce-mnz, Fe>Mg>Mn grt (or, Mg- or Ca-rich) in highly evolved peraluminous samples; rarer xenotime (frequently Ta-rich), rt, tur; U- and REE oxide minerals (e.g., uraninite, thorite, betafite) and fine-grained F ± Fe ± Nb-rich ttn in more evolved examples; fl (magmatic and secondary in pl), scheelite, pyrite, brl, uraninite, and molybdenite were reported in the Northern Dancer pluton; Ca>>Fe pyroxene and (or) ep (magmatic and secondary) in less evolved and more oxidized samples Zircon Textures well-developed oscillatory zoning is typical of magmatic zrn; slightly older, poorly to well-zoned, partially resorbed (1+ major resorption events) magmatic cores are observed in samples from the northwest CT; resorbed inherited cores and xenocrysts are nearly ubiquitous; inclusions of ap and mnz in magmatic zrn are typical References Noble et al. (1984); Lowenstern and Sinclair (1996); Driver et al. (2000); Ash (2001); Black et al. (2003); Mortensen et al. (2007); Colpron et al. (2011); Gordey (in press); this study. Anvil Patterns of Magmatism several small highly evolved plutons in the YTT inlier, and two laccolithic batholiths concordant to regional D2 foliation are separated by a metasedimentary horizon, with grt and staurolite in contact aureoles in the western SB; both regions lack miarolitic cavities (i.e., relatively deep level of emplacement) Macroscopic Textures massive to weakly foliated, fine- to medium-grained, and equigranular in the YTT inlier; moderately foliated (shallow-dipping deformation texture) and frequently Kfs megacrystic in the western SB Composition strongly reduced, ilm-bearing, bt ± ms monzogranite to ms ± tur leucogranite; cut by fine-grained hbl-(pl)-phyric diorite dykes in the western SB (Tay River suite?), or massive to highly strained aplites (with tur, and brl-bearing pegmatitic and qtz-rich segregations) and qtz-tur-(brl) veins in the YTT inlier Mineralogy and Microscopic Textures always bt-bearing, but with magmatic to sub-solidus ms; in the Anvil batholith ms is also present as ragged fine-grained sub-solidus grains and as clots with localized fibrolite (pseudomorphs of aluminous minerals?); in the Anvil batholith pl is oligoclase (An = 25–31) and displays oscillatory and normal zoning; minor myrmekite; near complete chloritization of bt, moderate sericitization in pl, and clay speckling Kfs Accessory Minerals ubiquitous zrn, ap, ilm (not observed so may not be abundant, but assumed to be present based on reduced character); Ce-mnz; Nb–Ta- rich rt (Square Lake pluton); late-stage aln, tur, and (rarer) grt (Anvil batholith); rare grt (Tsa Da Glisza pluton); grt and brl (white and green) locally (Tsa Da Glisza aplites) Zircon Textures large, resorbed and porous, highly fractured cores of unknown age with thin, oscillatory zoned magmatic rims (sample KR-07-13 only) References Pigage and Anderson (1985); Mortensen et al. (2000); Murphy et al. (2001, 2002); Neufeld (2004); Pigage (2004); Neufeld et al (2005); this study. Hyland Patterns of Magmatism thick batholithic laccoliths with mainly concordant contacts with schistose, gneissic, and migmatitic country rocks; staurolite to grt-grade contact aureoles (i.e., relatively deep level of emplacement) Macroscopic Textures typically massive (although sometimes weakly foliated), fine- to medium-grained Composition reduced to oxidized, bt granodiorite and monzogranite, with lesser granodioritic to dioritic pl-bt porphyry phases and enclaves Mineralogy and Microscopic Textures bt is the primary ferro-magnesian mineral; hbl not reported in any of the sampled intrusions or enclaves, but its presence has been noted in assessment reports (e.g., hbl granodiorite, qtz monzonite, and qtz monzodiorite phases as well as abundant mafic xenoliths and dioritic dykes); ms was also not observed in any of the samples, but has been reported in some small-volume leucocratic phases associated with W skarns Accessory Minerals zrn, ap; nd Zircon Textures nd References Roots et al. (1966); Kuehnbaum (1977); Brock (1978); Durgin (1979); Heffernan (2004); Hart et al. (2004a); Hart and Lewis (2006). Seagull Patterns of Magmatism narrow NNW-trending belt at the western margin of the CT; typically comprising small plutons or plugs, but batholiths occur at the south end of the belt; miarolitic cavities with tur and fl ± brl and toz, pegmatitic pods, and (or) qtz-tur orbicules with aplitic margins are abundant in equigranular bt granite of the eastern Seagull batholith (i.e., relatively shallowly emplaced) Macroscopic Textures fine-to coarse-grained, equigranular to Kfs porphyritic to megacrystic, to porphyry Composition reduced, ilm-bearing, one-mica (bt) leucogranite to monzogranite; volumetrically small, lepidolite–(zinnwaldite)–albite–fl–toz leucogranite stocks, high-level dykes, and sills adjacent to larger batholiths; minor, disaggregated, fine-grained to hbl-bt-phyric granodiorite (equigranular to Kfs-phyric) to tonalite–diorite enclaves (Thirtymile Stock); rare cross-cutting hbl-bearing lamprophyres 141 Mineralogy and Microscopic Textures bt is the only ferro-magnesian mineral; rare magmatic ms (KR-07-03, -04); local rapakivi and granophyric textures; myrmekite is present in several plutons; minor to moderate clay speckling in Kfs, minor sericitization of pl, and minor chloritization of bt; blue tur and (or) fl associated with Li enriched bt rims or after pl in the Seagull batholith; more intense but localized toz ± brl mineralized greisen veins and greisenization of megacrystic granite and exoskarn Accessory Minerals ilm, Ce-rich mnz, zrn, and ap are typical, except ap is not found in highly evolved phases where mnz >> zrn; ep and aln are present in rare granodioritic intrusions; also contain grt (Fe- or Ca-rich varieties), and Ce-rich aln; rarer xenotime (may be Ta-rich), Fe±Nb-bearing ttn, thorite, rt, toz (probably sub-solidus; Clarke, 1981), zinnwaldite, tur, pyrite, and REE oxide minerals (e.g., fergusonite); trace Ca-pyroxene and hbl (not seen in thin section) in the Little Salmon pluton Zircon Textures highly fractured and may or may not contain resorbed inherited cores; resorbed and well-developed oscillatory zoning in magmatic rims (with porous margins or overgrowths), and in antecrystic cores; ap ± mnz inclusions References Liverton (1990, 1992); Sinclair and Richardson (1992); Liverton and Botelho (2001); Colpron et al. (2003); Liverton et al. (2005); Liverton (2011, pers. comm.); M. Colpron (2011, pers. comm.); this study. Tay River (Western SB) Patterns of Magmatism spatially associated with, and cross-cut the Anvil plutonic suite Macroscopic Textures massive to strongly foliated, fine- to coarse-grained, equigranular to rarely Kfs-phyric or megacrystic Composition reduced (ilm > mag), bt ± hbl (± Ca-rich clinopyroxene), monzogranite and granodiorite batholiths and plutons, with significant granodioritic to monzogranitic qtz–feldspar porphyry bodies (typically bt–hbl-bearing) and abundant dykes and sills Mineralogy and Microscopic Textures pl is typically andesine (An = 34–53) and has oscillatory and normal zoning, although irregular calcic cores are abundant in hbl-bearing phases; chl, ttn, ep, and mag after hbl and bt, minor to intense sericitization ± saussuritization or calcification of pl (particularly more calcic cores), and minor to intense clay speckling in Kfs Accessory Minerals zrn, ap, ilm, aln (Ce-rich), tur, and mnz are typical accessory minerals Zircon Textures relatively large, have thick and oscillatory zoned magmatic rims around antecrystic cores, display evidence for 2–3 significant resorption events, and contain particularly abundant ap inclusions (sample 07M-150) References Wood and Armstrong (1982); Pigage and Anderson (1985); Mortensen et al. (2000); Pigage (2004); Gordey (in press); this study. South Fork Patterns of Magmatism a cluster of eight calderas that are spatially associated with the Tay River suite Macroscopic Textures densely welded to crystal lithic tuff pyroclastic rocks; syn-magmatic dropped-down blocks related to magma chamber evacuation Composition thick sequences (up to 950 m thick in the Connolly Caldera) of predominantly reduced, pl–qtz–bt–(hbl) lapilli tuff of dacitic composition; lesser basaltic-andesite and andesite flows and dacitic ash flow tuffs Mineralogy and Microscopic Textures pl ranges from An = 75–10, through basaltic andesite to rhyodacite flows; hypersthene is locally present; chl, ttn, ep, and mag after hbl and bt, minor to intense sericitization ± saussuritization or calcification of pl (particularly more calcic cores), and minor to intense clay speckling in Kfs Accessory Minerals zrn, ap, ilm > mag, Ce-rich aln; rare corundum (1 grain) Zircon Textures relatively large, have thick and oscillatory zoned magmatic rims around antecrystic cores, display evidence for 2–3 significant resorption events, and contain particularly abundant ap inclusions (sample 07M-151) References Wood and Armstrong (1982); Pigage and Anderson (1985); Mortensen et al. (2000); Pigage (2004); Gordey (in press); this study. Tay River (Eastern SB) Patterns of Magmatism small plutons to large batholiths; small to large tur-infilled miarolitic cavities with aplitic margins in 'evolved' phases (i.e., relatively shallowly emplaced); associated with a 'welt' of higher grade host rocks in southeastern SB; Coal River sub-suite: small, isolated plugs Macroscopic Textures massive, fine- to coarse-grained, equigranular to Kfs-phyric and megacrystic; Coal River sub-suite: rarely foliated, equigranular or fine- grained to porphyry-textured Composition typically moderately to highly oxidized (i.e., ttn > mag > ilm); coarser-grained and Kfs porphyritic to megacrystic, bt monzogranite and leucogranite that also contain appreciable magmatic ms ± tur, and mnz (referred to as 'evolved': e.g., Shelf Lake, Mount Appler, and Shannon Creek plutons, portions of the Hole-in-the-Wall and Coal River batholiths, and the Jox, Cream, Ivo, and Salivo plugs); associated dykes range from very fine-grained pl-(qtz)–phyric qtz monzodiorite, to pl–bt-phyric microgranite, to aplite and rare pegmatite; rare igneous fine-grained to Kfs porphyritic bt ± hbl-bearing xenoliths in "evolved" phases; Coal River sub-suite: pl ± hbl ± bt ± qtz porphyritic, granodiorite to qtz monzodiorite Mineralogy and Microscopic Textures myrmekite ± spherulitic qtz–feldspar intergrowths in "evolved" phases; pl commonly displays complex or oscillatory zoning in Tay River suite intrusions (skeletal textures in rare examples), and is sieve-textured in associated qtz monzodiorite dykes; chl, ttn, ep, and mag after hbl and bt; minor to intense sericitization ± saussuritization or calcification of pl (particularly more calcic cores), and minor to intense clay speckling in Kfs; Coal River sub-suite: one large dark honey-brown zoned elongate-prismatic phenocryst or xenocryst was observed (aln?; sample KR-05-08; pronounced alteration includes: chl–ttn–ep–mag after hbl and bt, minor to intense sericitization ± saussuritization or calcification of pl (particularly more calcic cores), and minor to intense clay speckling in Kfs Accessory Minerals zrn, ap, ttn, mag, ilm, and Ce-rich aln are typical; Ca- and Fe-rich grt, Ta-rich xenotime, wolframite (possibly fracture/vein-related), and a rare Nb–F–Y-rich ttn have also been observed; accessory clinopyroxene and hbl (inherited?) in the 'evolved' Shelf Lake pluton; Coal River sub-suite: zrn, ap, ttn, mag, ilm, and Ce-rich aln Zircon Textures relatively large and thickly oscillatory zoned magmatic rims around poorly to unzoned resorbed magmatic cores of the same age; individual grains commonly display evidence for 1–2 significant resorption events; no inherited cores were observed; Coal River sub-suite: zrn from at least two small qtz monzodiorite porphyry bodies are texturally similar to those described for the Tay River suite plutons (although magmatic cores are typically well-zoned), but also have well-zoned resorbed inherited cores References Anderson (1983); Mortensen et al. (2000); Heffernan (2004); Bezzola (2009); Pigage et al. (2011, and pers. comm.); this study. Tungsten Patterns of Magmatism typically rounded plugs (possibly cupolas) and small plutons; typically massive, but bt ± Kfs megacrysts may be foliated along contacts, or locally intensely mylonitized 142 Macroscopic Textures fine-grained and equigranular, medium- to coarse-grained and Kfs porphyritic or megacrystic; gradational internal contacts Composition ilm (reduced) and mnz-bearing bt monzogranite and leucogranite; dominated by a Kfs-phyric or megacrystic phase, with equigranular and fine-grained leucocratic ms ± bt, grt, and (or) tur-bearing phases (typically localized along the margins); abundant cross-cutting and satellitic, locally miarolitic, aplite ± pegmatite ± qtz–feldspar–(bt) porphyry dykes; rare cross-cutting kersantitic lamprophyres; rare micro- to porphyritic diorite (biotitic) enclaves; sheeted qtz ± tur veins or gossanous fractures/joints cut apical portions of intrusions Mineralogy and Microscopic Textures may have magmatic ms, but more typically as a minor ragged sub-solidus phase in the groundmass; qtz commonly forms large composite, smokey grey grains in coarse-grained samples (particularly those associated with W skarns); significant quantities of myrmekite ± spherulitic qtz–feldspar intergrowths may present in highly evolved felsic phases (felsic dykes, apical intrusions); Kfs megacrysts up to 13 cm in length, but typically 2–5 cm long; boxy-cellular and sieve-textured pl observed in the Cantung intrusions; local chloritization (± rt, ep, ilm) of bt, minor to moderate sericite and fine-grained ms and carbonate in pl cores, and minor to intense clay speckling in Kfs; pervasive secondary calcic mineral assemblages (e.g., mainly Ca-rich pl ± qtz, or ep-group minerals or hbl) have been observed in altered leucocratic dykes directly associated with W skarns Accessory Minerals ubiquitous zrn and Ce-rich mnz; less abundant ap; Ce-rich aln, xenotime (may be Ta-rich), F-rich ttn (particularly associated with W skarns), uraninite, and scheelite (possibly fracture/vein-related) may also be present; tur or grt in some marginal intrusive phases or late-stage felsic dykes; andalusite reported in satellitic intrusions of the Cac pluton Zircon Textures thick, well-zoned, and generally highly fractured magmatic rims on partially to strongly resorbed (1–3 or more resorption events) and poorly to well-zoned magmatic cores, although resorbed and often porous inherited cores are still common; ap inclusions are not as abundant as in other plutonic suites, but mnz inclusions are typical References Anderson (1982, 1983, 1988); van Middelaar (1988); Gordey and Anderson (1993); Mortensen et al. (2000); Hart et al. (2004a, 2005); Rasmussen (2004); Rasmussen et al. (2011); this study. Mayo Patterns of Magmatism small, homogenous to composite plutons Macroscopic Textures massive to foliated (mainly aligned bt along pluton contacts); fine- to coarse-grained, equigranular to weakly Kfs porphyritic to coarsely megacrystic Composition reduced to slightly oxidized (ttn ± mag bearing), hbl > bt (± trace clinopyroxene) qtz monzonite or monzodiorite to bt > hbl (± trace augite) monzogranite to granodiorite; some Kfs-crowded phases; lesser ms ± grt-bearing leucogranite; cut by aplite-pegmatite dykes, graphic qtz- tur veins, base metal mineralized veins and sheeted fractures; may contain tur-infilled vugs and up to 5% fine-grained bt ± hbl diorite xenoliths; abundant associated late mafic dykes, including: monzonite, diorite, and late lamprophyre dykes ranging from sub-alkaline kersantite or spessartite to alkalic minette Mineralogy and Microscopic Textures where clinopyroxene is present, it is typically Mg- and (or) Ca-rich augite and diopside that are respectively interpreted to be early- and late- crystallizing; zoning in pl is frequently complicated and oscillatory, and although rare, probable xenocrysts (or inherited cores) of pl and qtz are present in samples; hbl after clinopyroxene, chl after bt and hbl, minor to moderate sericite ± ms in pl cores, and minor to moderate clay speckling in Kfs Accessory Minerals ubiquitous ttn (particularly coarse-grained and abundant, and may be Nb-rich), Ce-rich aln, zrn, and ap; rarer mnz, thorite, Fe-rich grt, mag, and scheelite (possibly vein/fracture-related; very fine-grained ilm inclusions in mafic phenocrysts have been reported Zircon Textures typically thick, oscillatory zoned magmatic rims with well-zoned magmatic cores displaying up to 2 major resorption-growth episodes; resorbed and commonly porous inherited cores are abundant in some samples (e.g., 30–50% of the zrns examined in the Mount Wilson, Pelly River, and North and South Nahanni plutons) References Anderson 1(982, 1983, 1988); Gordey and Anderson (1993); Mortensen et al. (2000); Lang (2001); Hart et al. (2004a, 2005); Mair et al. (2006a, 2011); this study. Tombstone Patterns of Magmatism scattered, rounded and commonly concentrically zoned, composite plutons with gradational internal contacts Macroscopic Textures tur-infilled orbicule or miarolitic cavities are abundant in some intrusive phases; massive, medium- to coarse-grained Composition frequently become felsic towards the centre (e.g., Syenite Range, Mount Brenner, and Emerald Lake plutons); reduced + ilm or ttn-bearing (eastern SB), to slightly-moderately oxidized with mag > ttn (western SB); syenites dominated by Kfs (frequently crowded), qtz, or aegerine– augite are the most voluminous intrusive phases in the northwest SB; sub-alkaline, fine- to coarse-grained, hbl > bt or hbl > clinopyroxene, qtz monzonite and monzogranite are dominant in the eastern SB (although in the northeast, up to 40% perthitic Kfs with subordinate anhedral pl are observed); less abundant, sub-alkaline to alkaline, granite, monzonite, granodiorite, hbl diorite, and pyroxenite, (syeno)gabbro, and localized Si-undersaturated phases form small, often satellitic bodies (e.g., tinguate) across the SB; bt monzogranite and highly fractionated, peraluminous leucogranitic plutonic phases and dykes with magmatic ms and tur phases; mafic, round to elongate inclusions (up to 3%) are typically fine-grained bt ± hbl diorite, although hbl-pl porphyry, inclusions within inclusions, and other mineralogical and textural varieties may also be present; late aplite and pegmatite dykes, locally abundant miarolitic cavities, and sheeted mineralized qtz veins Mineralogy and Microscopic Textures aegerine–augite is the dominant mafic mineral in alkaline intrusive phases, whereas hbl (arfvedsonite) is abundant than bt or clinopyroxene in sub-alkaline phases (except in the more highly evolved monzogranitic intrusions); pl locally displays complicated zoning patterns; hbl after aegerine–augite (which may be partially preserved in hbl cores), hbl overgrowths on bt in alkaline intrusions (e.g., O’Grady batholith), chl ± ms after bt in sub-alkaline intrusions, minor to intense sericite ± ms in pl, and minor to moderate clay speckling in Kfs Accessory Minerals mag, ttn, zrn, Ce-rich aln, and ap are characteristic accessory minerals; tur is locally abundant (e.g., O'Grady batholith); grt with a Mg-rich core (pyrope?) and an Fe-rich rim (almandine) was detected in the O’Grady batholith; Ta-rich mnz, thorite, F-rich ttn, and Ti-poor ilm after sulphide may also be present; magmatic fl, ep, and carbonate are also possible in strongly alkalic intrusions, and melanitic grt + ne ± pseudoleucite are typical of the Si-undersaturated members of this suite Zircon Textures in the northwestern SB magmatic zrn is described as poorly zoned with no inherited cores, but in the northeastern to southeastern SB inherited cores range from absent to abundant depending on the intrusion and intrusive phase; in the eastern SB thick well-zoned magmatic rims are developed on older partially resorbed well-zoned antecrystic cores (evidence for at least 1 major resorption event) References Anderson (1983, 1988); Abercrombie (1990); Gordey and Anderson (1993); Murphy (1997); Duncan et al. (1998); Duncan (1999); Mortensen et al. (2000); Lang (2001); Coulson et al. (2002); Hart et al. (2004a, 2005); this study. Rancheria 143 Patterns of Magmatism a north-trending belt of small plutons east of the Cassiar batholith in southern Yukon Macroscopic Textures no evidence for post-crystallization deformation; medium- to coarse grained, equigranular to Kfs megacrystic; very coarse-grained megacrystic pods that may be highly enriched in uranium (up to 2.1% U3O8) and REE Composition highly reduced, to highly oxidized (in U-porphyry mineralized Allen Stock); bt–(ms) leucogranite to monzogranite; cut by aplite dykes with granophyric intergrowths, qtz veins, and rare biotitic qtz-absent (kersantic?) mafic dykes Mineralogy and Microscopic Textures ragged sub-solidus ms, chloritization of bt, and minor to moderate sericite in pl and clay speckling in Kfs; Gregory (2007) also reported linear zones of clay-altered monzogranite in the Allen Stock that were enriched in U and REE Accessory Minerals ubiquitous ap, zrn, Ce-mnz, with or without Ce-rich aln, Ca-rich grt, Ta-rich xenotime, Fe–(F or Mg) ttn, Ta–Nb-rich rt, Mg-rich pyroxene, and a Nb–Y–Ta–U–Dy–Yb–Th ± Ca oxide (possibly betafite or formanite-fergusonite); in the Allen Stock, Gregory (2007) also reported: Ce-mnz as one of the main hosts of U, along with several Fe–Ti oxides (e.g., rt, pseudort, ilm, and mag), minor ap, and abundant zrn in the Kfs megacrystic phase; Ce-mnz, thorite, uraninite, and rare kobeite present in the equigranular phase; and V–Nb-bearing rt and a Ti–U–Nb–Y– (Ca-Fe) oxide (possibly kobeite) in the aplite dykes Zircon Textures variably resorbed oscillatory zoned magmatic rims on rounded resorbed and poorly zoned inherited cores (sample KR-07-23) or oscillatory zoned and partially to significantly resorbed magmatic cores lacking evidence for inheritance; ap and mnz inclusions are typical of zrn from the Rancheria suite References Gregory (2007); this study. Black River Patterns of Magmatism two plutons near/adjacent to Tintina Fault, although similar-age magmatism is scattered across the SB and YTT inlier northeast of the Tintina Fault Macroscopic Textures no evidence for post-crystallization deformation, aside from a well-developed east-west trending jointing surface in the larger Black Lake pluton; sub-porphyry to medium-grained, Kfs porphyritic Composition reduced; bt (± trace ms) monzo- to leucogranite Mineralogy and Microscopic Textures sub-solidus ms, chloritization of bt, moderate sericitization of pl cores, and minor to moderate clay speckling in Kfs; clots of secondary or sub-solidus ms in the Tintina pluton may be pseudomorphs of earlier aluminous minerals Accessory Minerals ubiquitous ap, zrn, and Ce-rich mnz, as well as tur, Fe-rich grt (almandine), Ta–Nb-rich rt, Ta-rich xenotime, and uraninite Zircon Textures well-developed, partially resorbed and (or) porous oscillatory magmatic rims on variably resorbed and highly fractured magmatic cores; ap and mnz inclusions References this study. 144 Table 2.4. Details of geochemical discrimination plots used in Figure 2.4–2.13. Plot Namea Description Source a) Harker Variation Harker variation plots illustrate the extent of chemical variation within a plutonic suite (with respect to SiO2). Continuous linear trends are commonly attributed to magma mixing or restite-unmixing proccesses, whereas curvilinear trends commonly attributed to fractional crystallization and (or) assimilation processes. In instances where other geological evidence for these processes is lacking, however, chemical variation attributed to the range of compositions within a single heterogeneous melt source, the amount and type of peritectic mineral entrainment, and the conditions of partial melting (which will control the reactant minerals, melt fractions, etc.) n/a b) Fractionation Indices Rb/Sr, Eu/Eu* (the magnitude and sign of an Eu anomaly when REEs are normalized to chondritic representations of bulk Earth composition), and LaCN/YbCN (a measure of the steepness of slope from light to heavy REEs when the REEs are normalized to chondritic compositions) versus SiO2 are commonly used to provide information on the degree of magmatic fractionation experienced by a magma. These variables are also controlled by the relative role of plagioclase versus mica (mainly biotite) reactants in the melt source. These ratios are also influenced by the entrainment, or lack thereof, of peritectic plagioclase (formed during biotite and hornblende dehydration) and REE- bearing accessory minerals. n/a c) Chondrite- and Primitive Mantle- Normalized multi- element Multi-element profiles are also useful for comparing trace element and REE abundances, and identifying anomalies, that characterize specific plutonic suites. In this study, REEs are normalized to the average chondritic meteorite (which is representative of bulk Earth composition), and trace elements (mainly) are normalized to a reference for the primitive mantle prior to bulk differentiation of the Earth. Concentrations of trace elements and groups of trace elements provide information on the composition and type of melt source(s) that the magmas were derived from. Sun and McDonough (1989) d) High-K Calc- Alkaline- Normalized multi- element A multi-element normalization to the high-K calc-alkaline Yenchichi–Telabit liquid line of descent reference series is used to distinguish post-collisional high-K calc-alkaline and shoshonitic granitoids from post-collisional to post- orogenic/anorogenic alkaline granitoids. Normalization to a liquid line of descent removes, or significantly reduces, the effects of magmatic differentiation on composition by comparing elemental abundances on the basis of SiO2 content. To simplify these plots, only average compositions for each plutonic suite and sub-division therein are plotted; however, plots of all samples normalized to the Yenchichi–Telabit series are presented in Appendix D.4. Liégeois et al. (1998) e) QKP Ternary The International Union of Geological Scientists (IUGS) quartz–K-feldspar–plagioclase (QKP) ternary diagram for naming Si-saturated or granitoid rocks, and on which CIPW-normative orthoclase, quartz, and albite + anorthite compositions have been plotted for the plutonic suites. Streckeisen (1976) f) Shand's Index Molar Al / [Ca – 1.67P + Na + K] or "A/CNK" and Al / [Na + K] or "A/NK" provide information on aluminum saturation relative to the alkali elements (Zen, 1986). If A/CNK >1.0, then the rocks are considered to be peraluminous ("Pm") and corundum-normative; therefore in order to accommodate the excess aluminum they must crystallize an Al- bearing mineral in addition to feldspar (usually biotite at low levels of peraluminicity, but muscovite, garnet, an Al2SiO5 polymorph, cordierite, or corundum are typical in strongly peraluminous rocks). If A/CNK <1.0, the rocks are metaluminous ("M") and have excess Ca relative to the Al than what can be accommodated in the feldspars, calcic amphibole or pyroxene precipitate if A/NK >1.0. If A/NK <1.0 in this latter scenario, the rocks are peralkaline ("Pk") and have excess K+Na than what can be accommodated in feldspar, therefore Na must partition into ferro- magnesian silicates and, in cases of low A/NK, will precipitate the sodic amphiboles (riebeckite, arfvedsonite) and pyroxenes (e.g., aegerine) that characterize peralkaline rocks. Peralkaline rocks typically evolve to strongly peraluminous compositions by the fractionation of albite and/or K-feldspar. Shand (1943); Zen (1986) g) Modified Alkali- Lime Index MALI: (“Na2O + K2O – CaO”) is adapted from the original alkaline/calk-alkaline versus shoshonitic scheme of Peacock (1931) and is used to differentiate co-magmatic alkalic ("A") versus alkali-calcic ("A-C") versus calc-alkalic ("C-A") versus calcic ("C") granitic series. Frost et al. (2001) h) K2O versus SiO2 Shoshonitic ("sh") versus high-K calc-alkaline ("h-K") versus calc-alkaline or medium-K ("m-K") versus tholeiitic or low-K ("l-K") sub-divisions is used to sub-divide the evolutionary trends exhibited by calc-alkaline and shoshonitic volcanic suites with distinct parent magma compositions. Peccerillo and Taylor (1976) i) Na2O versus K2O The Na2O versus K2O plot for distinguishing I-type granites (derived from an infracrustal or igneous protolith), and S-type granites (derived from a supracrustal or sedimentary protolith). S-type granites are depleted in elements that are soluble during the weathering of feldspar to clay minerals at the Earth’s surface (e.g., Na, Ca, Sr, and Mg). S- type granites are also typically enriched in K2O (as well as Rb and Pb) relative to I-type granites. Partial melting of immature supracrustal sources (e.g., feldspathic and clay-poor greywackes, ± calcareous or amphibolitic lithologies), however, may produce a geochemical signature more akin to an infracrustal source (e.g., Clemens, 2003). Thus, Na2O versus K2O must be interpreted carefully before genetic assignments are made. Chappell and White (1974, 2001) j) normative corundum versus FeT CIPW-normative corundum versus FeT also disinguishes I- and S-type granitoids. S-type granitoids have >1 wt.% normative corundum in their most mafic phases, and corundum in these rocks will increase with maficity. I-type granitoids will have low to negligible normative corundum, although mnormative corundum may decrease with maficity. Chappell and White (1974, 1992) k) Ce versus 10000*Ga/Al used to distinguish A-type (A), or ‘anorogenic’ or ‘alkaline’, granitic suites from I- or S-type (I+S) granitic suites, where Ce represent the LREE enrichment and 10000*Ga/Al represents the enrichment in highly charged cations and depletion in Al that characterize A-type granites (e.g., Collins et al., 1982; Whalen et al., 1987; King et al., 1997). Highly fractionated I- and S-type magmas may also be elevated in the elements that define A-type magmas, such as Ga, Y, and Ce (e.g., London, 1995). Whalen et al. (1987) 145 l) Hf – Rb/30 – Ta*3 ternary plot Distinguishes pre-, syn-, and late to post-collisional granitic rocks . Pre-collisional volcanic-arc ("VA") rocks are enriched with LILEs and typical medium-K calc-alkaline, I-type, magmatic arc granitoids; syn-collisional ("SC") S- type leucogranites are enriched in Rb, F, B ± Ta and depleted in REE, Zr, and Hf, suggesting partial melting of over- thickened continental crust comprising metasedimentary rocks or crystalline basement; late to post-collisional ("PC") calc-alkaline suites of I-type affinity are similar to VA suites, although they are frequently even more enriched in incompatible elements than typical VA magmatism; and within plate ("WP") magmatism originates in the mantle and is attributed to decompression during/following orogenic collapse, the development of a back-arc rift basin, or extension related to pre- to post-collisional strike-slip faulting (Harris et al., 1986). Crustal contamination is thought to be responsible for the elevated Ta in late to post-collisional granitoids, although in some orogenic belts these rocks are not enriched in Ta and will plot in the VA field (Harris et al., 1986). Harris et al. (1986) m) Trace element tectonic discrimination diagram Trace element tectonic discrimination diagrams for granitoids emplaced into "VA" (including post-tectonic biotite +/- hornblende bearing I-type granites), "SC" (generally referring to muscovite-bearing peraluminous S-type granites), "WP" (typically A-type granites), and ocean ridge ("OR"; typically ‘plagiogranites’) settings. While these plots may indicate simple tectonic settings, they are frequently not diagnostic of the granite type or overall tectonic setting as different granite types may be emplaced into more than one tectonic setting. Post-collisional granitoids in particular can occupy either the VA, SC, and/or portions of the WP fields (broadly implying I-, S-, or A-type granite affinities, respectively), depending on the amount of Rb in the melt source, the amount of fluid available for partial melting (which is positively correlated with Rb-enrichment of the magma), and the composition of the source rocks (e.g., mixed mantle-crust origin). In the latter case, trace element discrimination plots may be more effective at identifying the tectonic setting of the protolith than the setting of the plutonic rocks themselves (e.g., Pearce et al., 1984; Harris et al., 1986; Arculus, 1987; Twist and Harmer, 1987; Roberts and Clemens, 1993; Ferré and Leake, 2001; Heffernan, 2004). Pearce et al. (1984) n) SNY versus SNX Separation of post-collisional granitoids from post-orogenic granitoids by comparing the sum of the average Rb–Th– U–Nb concentrations (SNY) with the sum of the average Y–Ce–Sm–Yb concentrations (SNX) on a bivariate plot, after all elements have been normalized to the post-collisional high-K calc-alkaline Yenchichi–Telabit reference series (Liégeois et al., 1998). This scheme distinguishes post-collisional high-K calc-alkaline and shoshonitic granitoids from post-collisional to ‘anorogenic’ alkaline magmatism. Typical medium-K calc-alkaline I- or S-type granitoids related to pre-collisional settings ("VA"–"SC") can also be distinguished from highly-fractionated I- or S- type granites (although the latter occupy a similar field as high-K calc-alkaline granitoids). A-type granites can also be sub-divided into aluminous-alkaline shoshonitic granitoids, and alkaline-peralkaline granitoids emplaced into a truly post-orogenic setting ("PO"; Bonin et al., 1998). Bonin et al. (1998); Liégeois et al. (1998) a letters correspond to the letters in Figures 4-13. 146 Table 2.5. Summary of geochemical observations from Figures 2.4-2.14 for syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline. “High-T” and “low-T” refer to the relative magma temperature for I-type magmas (Chappell et al., 1998, 2000). Cassiar Plutonic Suite (widespread in Cassiar Terrane) References for Compiled Geochemical Data Driver et al. (2000); Abraham et al. (2001); Colpron et al. (2002); Morris and Creaser (2008); Brand (2009); this study a) Harker Variation SiO2 down to 57 wt.% (high-T I-type, at least for hornblende-bearing); relatively tight, continuous, negative major element oxide trends with gentle downward inflections ca. 63 wt.% SiO2 (low-T I-type); K2O displays opposite trends; P2O5 ± Y decreases to ca. 73 wt.% SiO2 (I-type; apatite fractionation) then increases sharply (low-T I-type); BaO and Ce increase to ca. 67-69 wt.% SiO2 then decrease sharply (high-T I-type); relatively scattered Ce–Y due to heterogeneous melt source(s) (low-T I-type); Zr (XRF) decreases for SiO2 (low-T I-type); Cr (spinel/cpx), Co (olivine) are very low and/or decrease with increasing SiO2 (mineral fractionation); Th increases to ~67 wt.% SiO2, then diverges dramatically although the majority of samples decrease with higher SiO2; NOT SHOWN: coupled Th–La–Ce (allanite fractionation); Nb is relatively constant, but slightly enriched in hornblende-bearing b) Fractionation Indices Rb/Sr is typically < 2.5, but many samples with ~75 wt% SiO2 have Rb/Sr of 2-12; Eu/Eu* ranges fully from very strongly negative to slightly positive (0.09–1.14; average = 0.59) and decreases with increasing SiO2; LaCN/YbCN is low to high (2.83–58.99; average = 22.83), does not vary with Eu/Eu* (not shown), and increases up to ca. 68 wt.% SiO2, ranges widely between 68-74 wt.% SiO2, then decreases to very low values (monazite ± zircon) c) Chondrite- and Primitive Mantle- normalized relatively high TE and REE variation (heterogeneous); PM: relatively depleted in some HFSE’s (Ti ± Zr, P) and enriched in other HFSE’s (Ta, Nb ± Zr, P), LILE’s (K, Ba, Th, Rb), Sr, and LREE’s d) Average High-K Calc-Alkaline- Normalized Ca–Ti–Yb–(Na–Ce–Nd) depleted, variable U–Th–Ta–(Nb), and Sr enriched: average biotite and muscovite-bearing samples are nearly identical (except for depleted U and enriched Ta in muscovite-bearing phases); average hornblende-bearing samples are similar but more enriched in U–Th–Ta–Nb e-k) Geochemical Discrimination one linear trend from quartz diorite to monzogranite; weakly metaluminous to moderately peraluminous (A/CNK = 0.9– 1.2); CA–AC (with A–AC trend for least siliceous hornblende-bearing samples), or high-K CA; high Na/K (I-type; except some muscovite-bearing); corundum increases with Fe (I-type); moderate Ga/Al (slightly elevated) l-n) Tectonic Discrimination VA–PC/WP–(SC); highly fractionated VA–SC, or PC high-K calc-alkaline (to aluminous-alkaline) I-/S-type granitoids related to shearing Interpretation PC, syn-shear, high-K calc-alkaline (to aluminous-alkaline); >1 melt source, or heterogeneous melt source; mainly low- T I-type affinity, but significant supracrustal input from a feldspathic and clay-poor source may be responsible for this compositional signature (e.g., metagreywacke ± amphibolite/carbonate); high-T melt input from enriched underlying mantle or interlayered OIB-like mafic rocks is suggested for the alkaline hornblende-bearing phases; S-type melt source appears dominant Anvil Plutonic Suite (Yukon–Tanana Inlier and western Selwyn Basin) References for Compiled Geochemical Data Wood and Anderson (1982); Groat et al. (2002); Neufeld (2004) a) Harker Variation very small range of SiO2 (71–74 wt.%; low-T I- or S-type); tightly clustered major element variations that appear to decrease with increasing SiO2; Na2O shows opposite trend; Ce (allanite) decreases with increasing SiO2 (S-type); Zr may be highly enriched, but decreases with increasing SiO2; Cr and Co are variably elevated at Tsa Da Glisza (wallrock contamination from mafic metavolcanics: chlorite–plagioclase schist thought to be the source of Cr for associated emerald mineralization; Neufeld et al., 2004, 2005); Th decreases with increasing SiO2; NOT SHOWN: coupled Th–Ce–La ± P2O5 (monazite ± allanite fractionation); Sr is < 200 b) Fractionation Indices Rb/Sr typically < 2, but up to ~7; Eu/Eu* is strongly to moderately negative (0.15–0.44; average = 0.31) and decreases with increasing SiO2; LaCN/YbCN is moderate-high (11.38–37.86; average = 26.12, but typically > 30) and decreases with deeper negative Eu (not shown) c) Chondrite- and Primitive Mantle- normalized C: low overall REE (particularly HREE); PM: enriched Cs–Rb–U–Th–Pb, and depleted (Ba)–Nb–(Ta)–Sr–(P); very similar to Hyland plutonic suite, but has larger anomalies related greater fractionation or lesser entrained restite d) Average High-K Calc-Alkaline- Normalized generally depleted in all of the elements plotted (particularly those that partition into feldspar, biotite/K-feldspar, garnet/apatite, ilmenite, and possibly monazite) except for Rb–(Th)–U–Ta e-k) Geochemical Discrimination monzogranite; slightly to moderately peraluminous (A/CNK = 1.05–1.25); magnesian to ferroan; alkali-calcic/calc- alkaline, or high-K calc-alkaline; indeterminate Na/K; corundum decreases with Fe (S-type); high Ga/Al (A-type; all other indicators are S-type, probably due to fractionation or protolith composition) l-n) Tectonic Discrimination SC–(VA); highly fractionated SC, or PC high-K calc-alkaline S-type (with slight enrichment in SNY) granitoids related to shearing Interpretation low-T, minimum-melt, syn-shear S-type granites lacking entrained plagioclase-bearing restite (e.g., low-Sr leucogranites; Kemp and Hawksworth, 2003) Hyland Plutonic Suite (central–eastern Selwyn Basin) References for Compiled Geochemical Data Heffernan (2004) a) Harker Variation high SiO2 (66–73 wt.%; 2 intermediate dykes and 1 enclave are outliers ca. 57–59 wt.%) in all phases; major element trends are scattered (heteogeneous), although negative trends for Al2O3, TiO2, FeT, MgO, and CaO are present; Na2O shows the opposite trend; K2O initially decreases in ca. 58 wt.% SiO2 samples, then increases; P2O5, Y, and Zr have very a shallow negative slope that steepens ca. 68 wt.% SiO2 (I-type); BaO and Ce broadly increase to ca. 65 wt.% SiO2 then decrease (high-T I-type); Co increases with SiO2 (probably contamination); Th increases with SiO2, then decreases sharply at ~70 wt.% SiO2; NOT SHOWN: A/CNK decreases with increasing SiO2, Rb/Sr, and Eu/Eu* (S-type); tightly coupled Th–La–Ce (allanite fractionation) b) Fractionation Indices Rb/Sr typically < 1, but increases with decreasing peraluminosity (S-type); Eu/Eu* is slightly to moderately–strongly negative (0.30–0.84; average = 0.58) and decreases with increasing SiO2; LaCN/YbCN is low (5.38–18.70; average = 147 11.83), decreases with deeper negative Eu (not shown), and does not vary significantly with SiO2 c) Chondrite- and Primitive Mantle- normalized C: low overall REE; enclave and mafic dykes have smaller negative Eu anomaly; PM: Th–(Pb)–Sr are enriched, and (Rb)–(Ba)–Nb–(P)–Ti are depleted; more enriched in HREE than the Anvil suite; mafic dykes and biotite-rich enclave are very similar d) Average High-K Calc-Alkaline- Normalized Rb–Th–Nb–Sr are similar to the reference series, and all other elements are depleted; Ti is strongly depleted; the mafic enclave is very similar to the dykes (except for elevated Rb in the enclave), but compared to the average plutonic sample the enclave is particularly Rb–Th-rich and Nb–Sr–Ca-poor e-k) Geochemical Discrimination granodiorite to monzogranite; slightly metaluminous to strongly peraluminous; (A/CNK = 0.97–1.35); magnesian >> ferroan; calc-alkalic >> calcic, or med- to high-K calc-alkaline; high Na/K (I-type); corundum broadly increases with Fe, especially for the enclave (heterogeneous S-type); Ga/Al slightly anomalous for I-/S-type l-n) Tectonic Discrimination VA–(SC); highly fractionated VA–SC, or PC high-K calc-alkaline I-/S-type granitoids related to shearing Interpretation heterogeneous S-type source (e.g., immature plagioclase-rich sediments eroded from an arc) to mixed low-T S-/I-type source (e.g., interlayered quartzo-feldspathic metaseds and amphibolites typical of underlying NeoProterozoic and older? rocks) Seagull Plutonic Suite (western Cassiar Terrane) References for Compiled Geochemical Data Colpron et al. (2002); Liverton et al. (2005); Morris and Creaser (2008); Liverton (2011 pers. comm.); this study a) Harker Variation high SiO2 (70-80 wt.%; 2 cross-cutting lamprophyres ca. ~54 wt.%); major elements decrease with increasing SiO2; K2O and Na2O do not appear to vary systematically although K2O may decrease and Na2O may increase in the most siliceous samples; clusters of high Al2O3–Na2O and low TiO2–CaO–K2O–P2O5 are lepidolite topaz leucogranites (Liverton and Alderton, 1994; Liverton and Botelho, 2001); arguably lower Al2O3–CaO–MgO and higher FeT than other I-/S-type suites (e.g., Cassiar and Anvil suites) with equivalent SiO2 contents (A-type); trace elements display enriched vertical trends, particularly for Nb (not shown), Zr, Ce, Y, and Th (A-type); Cr and Co are very low and/or decrease with increasing SiO2; NOT SHOWN: poorly coupled Th–La–Ce (some allanite fractionation, but also possibly thorite or Th- oxides) b) Fractionation Indices Rb/Sr is anomalously high (i.e., > 2, but up to ~9500 for lepidolite topaz leucogranites); Eu/Eu* is moderate to VERY strongly negative (0.03-0.69, with one anomalous 1.15; average = 0.40) and decreases with increasing SiO2; LaCN/YbCN is VERY low (0.99–18.53; average = 7.69) and does not vary with SiO2 or Eu/Eu* (not shown) c) Chondrite- and Primitive Mantle- Normalized multi-element extreme enrichments and depletions; C: LREE, and for some samples, HREE are enriched; PM: Eu, Sr, Ba, P, Ti are VERY depleted, and Rb, Th are VERY enriched; Ta–(Nb) are not notably depleted d) Average High-K Calc-Alkaline- Normalized multi-element extremely depleted in Ba-Sr-Ti, moderately depleted in Ca-Na, and highly enriched in Ta–Nb in comparison to high-K calc-alkaline rocks examined here; mafic dykes and one felsic phase appear to be cumulates; heavier element (e.g., Zr to Yb) pattern very similar to that of late-shear, one-mica (biotite) alkaline-peralkaline rocks resulting from stress reversals along shear zones (Liégeois et al., 1998, and references therein) e-k) Geochemical Discrimination monzogranite >> granodiorite/tonalite (lamprophyres are quartz monzonite-monzodiorite); peralkaline (A-type) >> slightly peraluminous (A/CNK typically < 1.1); ferroan (A-type); alkalic-calcic to calcic with increasing SiO2, but generally high-K calc-alkaline (to shoshonitic); very high Ga/Al (A-type) l-n) Tectonic Discrimination PC–(VA) and WP–(SC) (A-type); relative to SNY vs. SNX: three clusters that sub-parallel the alkaline-peralkaline boundary from the shoshonitic field (PC) through the high-K calc-alkaline and fractionated I-/S-type field (h.-fr. VA–SC) into the calc-alkaline I-/S-type field (VA–SC) near the origin Interpretation peralkaline to aluminous A-type; associated high-K calc-alkaline to shoshonitic lamprophyres suggest mantle melt involvement, although these have not been dated and could be younger Tay River (TR) and South Fork (SF) (western Selwyn Basin) References for Compiled Geochemical Data Wood and Anderson (1982); Pigage and Anderson (1985); Morris and Creaser (2008); Gordey (in press); this study a) Harker Variation moderate to high SiO2 (62–74 wt.% in TR; 59–72 wt.% in SF: low-T); near-linear negative trends for most major elements (except Na2O, K2O, and BaO, which increase); a “chloritized biotite beta-granite” (Gordey, 2011, in press) is anomalously Al–Na–K-rich and Fe–Mg–Ca-poor; TiO2 (and P2O5 in SF) has a slight downward inflection ca. 66–67 wt.% SiO2; P2O5 (in TR), Ce (in TR), and Zr (ICP-muscovite) increase to ~67 wt.% SiO2, then decrease in TR (high-T I-type); Ce (in SF) and Th increase with SiO2 (although diffusely in TR; I-type); Y (TR) decreases to ~67 wt.% SiO2, then increases (I-type); Cr and Co are very low or decrease; NOT SHOWN: Nb is flat and slightly lower than in in the eastern SB for TR; Th–La–Ce are loosely coupled for TR (allanite fractionation), and are tightly linearly coupled over small compositional range (allanite fractionation?) b) Fractionation Indices Rb/Sr is < 2; Eu/Eu* is slight-moderately to moderate-strongly negative (0.33–0.79; average = 0.56) and generally decreases with SiO2 (some samples diverge); LaCN/YbCN is low (3.42–26.41; average = 13.40) and generally increases with deeper negative Eu (not shown) and SiO2 until ~67 wt.% SiO2 c) Chondrite- and Primitive Mantle- normalized C: slightly convex upward in HREE's for TR; PM: HFSE’s (Ta, Nb, P, Ti) and Sr-Ba-(Rb) are depleted; minimal TE and REE variation (essentially none in SF; very homogeneous) d) Average High-K Calc-Alkaline- Normalized very depleted Ca–Ti, variably small depletions in (Ta)–Nb–K (± REE), and slightly enriched U–Th; very similar to the TR-CR from the eastern SB, but the scale of the anomalies (e.g., Ca, Ti, REE) vary and particularly deep negative Ca is unique to these suites in the western SB e-k) Geochemical Discrimination granodiorite to monzogranite; slightly to moderately peraluminous (A/CNK = 1.0–1.2); magnesian; calcic to calc-alkalic, or high-K calc-alkaline; low Na/K in TR (S-type); corundum increases with decreasing Fe (I-type); low Ga/Al (I-/S-type) l-n) Tectonic Discrimination VA; highly fractionated VA–SC, or PC high-K calc-alkaline I-/S-type (with slight enrichment in SNY) granitoids related to shearing Interpretation low-T or high-T I-type chemistry with an unusually homogeneous and K-rich and Ca-poor protolith; associated dacitic volcanism and sub-volcanic intrusions suggest shallow emplacement and low initial H2O Tay River (TR), 'evolved', and Coal River (CR) (eastern Selwyn Basin) References for Compiled Geochemical Data Heffernan (2004); Bezzola (2009); Pigage et al. (2011, pers. comm.); this study 148 a) Harker Variation high SiO2 (60–77 wt.%; but combined with mafic dykes see an continuous range 53 wt.% SiO2); gently curvilinear negative trends for most major elements (P2O5: I-type); Na2O and K2O are more diffuse and increase with SiO2; BaO steadily decreases (low-T I-type); CR samples are depleted in K2O and BaO, and enriched in Na2O for their SiO2 content compared to other samples (notably, Na2O decreases with increasing SiO2); P2O5 and Y are decoupled; P2O5 may increase (and Y decreases) with increasing SiO2 in some 'evolved' samples (S-type); Ce, Zr (ICP- muscovite), and Th are scattered above ~67 wt.% SiO2 (low-T I-type); Ce–Th increase with SiO2 (I-type); Zr (ICP- muscovite) may increase (in CR samples) and then some samples suggest a general decrease (high-T I-type) with increasing SiO2; NOT SHOWN: Th–La–Ce generally coupled (allanite fractionation, ± monazite in 'evolved'), although P2O5 also for mafic dykes (apatite fractionation) b) Fractionation Indices Rb/Sr is < 2 except in 'evolved' (up to ~6); Eu/Eu* is shallow to very deep (0.02–0.85; average = 0.55 for TR and 'evolved' samples) and is shallowest for CR sub-suite (0.58–0.95; average = 0.76), then deepens substantially with increasing SiO2; LaCN/YbCN is low to moderate (4.25–55.72; average = 17.66), increases with SiO2 in mafic dykes and CR sub-suite, then varies but generally decreases with increasing SiO2 and with deeper negative Eu (not shown; S-type) c) Chondrite- and Primitive Mantle- normalized C: 'evolved samples are notably depleted in HREE, and mafic dykes lack notable Eu anomalies; PM: HFSE’s (Ta, Nb, P, Ti), and Sr and Ba are depleted, particularly in 'evolved', and mafic dykes and CR sub-suite are very similar but with smaller depletions/enrichments (notably, Cs–Rb–Th–U–La–Pb are less enriched, Ta is more depleted, and Sr–Ti are less depleted); overall though variation is not that significant, especially in CR samples (homogeneous) d) Average High-K Calc-Alkaline- Normalized TR, 'evolved', and CR samples have very similar patterns with negative Ca and Ti anomalies, although CR samples are typically more depleted overall (and more enriched in Sr); mafic dykes are enriched in U–Th (for their SiO2 content!) and Ba, and have a deeper Ti anomaly (also for their SiO2 content); overall very similar multi-element profiles to those exhibited by the Tay River suite in the western SB, but with larger negative Ti anomaly and slightly more variability in REE concentrations, and Ta is generally enriched e-k) Geochemical Discrimination quartz monzodiorite (CR and mafic dykes) to granodiorite to monzogranite; slightly metaluminous to moderately peraluminous (A/CNK = 0.98–1.25; highest in CR); magnesian; calc-alkalic, or medium-K (most CR) to high-K calc- alkaline; low Na/K (TR; S-type) to high Na/K (and Ca/Fe; not shown, for CR: I-type); corundum increases with decreasing Fe (I-type); minimal Ga/Al enrichment (I-/S-type), although Ce is higher in some samples l-n) Tectonic Discrimination VA–PC–(SC); highly fractionated VA–SC, or PC high-K calc-alkaline I-/S-type (with slight enrichment in SNY) granitoids related to shearing, with associated shoshonitic dykes Interpretation low-T or high-T I-type chemistry with an homogeneous and K ± Ca-rich protolith; more mafic members (CR sub-suite) have high Na/K and unusually high peraluminosity for their mineralogy and composition Tungsten (eastern Selwyn Basin) References for Compiled Geochemical Data Lang (2001); Rasmussen et al. (2011); this study a) Harker Variation high SiO2 (64–-76 wt.%; mafic dykes in SE SB ca. 53 wt.%); tight negative linear trends for most major elements (although mafic dykes ca. 53 wt.% SiO2 are relatively higher in MgO and lower in Al2O3); Na2O increases with SiO2; K2O is relative constant; in NE SB P2O5 and Y are relatively constant, but P2O5 decreases and Y increases in leucocratic samples; in SE SB P2O5 and BaO decrease with increasing SiO2 (low-T I-type) and Y is constant but has 2 anomalously depleted samples; all other trace elements decrease with increasing SiO2; in NE SB K2O, Th, and Nb (not shown) ± Ce are elevated and Al2O3 and Na2O are depleted relative to SE SB samples; NOT SHOWN: Rb < 500 in NE SB, but spread out relative to SiO2 (S-type), and Rb < 300 in SE SB (I-type); Th–La–Ce are tightly coupled, and loosely with P2O5 (monazite fractionation; apatite fractionation controls P2O5 in mafic dykes) b) Fractionation Indices Rb/Sr typically < 2, but up to ~6 (NE SB); shallow to very deep negative Eu/Eu* (0.15–0.72; average = 0.48; mafic dykes = ~0.9), and decreases with increasing SiO2; LaCN/YbCN is low-(moderate) (1.89–24.89; average = 14.22), decreases with increasing SiO2 and with deeper negative Eu (not shown) c) Chondrite- and Primitive Mantle- normalized C: steep REE slopes, with highly fractionated leucocratic samples from the northeast displaying more variability; PM: HFSE's (Nb–P–Ti ± Ta) and Ba are depleted, especially in leucogranites; minimal variation, except in highly fractionated leucocratic samples (homogeneous) d) Average High-K Calc-Alkaline- Normalized very similar to the Tay River and much of the Mayo (particularly HM) suites; samples from the northeast SB are more enriched in relatively incompatible elements (Rb, U, Th, Ta, Nb, Sm, Y, and Yb); lamprophyric dykes display pronounced U-Th and smaller Ba-Ce enrichments, and lower Ca-Ti relative to plutonic samples (as do the lamprophyre dykes associated with Mayo suite in the northeast SB, particularly with the HM dyke) e-k) Geochemical Discrimination monzogranite; slightly peraluminous (A/CNK = 1.0–1.1 in NE SB: minimum-T melt, but up to 1.2 in SE SB); magnesian; calc-alkaline, or very high-K calc-alkaline; lower Na/K in NE SB (S-type), and higher Na/K in SE SB (I-type); corundum decreases with Fe in NE SB (S-type), but increases with decreasing Fe in SE SB (I-type); low Ga/Al (possible S-type trend in NE SB) l-n) Tectonic Discrimination mainly SC–WP–(VA) or VA in NE SB, and VA–SC or VA–PC in SE SB; highly fractionated VA–SC, or PC high-K calc- alkaline I-/S-type (with slight enrichment in SNY) granitoids related to shearing, with associated shoshonitic dykes Interpretation S-type in NE SB, and low-T I-type in SE SB; geochemically very similar to Mayo NE suite (except for Th divergence), but require younger and/or more Pb-radiogenic source (e.g., relatively homogeneous Pb-rich metasediments in NE SB; homogeneous Pb-rich infracrustal source in SE SB) Mayo (northern Selwyn Basin) References for Compiled Geochemical Data Lang (2001); Morris and Creaser (2008); Mortensen (2009); Gordey (in press); this study a) Harker Variation low to high SiO2 (54–74 wt.%); negative linear trends for most major elements (often more scatter in mafic samples); Na2O and K2O (despite some Ksp-rich cumulates in NE SB) increase with SiO2; P2O5, BaO, and Y decrease with increasing SiO2 (high-T I-type), although BaO is scattered and anomalously enriched in NW SB (hypersolvus?); Ce, Zr (XRF), and Th (highly enriched in some samples from the NE SB) increase to ~67 wt.% SiO2, then decrease with increasing SiO2 (high-T I-type); Cr is slightly elevated but decreases with increasing SiO2; Co decreases with increasing SiO2, except mainly in NW SB and in half of HM sub-suite samples where it increases diffusely; NOT SHOWN: Rb < 300 regardless of P2O5 content (I-type); Th–La–Ce are coupled, and loosely with P2O5 (monazite fractionation; apatite fractionation controls P2O5 in mafic dykes and is more significant in NW SB samples) 149 b) Fractionation Indices Rb/Sr < 2; slightly positive (mainly NW SB) to moderately negative (mainly NE SB) Eu/Eu* (0.56–0.96; average = 0.81; rare 1.01–1.78), and depth of Eu/Eu* decreases with increasing SiO2 in NW SB (hypersolvus), but increases with SiO2 in NE SB (sub-solvus); LaCN/YbCN (2.15–34.03, one sample ~112; average = 19.11) is typically lowest in mafic dykes, decreases with shallower Eu/Eu* in the NW SB, and decreases or increases with increasing SiO2 c) Chondrite- and Primitive Mantle- normalized C: relatively steep REE slopes; PM: HFSE's (Nb–P–Ti ± Ta) have small depletions, Ba is typically positive, P–Ti depletions are more pronounced in samples with moderately negative Ba, generally LILE-rich and strong Cs–U–Pb enrichments relative to the other elements; mafic dykes and phases are very similar with only slightly more enriched heavy REE + Ba–Yb ± Th–U–P–Ti; HM sub-suite is generally less enriched d) Average High-K Calc-Alkaline- Normalized mainly a pattern typical of post-collisional shoshonitic magmatism: lighter elements (e.g., Th to Ba) are typically strongly enriched, heavier elements (e.g., Ti to Sm) are moderately enriched, and middle elements are roughly equivalent; northeast SB samples (including the HM sub-suite) are very similar to the Tungsten suite, but are typically more U–Th- rich and have variably Ta enrichments, are Nb-depleted, have bimodal but mostly depleted Ba-Sr, and some samples are middle REE ± K ± Ba enriched; northwest SB samples are U–Ta–Ba ± Sr rich and Ca-poor, but lack Nb- depletions, high Th, and the MREE-rich population; mafic dykes have larger positive and negative anomalies in general, but dykes from the northeastern SB are exceptionally enriched in U–Th–K e-k) Geochemical Discrimination mostly monzogranite, with 2 trends: dominant monzonitic (A-type) and subordinate dioritic (I-/S-type); metaluminous to moderately peraluminous (A/CNK = ~0.8–1.25), but 2 trends: metaluminous high and low A/NK converge ca. A/CNK = 1.1; magnesian; 2 trends: calcic to calc-alkaline and alkali-calcic to calc-alkaline, or high-K calc-alkaline to shoshonitic; low Na/K (S-type, or potassic A-type); corundum increases with decreasing Fe in SE SB (I-type); mainly low Ga/Al, but some Ce enrichment (I-/S-type) l-n) Tectonic Discrimination VA–SC–(WP), or VA (dominantly NE SB) to PC (dominantly NW SB); a continuous trend form shoshonitic (most mafic phases and late lamprophyre dykes) to PC high-K calc-alkaline I-/S-type (or highly fractionated syn-collisional I-/S-type) magmas Interpretation mainly high-T I-type, old (xenocrystic zircon ca. 1.6–1.8 Ga), and K-rich (e.g., felsic, LILE- and K-rich Precambrian lower crust); possibly contaminated by and mixed with shoshonitic magmas derived from an enriched mantle Tombstone (northern and eastern Selwyn Basin) References for Compiled Geochemical Data Duncan (1999); Lang (2001); Heffernan (2004); Morris and Creaser (2008); this study a) Harker Variation low to high SiO2 (54–74 wt.%); negative linear trends for most major elements (except for most K2O and Na2O samples) with large degree of scatter, particularly in more mafic samples; 2 main trends: a subordinate alkaline trend of mainly syenitic rocks (indicated by low SiO2 tinguaite end-members) of higher Al2O3, Na2O, and K2O and lower TiO2, FeT, MgO, CaO, P2O5, and BaO; and a dominant sub-alkaline trend (indicated by most of the shoshonitic mafic dykes and gabbroic end-members) with opposing enrichments and depletions; similarly scattered dual trends exhibited by Y, Ce, Zr, and Th (all highly enriched; A-type); some BaO enrichment in NW SB, but less pronounced than in Mayo suite; HM sub-suite is anomalously enriched in Y; Cr decreases with increasing SiO2, except in some Emerald Lake samples in the NE SB; Co increases with SiO2, except in some samples in the NE and SE SB; NOT SHOWN: Th–La–Ce loosely coupled with widely scattered P2O5 b) Fractionation Indices Rb/Sr < 2 (rarely up to ~7); possible positive, but mainly small to moderately deep Eu/Eu* (0.46–2.11; average = 0.78) that increases to ca. 63 wt.% SiO2, then decreases with SiO2 (particularly for southeast SB and HM sub-suite); relatively steep LaCN/YbCN (8.01–44.18; average = 23.18) that are not related to SiO2 content or the size of Eu/Eu* c) Chondrite- and Primitive Mantle- normalized C: generally steep REE slopes; PM: notably depleted Nb, P, Ti, Ba ± Sr, and enriched Cs, K, Rb, Th, U (generally more enriched than the Mayo suite): sub-alkalic metaluminous phases have negative Ba–Nb–Ta–P–Ti (and Sr in the southeast SB, tinguaite and alkalic phases are some of the most enriched (e.g., Lang, 2000; Coulson et al., 2002) and have some of the most pronounced anomalies (high Rb–Th–Ta–Nb–Pb–Zr–Hf and low Ba–P), and HM sub-suite is generally less enriched and more Eu- and Sr-poor (earlier plagioclase fractionation); mafic dykes and the least siliceous intrusions either have similar patterns with a similar range of enrichment and depletion to that exhibited by intermediate-felsic plutonic phases, or have much flatter profiles profiles with strong positive Ba–Pb anomalies; samples from the southeast SB are generally near the lower end of the range of enrichment observed d) Average High-K Calc-Alkaline- Normalized very similar to shoshonitic rocks (and the Mayo plutonic suite, however the light element enrichment is more pronounced, and the heavy element enrichment is absent or not significantly greater than that of the high-K calc- alkaline reference series); moderate variability between samples from the northwest and northeast SB, except that Ca- Ti are strongly depleted and Na is equivalent to the reference series; mafic intrusive phases in the northwest SB are very similar to more siliceous phases, but silica-undersaturated tinguaites have widely anomalous enrichments and depletions; samples from the southeast SB are notably different with slight light element enrichment, moderate middle element depletion lacking pronounced negative Ca–Ti, and typically slight heavy element depletion (unique and unlike any patterns from the other plutonic suites examined) e-k) Geochemical Discrimination mostly monzonite to monzogranite, 3 trends: syenitic (tinguaites; A-type ), monzonitic (shoshonites; A-type) and dioritic; peralkaline to moderately peraluminous (A-type), but 2 trends: high A/NK (A/CNK > 0.8) and low A/NK (A/CNK > 0.5) converge ca. A/CNK = 1.1, then A/NK increases steeply to A/CNK = 1.2 (NW SB); magnesian > ferroan; calcic to calc- alkaline (high-K calc-alkaline), alkali-calcic to calc-alkaline (shoshonites), and strongly alkalic to calc-alkaline (tinguaites); trend towards high Ga/Al and Ce (A-type) l-n) Tectonic Discrimination VA–SC–(WP), or VA–PC–(SC)–(WP); shoshonitic to high-K to medium-K calc-alkaline post-collisional magmatism, following a trend similar to that displayed by rapakivi topaz granites Interpretation a hybrid of peralkaline trend high-T A-type magmas contaminated by high-T I-type shoshonitic? melt (no xenocrystic zircon); I-type melt was probably LC (where A-type basalts pool and hybridize); high K/Na and LILE enrichment due to assimilation of felsic LC; in southeast SB I-type source was younger than rest of Tombstone-Mayo suites (1.8–1.9 Ga xenocrystic zircon) Rancheria (southeast Cassiar Terrane) References for Compiled Geochemical Data Gregory (2007); this study 150 a) Harker Variation high SiO2 (64–76 wt.%; 1 mafic dyke ca. 54wt.%); diffuse negative major element trends (except for Na2O, which has a shallow positive trend), although negative K2O trend is very shallow; more scatter for samples with < 70 wt.% SiO2; Al2O3 is relatively enriched, but MgO and CaO are relatively depleted (A-type) compared to other suites; Y decreases with increasing SiO2, although most siliceous samples may increase slightly to extremely; Ce, Zr, Th, and Nb (not shown) are elevated (A-type) and decrease with increasing SiO2; Cr is anomalously high and may increase with SiO2 in some samples; Co decreases with increasing SiO2; NOT SHOWN: Th–La–Ce are loosely coupled (allanite fractionation) except in mafic dyke (apatite fractionation) b) Fractionation Indices Rb/Sr increases diffusely with SiO2 up to ~ 3.5; moderate to deep Eu/Eu* (0.25–0.88; average = 0.49) that mainly decreases with increasing SiO2; LaCN/YbCN (0.48–50.54; average = 24.63) is steep, but decreases steeply with increasing SiO2, and does not vary with Eu/Eu* c) Chondrite- and Primitive Mantle- normalized C: typically steep REE profiles, but some samples are HREE (Allen Stock) enriched; PM: extremely enriched Th–U– HREE, but depleted Ba and Ti ± Sr, P, Zr, Hf, and Pb is low compared to all other suites; strongly negative Nb–Ta anomalies are not present d) Average High-K Calc-Alkaline- Normalized remarkably similar to Seagull suite, but more enriched in all elements; Rb to Nb are strongly enriched (especially U– Ta–Nb), Ca–Na–Ti are somewhat depleted, and a trend of increasing enrichment from Ce to Yb is obvious; mafic dyke has a similar pattern but with less anomalous variations and a generally decreasing (and depleted) trend for Ce to Yb) e-k) Geochemical Discrimination quartz monzonite to monzogranite (mafic dyke is quartz monzodioritic); slightly to moderately peraluminous (A/CNK = 1.04–1.25); magnesian; alkalic to alkali-calcic to calcic, or shoshonitic to high-K calc-alkaline; high Ga/Al and Ce (A- type) l-n) Tectonic Discrimination PC/WP; post-collisional shoshonitic magmatism with an evolutionary trend similar to that of the rapakivi topaz granites Interpretation aluminous A-type; no evidence for involvement of mafic melts (mafic dyke is frequently very different compositionally) Black River (eastern Cassiar Terrane) References for Compiled Geochemical Data Black River: this study; Ross Group: Neufeld (2004); Gordey (in press); Mount Reed: T. Liverton (2011, pers. comm.); McQuesten: Murphy (1997); Ting: Pigage (2009) a) Harker Variation no range in SiO2 for Black River suite (ca. 73 wt.%, including rhyolitic, Mt. Reed, and felsic McQuesten samples); intermediate McQuesten, Ting, and basaltic outliers range from ~47–63 wt.% SiO2; scattered but reasonably consistent, negative major element trends, although K2O and Na2O increase with SiO2 concentrations (Fig. 13). Minor and trace element plots are less consistent, although general trends are also evident. The majority of the samples, including the Black River suite, cluster ca. 74 wt.% SiO2 and are variably enriched relative to the less siliceous samples with the exceptions of Ba and Zn. b) Fractionation Indices Rb/Sr < 5 for Black River suite (rhyolites up to ~45); (moderate)-deep Eu/Eu* (0.22–0.31) that decreases with SiO2 with other siliceous suites (increases with SiO2 for other less siliceous samples); shallow LaCN/YbCN (10.56–11.21) that appears to decrease with SiO2 with other siliceous suites (increases with SiO2 for less siliceous samples) and with deeper Eu/Eu* for all suites c) Chondrite- and Primitive Mantle- normalized C: shallow slopes, moderate-large negative Eu anomalies, and generally elevated REE contents relative to other plutonic suites (closest to McQuesten and Ross suites, but lacking large Eu depletions); PM: strongly depleted Eu, Sr, Ba, P, and Ti; enriched in Rb and Th, moderately positive Pb, and only a small negative Nb anomaly (and no Ta anomaly) (most similar to Mt. Reed samples) d) Average High-K Calc-Alkaline- Normalized moderately to highly enriched in the light and heavy elements, but lack the massively depleted middle elements that characterize the Seagull plutonic suite and late-shear alkaline magmatism of the Tuareg Shield (Liégeois et al. (1998); most similar to McQuesten and Mt. Reed samples e-k) Geochemical Discrimination monzogranite (although all suites evolved along a dioritic trend); slightly peraluminous (A/CNK = 1.06); alkali-calcic to calc-alkaline, or high-K calc-alkaline; high Ga/Al and Ce (A-type) l-n) Tectonic Discrimination WP or PC; post-collisional aluminous-alkaline shoshonitic field, but sub-parallels its boundary with the post-orogenic alkaline-peralkaline field (similar to Seagull suite) Interpretation aluminous A-type; no evidence for involvement of mafic magma in the Black River suite; most similar to the Mt. Reed plutons and to the rhyolitic members of the bimodal (sub)volcanic rocks (that latter is likely to have fractionated from basaltic magmas) 151 3 CHAPTER: ISOTOPIC CHARACTERIZATION OF SYN- TO POST- ACCRETIONARY, EARLY CRETACEOUS TO EOCENE MAGMATISM IN THE CASSIAR TERRANE 3.1 INTRODUCTION The composition and the magma source(s) of syn- to post-accretionary magmatism in the Cassiar Terrane of the northern Cordilleran miogeocline, as well as the petrogenetic link between plutonism and mineralization, have been discussed by previous workers (e.g., Mortensen et al., 2000; Hart et al., 2002). Earlier interpretations were made mainly on the basis of geochemical arguments, albeit using limited data and (or) small sample sets, with poor geochronological control. Previous age constraints on syn- to post-accretionary igneous rocks in the northern Cordillera have relied heavily on K–Ar ages and Rb–Sr isochrons―no longer regarded as reliable―as well as 40Ar/39Ar datasets that may record cooling or re-heating ages rather than ages of magma emplacement. Furthermore, without accurate geochronological information, syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline could not be distinguished on the basis of geochemistry alone. The felsic to intermediate composition of most of this magmatism limits the diagnostic ability of geochemical discrimination diagrams (cf. Pearce et al., 1984). Unlike geochemical petrogenetic indicators, however, the radiogenic (Sm–Nd, Lu-Hf, Rb–Sr, Pb–Pb) and stable (S, O) isotopic systems are less strongly affected by anatectic and magmatic processes in the crust. As a result, the isotopic composition of a magma is largely inherited from is protolith rock(s). Few such isotopic datasets exist, however, for intrusive rocks in the Cassiar Terrane. Most of the available data are for the Rb–Sr isotopic system alone because it was formerly used for geochronological purposes. These previous isotopic studies of igneous rocks in the Cassiar Terrane (e.g., Driver et al., 2000; Morris and Creaser, 2008) have also been hampered by the poor age constraints, for two reasons. First, all ‘undeformed’ felsic plutonism in the Cassiar Terrane had been historically classified as the Cassiar plutonic suite (e.g., Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004a), which we now know to be an oversimplified interpretation (see Chapter 2). Second, radiogenic isotopic compositions should be calculated for the crystallization age of the igneous rock, particularly for the rapidly evolving Sm–Nd and Rb–Sr isotopic systems. Our understanding of the age and composition of syn- to post- accretionary magmatism in the northern Cordillera has, therefore, been largely restricted to simplistic interpretations that were limited by the lack of robust age constraints, the relatively felsic composition of the igneous rocks, and the rarity of radiogenic and stable isotopic datasets for plutonic rocks in this region. Several Early Cretaceous to Eocene plutonic suites in the Cassiar Terrane have now been identified and characterized temporally (using extensive new and compiled U–Pb ages in zircon) and geochemically (Chapter 2). The geochronological and geochemical classification of these magmatic suites provides the basis for a comprehensive radiogenic and stable isotopic study that includes whole rock Sm–Nd, Rb–Sr, and sulfur isotopic compositions, Pb isotopes from feldspar mineral separates, and oxygen isotopes from quartz separates. Together, these data not only help to better constrain the age, composition, and petrogenetic evolution of syn- to post- accretionary plutonism in the Cassiar Terrane, but also provide indirect information on the age and composition of 152 the unexposed basement to the Cassiar Terrane, if the magmas are sourced in the crust. 3.1.1 Isotopic Datasets as Petrogenetic Indicators Radiogenic isotopic compositions are the result of the radioactive decay of elements that are present in variable concentrations in the isotopic reservoirs that comprise the Earth’s atmosphere, crust, and mantle. There are two primary controls on the radiogenic isotopic composition of igneous rocks. First, the ratio of the parent to daughter elemental concentrations (e.g., Rb/Sr, Sm/Nd, U/Pb) in the source isotopic reservoir (i.e., magma source or protolith), which varies with the composition of the reservoir. Second, the amount of time that has passed since the extraction of the material from its original source (i.e., undepleted mantle). The combination of high parent:daughter ratios and prolonged periods of time cause the daughter isotope to become enriched relative to the parent isotope. The radiogenic isotopic composition of the protolith is, therefore, typically inherited by magma, providing information on the composition and age of the source rocks. Radiogenic isotopic compositions are expressed as the ratio of the radiogenic daughter isotope to a non-radiogenic isotope of the daughter element. Stable isotope variations result from their mass-related fractionation during chemical reactions, which can be significant for lighter elements (e.g., S, O, H, and C) at low or atmospheric temperatures (this fractionation is markedly less pronounced during higher temperature processes, including partial melting and magmatic differentiation; Taylor and Sheppard, 1986; Poulson et al., 1991; Wilson, 2007). Oxygen, which is abundant in all rocks, reflects the composition of the main melt source, unless a large amount of material with substantially different oxygen isotopic composition has been assimilated (as a solid) or mixed (as a magma or melt) with the parent magma. Conversely, the low sulfur concentrations of all magmas, and its relative volatility, makes sulfur isotopes a sensitive indicator of interactions between magmas and sulfur-rich rocks in the crust (e.g., organic-rich shales, evaporites, etc.; Poulson et al., 1991; DeYoung, 2007). The addition of a small amount of sulfur may dramatically affect sulfur isotopic composition, with little to no apparent change to oxygen and radiogenic isotopic compositions, in a magma (Poulson et al., 1991). Low sulfur abundances in the mantle (typically 20–50 ppm in peridotite, but up to ~140 ppm in some pyroxenites; Ionov et al., 1992) means that mantle-derived magmas will be strongly affected by the addition of sulfur from sulfur-rich crustal reservoirs. Sulfur isotopic compositions of igneous rocks have been under-appreciated in the past, and few studies of igneous rocks have combined sulfur data with the other isotopic systems to better understand the source of magmatic sulfur (e.g., DeYoung, 2007). Stable isotopic datasets nicely complement radiogenic isotopic datasets by providing additional, and in some cases more specific, information on the composition of melt source(s) and contaminants. Used separately or together, however, radiogenic and stable isotopic datasets are unparalleled tools for studying the petrogenesis of igneous rocks. 3.2 GEOLOGICAL BACKGROUND 3.2.1 Evolution and Composition of Northern Cordilleran Basement . Of vital importance to the study of magmas and their sources, particularly in regions where magmatism is thought to 153 be mainly crustally derived, is knowledge of the age and composition of basement rocks. The following summarizes what is known about the tectonostratigraphic evolution of the northern Cordilleran miogeocline and its basement. The tectonic evolution of the northern Cordillera involves a long history of extensional, contractional, and accretionary events that began as early as the early Mesoproterozoic and lasted through to the early Cenozoic (although tectonism continues today in parts of the northern Cordillera; e.g., Hyndman et al., 2005a). This past complexity is compounded by large, Early Cretaceous to Cenozoic, dextral offsets of several 100’s of km along strike-slip faults, which are responsible for the displacement of terranes and continental fragments (e.g., the Cassiar Terrane) from their original position along the western continental margin to North America. For the purposes of this study, the “Cassiar Terrane” is considered to be a post-Mesozoic feature that is located between the Teslin and Tintina faults (Fig. 3.1a). It is defined here as comprising three identifiable tectonostratigraphic and pericratonic entities underlain by Laurentian basement, including: the Cassiar Platform; eastern portions of the pericratonic terranes (located southwest of the Tintina Fault); and, the Klinkit Assemblage between the Cassiar Platform and Yukon–Tanana terrane (Fig. 3.1a). Rocks of oceanic affinity that structurally overlie portions of the Cassiar Terrane are thin klippen of the Slide Mountain terrane (Fig. 3.1a), a Mississippian ocean basin that separated the pericratonic terranes from the western margin of Laurentia. When Cretaceous to Cenozoic dextral displacements along regional structures are restored, it is apparent that the Cassiar Platform was located south of the Selwyn Basin, and west of the MacDonald Platform and Kechika Trough in Early Cretaceous time, which was approximately 430 km south of its current location in south–central Yukon (Fig. 3.2a; Gabrielse et al., 2006). 3.2.1.1 Precambrian Basement to the Cassiar Terrane The composition and western extent of North American basement is inferred from present-day reflection and refraction studies. The most widely accepted model for the formation of the basement to the northern Cordillera is that thick packages of late Paleo- to late Mesoproterozoic sedimentary rocks were deposited onto crystalline basement along the western margin of North America, probably during multiple continental rifting and (or) epicratonic extension events (e.g., Cook et al., 2004; Evenchick et al., 2005; Jones et al., 2005; Snyder et al., 2009). In south–central Yukon, these metasedimentary rocks and underlying crystalline basement are inferred on the basis of geophysical studies to thin to the west and terminate about halfway between the Teslin Tectonic Zone and the western margin of the Cassiar Platform (Fig. 3.1). The isotopic composition of Quaternary to Recent volcanic rocks and entrained mantle xenoliths corroborates the interpreted termination of Laurentian lithosphere at a position between the Tintina Fault and Teslin Tectonic Zone in south–central Yukon (Creaser et al., 1997; Abraham et al., 2001, 2005; Snyder et al., 2009). In northern British Columbia, Laurentian lithosphere and basement is thought to terminate near the Kutcho Fault (Fig. 3.1; Cook et al., 2004), although they may extend another 100 km farther west (Evenchick et al., 2005). The crystalline basement that is hypothesized to underlie the northern Cordilleran miogeocline is referred to as the Nahanni terrane (Fig. 3.2b; e.g., Hoffman, 1989), although it is of uncertain age and composition. Although the Nahanni terrane likely comprises multiple lithologies of varying ages, it is usually described as attenuated, 2.1–1.8 154 Ga North American crust (e.g., Jefferson and Parrish, 1989; Parrish, 1991; Villeneuve et al., 1991; Norris and Dyke, 1997; Cook et al., 1999). Possible components of this extended cratonic crust include: • the ~1.85 Ga Fort Simpson terrane, located immediately east of the Nahanni terrane (Fig. 3.2b; Cook et al., 1999; Aspler et al., 2003; Snyder et al., 2009); • analogs to the Paleoproterozoic basement domains identified further east in Alberta, such as the Talton, Buffalo Head, Chinchaga, Ksituan, and Nova domains (Fig. 3.2b; Thériault and Ross, 1991; Villeneuve et al., 1991; Ross et al., 1991; Ross, 2002); • Cordilleran gneisses similar to those exhumed in southeast British Columbia, such as the Malton, Monashee, and Frenchman Cap complexes (Vernon–Chamberlain, 1983; Armstrong et al, 1991; McDonough and Parrish, 1991; Parkinson, 1991). • a unique accreted terrane (possibly as young as ~1.0 Ga; Aspler et al., 2003); or • oceanic to transitional oceanic-continental crust formed during Proterozoic rifting events (Snyder et al., 2009). An additional possibility is that the lower crust that originally underlay the northern Cordillera is no longer present (or is not geophysically detectable) and overlying Proterozoic metasedimentary rocks now extend to MOHO depth of 33–36 km (Clowes et al., 2005; Snyder et al., 2009). All of the above possibilities are consistent with the low gravity (Cook et al., 1991, 1999) and magnetic (the “Nahanni low”; Pilkington et al., 2000; Aspler et al., 2003) geophysical signatures that characterize this enigmatic basement domain. The consensus, however, is that thinned Paleoproterozoic cratonic basement of Laurentian origin underlies the northern Cordilleran miogeocline. Major basin and (or) rift-forming events occurred at several times in the Proterozoic and led to the deposition of thick packages of supracrustal rocks in the northern Cordillera (Figs. 3.1b and 3.2a; e.g., Delaney, 1981; Aitken and McMechan, 1991; Gabrielse and Campbell, 1991; Rainbird et al., 1997; Ross et al., 2001; Cook et al., 2004; Evenchick et al., 2005; Thorkelson et al., 2005; Leslie, 2009; Furlanetto et al., 209; Medig et al., 2012). These include the: • <1.8 Ga Mesoproterozoic Muskwa Assemblage; • 1.72–1.2 Ga Wernecke Supergroup (~13 km thick); • <1.4–1.0 Ga Pinguicula Group (~3.5 km thick); • 1.2(0.88?)–0.78 Ga Mackenzie Mountain Supergroup (5–7+ km thick); and • 0.78–0.54 Ga Windermere Supergroup (composed of the 3–4.5+ km thick Ingenika Group in the Cassiar Platform). In the Cassiar Terrane, the Neoproterozoic Windermere Supergroup are the oldest exposed rocks (Figs. 3.1b and 3.2a), but are only inferred on the basis of seismic imaging and surface to extend to depths of 5–7 km below the current surface. Latest Paleo- to Mesoproterozoic metasedimentary rocks cropping out near the northern and eastern boundaries of the northern Cordillera (listed above), or their distal equivalents, are thought to underlie the widely exposed upper crustal Neoproterozoic rocks (Figs. 3.1b and 3.2a; e.g., Gordey and Makepeace, 2001; Cook et al., 2004; Clowes et al., 2005; Evenchick et al., 2005). A tapering wedge of geophysically ‘slow’ basement is defined by west-dipping reflectors (possible stratigraphic contacts) and forms the bulk of the middle to lower crust in the northern Cordillera (Cook et al., 2004; Clowes et al., 2005). This wedge has been linked to surface exposures of the 155 older metasedimentary rocks farther to the east is, therefore, interpreted to be of mainly metasedimentary origin. Multiple episodes of mainly rift-related mafic sills, dykes, and flows have intruded or were extruded onto Proterozoic clastic rocks (Figs 3.1; e.g., Souther, 1991; Dudàs and Lustwerk, 1997; Thorkelson et al., 2001, 2005; Schwab et al., 2004), and include the: • 1.71 Ga Bonnet Plume River intrusions; • 1.38 Hart River sills; • 1.27 Ga Bear River dykes (related to the Mackenzie Large Igneous Province; Schwab et al., 2004); and • 0.78 Ga Little Dal basalt flows and Tzetzotene sills (the Gunbarrel mafic magmatic event; Harlan et al., 2003). Basaltic rocks range in composition from potassic tholeiite to alkali basalt, both sourced in undepleted mantle. Buried sills and flows are another possible source for the high reflectivity in the west-tapering wedge of geophysically ‘slow’ metasedimentary rocks in the middle to lower crust (e.g., Evenchick et al., 2005). It is also conceivable that under- and intra-plating of basaltic magmas in the lower crust during these magmatic events has occurred locally. For example, thin, geophysically ‘fast’ lower crust could represent mafic underplating in both east– central Alaska (Beaudoin et al., 1992a) and below the Foreland Belt in northeast British Columbia (Welford et al., 2001). Geophysical surveys across the Cassiar Terrane, however, have not detected evidence for a preserved mafic lower crust (Cook et al., 2004; Hammer and Clowes, 2004; Ledo et al., 2004; Creaser and Spence, 2005; Evenchick et al., 2005; Fernández-Viejo et al., 2005; Jones et al., 2005). Overall, Mesoproterozoic siliciclastic packages(or their basin-ward equivalents) are thought to comprise most of the lower to middle Laurentian crust—which in turn probably overlie felsic to intermediate crystalline basement of late Paleoproterozoic age—below the Cassiar Terrane (e.g., Aitken and McMechan, 1991; Gabrielse and Campbell, 1991; Souther, 1991; Woodsworth et al., 1991; Cook et al., 2004; Evenchick et al., 2005; Snyder et al., 2009). 3.2.1.2 The Assembly of the Cassiar Terrane The rifting event along the western margin of Laurentia that triggered the deposition of the Windermere Supergroup also controlled the geometry of the rifted margin and the distribution of basinal and platformal miogeoclinal rocks. An upper plate rifted margin, typified by a sub-crustal detachment and steeply dipping normal faults that minimize crustal thinning during the rifting event (see Fig. 1.3, Chapter 1) was formed in the vicinity of the Cassiar Terrane (Hansen et al., 1993; Cecile et al., 1997). This contrasts with the “lower plate” rifted margin geometry that characterizes the Selwyn Basin, originally located north of the Cassiar Platform along the western margin of North America (Fig. 3.2a). Passive margin or miogeoclinal sedimentation on the upper plate rifted margin began in the earliest Paleozoic and led to the deposition of platformal (continental shelf) and basinal rocks. The Cassiar Platform (together with its northern extension, the McEvoy Platform; Figs. 3.1 and 3.2a) was deposited as a relatively large offshelf platform separated from the extensive MacDonald Platform by an apparently narrow basin (the Kechika Trough; Figs. 3.1a and 3.2a; e.g., Gordey and Anderson, 1993). Intermittent extension and subsidence in the Paleozoic led to the extrusion of sub-alkaline, OIB-like, basaltic magmas with geochemical affinities to ocean island 156 basalts (OIB) near basin–platform boundaries (e.g., the Cambro-Ordovician Groundhog and Looncry volcanic occurrences; Figs. 3.1b and 3.2a; Souther, 1991; Woodsworth et al., 1991; Goodfellow et al., 1995). Miogeoclinal sedimentation was interrupted in the Late Devonian to Early Mississippian by the deposition of locally derived turbiditic and cherty detritus. Back-arc extension and exhumation west of the Cassiar Platform (as the western portion of the Laurentian margin rifted from North America to form what are now known as the pericratonic terranes) led to the erosion and reworking of underlying strata (e.g., Gordey and Anderson, 1993; Creaser et al., 1999; Erdmer et al., 2005; Nelson et al., 2006; Murphy et al., 2006; Nelson and Colpron, 2007; Leslie, 2009; Beranek et al., 2010b). At this time, alkaline back-arc magmatism, exhalative barite and base metal mineralization, and block faulting were also occurring west of the Cassiar Platform (Mortensen, 1982; Gordey et al., 1987; Gordey and Anderson, 1993; Hunt, 2002; Erdmer et al., 2005; Mortensen et al., 2006a; Murphy et al., 2006; Gordey, 2012, in press). Basinal (east-derived) and platformal sedimentation in the Cassiar Terrane was re-established in the Carboniferous and persisted until at least the Triassic (e.g., Dusel–Bacon et al., 2006; Murphy et al., 2006; Nelson and Colpron, 2007). The Middle to Late Permian closure of the Slide Mountain ocean basin (Beranek and Mortensen, 2011) led to the collision and obduction of the pericratonic, island arc, Yukon–Tanana and Quesnellia terranes onto the continental margin. The Cassiar Platform was over- and under-thrust by a ~10 km-thick package of penetratively deformed the Yukon–Tanana terrane and Slide Mountain terrane rocks (Gabrielse, 1991; Gordey, in press). This was followed by the obduction and (or) tectonic wedging of the Cache Creek oceanic terrane and the pericratonic island arc terrane of Stikinia by mid-Jurassic time. Of the accreted terranes, only Stikinia appears to comprise a complete lithospheric block (e.g., Abraham et al., 2001; Clowes et al., 2005). Crustal scale seismic studies of the Yukon–Tanana (and Quesnellia) terrane near the western margin of the Cassiar Terrane indicate that these terranes form thin, mid- to upper crustal blocks that were structurally transposed onto North American basement (e.g., Cook et al., 2004; Snyder et al., 2009) and lithosphere (Abraham et al., 2001). Miogeoclinal sedimentation in the Cassiar Terrane ceased with mid- to Late Jurassic crustal thickening, uplift, and metamorphism resulting from one or more accretionary events (Gordey and Anderson, 1993; Murphy, 1997; Beranek and Mortensen, 2006; Nelson and Colpron, 2007; Beranek, 2009). During mainly mid-Jurassic contraction, miogeoclinal rocks were deformed into a northeastward-migrating, thin- skinned fold and thrust belt (Gordey and Anderson, 1993). Although obduction and tectonic wedging of pericratonic and oceanic terranes caused local, (south)west-verging structures (Price, 1986; Gabrielse et al., 1991), the majority of the observed 30–50% Mesozoic shortening in the Cassiar Terrane appears to have been accommodated by (north)east-verging imbrication on broadly folded thrust sheets with a shared upper crustal detachment (Gabrielse, 1991; Gabrielse and Yorath, 1991; Gordey and Anderson, 1993; Gordey, in press). Miogeoclinal rocks of mainly Paleozoic age are not thought to be structurally duplicated in the northern Cordillera based on mapping (e.g., Gordey and Anderson, 1993; Gordey, in press) and geophysical (e.g., Cook et al., 2004) studies. Rift-related structures of Proterozoic age―including transform faults, basin-bounding faults, and normal faults―may also have been reversely activated during compressional deformation (e.g., Norris, 1985; Gabrielse, 1991; Gordey and Anderson, 157 1993). The Gravina basin that separated the Insular Superterrane from Laurentia, may have closed sometime between the Late Jurassic to Early Cretaceous (e.g., Trop and Ridgway, 2007), but possibly as recently as the Late Cretaceous (e.g., Cole et al., 1999; Mezger et al., 2001). This accretionary event appears to have occurred at different times along the northern Cordilleran margin, although accretion seems to have migrated from south to north over time. An initial contact point in south–central Alaska (today’s coordinates; Trop and Ridgway, 2007) would have been outboard of the Cassiar Terrane in the Early Cretaceous if estimates of ~390 km of Cenozoic dextral displacement on the Denali fault are correct (Gabrielse et al., 2006). Although the timing of the accretion of the Insular Superterrane is poorly constrained, there are several possible times that it may have occurred. It may have caused a 10 m.y. magmatic lull in the northern Cordillera from ~135–125 Ma, which preceded a widespread, Early Cretaceous, syn- to post-accretionary magmatic event related to northeast-dipping subduction along the continental margin (e.g., Anderson, 1988; Woodsworth et al., 1991; Mihalnyuk, 1999; Hart et al., 2004a). Alternatively, a marked change in the composition of magmatism occurred at ~103 Ma in the Cassiar Terrane with the emplacement of the post-arc, A-type Seagull plutonic suite (and the termination of continental arc magmatism related to subduction farther west). This A-type plutonic suite is compositionally typical of post-collisional suites elsewhere in the world (Liégeois et al., 1998; Bonin et al., 1998) and suggests that an accretionary event (e.g., the Insular Superterrane) occurred prior to, or at, ~103 Ma. 3.2.2 Regional and Contact Metamorphism in the Cassiar Terrane Regional metamorphic grades in miogeoclinal rocks of the Cassiar Terrane are typically of greenschist facies (chlorite- to garnet zones in Proterozoic strata) to sub-greenschist facies (in Paleozoic and younger strata; e.g., Greenwood et al., 1991; Gordey and Anderson, 1993). However, there are some areas of the Cassiar Terrane, particularly in northern British Columbia, where uplifted Neoproterozoic rocks of the Windermere Supergroup display amphibolite metamorphic facies (e.g., the Horseranch, Sifton, and Swannell ranges; Fig. 3.2a; Evenchick 1988; Plint et al. 1992; Evenchick et al., 2005). Typically these high metamorphic grades occur in the cores of regional anticlinoria, or were exhumed along regional strike-slip structures that were active in the Eocene (Greenwood et al., 1991). In the Horseranch and Sifton ranges in northern British Columbia, the highest-grade Neoproterozoic rocks experienced Jurassic to Cretaceous peak metamorphism at amphibolite facies conditions (5–7 kbar and 500–650oC, roughly equivalent to 18–25 km of burial; Greenwood et al., 1991; Plint et al. 1992). Higher metamorphic grades are also present in contact aureoles associated with most of the plutonic rocks. Typically these higher grades are manifested in the form of biotite and andalusite, sillimanite, or cordierite-bearing mineral assemblages (e.g., Greenwood et al., 1991; Gordey and Anderson, 1993). The contact aureoles surrounding several Early Cretaceous batholiths were metamorphosed at amphibolite facies conditions (e.g., Greenwood et al., 1991). Some of the later, northwest-trending batholiths in the Cassiar Terrane, for example, have contact aureoles that are zoned over several km, from garnet-, to staurolite-, to andalusite-, to sillimanite ± cordierite-bearing schist, gneiss, and (or) migmatite, with increasing proximity to intrusive contacts (e.g., the Big Salmon–Nisutlin, Glenlyon, and 158 Cassiar batholiths; Greenwood et al., 1991; Mortensen and Hansen, 1992; Black et al., 2003). These amphibolite- grade and locally migmatitized contact aureoles are probably the result of heating during to the emplacement of these large magma bodies (e.g., Mortensen and Hansen, 1992). Alternatively, the migmatitic ‘halos’ may have been produced from the surrounding metasedimentary rocks by muscovite dehydration reactions, which produce lower temperature, smaller volume, and relatively immobile melts compared to biotite dehydration reactions (e.g., Patiño Douce et al., 1990). Migmatitic rocks are not necessarily cogenetic to voluminous granitic magmatism, which is derived at greater depths and higher crustal temperatures in rocks where biotite is an important reactant phase (e.g., Patiño Douce et al., 1990). 3.2.2.1 Burial Depths and Estimates of Cretaceous Crustal Thickness in the Cassiar Terrane The amphibolite-grade rocks in the Cassiar Terrane are either confined to the immediate contact aureoles to batholithic intrusions or to Neoproterozoic rocks that were locally uplifted along faults (Greenwood et al., 1991). Therefore, pressure–temperature estimates determined from the more widespread greenschist facies rocks are more representative of regional burial depths. Regional sub-greenschist to greenschist metamorphic grades prevail at the present surface in the Cassiar Terrane (Greenwood et al., 1991; Plint et al., 1992). Mineral assemblages from andalusite-bearing enclaves in the Cassiar batholith, and greenschist to sub-greenschist host rocks suggest that rocks at the Earth’s surface today were metamorphosed at no more than 4 kbar, corresponding to ~15 km depth (Driver et al., 2000). Given the estimated current depth to the MOHO of ~33–36 km below the northern Cordilleran miogeocline (e.g., Clowes et al., 2005), a maximum crustal thickness of 50 km would have been attained prior to (although not necessarily retained during) the Cretaceous magmatic event in the Cassiar Terrane. Conversely, exhumation of adjacent portions of the Yukon–Tanana terrane likely occurred in the Early Cretaceous (ca. 114–107 Ma; Mortensen, 1990; Murphy, 2004; Berman et al., 2007), which is coincident with Early Cretaceous plutonism throughout the northern Cordillera. Metamorphic grades in these exhumed rocks suggest that the Yukon– Tanana terrane may have attained a Jurassic crustal thickness of about 60 km, and maintained it for over 45 m.y. prior to Early Cretaceous exhumation, gravitational collapse, and syn- to post-accretionary magmatism. This estimate of crustal thickness and its duration only applies to the westernmost Cassiar Terrane, which is underlain by penetratively deformed and highly metamorphosed Yukon–Tanana terrane rocks. 3.2.3 Syn- to Post-Accretionary Magmatism At about 125 Ma, a widespread syn- to post-accretionary magmatic event was initiated, leading to the most voluminous and metallogenically important magmatism to have affected the northern Cordillera (Hart et al., 2004a). Early Cretaceous magmas are thought to be related to northeast-dipping subduction of an oceanic plate along the northwest Laurentian margin at this time, resulting from processes that occurred in both ‘outboard’ (i.e., margin- proximal) and ‘inboard (i.e., margin-distal) regions (e.g., Anderson, 1988; Woodsworth et al., 1991; Mihalnyuk, 1999; Hart et al., 2004a). Changing plate dynamics and the interruption of this subduction, however, produced more than one magmatic lull through the Cretaceous to the Eocene, and also led to magmatism that may not have been 159 specifically related to subduction at the continental margin (e.g., Nelson and Colpron, 2007). The following sections summarize the ‘outboard’ and ‘inboard’ magmatic suites in southern Yukon. 3.2.3.1 Magmatism ‘Outboard’ (west) of the Cassiar Terrane Early Cretaceous to Eocene magmatism located west of the Cassiar Terrane can be sub-divided into three main episodes, each related to tectonism along the continental margin and specific deformational regimes within the northern Cordilleran miogeocline. These are: the Early to ‘mid’-Cretaceous, the Late Cretaceous, and the latest Cretaceous to Eocene (Fig. 3.3). West of the Cassiar Terrane, in an outboard position, these igneous rocks are thought to be related to northeast-dipping subduction along the newly amalgamated western margin of the northern Cordillera (e.g., Mortensen et al., 2000; Hart et al., 2004a). The angle of convergence and rate of subduction, however, likely differed for each of the three magmatic episodes (e.g., Kelley, 1993; Nelson and Colpron, 2007). Inferred continental magmatic arcs are represented by: • 124–115 Ma Teslin plutonic suite in the westernmost Cassiar Terrane, and the 115–104 Ma Whitehorse plutonic suite and Mount Nansen volcanic suite in Yukon–Tanana terrane (Fig. 3.3a; Souther, 1991; Woodsworth et al., 1991; Hart, 1995, 1997; Gordey et al., 1998; Mihalnyuk, 1999; Mortensen et al., 2000; Hart et al., 2004a; Morris and Creaser, 2008), emplaced into a transpressive setting; • the 78–67 Ma Carmacks Group volcanic rocks and associated plutonism (e.g., the Windy Table and Prospector Mountain suites; Fig. 3.3b; Woodsworth et al., 1991; Souther, 1991; Selby and Creaser, 2001; Mortensen and Hart, 2010, unpub. data), emplaced into a transitional transpressive–transtensional setting; and • the circa 59–55 Ma Sloko–Skukum Group volcanic rocks and particularly voluminous, mainly 70–55 Ma plutonism of the Coast Plutonic Complex (Fig. 3.3b; Souther, 1971; Lambert, 1974; Smith, 1983; Gehrels et al., 1991; Souther, 1991; Currie, 1994; Mihalynuk et al., 1995; Mihalnyuk, 1999; Morris and Creaser, 2003; Simmons et al., 2006, 2007; S. Israel, 2012, pers. comm.) emplaced into a mainly transtensional setting. Although most researchers consider these episodes to reflect the development of three, I-type continental magmatic arcs related to subduction (e.g., Souther, 1991), other explanations for the latter two magmatic episodes include: lithospheric delamination, particularly for the effusive volcanism associated with the youngest Carmacks Group and Prospector Mountain rocks (e.g., Mortensen and Hart, 2010); and, subduction of one or more slab windows related to the Sloko-Skukum and some or all of the coeval Coast Plutonic Complex rocks (e.g., Madsen et al., 2006). Relatively juvenile initial isotopic compositions (εNdi ≈ +5.2 to –5.6 and Sri ≈ 0.701–0.710; Samson et al, 1991; Selby et al., 1999, 2001;Morris and Creaser, 2008) for these three magmatic events are consistent with a derivation for from mainly mantle and (or) juvenile mafic crustal rocks. Also associated, or coeval, with the arcs are two A-type plutonic suites west of the Cassiar Terrane (Fig. 3.3b): • the 84–78 Ma (U–Pb ages) Surprise Lake suite, exemplified by the Surprise Lake batholith17 and including 17 Although younger, the ~84 Ma Surprise Lake batholith is remarkably similar in terms of texture, mineralogy, chemistry, and metallogenic association to the 103–94 Ma Seagull plutonic suite (Chapter 2). 160 several intrusions forming a west–northwest-trending, 240 km-long belt associated with local graben structures (Fig. 3.3b; Anderson, 1988; Souther, 1991; Woodsworth et al., 1991; Mihalnyuk, 1999; Ray et al., 2000; Morris and Creaser, 2008; Smith, 2009); and • the circa 62–52 Ma (K–Ar ages) Nisling Range suite, exemplified by the Pattison pluton in the northern Coast Plutonic Complex (Fig. 3.3b; Ridgeway, 1973; LeCouteur and Tempelman–Kluit, 1976; Tempelman–Kluit, 1977a,b, 1979; Lynch and Pride, 1984; Anderson, 1988; Woodsworth et al., 1991; Souther, 1991). Anderson (1988) classified these two plutonic suites as “A-type” on the basis of textural, mineralogical, and geochemical criteria18. Both the Surprise Lake and Nisling Range suites have low initial strontium ratios (0.704– 0.708; Le Couteur and Tempelman–Kluit, 1976; Sinclair, 1995) that are consistent with a juvenile melt source. These suites appear to have been emplaced during the tectonically transitional Late Cretaceous and the increasingly transtensional Cenozoic, respectively. Their timing may also coincide with the accretion of exotic Insular and (or) Arctic terranes and the associated plate reorganization (e.g., Anderson, 1988). 3.2.3.2 ‘Inboard’ Magmatism within the Cassiar Terrane This study focuses on syn- to post-accretionary magmatism that occurred in a more ‘inboard’, or easterly, position relative to the continental magmatic arc(s). A large volume of intrusive rocks was emplaced into the Cassiar Terrane (Fig. 3.3). Prior to this study, available geochronological and compositional constraints on the igneous rocks in this region were limited, which had resulted in the inclusion of all apparently ‘post-tectonic’ granitoids in this region into the ‘mid’-Cretaceous Cassiar plutonic suite (e.g., Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004a). However, it is now clear that syn- to post-accretionary magmatism in the Cassiar Terrane comprises at least three distinct plutonic episodes, each separated by a lull: Early to ‘mid’-Cretaceous, which may be subdivided into Early (115–104 Ma) and ‘mid’- (103–94 Ma) Cretaceous; Late Cretaceous (82–75 Ma); and, Eocene (49–47 Ma). The composition and locus of this magmatism varied in the Cassiar Terrane, but the most pronounced change occurred at ~103 Ma with end of widespread voluminous plutonism and the onset of more localized and mainly small plutons associated with structural discontinuities. Using an extensive dataset of new (this study) and compiled (Appendices G.1 and G.2) geochronological and geochemical data four temporally, spatially, and geochemically constrained plutonic suites have been described in the Cassiar Terrane (Chapter 2). The main features characterizing these plutonic suites are summarized below from Chapter 2, although inferences regarding protolith compositions are based only on geochemical and mineralogical arguments (and not the new and compiled isotopic datasets presented later in this Chapter). The Cassiar plutonic suite of 115–104 Ma, high-K calc-alkaline, S > I-type granitoids comprises biotite ± 18 E.g., their reduced and highly siliceous compositions, apparent shallow levels of emplacement (indicated by abundant miarolitic cavities), the presence of magmatic fluorite and (or) high fluorine concentrations, and elevated rare earth elements (REEs) and other incompatible elements (e.g., Ce, La, U, Y, Yb, Nb, Rb in the Surprise Lake batholith; Ray et al., 2000) 161 muscovite- or hornblende-bearing batholiths and plutons with amphibolite-grade contact aureoles (Figs. 3.3a and 3.4). This suite was emplaced across most of the Cassiar Terrane (Fig. 3.3a), and many of these intrusions are intimately associated with Early Cretaceous movement along transpressional dextral structures (Gabrielse et al., 2006; Colpron et al., 2011). With proximity to strike-slip structures, for example, local highly strained to mylonitic textures become evident in some plutons (e.g., Cassiar batholith; Fig. 3.4); away from strike-slip structures, these intrusions are generally massive-textured with only minor foliation observed along plutonic contacts. The Cassiar plutonic suite was emplaced in a transpressive back-arc to the coeval Whitehorse continental arc. Reduced, felsic, biotite ± muscovite monzogranite intrusions dominate the Cassiar plutonic suite and have calc-alkaline compositions (cf. Frost et al., 2001). Oxidized, intermediate to felsic, biotite ± hornblende granodioritic to monzogranitic phases are less abundant and follow an alkaline to alkali-calcic evolutionary trend (cf. Frost et al., 2001). Associated mafic intrusive phases, including quartz monzodiorite to monzonite, and “lamprophyre” dykes are rarely observed (e.g., Cooke and Godwin, 1984; Driver et al., 2000; Ash, 2001). Ubiquitous unzoned xenocrystic zircons and resorbed poorly to unzoned old inherited zircon cores indicate that a significant crustal component makes up much, if not all, of the Cassiar suite magmas. Entrained metasedimentary to metaigneous rafts (quartzite, andalusite-bearing schist, hornfels, quartzo–feldspathic gneiss, and amphibolite) may provide examples of magma source rocks, although they are more likely represent contact aureole, roof pendant, and (or) mid-crustal xenoliths (Driver et al., 2000). At least two distinct melt sources have been proposed to account for the range of geochemical compositions and mineralogical variations exhibited by the Cassiar plutonic suite (Driver et al., 2000; Morris and Creaser, 2008; Chapter 2). Reduced, calc-alkaline, biotite ± muscovite monzogranite intrusions have heterogeneous trace element concentrations consistent with a derivation from a mixture of siliciclastic rocks (of average greywacke composition) ± interlayered chemical sedimentary rocks (e.g., carbonates or evaporites) and (or) mafic–intermediate igneous rocks. Oxidized, alkaline to alkali-calcic, biotite ± hornblende-bearing (and less commonly muscovite-bearing) monzogranite to granodiorite intrusions comprise magmas derived mainly from large-fraction partial melting of mafic crust, or small-fraction partial melting of an enriched mantle. The 103–94 Ma Seagull plutonic suite forms a peralkaline to peraluminous, alkali-calcic to calc-alkaline, A-type belt comprising highly siliceous, “one-mica” (biotite-bearing, although rare muscovite or lepidolite are observed) monzogranite intrusions. Most examples of the Seagull suite form small plutons. The Seagull and Hake intrusions are the only batholithic members of this suite, although some phases in the Quiet Lake and Nisutlin batholiths have also been included in this designation on the basis of age and composition. These intrusions are emplaced as a long, narrow belt mainly hosted by Snowcap assemblage rock units of the eastern Yukon–Tanana terrane in the westernmost Cassiar Terrane (Figs. 3.3a and 3.4). Despite the mineralogical similarity to many of the biotite-bearing Cassiar suite intrusions, the Seagull suite is geochemically distinct. Some extremely elevated trace elements (Ce, Ga, Nb, Y, Yb) that are not correlated with typical magmatic fractionation indices (e.g., Rb/Sr and Eu/Eu*)19 suggest that the protolith to these intrusions was geochemically enriched. Other notable features of this suite include abundant miarolitic cavities (Seagull batholith), partially developed rapakivi textures (locally in the Hake batholith; Liverton and Alderton, 1994; T. Liverton, 2012, pers. comm.), a strongly reduced oxidation state, and an association 19 Eu/Eu* = the magnitude and direction of the europium anomaly (Eu/Eu* > 1.0 is a positive anomaly; Eu/Eu* < 1.0 is a negative anomaly) relative to the middle REE samarium and gadolinium, after normalization of all three elements to chondritic compositions. 162 with widespread tin mineralization (south end of belt). This suite is similar to calc-alkaline to alkaline A-type granitoids emplaced in post-orogenic environments (Barbarin, 1990; Liverton and Botelho, 2001; Liverton et al., 2005). The Seagull suite trends from slightly less siliceous “alkali-calcic peralkaline” to highly siliceous “calc- alkalic peraluminous” ferroan (A-type) granitoid varieties of Frost and Frost (2011). Although A-type granites are poorly understood in terms of source rock composition (largely because their compositional range varies considerably; e.g., Frost and Frost, 2011) the dominant alkali-calcic, peralkaline composition exhibited by most samples of the Seagull plutonic suite suggests a primary derivation for these magmas from an undepleted upper asthenospheric to lower lithospheric mantle (e.g., Eby, 1992; Turner et al., 1992; Frost and Frost, 1997, 2011; Liégeois et al., 1998; Bonin et al., 1998; Chappell et al., 2000; Liverton and Botelho, 2001; Morris and Creaser, 2008). A less important high-K calc-alkaline, peraluminous component in several of the most siliceous samples of the Seagull suite may have resulted from coeval partial melting of a quartzo–feldspathic crustal protolith ± local mixing between the two felsic magmas (e.g., Frost and Frost, 1997, 2011). The 82–76 Ma Rancheria plutonic suite is a peraluminous, alkali-calcic (high-K calc-alkaline to shoshonitic), A-type suite with compositions that are typical of post-orogenic or post-collisional magmatism. It comprises small leucocratic biotite ± muscovite monzogranite plutons located east of the Cassiar batholith in the eastern Cassiar Terrane (Fig. 3.3b, Fig. 3.4). Two of the three plutons examined by this study are strongly reduced, but a third pluton is extremely oxidized (the Allen Stock) and hosts U-rich porphyry-style mineralization with anomalous heavy rare earth element (REE) enrichments (Gregory, 2007). The Allen Stock is also cut by late biotite-bearing and silica-undersaturated “mafic” dykes (Gregory, 2007). These three Rancheria suite plutons are located at the northern tip of a belt of Late Cretaceous magmatic, metamorphic, and hydrothermal K–Ar and Rb–Sr ages (Fig. 3.3b; Canadian Geochronological Knowledgebase, 2013). The Rancheria ‘belt’ is located north and west of the Kechika fault, and trends south–southwest into northern British Columbia where it may merge with the coeval, west– northwest-trending Surprise Lake plutonic suite (Fig. 3.3b). Several circa 84–74 Ma plutons have also been identified in the western portions of the Cassiar Terrane (e.g., the eastern lobe of the Deadman batholith20; Figs. 3.3b and 3.4), but are compositionally unlike the three plutons defining the Rancheria suite. Trace element enrichments (e.g., Th, U, Ta, Nb, and Sr) in the least fractionated samples of the Rancheria suite were probably imparted by either extreme differentiation or small-fraction partial melting of a mafic protolith (e.g., Whalen et al., 1987; Sylvester, 1989; Turner et al., 1992; Frost and Frost, 1997, 2011; Liégeois et al., 1998). The shoshonitic affinity exhibited by the Rancheria plutonic suite is typical of a continental lithospheric (either enriched sub-crustal, or crustal) magma origin (e.g., Eby, 1992). The 49–47 Ma Black River plutonic suite comprises only two known intrusions (the Tintina pluton and the Black River batholith) located immediately southwest of the Tintina Fault (Figs. 3.3b and 3.4). This plutonic suite, however, may be cogenetic to other compositionally and temporally similar plutons scattered through the Cassiar Terrane, such as the ~50 Ma Mount Haskin and Mount Reed plutons (Fig. 3.3b; Christopher et al., 1972) and the 20 These undated samples of the Deadman batholith comprise metaluminous, (typically) calc-alkaline, hornblende- rich quartz monzodiorite to hornblende-biotite-bearing granodiorite to biotite-bearing granite phases (Morris and Creaser, 2008). 163 ~42 Ma Balourdet pluton (Evenchick, 1988) in northern British Columbia. The Black River suite may represent just one of several expressions of Paleocene–Eocene magmatism found within the vicinity of crustal-scale structures (Fig. 3.3b). The Black River and Tintina plutons are highly siliceous, peraluminous, biotite- (± trace muscovite-) bearing monzogranites of high-K calc-alkaline composition (straddling the alkali-calcic and calc-alkaline divide of Frost et al., 2001) and post-collisional A-type affinity. These two plutons completely lack inherited zircon, which is consistent with their interpreted petrogenesis as highly fractionated derivatives of mafic parent magmas sourced in an undepleted mantle (e.g., Eby, 1992; Turner et al., 1992; Frost and Frost, 1997, 2011). The Black River suite represents within-in plate magmatism related to localized extension resulting from transtensional movement along the Tintina Fault during a lull in continental arc magmatism to the west (Fig. 3.3b). This setting is similar to that suggested for other coeval and compositionally similar magmas associated with large structures and scattered throughout the Canadian Cordillera (Gabrielse, 1991; Yorath, 1991; Woodsworth et al., 1991; Gabrielse et al., 2006). 3.2.4 Relevant Regional Studies on Igneous Rocks in the Cassiar Terrane21 Three previous petrogenetic, isotopic ± geochemical studies of regional scope have examined syn- to post- accretionary Cretaceous intrusions emplaced in the Cassiar Terrane (Dagenais, 1984; Driver et al., 2000; Morris and Creaser, 2008). None of these studies, unfortunately, employed robust geochronological methods to confirm temporal linkages between intrusive samples, and to allow for the accurate calculation of radiogenic isotopic ratios for the time of magma emplacement. Despite this limitation, each study contributed valuable information on the composition of plutonic rocks in the Cassiar Terrane. 3.2.4.1 The Cassiar Batholith Driver et al. (2000) completed a geochemical and isotopic study of the Cassiar batholith, the ‘inboard’ largest batholith in northern British Columbia and southern Yukon. A wide range of neodymium, strontium, and oxygen isotopic compositions (87Sr/86Sr100 = 0.706–0.734; εNd100 = –2.7 to –17.1; δ18O = +7.6 to +11.6), and a more limited range of lead isotopic compositions (206Pb/204Pb = 19.14–19.24; 207Pb/204Pb = 15.70–15.72; 208Pb/204Pb = 39.14– 39.26) characterize this composite intrusion. Two distinct melt sources, represented by end-member neodymium and strontium compositions, were proposed to explain the chemical and isotopic variation exhibited by this batholith: (1) a relatively isotopically primitive reservoir (e.g., high εNdi, low Sri); and, (2) a isotopically highly evolved reservoir (e.g., low εNdi, high Sri). Three possible scenarios for this dual-source interpretation were examined by Driver et al. (2000). The first possibility, whereby mantle-derived basalt interacted and mixed extensively with coeval crustally derived melts, was discarded because of the lack of coeval basaltic magmatism throughout the Cassiar Terrane, and the mechanical limitations associated with mixing mafic and felsic magmas with different rheological properties. 21 All compiled neodymium and strontium isotopic compositions in the following discussion are corrected to an age of 100 Ma. 164 Mathematical models of this scenario require that at least 50% of the Cassiar batholith was generated by mantle- derived magma (which was inconsistent with the range of isotopic compositions) or that the two distinct isotopic end-members had similar SiO2 contents (<10 wt.% difference). A second scenario involving the assimilation of solid crustal materials by a basaltic magma, assuming two crustal end-members: an average of Proterozoic paragneiss, and a hypothetical crustal composition adopted from the most isotopically evolved plutonic sample. High-temperature (ca. 900°C) basaltic assimilation of the crust was deemed possible if only the ratio of basalt to crust was less than 5:1. This model, however, was thought to be incapable of producing the siliceous geochemical compositions exhibited by most samples from the Cassiar batholith, without invoking extensive magmatic fractionation for which evidence (e.g., high Rb/Sr and low Eu/Eu*) lacked. Finally, the third possibility invoked mixing of melt derived from two compositionally and isotopically distinct crustal rocks appeared to produce the best match with the isotopic data set. The isotopically evolved end-member was interpreted to be Mesoproterozoic to Neoproterozoic metasedimentary packages and crystalline Paleoproterozoic basement rocks. The isotopically juvenile end-member was interpreted to be widespread but small volume, Proterozoic basaltic to dioritic sills/flows and mafic gneisses interlayered with the more evolved melt source. Between 10–80% of the melt in any given sample of the Cassiar batholith, would have been derived from the mafic–intermediate lithology. Based on the thermal modeling of Patiño Douce et al. (1990), anatexis in the Cassiar Terrane was attributed to intra- continental, syn-collisional, crustal thickening of up to 50% , which led to extensive melting of the continental crust 5–25 m.y. (Driver et al., 2000). In the Patiño Douce model, heat from the underlying mantle was limited to normal conductive exchange and additional heat was generated in the melt source, which was described as a hypothetical 10 km-thick mid-crustal metasedimentary package enriched in heat-producing elements. 3.2.4.2 A Regional Transect Morris and Creaser (2008) examined isotopic compositions of several plutonic suites along a transect from the Yukon–Tanana terrane in the southwest Yukon, to the Selwyn Basin in east–central Yukon, crossing the Cache Creek Terrane and the Cassiar Terrane. Several isotopic groupings were distinguished (from west to east): 1) Isotopically primitive (εNd100 = +2.3 to –1.2; 87Sr/86Sr100 = 0.705), mantle-sourced and variably crustally contaminated Early Cretaceous continental magmatic arc rocks of the 115–107 Ma Whitehorse plutonic suite emplaced into the accreted pericratonic rocks. 2) Isotopically primitive (εNd100 = +4.2 to +1.4; 87Sr/86Sr100 = 0.703–0.7035), but not subduction-related (i.e., post- orogenic A-type) intrusions derived from partial melting of mafic rocks―possibly their host Cache Creek terrane―including the Late Cretaceous Surprise Lake pluton and nearby undated plutons. 3) Slightly more evolved (εNd100 = +4.5 to –7.3; 87Sr/86Sr100 = 0.709–0.710) Early Cretaceous subduction-related mixture (Late Triassic to Early Jurassic volcano-sedimentary rocks of oceanic island arc affinity and underlying attenuated Laurentian basement) for the 123–118 Ma Teslin plutonic suite and the Quiet Lake batholith of the 165 Cassiar plutonic suite in the eastern Yukon–Tanana terrane. 4) A non-subduction-related and relatively primitive isotopic signature (εNd100 = +2.0 to –5.4; 87Sr/86Sr100 = 0.703– 0.707) similar to that of group (2) was identified for intrusions in the westernmost Cassiar plutonic suite (the Deadman batholith, and a plug adjacent to the Quiet Lake batholith22). 5) An evolved isotopic signature (εNd100 = –16.6 to –21.7; 87Sr/86Sr100 = 0.717–0.735) attributed to mixing between two distinct metasedimentary melt sources for plutons emplaced in the Cassiar Platform (and Selwyn Basin). These radiogenic isotopic subdivisions were correlated to changes in the basement composition across several tectonostratigraphic entities, although samples from several plutonic suites with different U–Pb ages were inadvertently grouped in the study. 3.2.4.3 Southwest to South–Central Yukon Dagenais (1984) examined whole rock oxygen isotopic compositions of plutonic rocks in Yukon–Tanana terrane, Cache Creek terrane, and the western Cassiar Platform, and compared them to compiled strontium isotopic compositions. The results are complementary to those of Morris and Creaser (2008). The Whitehorse and Teslin plutonic suites, emplaced in pericratonic terranes, are characterized by relatively primitive δ18O of +6.0 to +8.5‰ and 87Sr/86Sr100 (<0.706), which was consistent with their interpretation as mantle- derived continental magmatic arc rocks that experienced minimal crustal contamination (Groups 1 and 3; Section 3.2.4.2). The Surprise Lake batholith and nearby intrusions (ranging from Jurassic to Tertiary in age) emplaced into marine sedimentary rocks of the Cache Creek terrane in northern British Columbia have less primitive 87Sr/86Sr100 (>0.706) and variable δ18O (+6.8 to +10‰)—the latter becoming lower with increased proximity to mafic to ultramafic oceanic rocks at the margins of the Cache Creek terrane (Groups 2 and 4; Section 3.2.4.2). Finally, oxygen and strontium compositions for intrusions in the Cassiar Terrane were relatively high (δ18O = +8 to +11‰; 87Sr/86Sr100 > 0.710), probably coupled, and consistent with anatexis or assimilation of, or contamination by, craton- sourced sedimentary rocks. (Group 5; Section 3.2.4.2). A general lack of correlation between δ18O and 87Sr/86Sr for most of the intrusions led to the suggestion that 87Sr/86Sr was controlled by the tectonostratigraphic entity (as was also suggested by Morris and Creaser, 2008), but δ18O was more likely to be correlated to the composition of the specific host rocks (Dagenais, 1984). These features are particularly well exemplified by intrusions emplaced into the Cache Creek terrane (Group 2; Section 3.2.4.2). 3.2.4.4 The Mesozoic Mantle Information on the composition of the mantle presently beneath the Cassiar Terrane can be extracted from 22 This plug has been included in the Seagull plutonic suite on the basis of its geochemical composition (Chapter 2), although geochronological and isotopic data for this sample are lacking. 166 Quaternary to Recent alkaline volcanic centres and flows that bisect the Cassiar Terrane (Fig. 3.1). These basaltic magmas are interpreted to have been derived from an enriched mantle, through the mixing of small-volume and high-fraction partial melts from amphibole-bearing veins with high-volume and low-fraction partial melts from hydrous garnet peridotite. The basaltic rocks have elevated large ion lithophile element (LILE; e.g., Rb, Th, U, and Pb) concentrations, and high Rb/Sr and Th/U (Abraham et al., 2005)23, which has resulted in their relatively evolved strontium and lead isotopic compositions. Abraham et al. (2001) also noted that there is a gradual change in isotopic composition (that is not related to chemical composition) from evolved values suggestive of a derivation in old cratonic lithosphere, to juvenile values for lithosphere below the pericratonic terrane(s) southwest of the Teslin Fault. This implies that two distinct lithospheric mantles are juxtaposed, by an east-dipping lithospheric boundary or by a zone of lithospheric “mixing” (Abraham et al., 2001). Isotopically evolved ‘basalts’ in the Cassiar Terrane suggest that the underlying lithosphere was enriched during Mesozoic subduction, prior to, during, and following the accretion of Yukon–Tanana terrane (Abraham et al., 2001). The isotopic compositions of Quaternary to Recent ‘basalts’, therefore, provide a proxy for that of the enriched lithospheric mantle that underlay the Cassiar Terrane during the Early Cretaceous to Eocene magmatic events. Therefore, the lithospheric mantle below the Cassiar Terrane may have had a similar overall isotopic composition, with radiogenic strontium (87Sr/86Sr100 = 0.704–0.705) and relatively non-radiogenic neodymium (εNd100 = +5.8 to +1.8; Abraham et al., 2001) compared to typical depleted mantle compositions. The lead isotopic compositions of the ‘basalts’ (206Pb/204Pb = 19.02–19.41; 207Pb/204Pb = 15.56–15.67; 208Pb/204Pb = 38.43–39.43), however, have been nearly homogenized to upper crustal compositions (cf. Zartman and Haines, 1988) due to contamination from the overlying continental crust or subducted pelagic to terrigenous rocks (Abraham et al., 2001). 3.3 METHODS AND RESULTS Sample(s) inferred to be representative of the general composition of each intrusion (more than one sample was taken from several batholiths) were chosen for isotopic analysis. These samples were typically a medium- to coarse- grained, K-feldspar-phyric, phase of intermediate to felsic composition. Basic descriptions, as well as U–Pb crystallization ages, trace element ratios, and geochemical parameters presented in Chapter 2 for all samples that underwent isotopic analysis are summarized in Table 3.1, and are located in Figure 3.4. Prior to crushing, hand samples and thin sections were examined and sample material containing quartz ± sulphide bearing veins or fractures, or pervasively altered samples, were avoided. 3.3.1 Radiogenic Isotopes All sample preparation (with the exception of crushing and grinding, which were done at ALS Chemex, Vancouver, for whole rock isotopic samples), geochemical separation, isotopic analysis, and data correction and reduction for radiogenic isotopic analyses were done at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the 23 Crustal contamination of these basaltic magmas is thought to have been limited due to their rapid ascension and eruption in an extensional setting (Abraham et al., 2001). 167 University of British Columbia (UBC). 3.3.1.1 Sm–Nd and Rb–Sr Sm–Nd and Rb–Sr24 isotopic analysis was done on twelve whole rock powders, along with one sample duplicate, one standard (GS-2), and one standard replicate (per sample set). Detailed methods for dissolution, ion exchange chemistry, and mass spectrometry for neodymium and strontium isotopic analysis are described by Weis et al. (2006), and detailed sample digestion and mass spectrometry methods for trace element analysis (Sm, Nd, Rb, Sr) are described by Pretorius et al. (2006). Present-day neodymium and strontium isotopic compositions, and calculated neodymium and strontium isotopic compositions for 100, 80, and 50 Ma are reported with 2-sigma errors in Table 3.2 and are summarized in Figure 3.5a. Compiled isotopic data are tabulated in Appendix G.3 and are also presented and referenced in Figure 3.5. Of the data resulting from this study, biotite-bearing samples (Cassiar-bt” in Fig. 3.5) display the widest range of both ɛNd100 and 87Sr/86Sr100, from relatively isotopically juvenile values of –7.6 and 0.709 to relatively isotopically evolved values of –16.2 and 0.721, respectively (Fig. 3.5a). Muscovite ± biotite-bearing samples (“Cassiar-ms” in Fig. 3.5) have an overlapping, but much more limited range of –11.2 and 0.712 to –14.3 and 0.718 (Fig. 3.5a). Supplementary data compiled for the Cassiar plutonic suite establishes slightly larger ɛNd100 and 87Sr/86Sr100 ranges of –3.2 to –19.88 and 0.706–0.736 for biotite-bearing samples, and –3.7 to –17.33 and 0.707–0.734 for muscovite- bearing samples (Fig. 3.5a). No hornblende ± biotite-bearing phases (“Cassiar-hbl” in Fig. 3.5) sampled by this study were analyzed for radiogenic isotopic compositions, however, compiled isotopic data establish ɛNd100 and 87Sr/86Sr100 ranges of –2.8 to –6.1 and 0.706–0.711 for three such samples (Fig. 3.5a). Compiled 87Sr/86Sr100 for the Seagull plutonic suite range between 0.706–0.708 (Fig. 3.5a)25, but neodymium isotopic data are lacking. Samples from the two A-type plutonic suites (Rancheria, and Black River) have relatively restricted neodymium and strontium isotopic compositions, although this may be a consequence of the limited number of samples analyzed from these plutonic suites. At 80 Ma, two samples (and one duplicate) for the Rancheria plutonic suite have a ɛNd100 and 87Sr/86Sr100 range of –8.1 to –11.3 and 0.708–0.709 (Fig. 3.5a; Table 3.2). At 50 Ma, one sample of the Black 24 While neodymium isotopic data are conventionally expressed relative to the chondritic uniform reservoir (CHUR) representing the composition of bulk Earth at a specified age (ɛNdt, where t = age), strontium isotopic data are more conventionally expressed as the initial strontium ratio (87Sr/86Srt). Here, neodymium and strontium isotopic compositions calculated for an initial age of 100 Ma are denoted by: • ɛNd100 ([144Nd/143Nd100 / 144Nd/143NdCHUR@100Ma] – 1 x 10000): increasingly negative values are associated with less radiogenic neodymium than CHUR (or bulk Earth), but more isotopically “evolved” (and less isotopically “juvenile”) magmas; and • 87Sr/86Sr100: higher values are associated with more radiogenic strontium in general, and more isotopically “evolved” (and less isotopically “juvenile”) magmas. In most cases, crustal rocks are more isotopically evolved than the mantle and bulk Earth. This is a function of the different partitioning behaviour of Sm–Nd and Rb–Sr during melt extraction from the mantle, as well as the concentration of rubidium into crustal rocks by magmatic and weathering processes. The division between these two broad reservoirs (mantle and crust) are generally (but not always) taken as the ɛNd value for CHUR (always “0”) and by a Sri (initial) value of 0.706. 25 The strontium isotopic data come from five samples of the Thirtymile pluton (Liverton et al., 2001), as well as one sample of a small plug near the Quiet Lake batholith (Morris and Creaser, 2008) thought to be a cumulate phase of the Seagull plutonic suite (on the basis of geochemical interpretations; Chapter 2). 168 River plutonic suite has a ɛNd100 and 87Sr/86Sr100 composition of –12.9 and 0.718 (Fig. 3.5a; Table 3.2). Despite the wide range of neodymium and strontium isotopic compositions exhibited by the Cassiar plutonic suite, the average εNd100 and 87Sr/86Sr100 for biotite-bearing (16 and 17 samples, respectively) and muscovite-bearing (12 samples) samples of this suite are virtually identical (Fig. 3.5a; Table 3.2;Appendix G.3). The hornblende-bearing samples from the Cassiar suite are distinctly less isotopically evolved as a group (Fig. 3.5a; Table 3.2; Appendix G.3), although this could be partly to entirely an effect of the limited number of hornblende-bearing samples with existing isotopic data. 3.3.1.2 Pb–Pb Relatively fresh, inclusion-free feldspar grains were analyzed from a total of thirteen samples for lead isotopic analysis. This material was hand-picked from crushed and ground light material collected during separation of heavy and light mineral fractions on a wet shaker table. Wherever possible, K-feldspar was selected in preference of plagioclase due to its greater resistance to sub-solidus and hydrothermal processes. Sample preparation and lead isotopic analysis followed standard techniques, as described in detail by Heffernan (2004). The measured ratios were corrected for instrumental mass fractionation of 0.12% per mass unit based on repeated measurements of the NBS SRM 981 standard and the values recommended by Thirwall (2000). Mass fractionation and analytical errors were numerically propagated throughout all calculations and are reported at the 2-sigma level. The lead isotopic ratios have not been age-corrected due to the very low concentrations of uranium relative to lead incorporated into feldspar (and sulphide) minerals, and to the Phanerozoic age of the samples (cf. Tosdal et al., 1999). Lead isotopic data are reported in Table 3.2 and summarized in Figure 3.5. Compiled isotopic data are tabulated in Appendix G.3 and are also presented and referenced in Figure 3.5. Lead isotopic compositions also range widely for the Cassiar plutonic suite. 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb have respective ranges of: 19.32–20.17, 15.69–15.82, 39.26–39.82 for biotite-bearing samples; and 19.24–19.31, 15.70–15.74, and 39.172–39.36 for muscovite-bearing samples (Fig. 3.5b). Additional compiled isotopic data lower the 206Pb/204Pb and 208Pb/204Pb ranges to 19.22 and 39.22 for biotite-bearing samples, and to 19.14 and 39.14 for muscovite-bearing samples (Fig. 3.5b). The only analysis of a hornblende-bearing sample has a composition of 19.24, 15.72, and 39.20 (206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb, respectively; Fig. 3.5b). The Seagull plutonic suite has not analyzed for lead isotopic compositions. Respective 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb values for 19.60–19.75, 15.72–15.74, and 39.77–39.93 for the Rancheria suite (2 samples and 1 duplicate sample), and 19.48, 15.70, and 39.69 for the Black River suite (1 sample; Fig. 3.5b). 3.3.2 Stable Isotopes Thirty samples underwent stable isotopic analysis for sulfur (whole rock powders, prepared by ALS Chemex, Vancouver) and oxygen (hand-picked quartz separates) at the Nevada Stable Isotope Laboratory, University of Nevada-Reno. Stable isotopic methods and data used for plotting here are also reported and the data have previously 169 been briefly discussed in Rasmussen et al. (2010). 3.3.2.1 Oxygen Petrographic evidence for water-rock interaction may be used to selectively discriminate against hydrothermally- altered rocks on a first-order basis. Quartz mineral separates (which are more likely to retain magmatic isotopic compositions than mica or feldspar minerals during hydrothermal alteration; Gregory and Criss, 1986; DeYoung, 2007) were hand-picked from a wet shaker table light fraction of crushed and ground whole rock material from unaltered and vein-free samples. Duplicate quartz separates were picked for 10% of the samples. Quartz grain samples weighing approximately 1000 µg were sealed into tin cups, then reacted at temperatures as high as 3000°C in the presence of graphite in a graphite electrode furnace (i.e., inert gas fusion analysis). This high-temperature carbon reduction technique liberates silicate oxygen by reacting with graphitic carbon to produce CO (Arehart and Poulson, 2006). The resultant CO gas is then introduced by a stream of helium gas to a very high temperature elemental analyzer – mass spectrometer operating in continuous-flow mode. This method closely reproduces the oxygen isotopic compositions and uncertainties of a number of standard materials analyzed by traditional fluorination methods. Replicate analyses of standards NCSU, NBS-28, and ARQ indicated a precision of ± 0.2‰, and δ18O is reported relative to an NBS-28 value of +9.6‰. All oxygen isotopic data and lab standards are reported in standard ‘del’ (δ) notation relative to the Vienna Standard Mean Ocean Water (Coplen, 1994). Oxygen isotopic data are reported in Table 3.2 and summarized in Figure 3.5. Compiled isotopic data are tabulated in Appendix G.3 and are also presented and referenced in Figure 3.5c. All intrusions in the Cassiar Terrane display a normal distribution of highly positive δ18O values, with a median value of +12.8‰ and a range of +10.6 to +16.5‰ (Fig. 3.5c). Oxygen isotopic compositions from quartz separates for the Cassiar plutonic suite have a wide range of high positive values (+10.6 to +16.5‰), although data for whole rock samples are typically several per mil lower (Fig. 3.5c). Biotite-bearing samples are the most variable, with δ18O ranging from +11.8 to +15.7‰ (Fig. 3.5c). Muscovite-bearing samples have a similar but more limited range of +11.5 to +14.4‰ (Fig. 3.5c). Compiled data from quartz separates extend the bottom range to +10.8 for biotite- bearing samples, and +10.9‰ for muscovite-bearing samples (Fig. 3.5c). The hornblende-bearing phase sampled by this study has a δ18O of +13.2‰ (Fig. 3.5c). The lowest δ18O value obtained from the analysis of quartz is +10.6‰ (Dagenais, 1984), but the mineralogy of this sample was not specified (it is plotted as biotite-bearing by default in Fig. 3.5c). Samples from the three A-type plutonic suites (Seagull, Rancheria, and Black River) have oxygen isotopic compositions that mutually overlap: +12.5 to +15.4‰ for the Seagull suite, +11.6 to +16.5‰ for the Rancheria suite, and +12.5 to +14.0‰ for the Black River suite (Fig. 3.5c). These oxygen isotopic compositions are within the range exhibited by the Cassiar plutonic suite (with the exception of the Gravel Creek Stock; KR-07-23). The oxygen isotopic composition of the Seagull plutonic suite just barely overlaps with that of the Surprise Lake plutonic suite (+9.5 to +12.6; indicated by the red dotted box in Fig. 3.5c), which is also plotted for comparative purposes. 170 3.3.2.2 Sulfur Sulfur was extracted from 30–50 grams of whole rock powder using the Kiba method (Sasaki et al., 1979) for intrusive samples that did not display evidence for the addition or loss of sulphides or sulphates (e.g., hydrothermal alteration). A mixture of rock powder and ~50 mL of Kiba reagent was heated under a nitrogen gas stream to ~280oC. The extracted H2S gas was bubbled through ~15 mL of zinc acetate solution, precipitating ZnS, which was then reacted with AgNO3 to ultimately precipitate Ag2S. Approximately 10% of the samples underwent a duplicate sulphide extraction. About 380 µg of Ag2S was weighed into a tin cup with about 200 µg V2O5 (to aid combustion), sealed, and combusted to SO2 gas in a high-temperature Eurovector Elemental Analyzer. The resulting SO2 gas was introduced in continuous-flow mode to a Micromass Isoprime Stable Isotope Ratio Mass Spectrometer, following the method of Fry et al. (2002). Approximately 30% of the samples, including half of the duplicate samples, also underwent replicate analysis. The GSL (sphalerite: δ34S = –2.92‰) and MIC (marcasite; δ34S = +17.90‰) standards were used to correct for the presence of 32S16O18O (which has the same isotopic weight as 34S16O16O) and for machine drift. Typical analytical precision for the reported analyses is ± 0.3‰ or better, and all data and standard values are reported in δ notation relative to the Vienna Cañon Diablo Troilite (Ding et al. 2001). Several sulfur-poor samples with small analytical peaks were corrected on the basis of analyses of similarly small quantities of the GSL standard, which produced smaller peaks and lower δ34S values. The correction factor for the unknown samples is based on the known δ34S for GSL. Sulfur isotopic data are reported in Table 3.2 and are summarized in Figure 3.5. Compiled isotopic data are tabulated in Appendix G.3 and are also presented and referenced in Figure 3.5. Throughout the Cassiar Terrane study area, δ34S values for syn- to post-accretionary intrusions range from +2.0 to +10.2‰, although the majority of the intrusions have δ34S values greater than +8‰ (Fig. 3.5c). As with the other isotopic systems, sulfur compositions for the Cassiar plutonic suite also display a relatively wide range of values. Biotite-bearing samples display the widest range of δ34S, from +2.0 to +10.1‰, whereas muscovite-bearing samples show a similar but more limited range of +4.5 to +9.0‰ (Fig. 3.5c). The hornblende-bearing phase sampled by this study has a δ34S value of +8.1‰ (Fig. 3.5c). Samples from the three A-type plutonic suites (Seagull, Rancheria, and Black River) have relatively restricted δ34S compositions, although this could simply reflect the limited number of samples from these suites. Sulfur isotopic compositions increase from +3.5 to +4.9‰ for the Seagull suite, to +6.0 to +6.3‰ for the Rancheria suite, to +9.0 to +10.2‰ for the Black River suite (Fig. 3.5c). If supplementary data from the south end of Rancheria belt are included, this range can be extended to +3.7 to +9.4‰. Like oxygen, the sulfur isotopic compositions of the A-type plutons are within the range exhibited by the Cassiar plutonic suite. 3.4 ISOTOPIC COMPOSITIONS AND COMPARISONS 3.4.1 Radiogenic Isotopic Datasets Additional neodymium, strontium ± lead isotopic data for the Cassiar plutonic suite (Driver et al., 2000; Abraham et al., 2001; Morris and Creaser, 2008) and strontium isotopic ranges for the Seagull plutonic suite (Liverton et al, 2001; Morris and Creaser, 2008) are plotted together with data generated by this study in Figure 3.5. Compositions 171 of Early Cretaceous to Eocene I-type continental arc magmatic rocks west of the Cassiar Terrane (the Early Cretaceous Teslin26 and Whitehorse plutonic suites; the Late Cretaceous Prospector Mountain plutonic suite associated with later Carmacks Group volcanism; and the Paleocene–Eocene Coast Plutonic Complex; Samson et al., 1991; Selby et al., 1999; Morris and Creaser, 2008; Bineli Betsi and Lentz, 2011) are plotted for comparative purposes. Limited neodymium and strontium isotopic data for the (A-type) Surprise Lake plutonic suite are also included for comparisons (Mihalynuk, 1999; Morris and Creaser, 2008). Radiogenic isotopic data for Quaternary to Recent basalt erupted in the Cassiar Terrane (Fig. 3.1) provide a compositional proxy for the enriched mantle that is inferred to have underlain the Cassiar Terrane throughout the syn- to post-accretionary magmatic events (Abraham et al., 2001; see above discussion). 3.4.1.1 87Sr/86Sr100 versus εNd100 All of the four syn- to post-accretionary plutonic suites studied in the Cassiar Terrane have initial neodymium isotopic compositions that are more evolved than the bulk Earth, or the chondritic uniform reservoir (CHUR), and have initial strontium isotopic compositions more evolved than the mantle (i.e., 87Sr/86Sr100 > 0.706; DePaolo and Wasserburg, 1979; Faure, 2001), including underlying enriched mantle (Fig. 3.6a; Abraham et al., 2001). The Cassiar plutonic suite forms a relatively continuous curvilinear array in a plot of 87Sr/86Sr100 versus εNd100 (Fig. 3.6a). This classic “mixing curve” is seen in most regional studies of plutonism, although rarely over such a wide range of isotopic compositions within a single suite. The more isotopically evolved members of the Cassiar plutonic suite have compositions that are typical of older and (or) felsic rocks (DePaolo and Wasserburg, 1979), such as cratonic basement or sedimentary rocks derived from the erosion of such basement. The least isotopically evolved samples have moderately juvenile compositions that are typical of a younger and (or) mafic rocks that are only slightly more evolved than enriched mantle compositions (DePaolo and Wasserburg, 1979). The Cassiar suite sample array is disrupted by two perturbations, at which samples appear to diverge towards higher 87Sr/86Sr100, occurring at: εNd100 ≈ –5 and 87Sr/86Sr100 ≈ 0.707, and εNd100 ≈ –13 and 87Sr/86Sr100 ≈ 0.72 (herein referred to as the “upper” and “lower” perturbations, respectively; Fig. 3.6a). Given the overall continuity and density of data, and the extremely small errors on the isotopic measurement (i.e., smaller than the symbols in Figure 3.6), these perturbations appear to be a real isotopic feature of these magmas. There are a number of variables related to source composition, partial melting, and magmatic processes that could create these disruptions to the otherwise continuous curvilinear sample array. These two perturbations could also represent the average end- member isotopic compositions of the crustal reservoirs that underwent partial melting in the Cassiar Terrane. Samples that plot above the upper perturbation, or the least isotopically evolved samples included in the Cassiar suite, range up to relatively unevolved compositions of εNd100 ≈ –3 and 87Sr/86Sr100 ≈ 0.706 (Fig. 3.6a). These samples are considered to be isotopically juvenile compared to the rest of the Cassiar plutonic suite. The isotopically 26 Four samples from the northern tip of the eastern lobe of the Deadman batholith (Morris and Creaser, 2008) are included, here, in the Early Cretaceous Teslin plutonic suite, although these samples from the Deadman batholith may well be Late Cretaceous in age (Fig. 3.4). However, they plot within the εNd100 and 87Sr/86Sr100 isotopic range of other known Teslin suite samples. 172 juvenile phases comprise several muscovite-, biotite- and (or) hornblende-bearing samples collected by Driver et al. (2000) and Morris and Creaser (2008)27. These samples appear to be isotopically more akin to the Teslin plutonic suite, a 123–115 Ma continental magmatic arc developed along the western Cassiar Terrane and the eastern Yukon– Tanana terrane (e.g., Hart, 1997; Mortensen et al., 2000; Morris and Creaser, 2008). Given that no age determinations are available for these specific samples, a viable possibility is that the isotopically juvenile intrusions belonging to the Cassiar plutonic suite may have had a petrogenetic evolution more similar to that of the older Teslin plutonic suite. Additional work could determine whether or not these samples are coeval to the 115–104 Ma Cassiar plutonic suite or should instead be considered as belonging to the 123–115 Ma Teslin plutonic suite. Samples that plot immediately below the upper perturbation are biotite-bearing phases of the Marker Lake batholith and Cabin Creek pluton of the Cassiar plutonic suite (two samples of the Late Cretaceous Rancheria plutonic suite also plot below the upper perturbation; Fig. 3.6a). These intrusions cluster geographically near the northern tip of the Rancheria belt (Fig. 3.4). The most isotopically evolved samples of the Cassiar plutonic suite are from the biotite- bearing Big Salmon and Nisutlin batholiths in the central study area, and one muscovite-bearing sample from the Cassiar batholith (εNd100 < –15; Fig. 3.6a). The Big Salmon and Nisutlin batholiths are partially hosted by what has been described as a “migmatitic complex” comprising 111 Ma granodioritic orthogneiss (Mortensen and Hansen, 1982), or sheeted granodioritic dykes separated by amphibolite-grade metasedimentary rocks (Tempelman–Kluit, 1977a). The orthogneiss is thought to be the product of either deep equivalents to the Cassiar suite (Mortensen and Hansen, 1982), or in situ melting of the quartzo–feldspathic Ingenika Group (Windermere Supergroup) host rocks (Morris and Creaser, 2008). A granitic “migmatite” sampled from the orthogneiss (Morris and Creaser, 2008) has one of the most ‘evolved’ isotopic compositions in Figure 3.6a. Biotite- and (or) muscovite-bearing plutonic phases from the Cassiar suite trend from base of the upper perturbation towards the highly evolved “migmatitic” end- member sample (Fig. 3.6a). The isotopic compositions determined for the small number of samples from the Rancheria (at 80 Ma) and Black River (at 50 Ma) plutonic suites plot mainly within the range of compositions exhibited by the Cassiar plutonic suite (Fig. 3.6a). 87Sr/86Sr80 for the 82–75 Ma Rancheria plutonic suite is low (0.709–0.710), but by 50 Ma both ɛNd50 and 87Sr/86Sr50 for the Rancheria suite have evolved compositions similar to that of the 47 Ma Black River sample (0.718). An initial strontium isotopic composition for the Late Cretaceous Storie (or Cassiar) pluton at the eastern end of the Surprise Lake belt (Fig. 3.3b) is similar to that of the Rancheria plutonic suite (Fig. 3.6a). An initial strontium isotopic composition for the Mount Hansen pluton (Fig. 3.3b)—which was tentatively related to the Black River magmatic event on the basis of age and apparent composition in Chapter 2—is less isotopically evolved than the Black River suite but still within the range exhibited by the A-type plutonic suites (Fig. 3.6a). Like the Rancheria plutonic suite at its time of emplacement, 87Sr/86Sr100 for the Seagull plutonic suite is low and tightly constrained 27 These isotopically juvenile samples include all of the hornblende-bearing intrusive phases for which there are isotopic data, as well as several muscovite- and biotite-bearing samples from the Quiet Lake (Morris and Creaser, 2008) and the Cassiar (Driver et al., 2000) batholiths. There is no geochronological control on these samples, so they are inferred to be members of the Cassiar plutonic suite based on their spatial association with the Cassiar suite. It is possible, however, that the isotopically juvenile samples would be better associated with the older Teslin plutonic suite (especially the Quiet Lake batholith) or with Late Cretaceous plutonism associated with the Carmacks Group volcanic episode, both found west of most of the Cassiar suite magmatism in the Cassiar Terrane (Fig. 3.4). 173 (87Sr/86Sr100 = 0.706–0.708; Fig. 3.6a). The Seagull suite has 87Sr/86Sr100 values that are slightly more evolved with respect to the enriched mantle represented by Quaternary to Recent basaltic volcanics (Fig. 3.6a) and the upper Sri limit of 0.706 for enriched mantle in general (e.g., DePaolo and Wasserburg, 1979). For comparative purposes, 87Sr/86Sr values for the Seagull suite calculated at 80 and 50 Ma are also more radiogenic, and overlap with the Rancheria and Black Lake plutonic suite samples. As will be demonstrated later, the A-type plutonic suites typically form linear trends on lead and sulfur isotope plots, which may allow for εNd100 of the Seagull plutonic suite to be tentatively approximated at values of –5 to –7 is indicated by the red boxes in Figure 3.6a. The strontium isotopic composition of the A-type plutons clearly evolves towards more radiogenic compositions faster than the Cassiar suite samples (with the exception of the granitic “migmatite”). This is probably a result of the high Rb/Sr that characterizes these plutons (Chapter 2), and this type of magmatism in general. 3.4.1.2 147Sm/144Nd versus εNd100 On plots of εNd100 against 147Sm/144Nd (Fig. 3.6b), the Cassiar plutonic suite, continental magmatic arc rocks of the Whitehorse and Teslin plutonic suites, and proxies for the composition of the Cretaceous mantle below the Cassiar Terrane (Quaternary to Recent alkaline ‘basalts’) form a crescent-shaped array. The upper portion of the crescent is characterized by the least isotopically evolved continental magmatic arc and enriched mantle samples, including the isotopically juvenile Cassiar suite samples (i.e., those plotting above the upper perturbation in Fig. 3.6a). This is not unexpected; 147Sm/144Nd is higher in rocks derived from the mantle than in most felsic crustal rocks. The continental arc magmas likely incorporated some mantle component from their protolith(s), and the Quaternary to Recent alkaline ‘basalts’ are known to be mantle-derived. In comparison, the majority of the Cassiar suite samples and the limited number samples available for Rancheria and Black River suites have lower 147Sm/144Nd values consistent with their lower εNd100 and felsic compositions. This is consistent with the composition of most upper crustal rocks (e.g., granite, felsic gneiss, and shale), which have lower Sm/Nd ratios than mantle-derived rocks. The lower third of the 147Sm/144Nd versus εNd100 crescent, however, is defined mainly by high 147Sm/144Nd and the most isotopically evolved samples of the Cassiar plutonic suite (e.g., εNd100 < –16 and 87Sr/86Sr100 > 0.72). Although the sample with the highest 147Sm/144Nd and lowest εNd100 is the granitic “migmatite”, several (4 of 8) biotite- bearing samples from the nearby Big Salmon and Nisutlin batholiths have anomalously high 147Sm/144Nd, as do several muscovite-bearing phases of the Cassiar batholith (circled with blue dashed line; Fig. 3.6b). There is the possibility that this unusually elevated Sm/Nd is the result of an analytical issue, perhaps caused by incomplete crushing or digestion of accessory minerals that host the light to middle REEs. However, the affected samples are from three sample sets analyzed by the same methods in two laboratories (Driver et al., 2000; Morris and Creaser, 2008; this study). If this is an analytical issue, it must have occurred in both laboratories but only affected a small proportion of samples from three large sample sets28. An analysis of the data quality (Appendix D.3) does not 28 That said, a close examination of Fig. 3.4b that many of the “moderate” Cassiar suite samples (i.e., samples with εNd100 < –9) do deviate slightly towards, rather than away from (as would have been expected), elevated 147Sm/144Nd with decreasing εNd100. 174 indicate any obvious problem with REE analyses for this study29. It would seem that there is a mineralogical, geochemical, or other factor shared by these samples that has affected their measured light to middle REE contents. An ultramafic to mafic, calcareous melt source may produce high Sm/Nd, but both are unlikely given the high degree of isotopic evolution for these samples; therefore, samarium and neodymium were most likely decoupled during magmatic or analytical processes in these samples. If unusually coarse-grained monazites (which contain the bulk of Sm and Nd in granitoids; Ayres and Harris, 1997) were present in these isotopically highly evolved samples, incomplete dissolution of these grains might affect the whole rock isotopic composition. Alternatively, if monazite was present in the magma protolith or residuum, isotopic disequilibrium between monazite and the extracted magma might also result in unusually high Sm/Nd because of the greater relative dissolution of samarium-rich apatite compared to monazite under fluid-absent partial melting conditions (Ayres and Harris, 1997; Zeng et al., 2005). Therefore, although the 147Sm/144Nd enrichment in these granitoids is unusual given their felsic composition, it is not improbable if anatectic conditions did not allow equilibration between magma and restite, or if monazites were incompletely dissolved during chemical analysis. 3.4.1.3 TDM versus εNd100 Depleted mantle model ages (TDMs) based on the Sm–Nd isotopic system approximate the age of neodymium extraction from a depleted mantle isotopic reservoir, and provide an average crustal residence age for magma protoliths. If the protolith is infracrustal (igneous or metaigneous), was derived directly from a depleted mantle, did not undergone any crustal recycling (i.e., partial melting), did not experience crustal contamination, and remained a closed isotopic system, then the TDM should be the same as the age of the rock. Rarely are all of these criteria fulfilled, however, and TDMs are usually older due to the incorporation of material from the neodymium enriched crust by contamination or assimilation. If the protolith is supracrustal, plutonic TDMs represent the average crustal residence age (not the depositional age) for the materials that melted to form the plutonic body. TDMs for all of the Cassiar suite samples ranges from 0.9–3.75 Ga (Fig. 3.6c). The majority of the Cassiar suite samples, as well as all of the few Rancheria and Black River samples, have TDMs between those of the upper and lower perturbations (i.e., ~1.2 to 1.9 Ga). Isotopically evolved samples with anomalously elevated 147Sm/144Nd values (circled; Fig. 3.6c) have TDM estimates ranging from 1.9–3.75 Ga. The positive correlation between 87Sr/86Sr100 and TDM (not shown, but implied by in Fig. 3.6a,c) is consistent with the generalization that rocks with longer crustal residence ages are more isotopically evolved (e.g., DePaolo and Wasserburg, 1979). If the TDMs in Figure 3.6c represent averages of the age of the crust or mantle from which the intrusions were derived, they imply that the anomalously elevated 147Sm/144Nd samples were derived from 1.9–3.75 Ga crust. High 147Sm/144Nd, however, produces anomalously old TDMs in rocks that have lower 143Nd/144Nd (or strongly negative εNd)30, making 29 Residual standard deviations (%) for samarium and neodymium trace element analysis by ICP–MS determined for this study: • Sm: 1.8 and 3.1 for two reference materials, and 2.3, 4.7, and 12.3 for three duplicate samples; and • Nd: 2.6 and 3.3 for two reference materials, and 2.3, 3.6, and 12.7 for three duplicate samples. 30 Based on the equation: TDM = 1/λ147Sm * ln[ (143Nd/144Ndm – 143Nd/144NdDM) / (147Sm/144Nd – 147Sm/144NdDM) +1]. 175 TDMs in these samples invalid. Therefore, none of the Cassiar suite samples are likely to have been derived from rocks with an average residence age that is much older than 1.9 Ga. Except for the most isotopically juvenile samples of the Cassiar suite, all four syn- to post-accretionary plutonic suites identified in the Cassiar Terrane appear to have incorporated crust with an average crustal residence age of about 1.2–1.9 Ga. 3.4.1.4 206Pb/204Pb versus 207Pb/204Pb and 208Pb/204Pb Changes in the uranogenic lead ratios for Phanerozoic rocks are a function of the time of lead extraction, as recorded by 206Pb/204Pb, and the input of young versus old crustal lead, as recorded by 207Pb/204Pb (Fig. 3.7a). When compared to lead evolutionary curves, elevated 207Pb/204Pb and 206Pb/204Pb in plutonic rocks are consistent with a derivation from a highly radiogenic (e.g., old, felsic, and (or) U-rich) crustal lead reservoir. The greater abundance of 235U compared to 238U in the Archean and the shorter half-life of 235U leads to high 207Pb/204Pb for Archean rocks, and low 207Pb/204Pb for younger or less evolved rock with low U/Pb (e.g., mafic rocks derived from the mantle, or felsic rocks metamorphosed to high grades shortly after their deposition). Samples that are anomalously radiogenic (particularly with respect to 206Pb/204Pb for Phanerozoic rocks), however, may have been affected by post- emplacement re-heating (Tosdal et al., 1999). The most 206Pb/204Pb-radiogenic sample from the Cassiar suite is the Buttle Creek pluton (labeled “BC” in Fig. 3.7), which exhibits several features that are compatible with post- emplacement disruption to its lead isotopic system31. Thorogenic lead plots (208Pb/204Pb vs. 206Pb/204Pb; Fig. 3.7b) provide some details concerning the type of crustal melt source. This is a result of the higher mobility of uranium relative to thorium in fluids (e.g., Jenner, 1996). The decay of 232Th to 208Pb is responsible for the 208Pb/204Pb composition; therefore, the 208Pb-rich area above upper crustal growth curves on the thorogenic plot is regarded as characteristic of a highly metamorphosed (i.e., granulite-grade) lower crustal reservoir (e.g., Zartman and Haines, 1988). Conversely, samples that plot below upper crustal growth curves on thorogenic plots may contain lead from reservoirs that interacted with seawater, which has high U/Pb (e.g., ocean island arc and (or) chemically precipitated sedimentary rocks; Taylor and McLennan, 1985; Tosdal et al., 1999). Lead isotopic compositions for sedimentary rocks―and presumably granitoids derived from sedimentary rocks―can be heterogeneous depending on the variation in age and composition of detrital sources as well as degree of sedimentary mixing (Tosdal et al., 1999). Greater relative variation in 207Pb/204Pb relative to 206Pb/204Pb characterizes most of the Cassiar plutonic suite, which is consistent with significant variation in the age of the lead source(s) for these intrusions (Fig. 3.7a). This is compatible with a dominantly supracrustal magma source, containing detritus eroded from terranes characterized by a wide range of ages, for these plutons. The Cassiar plutonic suite generally parallels the lead evolutionary path of northern Cordilleran upper crust in Figure 3.7―approximated by the “Shale Curve” (Godwin and Sinclair, 1982)―implying that the Cassiar suite was derived 31 The Buttle Creek pluton was emplaced along the Tintina Fault (a strike-slip structure active mainly in the Cenozoic), is strongly foliated, has a high degree of K-feldspar ordering (e.g., no tartan twins or perthite), and was sampled near several cross-cutting and undeformed aplite dykes. These features are compatible with sub-solidus re- heating of the sample by later deformation and (or) magmatism. 176 from sources that evolved along a similar lead curve as northern Cordilleran upper crust. A cluster of samples for this plutonic suite32 plot mainly on or near the Shale Curve (circled in Fig. 3.7a) at a position broadly corresponding to a Jurassic model age, which is reasonable for Early Cretaceous magmas given the large errors associated with lead growth curves. Two samples of less radiogenic biotite and (or) muscovite-bearing examples of the Cassiar plutonic suite (also circled; Fig. 3.7a) may trend away from the sample cluster on the Shale Curve towards less radiogenic lead compositions that are more typical of the enriched mantle (represented by the Quaternary to Recent ‘basalts’; Abraham et al., 2001). These samples of the Cassiar plutonic suite may have evolved by mixing between lead sourced from the enriched mantle, or a similarly unevolved (i.e., relatively young and U–Th-poor) crustal reservoir, and a more radiogenic reservoir similar to that of northern Cordilleran upper crust. Mafic sills and flows observed throughout Proterozoic metasedimentary basement to the northern Cordillera are one potential source of this juvenile component. Several other samples that lack evidence of sub-solidus re-heating or fluid flow have more radiogenic compositions than the main cluster of Cassiar suite intrusions (Fig. 3.7). Errors associated with isotopic fractionation or the measurement of 204Pb cannot explain these trends towards more lead radiogenic compositions on the uranogenic plot (although the trend of these samples do sub-parallel the fractionation and error lines on the thorogenic plot; Fig. 3.7a,b). Lead from a radiogenic source may have been added to these samples, perhaps by contamination or assimilation processes during magma ascension and emplacement (Fig. 3.7). Most of the samples that plot on the Shale Curve (circled; Fig. 3.7) are from the Cassiar batholith, which may have largely escaped post-emplacement contamination with respect to radiogenic lead because of the large size of this intrusive body (Figs. 3.3a and 3.4). Linear sample arrays on the uranogenic and thorogenic plots that do not parallel fractionation and 204Pb error lines may have geochronological or petrogenetic significance33. The Rancheria and Black River samples plot as a linear array that does not parallel the upper crustal Shale Curve in Figure 3.7. The slope34 of this linear array on the uranogenic plot (indicated by the red dotted line; Fig. 3.7a) corresponds to an isochron age of ~2.4 Ga, although 207Pb/204Pb for these A-type suites is too low to be compatible with a (near-) Archean lead source. Instead, the linear array of lead compositions exhibited by the A-type plutons appears to be petrogenetically controlled, and probably represents mixing of lead from two distinct reservoirs. Despite radiogenic 206Pb/204Pb, the A-type plutons are only slightly more radiogenic than enriched lithospheric mantle and (or) depleted asthenospheric mantle with respect to 207Pb/204Pb (Fig. 3.7a), and they form a linear trend towards lower crustal lead compositions or another thorium depleted reservoir on the thorogenic plot (Fig. 3.7b). The less radiogenic (206Pb/204Pb and 207Pb/204Pb) lead reservoir defining this inferred mixing trend could be compatible with a mafic lower crustal, or even a mantle melt source for 32 All muscovite-bearing (and one hornblende-bearing) samples from the Cassiar batholith. 33Assumptions regarding initial lead isotopic compositions, U/Pb and Th/Pb ratios, and a closed system are also required in the interpretation of sample arrays in plots involving common lead (e.g., Faure, 2001). 34 The slope of this line (0.1543) is defined by 3 samples from the Rancheria and Black River plutonic suites, and one duplicate sample for the Rancheria plutonic suite. 177 the A-type plutons. The trend towards more radiogenic 206Pb/204Pb would, however, be the result of interactions (i.e., assimilation or contamination) between these magmas and upper crustal (meta)sedimentary rocks or deeper Paleoproterozoic gneisses (see insets; Fig. 3.7a,b). Lead isotopic contamination of the A-type magmas (or of their protolith) might have been facilitated if the original isotopic reservoir and (or) the magmas derived from this reservoir were relatively uranium- and thorium-poor. Such magmas may be relatively easily contaminated by radiogenic lead-rich crust during ascension and (or) emplacement. With this consideration in mind, the mafic melt source suggested for the A-type plutonic suites in Chapter 2 is not necessarily incompatible with their lead isotopic compositions as magmas derived from a U–Th-poor mafic reservoir may be easily contaminated by lead-rich crust. 3.4.1.5 207Pb/206Pb versus 208Pb/206Pb The 207Pb/206Pb versus 208Pb/206Pb plot (Fig. 3.7c) eliminates the error associated with the measurement of 204Pb, limits the effects of lead isotopic fractionation during analysis, and allows the data to be examined simultaneously in both uranogenic and thorogenic space. This comparison also allows populations of granitoids and mixing between lead reservoirs to be better distinguished. In Figure 3.7c, the Cassiar plutonic suite forms a broad array of samples that sub-parallels the upper crustal Shale Curve, implying that the main melt source for the Cassiar plutonic suite was a reasonably well-homogenized or well-mixed crustal reservoir (represented by the circled samples in Fig. 3.7) that evolved to modern lead compositions. These data are compatible with a mid- to upper crustal, clastic metasedimentary protolith for the majority of the Cassiar suite magma(s), although some input from a more juvenile melt source implied by two samples on the uranogenic plot (Fig. 3.7a) may also be important. The Late Cretaceous to Eocene A-type granites still form a linear trend in Figure 3.7c, which reinforces previous interpretations that these plutons plot along a mixing line between two lead reservoirs. This array also has a shallower slope than the broader sample array exhibited by the Cassiar plutonic suite, which supports the hypothesis that the Cassiar suite was derived from a different melt source than the A-type suites. The trend towards more radiogenic lead for the Cassiar suite, the A-type suites, and even some of the Quaternary to Recent basalt samples in Figure 3.7c is probably a result of the addition of radiogenic (upper crustal) lead through contamination or assimilation in the melt source, during ascension, or (in the case of the plutonic rocks) at the level of magma emplacement. 3.4.2 Stable Isotopic Datasets Unfortunately, there are very few stable isotopic datasets in the northern Cordillera with which to compare the results of this study. Limited whole rock oxygen (± Sr) isotopic ranges for the Whitehorse, Carmacks, and Coast Plutonic Complex magmatic events to the west (LeCouteur and Tempelman–Kluit, 1976, Dagenais, 1984; Morris and Creaser, 2008) are plotted with data from this study and considered in the following discussion. Oxygen– neodymium data ranges for the non-subduction-related A-type Surprise Lake plutonic suite (Dagenais, 1984; Morris 178 and Creaser, 2008) and sulfur–strontium isotopic compositions for the Kuhn Stock (adjacent to the Storie or Cassiar pluton; Fig. 3.3b; Cook and Godwin, 1984)―thought to belong to the Rancheria plutonic suite on the basis of age and location―are also plotted. 3.4.2.1 δ18O Oxygen isotopic compositions in magmatic rocks reflect those of the magma source(s) (e.g., Rollinson, 1993). The high positive oxygen isotopic compositions (commonly referred to as ‘heavy’) exhibited by all four plutonic suites in the Cassiar Terrane (Figs. 3.5c and 3.8a)35 require that the light isotope (16O) was depleted in the source material. 16O is lacking in rocks containing diagenetic, authigenic, or alteration minerals formed in equilibrium with surface waters (typically seawater; e.g., Taylor and Sheppard, 1986). Granitoids derived by the anatexis of supracrustal rocks, for example, have δ18O values greater than +8‰ (Ishihara, 2004)—although δ18O for plutonic rocks derived from infracrustal melt sources contaminated by supracrustal rocks can have similar values (Taylor and Sheppard, 1986). Altered oceanic ‘basalts’, or magmas derived from such rocks, are another example in which high positive δ18O (+6.5 to +10‰) is found in igneous rocks (e.g., Taylor and Sheppard, 1986). Although protolith compositions are the primary control on δ18O (Taylor, 1980; Rollinson, 1993), slow cooling and (or) extensive fractionation from basaltic to rhyolitic compositions help to increase plutonic δ18O, but only by several per mil (e.g., Taylor and Sheppard, 1986). The high positive oxygen isotopic compositions (i.e., > +10.6‰; Fig. 3.5) that characterize all four plutonic suites in the Cassiar Terrane, therefore, are most compatible with a supracrustal protolith. Globally, there appears to be upper δ18O limit of about +14.5 for quartz from plutonic rocks. Extensive assimilation of, or contamination by, crust with high positive oxygen isotopic compositions, however, can produce an increase of up to 4‰ in δ18O for small intrusions (plugs, dykes), or within 300 m of plutonic margins for larger intrusions (Clemens and Wall, 1981; Taylor and Sheppard, 1986). Three of the four plutonic suites in the Cassiar Terrane have one or more samples with a δ18O of +14.5‰ or more (Figs. 3.5c and 3.8a). An examination of the samples with δ18O greater than +14.5‰ (Table 3.2; Fig. 3.5c) demonstrates that, with the exception of samples KR-07-02 and KR-07- 24 from the Quiet Lake and Cassiar batholiths, respectively, all of these samples with high positive δ18O were collected either near plutonic margins or from very small intrusions (Fig. 3.4). This is consistent with a role for host rock contamination in imparting high positive oxygen compositions to syn- to post-accretionary intrusions in the Cassiar Terrane (e.g., Dagenais, 1984). Broad correlations between δ18O and 87Sr/86Sr100, εNd100, and 208Pb/206Pbdemonstrate that the oxygen isotopic system is coupled to the radiogenic isotopic systems (Figs. 3.9a and 3.10a) 36. Oxygen isotopic compositions that vary systematically with other isotopic or geochemical parameters are likely inherited from source rocks containing a mixture of isotopically contrasting materials (Taylor, 1980; Clemens et al., 2011). As oxygen isotope values 35 Note that Figures 3.9 and 3.10 do not show the full range of stable isotopic compositions because there are not always other complementary isotopic data available for each sample; see Figure 3.5c and Table 3.2 for the full range of stable isotopic compositions. 36 A very similar trend is seen for 207Pb/206Pb versus δ 18O. 179 increase, radiogenic isotopic compositions become more evolved, both of which suggest that supracrustal rocks are the source of the more evolved isotopic signature. Most of the samples plot in the “greywacke” field (after Rollinson, 1993) with respect to oxygen and strontium compositions (Fig. 3.10a). Although positively coupled to strontium, δ18O above the greywacke field (Fig. 3.10a) are not particularly enriched in radiogenic strontium (i.e., have low Rb/Sr), which is compatible with isotopically ‘heavy’ oxygen sourced from carbonate rocks in the melt source (e.g., Bernard–Griffiths et al., 1991). Three quartz analyses for muscovite-bearing samples of the Cassiar plutonic suite, however, form a shallow, positively sloping trajectory towards the slightly ‘heavier’ δ18O and more radiogenic 87Sr/86Sr100 that characterize metapelitic rocks from the southeast Canadian Cordillera (Figs. 3.9a and 3.10a; Bowman and Ghent, 1986; Kyser et al., 1994). Neodymium and lead isotopic compositions do not distinguish between input from different melt source lithologies as clearly as 87Sr/86Sr100 does because Sm–Nd and U–Th are not as easily fractionated by lithological or mineralogical variations. The negative correlation between 208Pb/206Pb and δ18O in Figure 3.9a, however, is consistent with the influence of calcareous lithologies in the protolith (cf. Tosdal et al., 1999). Overall, the oxygen isotopic patterns exhibited by the Cassiar plutonic suite in Figures 3.9a and 3.10a suggest derivation primarily from quartzo–feldspathic lithologies, with input from both pelitic and calcareous rocks. There are no oxygen isotopic constraints for middle to lower crustal rocks in the Cassiar Terrane in the northern Cordillera. Ortho- and paragneissic rocks, including sheared igneous rocks at the base of the Windermere Supergroup, exhumed in the Malton Complex of southeast British Columbia (Vernon–Chamberlain, 1983) provide the one of closest proxies for the age and composition of this basement. Neoproterozoic gneisses from the Malton Complex have radiogenic 87Sr/86Sr100 (0.714–0.809) and crustal δ18O values (+6.2 to +10.9‰), whereas Archean to Paleoproterozoic orthogneisses in the Malton Complex have evolved 87Sr/86Sr100 (0.716–0.819) and a wider range of crustal to mantle δ18O values (+4.7 to +10.1‰; Fig. 3.10a; Vernon–Chamberlain, 1983). If the basement to the Cassiar Terrane is isotopically similar to the predominantly orthogneissic basement in southeast British Columbia, it is not the source of the high positive oxygen isotopic compositions exhibited by the syn- to post-accretionary plutonic rocks. Most of the syn- to post-accretionary igneous rocks in the Cassiar Terrane are hosted in upper crustal, miogeoclinal rocks of Neoproterozoic to Triassic age (Gordey and Makepeace, 2001), which have a whole rock δ18O range of +5.7 to +12.7‰ (Fig. 3.10a; Dagenais, 1984). Higher δ18O is be typical of chemically precipitated lithologies (e.g., Neoproterozoic to Lower Cambrian carbonate rocks in the Cassiar Terrane; δ18O ≈ +17 to +18‰; Dagenais, 1984) and quartz-rich lithologies (e.g., phyllitic quartzite of the Neoproterozoic Hyland Group, northwest Selwyn Basin; δ18O = +13.8 to +16.5‰; whole rock values; Mair et al., 2006a), as is evident in Figure 3.10a. Since the heavy oxygen isotope (18O) concentrates in later crystallizing minerals, δ18O is also 1–2‰ higher in quartz separates than in whole rock powders37, depending on the lithology and temperature of crystallization (Taylor and Sheppard, 1986). Thus, the range of δ18Oqtz for granitoids in the Cassiar Terrane largely overlaps with the range of δ18OWR for upper crustal miogeoclinal rocks (Table 3.2; Fig. 3.10a38). The miogeoclinal rocks are an unlikely melt source given 37 The 1–2‰ difference between quartz and whole rock oxygen isotopic compositions is evident when δ18O from plutonic whole rock and quartz samples are plotted together (Fig. 3.5c). 38 Note that Figures 3.9 and 3.10 do not show the full range of stable isotopic compositions given that there are not 180 their widespread (sub)-greenschist facies metamorphic grades (e.g., Greenwood et al., 1991; Gordey and Anderson, 1993) and position in the upper crust (e.g., Cook et al., 2004). They were, however, derived by reworking of older basement (Gabrielse and Campbell, 1991; Gehrels and Ross, 1998; Ross et al., 2005; Lemieux et al., 2007; Beranek, 2009; Leslie, 2009), so their oxygen compositions may also be representative of deeper metasedimentary or crystalline rocks. It is also possible that the miogeoclinal rocks are present at greater depths in the northern Cordillera (e.g., Evenchick et al., 2005). The limited isotopic data available for the Rancheria plutonic suite generally overlap with range of isotopic compositions exhibited by the Cassiar plutonic suite (Fig. 3.9a). This implies that the composition of the protolith rocks also controls the oxygen isotopic composition of the Rancheria suite. Given that data are available for only three Rancheria suite samples, however, it is still possible that the protolith to this plutonic suite differs from the metasedimentary magma source suggested for the Cassiar suite. The Seagull plutonic suite (for which only O and Sr data are available) has a near-vertical slope in Figure 3.10a. The extreme difference between the limited variation in 87Sr/86Sr100 and more extensive variation in δ18O for the Seagull suite implies that these two isotopic systems are not coupled. This suite had a protolith with a limited range of low strontium isotopic composition, and a large range of oxygen isotopic compositions (unless the A-type magmas were contaminated by rocks with high positive oxygen and low strontium isotopic compositions). The former might be consistent with a low-temperature hydrothermally altered, and isotopically juvenile protolith for the Seagull plutonic suite. One possible magma source for the A-type plutons in the Cassiar Terrane, therefore, is oceanic crust of mainly, if not entirely, infracrustal origin (since most supracrustal oceanic rocks would also contribute radiogenic Sr). 3.4.2.2 δ34S Sulfur isotopes are more sensitive than oxygen and radiogenic isotopes to contamination by supracrustal lithologies, particularly sulfur-rich rocks like evaporitic, organic-rich (coaly, graphitic, or pyritic), and—to a lesser degree— carbonate lithologies (Wedepohl, 1969–1978; Ohmoto et al., 1990). Sulfur contamination is particularly noticeable if the parental magma was derived from a sulfur-poor melt source, as is the case for igneous rocks originating in the mantle (e.g., Ionov et al., 1992). The addition of even the smallest quantities of a relatively sulfur-rich rock to a sulfur-poor magma will, therefore, largely control its sulfur content and isotopic composition. This contamination may not noticeably affect the oxygen composition of the magma; therefore, these two stable isotopic systems are commonly decoupled. In the Cassiar Terrane, δ34S for the granitoids is not correlated to δ18O, either between the four plutonic suites or within individual plutonic suites (Fig. 3.8a). This lack of O–S coupling is compatible with variable additions of sulfur to all of the magmas from sulfur-rich lithologies. Intrusions in the Cassiar Terrane have low sulfur concentrations for granitic rocks (e.g., Wallace, 2001, 2005), ranging from 3–110 ppm, and typically below 20 ppm (Table 3.2; Fig. 3.8b). The Seagull, Rancheria, and Black River plutonic suites have some of the lowest sulfur concentrations (<11 ppm), although sulfur content does appear always complementary data for other isotopic or geochemical datasets available for all plutonic samples; see Figure 3.5c and Table 3.2 for the full range of stable isotopic compositions from this study and compiled data. 181 to increase slightly with δ34S if these three A-type suites are considered as a group (Fig. 3.8b). Cassiar suite samples with low δ34S also have similarly low sulfur contents (Fig. 3.8b), although when δ34S exceeds +7‰ in the Cassiar suite sulfur concentrations increase to 66 ppm (Fig. 3.8b). The low sulfur concentrations in all of the plutons that were examined in the Cassiar Terrane are consistent with a derivation from a sulfur-poor melt source, and with limited interactions between the magma and isotopically ‘heavy’ sulfur-rich rocks in the crust. As sulfur concentrations increase during contamination, the sulfur isotope composition becomes slightly higher (Fig. 3.8b), indicating that a sulfur source with a high positive isotopic composition is responsible for the broad trend towards higher δ34S in both the Cassiar and A-type suites. Muscovite-bearing phases of the Cassiar plutonic suite typically have lower sulfur contents than biotite-bearing phases (Fig. 3.8a,b). Sulfur concentrations, however, may be linked to the magnetic susceptibility of the sample, and the sulfur-poorer muscovite-bearing magmas are also slightly more reduced than the biotite-bearing magmas as a group (Fig. 3.8c). The differences in muscovite- and biotite-bearing samples are probably a result of the lower sulfur solubility in reduced magmas (Poulson and Ohmoto, 1990; Poulson et al., 1991; Yang, 2012) rather than differing petrogenetic histories for these mineralogical variations of the Cassiar plutonic suite. The Cassiar suite samples with the heaviest sulfur isotopic compositions (δ34S > +7‰) are weakly correlated to more evolved 87Sr/86Sr100, ɛNd100, and lead (e.g., 208Pb/206Pb) isotopic compositions—although not as tightly as the correlation between δ18O and the radiogenic isotopes (Figs. 3.9 and 3.10). This suggests that either sulfur is sourced from the protolith rocks, or the radiogenic isotopic compositions were also affected by contamination or assimilation during ascension and emplacement of the magma. Samples with the lowest sulfur isotopic composition (e.g., δ34S < +7‰) from the Cassiar plutonic suite include the Cabin Creek and Marker Lake plutons, as well as two samples from the Cassiar batholith (KR-05-16, -24; Fig. 3.8b), and are also relatively sulfur-poor (Fig. 3.9b).39 The Cassiar suite samples with low positive δ34S plot within the sub-vertical sample arrays defined by the A-type suites (Fig. 3.9b), and immediately below the upper perturbation in 87Sr/86Sr100 versus εNd100 space where the Rancheria (and possibly Seagull) plutonic suite also plots (Fig. 3.6a). These plutonic phases were sampled in the vicinity of the Rancheria suite samples (Fig. 3.4), which have similarly low sulfur isotopes and concentrations (Fig. 3.9b). The simplest explanation for the range of sulfur isotopic compositions exhibited by the Cassiar suite is, therefore, that most of the intrusions with δ34S below +7‰ were derived from sulfur-poor rocks with moderately evolved neodymium and strontium isotopic compositions. The high positive sulfur isotopic compositions of most of the Cassiar suite samples (δ34S > +7‰) may have been derived from an isotopically ‘heavy’ protolith with more evolved neodymium and strontium isotopic compositions. Basement sulfur isotopic compositions vary so widely that they are not useful in identifying specific sources of magmatic sulfur in the Cassiar Terrane (Fig. 3.10b). Metaigneous basement (e.g., felsic basement orthogneisses in the Malton Complex, southeast British Columbia; Vernon–Chamberlain, 1983) can be ruled out as a possible source 39 The relative sulfur-enrichment in the Marker Lake sample is anomalous compared to the other Cassiar suite samples with isotopically light sulfur compositions. There are is more than one possible reasons for this, including: contamination by a sulfur-rich, but isotopically light reservoir (e.g., pyritic shales with low positive to negative δ34S), and (or) the presence of hydrothermal sulphides not detected by hand specimen and thin section examination. 182 of sulfur due to the low sulfur contents in intermediate–felsic rocks (e.g., Wedepohl, 1969–1978; Wallace, 2001, 2005; DeYoung, 2007; Yang, 2012). Instead, the sulfur and strontium isotopic trends exhibited by the Cassiar suite samples are much more consistent with a sulfur-rich supracrustal lithology. Typically pyritic shales have negative δ34S (Rollinson, 1993), however, in the northern Cordillera bacterial reduction of sulphate in an anoxic environment produced a wide range of negative to positive δ34S values for pyrite in Neoproterozoic to Paleozoic miogeoclinal rocks (e.g., –30 to +35 ‰; Goodfellow et al., 1987; Ross et al., 1995). The slight positive correlation between δ34S and 87Sr/86Sr100 for most of the Cassiar suite samples (i.e., those with δ34S > +7‰) is consistent with sulfur contamination by a clay-rich (pelitic) lithology with high positive δ34S overall. A steeper array, or even a negative correlation, would be expected in Figure 3.10b if sulfur was derived from a strontium-rich lithology (e.g., carbonate and evaporite rocks, which have low 87Sr/86Sr100; Bernard–Griffiths et al., 1991; Rollinson, 1993). It is unlikely that the supracrustal source rocks for the Cassiar plutonic suite contained abundant sulphides, given the low sulfur concentrations in samples from this suite. However, the possibility of minor pyritic layers in a relatively strontium-radiogenic metasedimentary lithology, such as a metapelite, is a viable source of at least some of the sulfur responsible for the high positive sulfur isotopic composition of the Cassiar suite plutons. The oxygen and sulfur isotopic compositions, therefore, were derived from distinct lithologies in the magma source. δ18O was sourced mainly from isotopically ‘heavy’ quartz-rich (and, in some cases, calcareous) lithologies, and δ34S reflects input from isotopically ‘heavy’ clay-rich lithologies containing minor pyrite. The decoupling between δ18O and δ34S, and the poorer coupling between δ34S and the radiogenic isotopic compositions for the Cassiar plutonic suite, however, suggest that some sulfur was also added by contamination or assimilation involving pyritic upper crustal rocks during the ascension and emplacement of these magmas. When considered as a group, the A-type samples plot as linear sample arrays40 when radiogenic isotopic compositions are compared to δ34S (Fig. 3.9b). As in lead isotopic space (Fig. 3.7), the orientation of the sample arrays defined by A-type plutons does not always parallel trends exhibited by the Cassiar plutonic suite (e.g., Sr, Pb; Fig. 3.9b), allowing for a different petrogenetic interpretation for the A-type plutons. Melt from mafic crust (i.e., underplated basalt or oceanic crust) or the mantle may be inferred on the basis of low strontium isotopic compositions to have been the main source of the A-type plutonic suites. The mantle, and its derivative melts, are typically characterized by low sulfur concentrations and low positive sulfur isotopic compositions (20–50 ppm S and δ34S = 0 ± 3; Chaussidon and Lorand, 1990; Ionov et al., 1992). Several of the A-type samples, particularly those that are the most sulfur-depleted, have some of the lowest sulfur compositions observed in the Cassiar Terrane (e.g., the Seagull and Rancheria plutonic suites), although as a group these suites have a wide range of δ34S (Fig. 3.8b). The large range of δ34S, combined with a small range of 87Sr/86Sr100 are more suggestive of sulfur contamination by supracrustal rocks with high positive δ34S, during ascension and emplacement of the A-type magmas. 40 If the linear trend exhibited by all A-type suites with respect to δ34S and 87Sr/86Sr100 is representative of the relationship between δ34S and radiogenic isotopic compositions, then extrapolation of the linear sample array of the Rancheria and Black River samples to the δ34S range of the Seagull plutonic suite suggests that the Seagull plutonic suite could have an approximate εNd100 range of –5 to –7 and 208Pb/206Pb range of 2.01–2.02. 183 3.4.3 Isotopic Comparisons to Coeval Intermediate–Felsic Intrusive Rocks in the Northern Cordillera 3.4.3.1 The Northern Cordillera Widespread Early Cretaceous to Eocene magmatism emplaced into the Yukon–Tanana Uplands in east–central Alaska (Fig. 3.1) is similar in age to Early Cretaceous intrusions in the Cassiar Terrane (e.g., Mortensen et al., 2000; Hart et al., 2004a; Mortensen and Hart, 2010). Prior to ~430 km of Cenozoic offset along the Tintina Fault, this region would have been located north of the Cassiar Terrane (i.e., west of the Selwyn Basin) in the Early Cretaceous. Early Cretaceous magmatism in east–central Alaska comprises the Yukon–Tanana Uplands plutonic suite (Fig. 3.3a; Hart et al., 2004a; Dilworth et al., 2007) and is known to range from 109–103 Ma (but is as old as 115 Ma; Aleinikoff et al., 1984). ‘Mid’-Cretaceous magmatism in east–central Alaska from about 94–90 Ma comprises the Fairbanks–Salcha plutonic suite (Fig. 3.3a; Aleinikoff et al., 2000; Dilworth et al., 2007; Hart et al., 2004a). Plutonic rocks in east–central Alaska were emplaced at ~1.5–2.5 kbar as sheets or sills that parallel a shallow-dipping fabric created during extensional ductile deformation (Dilworth et al., 2007). Small occurrences of felsic and bimodal magmatism in this region are poorly constrained at ~87–65 Ma and ~60–50 Ma, respectively. It is important to note that exotic components of east–central Alaska (e.g., Triassic to Jurassic arc-related rocks, and over-thrust oceanic klippen) that are mapped at the surface are thought to be limited to the uppermost crust only; North American rocks of the continental margin are believed to comprise most of the unexposed basement in the Yukon–Tanana Uplands. Although the age and composition of this Laurentian basement is unknown, there is no reason to suppose it is substantially different than unexposed basement in the Cassiar Terrane. Plutonic lead isotopic compositions in the Cassiar Terrane generally overlap with data for coeval intrusive rocks in east–central Alaska (Fig. 3.11a; Appendix G.3), implying similar basement below both regions. The bulk of samples for Early Cretaceous magmatism in east–central Alaska, however, have higher 207Pb/204Pb than most of the coeval Cassiar suite―typically plotting above the upper crustal Shale Curve―and higher 206Pb/204Pb than about half of the Cassiar suite samples. ‘Mid’-Cretaceous, Late Cretaceous, and Paleocene plutons in Alaska have lower 206Pb/204Pb than coeval suites in the Cassiar Terrane, ranging from compositions similar to those of the enriched mantle to the upper crustal Shale Curve. Dilworth et al. (2007) interpreted the difference between Early and ‘mid’-Cretaceous plutonic lead compositions to reflect a derivation from old continental crust (or siliciclastic rocks derived from such crust) for the older plutons. Less evolved lead isotopic compositions for ‘mid’-Cretaceous intrusions, therefore, were interpreted to reflect the input of a larger proportion of magma derived from a more mafic and primitive lead source (either a mantle-derived, or a mafic crustal, component). Late Cretaceous and Paleocene plutons continue the trend towards less radiogenic lead compositions with younger magmatism (Aleinikoff et al., 2000). These trends were also reflected in the whole rock composition of the plutons, with peraluminous older intrusions and increasingly metaluminous younger intrusions (Aleinikoff et al., 2000). The plutonic rocks in east–central Alaska also become progressively less evolved with respect to neodymium and strontium isotopic compositions, from the Early to Late Cretaceous (Fig. 3.12a; Appendix G.3). Like several of the most isotopically evolved samples of the Cassiar suite, 147Sm/144Nd in Early Cretaceous samples of the Yukon–Tanana Uplands suite also trend towards anomalously 184 elevated values (Fig. 3.12a; Appendix G.3). Given the anomalous 147Sm/144Nd, it is likely that the only estimates of TDM (Fig. 3.12a; Appendix G.3) are too old to be useful for these samples. Limited oxygen isotopic compositions from plutonic quartz suggests that δ18O is slightly more primitive in Late Cretaceous samples (+8.1 to +9.8 ‰) than it is in Early and ‘mid’-Cretaceous samples (+8.4 to +11.1‰; Appendix G.3) in the Yukon–Tanana Uplands. Both of these ranges are, however, lower than the δ18O range exhibited by the majority of the Cassiar plutonic suite (Figs. 3.5c and 3.8a; Table 3.2). Early to Late Cretaceous intrusions in east–central Alaska are hosted mainly by a pre–Devonian package of Laurentia-derived continental margin siliciclastic rocks―the Lake George assemblage, similar to mainly Paleozoic miogeoclinal Selwyn Basin sedimentation in Yukon―that was intruded by peraluminous granitoids of Devono– Mississippian age (e.g., Dusel–Bacon et al., 2002, 2006; Colpron et al., 2006; Piercey and Colpron, 2009). Early Jurassic over-thickening of these rocks related to compressional deformation (e.g., Dusel–Bacon et al., 2002) is interpreted to have led to the convective removal of a gravitationally unstable, over-thickened mantle lithosphere (cf. Houseman et al., 1981), and a slow rise in the crustal geothermal gradient until the Early Cretaceous (Mair et al., 2006b). The result was thermally weakened crust and regional uplift that drove ductile extension and cooling (135– 108 Ma; Dusel–Bacon et al., 2002), exhumation of regional amphibolite-grade rocks, and extensive Early Cretaceous crustal anatexis (Pavlis et al., 1993; Mair et al., 2006a; Berman et al., 2007). Thus, in east–central Alaska, Early Cretaceous intrusions were emplaced during a period of widespread extension and uplift (e.g., Dilworth et al., 2007). Aleinikoff et al. (2000) suggested that the Late Proterozoic to Paleozoic host rocks (the Lake George assemblage) were possible sources of isotopically evolved magmas of the Early Cretaceous magmatic event (and isotopically evolved contributions to the later magmas). Dilworth et al. (2007) instead argued for a melt source in the older and deeper Laurentian basement that is thought to underlie this region. Late Cretaceous to Eocene within-plate (A-type) granitoids are consistent with emplacement during transtensional tectonism, and followed by rapid cooling and exhumation of the Yukon–Tanana Uplands in the Eocene (e.g., Aleinikoff et al., 2000; Dusel– Bacon and Murphy, 2001; Dusel–Bacon et al., 2002; Dilworth et al., 2007). 3.4.3.2 The Southern Cordillera The White Creek, Bugaboo, Fry Creek, and Horsethief Creek intrusions in southeast British Columbia are also Early Cretaceous in age (115–106 Ma; Brandon and Lambert, 1993, 1994). Plutons in southeast British Columbia were emplaced at pressures of 2.3–3.8 kbar and are hosted by (sub-)greenschist grade rocks of the Mesoproterozoic Purcell Supergroup in the Purcell Anticlinorium. The plutons are composed of multiple intrusive phases spanning a similar range of compositions as the Cassiar plutonic suite (e.g., slightly metaluminous hornblende–biotite quartz monzodiorite and granodiorite, to peraluminous biotite–muscovite granite; Brandon and Lambert, 1993, 1994). They have historically been grouped together as the Bayonne plutonic suite (e.g., Woodsworth et al., 1991); herein, however, they are referred to as plutons in the Purcell Anticlinorium of southeast British Columbia. Compared to feldspar lead compositions for the Cassiar plutonic suite, whole rock lead isotopic compositions for the Early Cretaceous Purcell Mountains intrusions vary widely (Fig. 3.11b; Appendix G.3). The whole rock lead 185 isotopic compositions for the intrusive rocks in the Purcell Mountains have undoubtedly been disrupted by lead mobilization, which has resulted in the widely variable compositions. 87Sr/86Sr100 and εNd100 are less likely to have been disrupted in the intrusive rocks, and provide a better basis for compositional comparisons between Early Cretaceous magmatism in the Purcell Mountains and the Cassiar Terrane. The remarkable similarity between plutonic rocks in the Purcell Anticlinorium and the Cassiar Terrane is somewhat surprising considering the distance separating these two regions. Furthermore, samples from southeast British Columbia appear to parallel (and even extend) the isotopic trends towards radiogenic strontium that emanate from two perturbations in the Cassiar suite array in Figure 3.12b. Several Fry Creek samples, for example, stem from the lower perturbation exhibited by the Cassiar suite, and trend towards more evolved 87Sr/86Sr100 (circled; Fig. 3.12b). In southeast British Columbia, the least isotopically evolved samples are the hornblende-bearing phases of the White Creek batholith and biotite- bearing phases mainly of the Bugaboo batholith (Fig. 3.12b; Appendix G.3). The most isotopically evolved samples are aplitic or pegmatitic dykes and “two mica” granitic phases (Fig. 3.12b; Appendix G.3). It should be noted the most isotopically evolved samples in the Purcell Mountains have anomalously high 147Sm/144Nd (Fig. 3.12b; Appendix G.3). Given the anomalous 147Sm/144Nd, the estimates of TDMs for several of the muscovite-bearing phases (Fig. 3.12b; Appendix G.3) are likely anomalously old estimates of the average crustal residence age. Oxygen isotopic compositions for quartz from Cretaceous intrusions in the Purcell Mountains increase slightly from +10.0 to +10.2‰ in the least siliceous intrusive phases, to +10.3 to +11.2 ‰ in the most siliceous intrusive phases (Appendix G.3; Brandon and Lambert, 1994). The latter are within, but at the low end of the range of δ18O exhibited by the Cassiar plutonic suite (Figs. 3.5c and 3.8a; Table 3.2). Early Cretaceous plutons in southeast British Columbia are ascribed an entirely crustal melt source (Brandon and Lambert, 1993, 1994). The geochemical and isotopic variations in plutons in the Purcell Anticlinorium were attributed to partial melting of interlayered mafic and tonalitic gneisses similar to exhumed crystalline basement in southeast British Columbia (e.g., the Malton complex; Vernon–Chamberlain, 1983) to produce the least siliceous magmas (hornblende and biotite-bearing quartz monzodiorite and granodiorite). Partial melting of Meso- to Neoproterozoic pelitic rocks in overlying metasedimentary packages produced the most siliceous magmas (“two mica” granites). Thickening of the crust up to 60 km in southeast British Columbia was thought to have been important in generating biotite and muscovite dehydration reactions about 10–40 m.y. following the attainment of maximum crustal thickness during terrane accretion and thickening. It was noted, however, that the less siliceous magmas would have required amphibole dehydration in the mafic to tonalitic rocks, which implied greater heat than what is attained by crustal thickening alone (Brandon et al., 1994). The absence of associated mafic plutonic rocks implied that the additional heat was not in the form of mantle-derived magmas intruding the crust, so a higher crustal geothermal gradient (related to higher heat production in the basement), thermal preconditioning, or basaltic underplating were proposed as possible sources of the additional heat (Brandon et al., 1994). 186 3.5 DISCUSSION 3.5.1 Isotopic and Geochemical Variation 3.5.1.1 Sm/Nd, La/Sm, and Ba/Nb Trace element ratios provide complementary petrogenetic information regarding the type of protolith and depth of melting involved in the petrogenesis of a magma. The significance of Sm/Nd, for example, was discussed in Section 3.4.1.2 in terms of the overall composition of the melt source, although the fractionation of Sm–Nd during anatexis may also be an equally important control on Sm/Nd in a magma (and its neodymium isotopic composition; e.g., Ayres and Harris, 1997; Zeng et al., 2005). An incompatible element enrichment relative to the mantle characterizes the overall crust, but particularly the middle and upper crust (e.g., Rudnick and Gao, 2003). Crustal contributions to the genesis of a mantle-derived magma are commonly interpreted from elevated trace element ratios, such as La/Sm, La/Yb, La/Ta, and Ba/Nb (e.g., Thompson et al., 1982; Rollinson, 1993). The comparison of Sm/Nd, La/Sm, and Ba/Nb with neodymium and strontium isotopic compositions, produces relatively continuous, if scattered, arrays of data that trend between three apparent compositional end-members (calculated at 100 Ma for all samples): 1) a high Sm/Nd, and low La/Sm and Ba/Nb end-member with the least evolved isotopic compositions, which is characterized by the composition of Quaternary to Recent alkaline ‘basalts’ derived from an enriched mantle (Fig. 3.13a); 2) a low Sm/Nd, and high La/Sm and Ba/Nb end-member with a moderately evolved isotopic composition, which is characterized by several examples of the Cassiar plutonic suite and the most evolved continental arc samples (circled; Fig. 3.13a); and, 3) a high Sm/Nd, and low La/Sm and Ba/Nb end-member with very highly evolved isotopic compositions, which is characterized by the “migmatitic” sample (indicated by the grey filled box with a black “x”; Fig. 3.13a) and other highly evolved leucogranitic examples of the Cassiar plutonic suite (samples with anomalously high 147Sm/144Nd are outlined with dashed blue line; Fig. 3.13a). The approximate εNd100 and 87Sr/86Sr100 values for the upper and lower Cassiar suite perturbations from Figure 3.6a are indicated by grey arrows (next to the x-axis; Fig. 3.13a). The first isotopic end-member represents the composition of the underlying enriched mantle, which is the source of the Quaternary to Recent basaltic rocks (Abraham et al., 2001). The other two signatures reflect two reservoirs of probable crustal origin given their relatively evolved compositions. The first of these broadly corresponds to the upper perturbation (Fig. 3.13a) and is referred to here as “isotopically moderately evolved”. The second of these is referred to here as “isotopically very highly evolved”. A large proportion of the Cassiar plutonic suite samples plot between the two ‘crustal’ compositions and will are referred to here as “isotopically highly evolved” (outlined by the black box; Fig. 3.13a). Although the data as a whole appear to form a continuous array between the two crustal end-member compositions, the trace element ratios for the isotopically highly evolved samples do not vary systematically (Fig. 3.13a). 187 Trace element trends relative to stable isotopic compositions are scattered (Fig. 3.13b), but this is probably partly because the majority of samples analyzed for stable isotopic composition are of the isotopically highly evolved group (where a lack of systematic variation was also observed in Fig. 3.13a). Exceptions include the lowest δ18Oqtz value, which is from an isotopically juvenile sample (labeled in Fig. 3.13b), and limited δ18O and δ34S for the very isotopically highly evolved samples (circled with dashed blue lines; Fig. 3.13). The lack of systematic trace element variation with respect to oxygen and sulfur isotopic composition is, therefore, similar to the lack of systematic radiogenic isotopic variation exhibited by the same samples in Figure 3.13a. Samples from the A-type plutonic suites (Seagull, Rancheria, and Black River) are only distinct from the Cassiar plutonic suite in Figure 3.13a with respect to their consistently low Ba/Nb. Early Cretaceous and younger continental magmatic arc samples emplaced into the Yukon–Tanana terrane west of the Cassiar Terrane range from compositions that are typical of the enriched mantle underlying the Cassiar Terrane to the moderately evolved isotopic end-member (εNd100 ≈ –5 and 87Sr/86Sr100 ≈ 0.707) in Figure 3.13a. Although these arc-related samples must contain a significant amount of enriched mantle-derived magma, the low Sm/Nd, and high La/Sm and Ba/Nb that characterize the more evolved samples reflect an important contribution from the continental crust to some of these magmas. The observed coupling between an incompatible element enrichment and isotopic evolution is likely due to the incorporation of small, but highly enriched melt fractions from crustal rocks in contact with mafic mantle-derived magmas (e.g., Thompson, 1982). This process probably occurs at the base of, and within the lowermost crust (e.g., Hildreth and Moorbath, 1988; Annen et al., 2006), and is thought to drastically alter isotopic and trace element compositions, with very little associated change to the bulk magma composition (e.g., DePaolo et al., 1992; Schaltegger and Corfu, 1992; Davis and Hawkesworth, 1993; Reiners et al., 1995). An origin in enriched mantle for the arc-related intrusions is also consistent with their non-radiogenic 207Pb/206Pb and 208Pb/206Pb isotopic compositions (e.g., the Late Cretaceous Carmacks suite; Fig. 3.7) and low oxygen isotopic compositions (overlapping with mantle values; Fig. 3.9a). Like similar continental magmatic arc suites worldwide, and in accordance with models for the subduction of an oceanic plate and the resulting hydration and melting in the overlying mantle wedge and lower crust (e.g., Tatsumi et al., 1983; Hildreth and Moorbath, 1988; Annen et al., 2006), the parental magmas to the continental arc plutonic suites are sourced from the mantle (e.g., Roberts and Clemens, 1993). Several isotopically juvenile to moderately evolved samples in Figure 3.13a—including all of the hornblende- bearing samples, as well as several biotite or muscovite-bearing samples with εNd100 > –5 and 87Sr/86Sr100 < 0.71— are members of the Cassiar plutonic suite (circled; Fig. 3.13a). Based on the above discussion, this might suggest that the parental magmas to these Cassiar suite intrusions were also rooted in the mantle. As with most continental subduction zones, some component of lithospheric mantle-derived melt is also thought be involved in the genesis of some high-K calc-alkaline magmatic provinces, not unlike the Cassiar Terrane (e.g., the West African and Variscan orogenies; Wall et al., 1987; Roberts and Clemens, 1993; Liégeois et al., 1998; Barbarin, 1999; Ferré and Leake, 2001). The lack of abundant cogenetic mafic phases in the Cassiar suite, however, argue to the contrary. Furthermore, the difficulties inherent to magma mixing and crustal assimilation models for the production of voluminous intermediate to silicic granitoids (e.g., Clemens et al., 2009; Clemens and Stevens, 2012) are compatible 188 with a mixture of crust and mantle-derived magmas. The isotopically juvenile Cassiar suite intrusions probably have a dominantly, if not entirely, crustal origin even though none of the geochemical proxies for average crustal compositions in Figure 3.13 have such elevated La/Sm and Ba/Nb. These geochemically ‘enriched’ magmas may, therefore, have experienced a crustal anatectic (e.g., Clemens et al., 2009) or magmatic process that has enriched them in incompatible (light REE and LILE) element relative to even the upper crust. A protolith of intermediate bulk composition is consistent with the quartz monzonitic to granodioritic compositions exhibited by the least isotopically evolved members of this suite (Chapter 2; Roberts and Clemens, 1993; Clemens et al, 2011; Clemens and Stevens, 2012). Although the bulk of the Cassiar suite magmas are apparently of supracrustal origin (Chapter 2), the oxidized and alkaline to alkalic-calcic composition of the least isotopically evolved Cassiar suite samples (Chapter 2) could have been imparted from mixtures of melts from alkaline (ocean island basalt-, or OIB-, like) flows or sills interlayered with the supracrustal package (e.g., Driver et al., 2000). Isotopically highly evolved samples of the Cassiar plutonic suite outlined by the black box plot between the moderately evolved isotopic end-member (εNd100 ≈ –5 and 87Sr/86Sr100 ≈ 0.71)and the very highly evolved isotopic end-member (εNd100 ≈ –15 and 87Sr/86Sr100 ≈ 0.72; Fig. 3.13a). These biotite- and muscovite-bearing samples are characterized by moderate Sm/Nd, La/Sm, and Ba/Nb, and lack systematic trace element–isotopic variations (Fig. 3.13a). The combination of these trace element ratios, the isotopic compositions, and the intermediate–felsic composition of these intrusions are compatible with an entirely crustal origin for these members of the Cassiar plutonic suite (Clemens et al, 2011; Clemens and Stevens, 2012). In Chapter 2, these reduced and calc-alkaline examples from the Cassiar plutonic suite were interpreted be of supracrustal origin, particularly on the basis of geochemical heterogeneity (Chapter 2). The range in isotopic composition, however, requires that the supracrustal melt source contained a mixture of isotopically juvenile (mafic) and isotopically evolved (felsic) components. The very isotopically highly evolved samples (i.e., those with εNd100 > –15 and 87Sr/86Sr100 > 0.72) are characterized by higher Sm/Nd, and lower La/Sm and Ba/Nb, all of which overlap with trace element ratios in the enriched mantle-derived rocks (Fig. 3.13a). Most of the samples with anomalously elevated 147Sm/144Nd (circled with dashed blue line; Figs. 3.6b and 3.11) can be considered very isotopically highly evolved as well. Assuming these trace element ratios are not due to incomplete sample digestion, they might be interpreted to be indicative of a mafic or mantle melt source. A mafic protolith need not be invoked, however, if the REEs are fractionated as a result of disequilibrium between melt and restite during anatexis (e.g., Ayres and Harris 1997; Zeng et la., 2005; Farina and Stevens, 2011). Sm/Nd and La/Sm in intermediate to felsic rocks are controlled by light REE-bearing accessory minerals in the protolith. Low magmatic La/Sm implies that light REE-bearing minerals (e.g., monazite, allanite ) were less important reactants in the protolith, and (or) were less likely to be among the entrained phases in the extracted magma for the very isotopically highly evolved phases of the Cassiar plutonic suite (e.g., Ayres and Harris, 1997; Zeng et al., 2005). The extraction of a melt that has not fully equilibrated with the light REE-rich minerals in its source rocks creates an opportunity for middle REE-rich minerals (e.g., apatite) to control these trace element ratios (and the resulting Nd isotopic composition; Ayres and Harris, 1997; Zeng et al., 2005) in the magma. Late-stage magmatic fractional crystallization can also produce a similar effect (Ayres and Harris, 1997), although this is unlikely to be important except in the most highly fractionated rocks (e.g., aplite or pegmatite phases; 189 Clemens et al., 2009). Thus, the very isotopically highly evolved samples comprise magmas that were not in equilibrium with their isotopically evolved crustal protolith. The three A-type plutonic suites in the Cassiar Terrane are best distinguished from the Cassiar plutonic suite by their consistently low Ba/Nb, regardless of isotopic composition (Fig. 3.13). This is typical of A-type magmas, which are characteristically high field strength element (HFSE; e.g., Nb, Ta) enriched and LILE (e.g., Sr, Ba) depleted (Collins et al., 1982; Bonin et al., 1998; Frost and Frost, 2011; Chapter 2). Like the Cassiar plutonic suite, oxygen and sulfur isotopic compositions for the A-type suites do not display any definitive trends relative to their immobile element ratios (Fig. 3.13b). The only feature of note is that the Seagull and Black River samples have low and high sulfur isotopic compositions, respectively, but similar trace element ratios (Fig. 3.13b). Conversely, the Rancheria suite has a moderate sulfur isotopic composition and higher La/Sm and Sm/Yb (Fig. 3.13b). The generally low and restricted range of strontium isotopic compositions (despite anomalously elevated Rb/Sr) and low Ba/Nb (Fig. 3.13) in all three A-type suites is consistent with a less radiogenic melt source. 3.5.1.2 Sm/Yb Heavy REEs have higher partition coefficients for specific minerals, and magmatic Sm/Yb varies depending on the ferro-magnesian mineral(s) present when the magma was extracted from the protolith. Intermediate (quartzo– feldspathic) to mafic rocks are thought to comprise most of the middle to lower crust (e.g., Wedepohl, 1995; Rudnick and Gao, 2003), and different ferro-magnesian minerals are stable in these lithologies at different pressures. Plutonic Sm/Yb has been used to differentiate between melts extracted from pyroxene, hornblende, or garnet- bearing residual mineral assemblages, which in turn may allow estimates to be made regarding the composition of the melt source and even the approximate depth of partial melting (e.g., Hildreth and Moorbath, 1988). In most mafic to intermediate amphibolites, garnet will be stable at pressures that characterize the lower crust of over- thickened orogens (e.g., 12–15 kbar, or about 36–45 km depth; Rushmer, 1991). However, garnet is also a common mineral in aluminous amphibolitic to pelitic rocks metamorphosed at mid-crustal (amphibolite facies) conditions, and even garnet-zone greenschist-grade upper crustal rocks that are locally present in exposed Neoproterozoic rocks in the northern Cordillera (Greenwood et al., 1991). An additional complication is that clinopyroxene ± garnet are common peritectic minerals that form during the dehydration of hornblende ± biotite in the fluid-absent production of I-type magmas (Clemens et al., 2011), whereas garnet ± orthopyroxene are the main peritectic minerals resulting from the dehydration of biotite ± muscovite in the fluid-absent production of S-type magmas (Stevens et al., 2007). Clearly the presence or absence of garnet is not solely related to depth of partial melting; therefore, interpretations of Sm/Yb in intermediate to felsic (i.e., crustally derived) granitoids may be complex. In general, however, higher Sm/Yb is consistent with pre-anatectic or syn-anatectic (i.e., peritectic) garnet in the melt source that was not entrained in significant abundances in the magma. Lower Sm/Yb is consistent with either less garnet in the melt source (for metamorphic or compositional reasons) or greater entrainment of pre-anatectic (restitic) or syn-anatectic (peritectic) garnet in the magma. The Cassiar plutonic suite samples are generally scattered with respect to Sm/Yb, showing no well-defined trends 190 with isotopic composition (Fig. 3.13). The only correlation between Sm/Yb and isotopic composition for this suite is a poorly defined decrease in Sm/Yb with lower εNd100, higher 87Sr/86Sr100, and lower δ18O (Fig. 3.13). This suggests that a higher metamorphic grade and (or) a more aluminous (i.e., garnet-rich) composition might characterize a protolith with generally lower oxygen and more evolved radiogenic isotopic compositions. One possibility is that this garnet-rich rock was pelitic, which is also compatible with higher 87Sr/86Sr100 and lower δ18O. Several Cassiar suite samples, particularly some of the isotopically highly evolved samples (εNd100 = –14 to –15, 87Sr/86Sr100 = 0.716–0.720), have elevated Sm/Yb values that are within the range for garnet stability in the magmatic residuum (Fig. 3.13a). This residual garnet may have existed prior to anatexis in relatively aluminous rocks displaying a range of metamorphic grades or in granulite-grade metamorphic rocks, or it may have formed as a peritectic mineral produced by the dehydration of hydrous ferro-magnesian minerals like biotite during anatexis (e.g., Stevens et al., 2007). Low Sm/Yb for the Seagull and Black River plutonic suites is reflective of the unusually high heavy REE content in these A-type magmas, which is less pronounced in the Rancheria plutonic suite (Fig. 3.13a; Chapter 2). 3.5.1.3 Rb/Sr, Eu/Eu*, and A/CNK In addition to trace element ratios, several basic geochemical parameters (introduced in Chapter 2) are also be compared to isotopic compositions for the characterization of igneous rocks (e.g., Kemp and Hawkesworth, 2003). In granitoid suites, SiO2 and Rb/Sr increase and Eu/Eu* decreases as a magma approaches the ternary minimum composition of quartz–albite–K-feldspar (Tuttle and Bowen, 1958; Zen, 1986; Chappell and White, 1992; London, 1995; Kemp and Hawkesworth, 2003; Clarke et al., 2004). This occurs as melt fractions become smaller, or as magmas become more differentiated. Conversely, the aluminum saturation index, or A/CNK (Al / Ca+Na+K), in granitic magmas is typically inherited from the source rocks; in general, metaluminous infracrustal rocks will produce metaluminous magmas (A/CNK < 1) and peraluminous supracrustal rocks will produce peraluminous magmas (A/CNK > 1). Smaller melt fractions and (or) the removal of restite from a magma will cause A/CNK to approach a value of 1.0, regardless of whether or not the magma was originally of metaluminous or peraluminous composition (Chappell and White, 1992). Late-stage magmatic processes with the potential to also affect A/CNK may also occur locally and include wall rock contamination or assimilation, the partitioning of the alkalis into a fluid (Clarke et al., 2004), and mineral fractionation (Cawthorn and Brown, 1976; Zen, 1986; Chowdhury and Lentz, 2010). However, geochronological and geophysical (magnetic susceptibility and aeromagnetic) studies of intermediate to felsic granitoids have shown that these are typically composite bodies formed by the sequential emplacement of small batches of magma (e.g., Liégeois et al., 1998; Ferré et al., 2002; Vigneresse, 2004, 2006, 2007; Glazner et al., 2004; Žák and Paterson, 2005; Michel et al., 2008; Štemprok et al., 2008; Clemens et al., 2009; Farina et al., 2010). Smaller magma batches are thought to ascend as narrow (i.e., <20 m wide) dykes rather than as large diapiric volumes of magma (e.g., Clemens and Mawer, 1992; Petford et al., 1993, 2000, 2004). Together, the concepts of smaller magma volumes and narrower magma conduits have changed perceptions of how chemical variation occurs in intermediate to felsic magmas, particularly with respect to the effects of differentiation (Clemens and Stevens, 2012). Instead, aside from the obvious control of protolith composition, anatectic processes are likely 191 to be a major control on the overall compositional variation within a single intrusion, or a suite of intrusions derived from similar source rocks (e.g., Clemens et al., 2009; Clemens and Stevens, 2012). Anatectic processes, therefore, are the main control on the above-listed geochemical parameter in intermediate–felsic igneous rocks. The composition of crustally derived magmas, therefore, is controlled by the mixture of melt and peritectic crystals ± entrained restite that each magma contains, which is in turn affected by anatectic conditions. Magmas derived from the largest melt fractions (i.e., at higher temperatures and (or) with slower extraction rates) are more likely to incorporate the most entrained material. In this way, magmas formed from large melt fractions and contain abundant entrained peritectic and restitic material will have chemical and isotopic compositions that more closely reflect the composition of the melt source (Chappell et al., 1987, 2012; Knesel and Davidson, 1999; Clemens and Stevens, 2011; Stevens et al., 2011). For example, entrained peritectic garnet is largely responsible for the characteristically high A/CNK of S-type magmas (Stevens et al., 2007), whereas entrained peritectic clinopyroxene is likely responsible for the characteristically low A/CNK of I-type magmas (Clemens et al., 2011). As melt fractions decrease in size and approach minimum melt compositions (i.e., the quartz–albite–K-feldspar ternary minimum; Tuttle and Bowen, 1958), the geochemical and isotopic composition will less and less reflect the composition of the melt source. This is partly due to a presumed lower proportion of entrained materials in these smaller melt batches, as well as the reduced number of reactant minerals during anatexis. Instead, A/CNK in magmas formed from smaller melt fractions will approach a value of one (as the melt fraction becomes smaller), and the isotopic composition will be controlled mainly by the reactant minerals in the melt source (e.g., Chappell et al., 1987, 2012; Kemp and Hawkesworth, 2003; Farina and Stevens, 2011). Plagioclase and other anhydrous minerals are not involved in fluid- absent, small-fraction partial melting, so a minimum melt inherits only the geochemical and isotopic composition of the hydrous minerals (less any components that remain in the restite as the peritectic mineral assemblage). The rubidium enrichment in micas (particularly biotite), however, is largely transferred to the melt, producing more radiogenic strontium isotopic compositions in the melt than in the protolith (Knesel and Davidson, 1999; Zeng et al., 2005; Farina and Stevens, 2011). An intrusive phase comprising magmas derived mainly from small (or even true ‘minimum’) melt fractions, therefore, will be characterized by elevated Rb/Sr (and lower Eu/Eu*41), high initial strontium isotopic compositions, and A/CNK ≈ 1. If fluid is present in the melt source or partial melting occurs at higher temperatures, however, plagioclase and (or) hornblende in the melt source may become reactants. These minerals add non-radiogenic strontium to the melt, making it less radiogenic than the protolith with respect to strontium (e.g., Patiño Douce and Harris, 1998; Knesel and Davidson, 1999; Zeng et al., 2005; Farina and Stevens, 2011), and reducing its A/CNK and Rb/Sr (as well as increasing Eu/Eu*). The same logic applies with respect to the type of minerals (both peritectic and restitic) that are entrained by the magma. Clearly the reactant phases involved in partial melting, which are dictated by the conditions of partial melting and the composition of the melt source, can produce significant isotopic and geochemical variations in a suite of magmas if there is disequilibrium between the melt and the restite. The lithology of the melt source is also important as this will determined which minerals are available to react (e.g., calcareous rocks will contribute non-radiogenic strontium to magmas; Bernard–Griffiths et al., 1991). 41 If plagioclase is absent in the protolith, restitic monazite and apatite can also produce an Eu/Eu* of less than 1.0 in a magma (Ayres and Harris, 1997). 192 For the Cassiar plutonic suite, there are at least two main trends where Rb/Sr and A/CNK increase with 87Sr/86Sr100 (Fig. 3.14a). Corresponding decreases in Eu/Eu* also occur (Table 3.1), although Eu/Eu* is so closely correlated, inversely, to Rb/Sr that it is not shown in Figure 3.14. However, the very isotopically highly evolved samples (which include most of the high 147Sm/144Nd samples outlined with dashed blue line; Fig. 3.14a) trend towards lower A/CNK with increasing 87Sr/86Sr100. In cases where geochemical parameters vary but isotopic compositions are relatively static, the removal of restite from the magma and (or) late-stage mineral fractionation are likely responsible for the chemical variation. In cases, however, where both geochemical and isotopic composition vary significantly (which is evident in the two Cassiar suite trends indicated in Figure 3.14a), the protolith composition and anatectic conditions should be the dominant controls (e.g., Farina and Stevens, 2011; Clemens and Stevens, 2012). Farina and Stevens (2011) have shown that 87Sr/86Sr can range up to ~0.03 in granitic rocks that are derived by biotite dehydration in a geochemically heterogeneous (greywacke) rock with a homogeneous isotopic composition. In other words, a range in the initial strontium ratio similar to that exhibited by the Cassiar plutonic suite (0.706 to 0.736; Fig. 3.14a) could be produced by partial melting of a single, biotite-bearing, lithology with a homogeneous isotopic composition. Magma mixing or crustal assimilation processes, therefore, do not need to be invoked to produce this range of strontium isotopic composition. The range in εNd100 for the Cassiar plutonic suite (e.g., Fig. 3.5), however, is less likely to be produced by partial melting of a single, isotopically homogeneous melt source. Although magmatic neodymium isotopic compositions can vary from those of the protolith, depending on the fusion or entrainment (or lack thereof) of light REE-bearing minerals in the melt source, the amount of variation produced appears to be only on the order of several epsilon points (Ayres and Harris, 1997; Zeng et al., 2005). The larger range of εNd100 exhibited by the Cassiar plutonic suite necessitates an isotopically heterogeneous melt source, at least with respect to neodymium. A supracrustal protolith consisting of two or more interlayered lithologies with variable neodymium ± strontium isotopic compositions, or containing poorly mixed detritus eroded from isotopically distinct terranes, could explain the neodymium isotopic range of the Cassiar plutonic suite. The geochemical interpretation in Chapter 2 is strongly supportive of a geochemically heterogeneous supracrustal melt source of bulk greywacke composition, although lesser mafic or calcareous rocks in this supracrustal package were not ruled out as possible contributors of melt to this plutonic suite. Trends exhibited by the isotopically highly evolved intrusions in the Cassiar plutonic suite are most simply attributed to lithological and detrital variations likely to exist in supracrustal rocks (e.g., Clemens and Stevens, 2012), combined with differing degrees of disequilibrium between magma and restite (e.g., Ayres and Harris, 1997; Zeng et al., 2005; Farina and Stevens, 2011). Pelitic or clay-rich lithologies are likely to be more isotopically evolved with respect to strontium than quartzo–feldspathic lithologies, given the higher mica (especially biotite) to feldspar ratio in the former. An isotopically and lithologically heterogeneous supracrustal package may consist of interlayered pelitic and quartzo–feldspathic lithologies may account for the two trends highlighted in Figure 3.14a. The samples with lower 87Sr/86Sr100 may be magmas sourced from a dominantly quartzo–feldspathic protolith (with input from mafic infracrustal rocks in the samples with the lowest 87Sr/86Sr100; see Section 3.5.1.1) and the samples with higher 87Sr/86Sr100 may be magmas sourced from a dominantly pelitic protolith (Fig. 3.14a). The geochemical and isotopic variations attributed to the interplay between the mineral reactants involved in partial melting (i.e., plagioclase ± hornblende versus mica) in each of these lithological end-members are illustrated in 193 Figure 3.14a. The very isotopically highly evolved intrusive phases belonging to the Cassiar suite (e.g., 87Sr/86Sr100 > 0.72; Fig. 3.14a) and the “migmatite” sample (Morris and Creaser, 2008) deserve special mention. These samples display a high degree of strontium isotopic variability (Fig. 3.13a), and approach an A/CNK value of ~1.0 with increasing Rb/Sr and 87Sr/86Sr100 (as well as SiO2; Table 3.1) in Figure 3.14a. Both of these features are compatible with a derivation by low-temperature, minimum-fraction partial melting. If minimum melting conditions acted to increase Rb/Sr, then the strontium isotopic composition of these samples is likely to be more evolved than that of the protolith (e.g., Farina and Stevens, 2011). As the increase in strontium variability (and the appearance of samples derived from minimum-melting processes) roughly coincides with the location of the lower perturbation in Figure 3.6a, or 87Sr/86Sr100 ≈ 0.72. A isotopically highly evolved crustal end-member, inferred above to have a pelitic composition, therefore, appears to have had an average 87Sr/86Sr100 of about 0.72 (Fig. 3.14a). The very isotopically highly evolved samples also have anomalously elevated 147Sm/144Nd, which may be related to REE disequilibrium during minimum partial melting processes (Section 3.4.1.2; e.g., Brandon and Lambert, 1993, 1994; Ayres and Harris, 1997)—even this does not produce the same degree of isotopic variability exhibited by strontium. Samarium and neodymium in intermediate to felsic crust are concentrated in monazite (Ayres and Harris, 1997; Zeng et al., 2005), making it difficult to fractionated samarium from neodymium during anatexis when equilibration between magma and restite are attained (similar arguments also apply for magmatic fractionation). During crustal anatexis, however, isotopic disequilibrium between magma and restite is likely, particularly at the conditions that produce minimum melts (i.e., in the absence of fluid) and (or) when rapid heating and melt extraction occurs (Ayres and Harris, 1997; Zeng et al., 2005; Farina and Stevens, 2011). This allows apatite, which contains less REEs overall but has higher Sm/Nd than monazite, to control this ratio―and the neodymium isotopic composition―of the resulting magma (Zeng et al., 2005). The depletion in La/Sm for the very isotopically highly evolved samples is consistent with disequilibrium between monazite and the extracted magma. A similar disequilibrium involving HFSE-rich accessory minerals might also be responsible for low Ba/Nb in the same samples (Fig. 3.13a). Although some degree of disequilibrium during partial melting has probably affected most of the Cassiar plutonic suite (given the overall trend towards lower La/Sm and Ba/Nb, and higher Sm/Nd with increasing degree of isotopic evolution; Fig. 3.13a), only the very isotopically highly evolved samples have Sm/Nd that is notably higher than proxies for average crustal compositions (Fig. 3.13a). Estimates of TDM (>1.9 Ga) for the isotopically highly evolved samples are already regarded as suspect because of the anomalous 147Sm/144Nd values (Section 3.4.1.3). It is likely that εNd100 values below about –15 (corresponding to samples with TDM > 1.9 Ga; Fig. 3.6c) were also affected by REE disequilibrium during anatexis. Sulfur isotopic compositions for the Cassiar suite do not display any correlation with A/CNK (Fig. 3.14b) because the variation in δ34S should not noticeably affect major element concentrations. Higher positive δ34S (which characterizes most of the Cassiar plutonic suite), however, is generally associated with trace element variations, including higher Rb/Sr (and lower Eu/Eu*) in Figure 3.14b. The range in sulfur compositions is too large to be the result of magmatic fractionation (Ueda and Sakai, 1984); instead, the correlation between Rb/Sr (and lower Eu/Eu*; 194 Table 3.1) implies that a plagioclase-poor and clay-rich (i.e., pelitic) melt source is the responsible for the high positive values characterizing most of the Cassiar plutonic suite. Although the lower sulfur isotope compositions for several Cassiar suite samples might have been produced by magmatic degassing (Section 3.4.2.2), the lower Rb/Sr in these rocks suggests instead that these samples were less contaminated with sulfur from sulphide-bearing pelitic rocks. Unlike the Cassiar plutonic suite, the three A-type plutonic suites have very restricted initial strontium isotopic compositions at 100 Ma (Figs. 3.13a and 3.14a). For this reason (as well as a limited amount of data) correlations between radiogenic and geochemical compositions are not apparent. However, Rb/Sr is consistently higher (and Eu/Eu* is consistently lower; Table 3.1) in the A-type samples when compared to Cassiar suite samples with similar δ34S (Fig. 3.12b). Rb/Sr (and Ba/Nb) are, therefore, not controlled by contamination of the A-type magmas through interactions with sulfur-bearing, rubidium-rich (pelite) or barium-rich (carbonate/evaporite) rocks. These distinctive ratios were either inherited from the melt source or are the result of anatectic or magmatic processes, such as extremely high degrees of magmatic fractionation or extremely small degrees of partial melting. 3.5.2 Sources of Isotopic Variation in the Back-Arc Cassiar Plutonic Suite The range of εNd100 and 87Sr/86Sr100 for the Cassiar plutonic suite (εNd100 ≈ –3 to –20 and 87Sr/86Sr100 ≈ 0.706 to 0.735) extends from compositions typical of continental arc magmatism to the west to compositions that approach (but do not attain) averages of the Proterozoic felsic orthogneissic to paragneissic Laurentian basement composition (Fig. 3.15b). Similar trends in isotopic space have been interpreted to be the result of either assimilation of felsic crust by mantle-derived magma (e.g., Patiño Douce, 1995) or mixing between mantle-derived magma and crustally derived magma in the lower crust (e.g., Clemens et al., 2009). There is no evidence for widespread magma mixing between two compositionally distinct magmas in these plutons. Mafic magmatic enclaves or globules and disaggregated mafic magma textures have not been observed by this study, or reported by others, in the Cassiar suite intrusions. Even if they are present in some intrusions, local signs of magma mingling or mixing are not evidence for extensive magma mixing processes (Clemens et al., 2009; Clemens and Farina, 2011; Clemens and Stevens, 2012). The difference in rheological properties between compositionally distinct magmas present difficulties for large-scale magma mixing, particularly when a viscous felsic magma is involved (Watson, 1982; Sparks and Marshall, 1986; van der Laan and Wyllie, 1993; Kay et al., 1992; Glazner, 2011). It is largely for this reason that magma mixing is an unlikely (although not impossible; Clemens et al., 2009) source of isotopic variation in most intermediate–felsic plutonic suites (Clemens et al., 2010; Glazner, 2011; Clemens and Stevens, 2012). Furthermore, modeling of magma mixing and assimilation processes between basaltic magma and felsic magma or crust required unrealistic parameters and (or) were not able to reproduce the observed range of geochemical compositions in the Cassiar batholith (Driver et al., 2000). This discrepancy is common, but rarely addressed, in many studies invoking mixing processes (Chappell et al., 1987; Glazner, 2007; Clemens et al., 2009). It is improbable that widespread mixing, assimilation, or contamination processes were capable of efficiently altering the composition of the large volumes of silicic magma emplaced across the Cassiar Terrane. Recent interpretations regarding the genesis of intermediate to felsic granitoid provinces instead suggest that the chemical variation in these granitic rocks primarily reflects the 195 variation inherent to their crustal melt source (e.g., Kemp and Hawkesworth, 2003; Clemens et al., 2009; Clemens and Stevens, 2012). Heterogeneity within the melt source, combined with differing conditions of partial melting (e.g., temperature, pressure, rate of heating, the presence or absence of fluid, etc.) and timings of melt extraction (e.g., prior to, or after equilibration with restite is achieved), further complicate the chemical and isotopic variation exhibited by granitoid suites (e.g., Zeng et al., 2005; Stevens et al., 2007; Clemens et al., 2009; Farina and Stevens, 2011; Clemens and Stevens, 2012). Several petrogenetic scenarios involving the mixing of distinct isotopic reservoirs prior to anatexis might better explain the range of isotopic compositions—as well as the two perturbations interrupting the isotopic continuum— exhibited by the Cassiar plutonic suite (Figs. 3.6a and 3.15). Isotopically relatively juvenile samples plotting above the upper perturbation comprise the few hornblende-bearing phases sampled from the Cassiar suite, and several biotite- and muscovite-bearing phases (Figs. 3.6a and 3.15a). Together these form a (commonly oxidized) alkaline to alkali-calcic magma series (Chapter 2). Isotopically moderately to highly evolved samples plotting below the upper perturbation comprise the majority of biotite- and muscovite-bearing phases of the Cassiar suite (Figs. 3.6a and 3.15a) are typically reduced and calc-alkaline (Chapter 2). Distinct trace element ratios are also broadly associated with these isotopic perturbations (Figs. 3.13a and 3.15a). The four compositional groupings illustrated in Figure 3.15a reflect differences in either the protolith composition or the conditions of anatexis. 3.5.2.1 Possible Isotopically Juvenile Protoliths The isotopically moderately evolved component was interpreted in Sections 3.5.1.1 and 3.5.1.3 to be mainly supracrustal rocks with bulk intermediate, quartzo–feldspathic composition. The transition from alkaline to alkali- calcic (to calc-alkaline in the most siliceous examples) compositions for the isotopically juvenile samples of the Cassiar plutonic suite (Chapter 2) is suggestive of some form of isotopic and chemical mixing. The most likely crustal sources of the isotopically juvenile component (labeled “IJ-qp”; Fig. 3.15) are: a) partial melts originating in mafic–intermediate crust derived from the mantle, such as juvenile island arc crust (Morris and Creaser, 2008) b) partial melts derived from OIB-like igneous rocks interlayered with supracrustal rocks in the crust (Driver et al., 2000); or c) partial melting of sedimentary rocks containing a large proportion of detritus derived from mafic volcanic or intrusive rocks. The first scenario requires the presence of mafic–intermediate rocks in the middle to lower crust of the Cassiar Terrane, either in the form of an old continental magmatic arc in the Cassiar Terrane, under-thrust Paleozoic pericratonic to oceanic terranes, or mafic flows and sills. There is evidence for local Devono-Mississippian continental arc magmatism in the western Cassiar Terrane (e.g., Erdmer et al., 2005); however, the extent of this arc has not been evaluated and it cannot be inferred to underlie the entire Cassiar Terrane. Nor have geophysical surveys detected the presence of any mafic–intermediate igneous basement that could be related to a continental magmatic arc, pericratonic or otherwise, below the Cassiar Terrane (e.g., Clowes et al., 2005). Finally, the alkaline to alkali- 196 calcic compositional signature of the isotopically juvenile Cassiar suite samples is unlikely to have been inherited from a typical continental magmatic arc of predominantly calc-alkaline composition. Tholeiitic (sub-alkaline) to OIB-like (alkaline) flows and sills, however, have been observed in all of the exposed Proterozoic metasedimentary packages in the northern Cordillera, and OIB-like Paleozoic volcanic rocks and sills are locally exposed in the Cassiar Terrane near the Tintina Fault (Fig. 3.1b; e.g., Goodfellow et al., 1995; Driver et al., 2000; Thorkelson et al., 2005). High-velocity horizons detected seismically throughout Laurentian components of northern Cordilleran crust might be mafic sills or flows present at mid-crustal levels (e.g., Evenchick et al., 2005). Biotite-bearing pelitic and quartzo–feldspathic metasedimentary rocks interlayered with hornblende-bearing amphibolitic rocks may generate large volumes of melt at roughly the same mid- to lower crustal temperatures (Patiño Douce and Beard, 1995; Skjerlie and Patiño Douce, 1995; Thompson, 2001)42. There would also be fewer rheological barriers to mixing intimately associated granitic to granodioritic melts derived from interlayered metasedimentary and amphibolitic rocks (e.g., Rushmer, 1991; Patiño Douce and Beard, 1995). Finally, interlayered metasedimentary and amphibolitic rocks may mutually enhance anatexis through localized alkali exchange along their contacts (Skjerlie and Patiño Douce, 1995). However, depleted mantle model ages for mafic–intermediate sills and flows in the exposed Proterozoic supracrustal packages are older than those of the Cassiar suite plutons (e.g., >2 Ga; Fig. 3.16a), making these improbable sources of isotopically juvenile material in the plutonic rocks. Instead, OIB-like Paleozoic sills with younger TDMs (e.g., 0.4–0.9 Ga; Fig. 3.16a), appear to be a more appropriate source of juvenile melt, but only if they are present in the main supracrustal source region for this plutonic suite. The absence of low, near mantle-like oxygen isotopic compositions (+7.5 ‰ or less; Gregory and Taylor, 1981) in the Cassiar plutonic suite requires that any igneous source of the juvenile isotopic signature must have been either hydrothermally altered at extremely low temperatures (i.e., <150°C by tropical seawater; G. Arehart, 2012, pers. comm.), or significantly contaminated by supracrustal rocks. As none of the Cassiar suite samples overlap with the isotopic range of the Paleozoic basaltic rocks, it may be inferred that supracrustal materials were also incorporated into the isotopically juvenile intrusive phases. An infracrustal juvenile melt source, however, did not comprise the bulk of any of the Cassiar suite magmas. The geochemical (Chapter 2) and oxygen isotopic compositions of the Cassiar plutonic suite strongly suggest that it is mainly, if not entirely, of supracrustal origin. A third possible source for the isotopically juvenile component is lithic detritus eroded from mafic igneous rocks and mechanically mixed during deposition with detritus eroded from more isotopically evolved rocks. This implies that the moderately evolved isotopic end-member is actually that of a supracrustal rock that may contain a significant component of isotopically juvenile detritus. If the supracrustal rocks comprise detritus sourced from compositionally and isotopically variable basement, differing proportions of each material combined with varying 42 It should be noted that these experimental studies used high starting temperatures of 900°C and 850°C, respectively, which are unrealistic for crustal anatexis (unless rapid heating occurred because of basaltic interplating; Petford and Gallagher, 2001). These experimental partial melting studies of interlayered metasedimentary and metabasaltic rocks, therefore, may over-estimate the fertility muscovite- and biotite-bearing lithologies. If heating is not rapid, it might be postulated that micaceous supracrustal lithologies would be somewhat depleted of melt by the time interlayered hornblende-bearing lithologies begin to undergo extensive anatexis in the mid- to lower crust. Higher temperature (and later) anatexis of hornblende-bearing lithologies is consistent with the later emplacement of hornblende-bearing intrusions in the Cassiar Terrane (see Section 3.5.2.3), and makes models requiring mixing of melt or magma derived from such compositionally distinct rock types less practical. 197 degrees of mixing during transport, deposition, re-working and re-deposition could account for a large range of isotopic compositions in plutonic rocks derived from heterogeneous, but sedimentary, rocks. This suggestion is supported by the findings of Farina and Stevens (2011), who demonstrated that a large range of strontium isotopic compositions can be produced by fluid-absent partial melting of biotite with varying Mg / (Mg + Fe) from a single heterogeneous, but isotopically well-mixed greywacke protolith. The possibility of larger magmatic strontium isotopic variations were alluded to by Farina and Stevens (2011) for isotopically heterogeneous supracrustal protoliths (e.g., a package of two or more lithologies with isotopic composition controlled by the relative proportions of detritus derived from isotopically distinct terranes). Unlike the above two scenarios involving mafic infracrustal rocks, a sedimentary protolith containing a large proportion of isotopically juvenile detritus is also more compatible with the high positive δ18O exhibited by the Cassiar plutonic suite. Depositional mixing of isotopically juvenile and isotopically evolved detritus in one package of sedimentary rocks is a reasonable explanation for the source of the isotopically juvenile component in the Cassiar suite magmas. The mafic material would have to have been derived from the erosion of OIB-like mafic rocks (of possible Paleozoic age) in order to impart the alkaline composition to isotopically juvenile hornblende-bearing Cassiar suite samples. An alternate scenario involving relatively young (i.e., Paleozoic) interlayered OIB-like flows or sills is also a viable model, if they did not form the bulk of the magma in any of the Cassiar suite intrusions. Both models (e.g., an infracrustal or a detrital source of isotopically juvenile material) fit the available data, and are not mutually exclusive. The model invoking interlayered mafic sills or flows may be less practical due to the necessary mixing of melts, and issues associated with the temperature of anatexis—and, consequently the timing of melt extraction from the supracrustal and infracrustal lithologies. 3.5.2.2 Possible Isotopically Highly Evolved Protoliths A large proportion of samples from the Cassiar plutonic suite have highly evolved neodymium and strontium isotopic compositions (εNd100 ≈ –7 to –15 and 87Sr/86Sr100 ≈ 0.71 to 0.72; Fig. 3.15a)43, and a large range of TDM (~1.2–1.9 Ga; Figs. 3.6c and 3.16). In Section 3.5.1.3, these calc-alkaline samples were interpreted to have been derived from the anatexis of a package of isotopically highly evolved metapelitic rocks (labeled “HE-p”; Fig. 3.15) interlayered with isotopically moderately evolved quartzo–feldspathic metasedimentary rocks (labeled “HE-qp”; Fig. 3.15). The difference in the strontium isotopic composition of these rocks was attributed to different proportions of plagioclase and mica (especially biotite) in these two lithological end-members. The geochemical and strontium isotopic variation resulting from variable mixtures of melt derived from both lithological end-members was attributed to the interplay between the mineral reactants involved in partial melting (i.e., plagioclase ± hornblende versus biotite ± muscovite). The large neodymium isotopic variations in particular were interpreted to reflect the lithological variations and, thus, varying proportions of detritus from isotopically distinct sources (Section 3.5.1.3). A range of about 10 εNd100 points, for example, is present between more evolved pelitic and less evolved quartzose rocks in the limited neodymium isotopic data available for the Mackenzie Mountains Supergroup (Fig. 3.16a; 43 Samples with more evolved isotopic compositions in the Cassiar plutonic suite (referred to as “very highly isotopically evolved”) are only apparently more isotopically evolved than their melt source because of Sm/Nd and Rb/Sr fractionation during anatexis; these samples are not discussed further. 198 Rainbird et al., 1997). The neodymium isotopic composition of the Cassiar suite may also have been largely controlled by similar lithological variations in the magma source rocks. An isotopically variable package of supracrustal rocks, therefore, may comprise two or more distinctly different interlayered lithologies (e.g., feldspathic arenite and pelite) with different neodymium ± strontium isotopic compositions, or one poorly mixed lithology (e.g., greywacke) containing detritus eroded from isotopically distinct terranes. Either of these scenarios could explain the isotopic variation exhibited by the Cassiar suite. The low sulfur concentrations in the Cassiar suite magmas require a relatively sulfur-poor melt source, in which evaporite and carbonaceous rocks are absent, and abundant carbonate rocks are lacking. The minimal amounts of sulfur in these plutons are thought to have been sourced from minor pyritic shales in the protolith, and (or) by variable contamination by (or assimilation of) similar sulfur-rich rocks in the crustal column (Section 3.4.2.2). Most of the exposed stratigraphic packages in the northern Cordillera contain mixtures of rocks deposited in relatively shallow marine environments (i.e., continental shelves; e.g., Thorkelson et al., 2005). Deeper marine sedimentary rocks of the Selwyn Basin typically contain pyritic or graphitic shales, although calcareous rocks are less common and evaporitic rocks are absent (e.g., Gordey and Anderson, 1993). The melt source to the Cassiar plutonic suite would have been deposited in a similarly deep basin, although perhaps a more distal and oxidized setting lacking rocks typical of shallower settings. This melt source may have been basin-ward stratigraphic equivalents of one or more of the shallow-water supracrustal packages fringing the northern and eastern margins of the Selwyn Basin (Fig. 3.2a) or deeper rocks deposited in the Selwyn Basin. A heterogeneous package of interlayered siliciclastic lithologies deposited in a relatively deep basinal (if not an open marine) setting is envisioned as the most likely source of isotopic and chemical variation in moderately to isotopically highly evolved samples of the Cassiar plutonic suite. 3.5.2.3 Mineralogical ± Isotopic Variation with Time The average neodymium and strontium isotopic compositions of the biotite-bearing and muscovite-bearing Cassiar suite intrusions are identical, despite the wide range of isotopic compositions exhibited by these two mineralogical variants (Fig. 3.5a). The biotite-bearing samples (average U–Pb age = 109.1 ± 2.3 Ma; 19 samples), however, are typically younger than the muscovite-bearing samples (average U–Pb age = 113.0 ± 2.0 Ma; 8 samples; Table 3.1; Appendix G.1). The older average age of the muscovite-bearing intrusive phases demonstrate that they are not the product of fractionation and (or) differentiation of the biotite-bearing magmas, despite their nearly identical average isotopic compositions. Instead, differences between these two mineralogical variations might be related to differences in the degree of partial melting, the timing of melt extraction, the temperature of anatexis, and the degree of prior melt depletion and (or) metamorphic dehydration in the same protolith rocks. Earlier muscovite-bearing magmas are more likely to have been produced by partial melt fractions from undepleted or previously unmelted or dehydrated (i.e., amphibolite-grade or lower) metasedimentary rocks, at low temperatures of partial melting where the muscovite ± biotite solidi are intersected (e.g., 750–850°C). The oldest intrusions in the Cassiar Terrane, which are mostly muscovite-bearing, tend to have the lowest Eu/Eu* (Table 3.1), which is 199 consistent with a lack of plagioclase involvement in partial melting. Later magmas in the Cassiar Terrane (typically biotite- and hornblende-bearing phases), become less geochemically evolved as temperatures increase, muscovite in the melt source is exhausted, and biotite and plagioclase ± hornblende become more important reactant phases. The biotite-bearing samples with the lowest 87Sr/86Sr100 were probably formed from higher temperature partial melting of a quartzo–feldspathic lithology, in which plagioclase ± minor hornblende may also have been important reactants. Given the range of neodymium isotopic compositions exhibited by these two mineralogical variations of the Cassiar plutonic suite, these source rocks must have been isotopically heterogeneous. If crustal geotherms continued to rise with time, the latest and highest temperature partial melting may have started to affect hornblende-bearing or more ferro-magnesian lithologies with distinctly less evolved isotopic compositions, such as interlayered mafic rocks. Partial melting of interlayered biotite-bearing quartzo–feldspathic rocks and hornblende-bearing mafic flows or sills would have produced voluminous melts at temperatures between 850– 930°C (e.g., Patiño Douce and Beard, 1995). This may have been responsible for generating the alkaline to alkali- calcic hornblende-bearing phases of the Cassiar suite with an average U–Pb age of 104.2 ± 1.3 Ma (and possibly related, isotopically juvenile, alkalic to alkali-calcic biotite- and muscovite-bearing phases for which geochronological control is lacking), and with mixed infracrustal and supracrustal protoliths. Lithic detritus in the supracrustal rocks may also contribute larger quantities of less isotopically evolved melt at higher temperatures. 3.5.3 Isotopic Comparisons between the Cassiar Plutonic Suite and Northwest Laurentian Basement Due to the differences in fertility during fluid-absent partial melting in amphibolite-grade metasedimentary rocks that are rich in micas, compared to granulite-grade lower crustal rocks (e.g., Patiño Douce et al., 1990), the crustally derived Cassiar plutonic suite does not provide insight into the age and composition of any lower crustal basement to the Cassiar Terrane. This plutonic suite is almost certainly entirely derived from supracrustal rocks deposited in a deep basinal setting, rather than from the crystalline basement on which the sediments are thought to have been deposited. The S-type affinity of the Cassiar plutonic suite is compatible with the dominant geophysical view that a thick package of deep-water metasedimentary rocks underlies the northern Cordilleran miogeocline (e.g., Cook et al., 1999, 2004; Welford et al., 2001; Snyder et al., 2002, 2009; Hammer and Clowes, 2004; Clowes et al., 2005; Creaser and Spence, 2005; Fernández-Viejo et al., 2005; Cook and Erdmer, 2005). Although the Neoproterozoic to Paleozoic sedimentary rocks at the surface are isotopically similar with respect to neodymium, strontium, and oxygen compositions, the low sulfur content in the Cassiar suite magmas might be more consistent with derivation from a deeper (or depositionally more distal) package of lithologically variable siliciclastic rocks. The highly resistive geophysical properties of the inferred mid-crustal wedge of metasedimentary rocks requires an absence of interconnected graphite, carbon, metallic oxides, or sulphides (Wennberg et al., 2002; Ledo et al., 2004; Jones et al., 2005), which is consistent with the low sulfur contents in the Cassiar suite magmas. This study provides indirect information regarding the possible age(s) and composition of the supracrustal rocks that are located at lower–middle crustal depths in the Cassiar Terrane. 200 3.5.3.1 An Overview of the Isotopic Composition of Northwest Laurentian Basement Lead, oxygen, and strontium isotopic compositional data for Cassiar Terrane basement rocks are limited. These isotopic systems are typically avoided when examining metamorphic and sedimentary rocks, as they are easily disrupted by fluid flow in the crust and (or) at the Earth’s surface. The Sm–Nd system is one of the most robust isotopic systems, thus, the isotopic data for available basement rocks is nearly entirely confined to neodymium isotope datasets. A summary of the data is provided below as a basis for the following discussion on possible melt sources for the post-accretionary plutonic suites.44 These isotopic data for northwest Laurentian basement are also compiled in Appendix G.4. The analysis of undeformed granitic rocks in drill core intersections in the Fort Simpson magnetic high (Fig. 3.2b) suggest that it is a Paleoproterozoic (~1.85 Ga) continental magmatic arc developed on the Nahanni Terrane (Hoffman, 1989; Villeneuve et al, 1991). Maps of terranes, basement domains, and geophysical properties (e.g., Ross, 1991; Aspler et al., 2003) would suggest that the Fort Simpson magnetic high does not underlie the Cassiar Terrane in central Yukon—although it could conceivably be extrapolated to Cordilleran basement in northern British Columbia (Fig. 3.2b). Highly evolved neodymium (ɛNd100 ≈ –23 to –21; Fig. 3.16a) and older depleted mantle model ages (2.14 and 2.45 Ga) for these granites indicate they were derived by anatexis of their host terrane; if they are hosted by the Nahanni terrane, then fertile portions of this unexposed domain should range from 1.85 Ga to at least 2.45 Ga (and likely older; Villeneuve et al., 1991). These TDMs from the Fort Simpson magmatic high provide perhaps the only constraints on the age of crystalline basement below the northern Cordilleran miogeocline. Other proxies for the age and composition of crystalline basement below the Cassiar Terrane stem from two main sources: (a) exhumed core complexes in southeast British Columbia (e.g., the Malton, Monashee, and Frenchman’s Cap complexes); and (2) drill core intersections in basement domains beneath sedimentary cover in northern Alberta (e.g., the Buffalo Head, Nova, Taltson, and Chinchaga domains; Fig. 3.2b). Although these basement domains cannot necessarily be extrapolated to the Cassiar Terrane in the northern Cordillera (Fig. 3.2b), they do provide an indication of what ages and types of rocks could underlie portions of the Cordillera, particularly if the Nahanni terrane is an accreted Paleoproterozoic crustal terrane (Fig. 3.2b; Ross, 1991). Felsic examples of Archean to Paleoproterozoic (3.0–1.8 Ga) orthogneissic and paragneissic rocks typically have very evolved isotopic compositions due to their age and felsic composition (ɛNd100 ≈ –20 to –30; Fig. 3.16a). Coeval mafic orthogneissic rocks in these regions have a much wider range of neodymium compositions (ɛNd100 ≈ –28 to –2; Fig. 3.16a), which is probably due to variable crustal contamination. Paleoproterozoic rocks in the northern Cordillera for which isotopic data are available are still limited, although studies that should rectify this have recently been initiated (e.g., Furlanetto et al., 2009; Medig et al., 2010, 2012; Turner, 2010). Some of the oldest basement exposed in the northern Cordillera are rocks of the ~1.7 Ga, primarily siliciclastic Wernecke Supergroup (Fig. 3.2a). The siliciclastic rocks have highly evolved isotopic compositions (ɛNd100 ≈ –24 to –18, and –10 for a dolostone; Fig. 3.16a) that are consistent with detritus sourced from Archean to 44 Complete sources of compiled isotopic data for basement rocks are detailed in the footnotes to Figure 3.16. 201 Paleoproterozoic domains across the Laurentian craton (Furlanetto et al., 2009). Gabbroic to syenitic samples of the 1.7 Ga (mainly mafic) Bonnet Plume River plutons extracted from the mantle as undepleted continental tholeiitic magmas (e.g., Thorkelson et al., 2001) are emplaced into the Wernecke Supergroup. These mafic rocks are more neodymium-radiogenic, but still relatively evolved compared to younger mafic magmatism (ɛNd100 ≈ –17 to –8; Fig. 3.16a). The Paleo- to Mesoproterozoic Muskwa Assemblage and the Meso- to Neoproterozoic Mackenzie Mountains Supergroup in the northern Cordillera (Fig. 3.2a) are both thought to be aerially restricted on the basis of their depositional environment (e.g., Rainbird et al., 1996; Evenchick et al., 2005). However, it is possible that distal equivalents of these successions extend farther west and make up much of the thick, westward-tapering package of reflective rocks inferred by Cook et al. (2004; and others) to represent buried Mesoproterozoic strata below the Cassiar Terrane. The Muskwa Assemblage has particularly evolved isotopic compositions (ɛNd100 ≈ –28 to –24; Fig. 3.16a) that are similar to Archean to Paleoproterozoic crystalline domains, whereas the Mackenzie Mountains Supergroup is much less evolved overall (ɛNd100 ≈ –10 to –19; Fig. 3.16a). Neoproterozoic continental ‘tholeiites’ capping the Mackenzie Mountains Supergroup mark the earliest stages of the Cordilleran-wide rifting event that led to the deposition of the Windermere Supergroup (Dudás and Lustwerk, 1997). These rift-related ‘basalts’ have unusually radiogenic 87Sr/86Sr100 (0.712–0.717; see Fig. 3.15a) for their more juvenile range of ɛNd100 (–0.8 to –4; Figs. 3.15a and 3.16a). In the Cassiar Terrane the oldest exposed rocks are the Neoproterozoic Ingenika Group of the Windermere Supergroup (Fig. 3.2a; e.g., Gordey and Makepeace, 2001). Although some of the older rift-related sedimentary packages may have had a relatively limited geographic distribution (Evenchick et al., 2005), the Windermere Supergroup was deposited along the length of the Canadian Cordillera during the breakup of the Rodinia supercontinent. These rocks are dominantly siliciclastic (typically quartzo–feldspathic grits), with thin glaciogenic and carbonate horizons, and are intruded or overlain by mafic sill and (or) volcanic rocks (Gabrielse and Campbell, 1991; Souther, 1991; Woodsworth et al., 1991; Thorkelson et al., 2005). Their neodymium compositions (ɛNd100 ≈ – 20 to –26, although a few samples have εNd100 up to –13; Fig. 3.16a) are consistent with relatively proximal detrital sources in the uplifted, Paleoproterozoic to Archean, granitic basement to the east (e.g., Gabrielse and Campbell, 1991; Gehrels and Ross, 1998; Ross et al., 2005; Lemieux et al., 2007; Leslie, 2009). Compared to the Proterozoic supracrustal rocks, the Paleozoic sedimentary rocks in northwest Laurentia are typically more radiogenic with respect to neodymium, due to their younger age and (or) less siliceous composition (Fig. 3.16b). Paleozoic miogeoclinal rocks are inferred to have been derived by reworking of older rocks and (or) from crystalline basement to the east. However, they also appear to have incorporated a significant proportion of isotopically juvenile detritus, from either coeval OIB-like basaltic flows and diatribes or from older mafic rocks like erupted portions of the tholeiitic Mackenzie Large Igneous Province (e.g., the Coppermine flood ‘basalts’; Creaser and Erdmer, 1997). Cambrian to Devonian miogeoclinal sediments are moderately to isotopically highly evolved (ɛNd100 ≈ –8.8 to –17.6; Figs. 3.1 and 3.16b). Upper Paleozoic to Triassic miogeoclinal rocks are less isotopically evolved overall (ɛNd100 ≈ –7.3 to –13.5; Fig. 3.13b), and the mid-Paleozoic clastic influx of sediments overlaps with 202 both packages of miogeoclinal rocks (ɛNd100 ≈ –7.0 to –19.1; Fig. 3.16b). Lower to mid-Paleozoic mafic volcanic rocks and diatribes (Fig. 3.2a) are much more isotopically juvenile (ɛNd100 ≈ –2.1 to +6.0; Fig. 3.13b). 3.5.3.2 Possible Crustal Magma Sources for the Cassiar Plutonic Suite Morris and Creaser (2008) suggested that migmatitized Neoproterozoic metasedimentary host rocks may be a suitable melt source for the highly evolved Cassiar plutonic suite, whereas Driver et al. (2000) preferred deeper Neoproterozoic to Mesoproterozoic metasedimentary (± Paleoproterozoic felsic orthogneissic) melt sources (at least for the Cassiar batholith). Based on the isotopic comparisons presented in Figure 3.16, neither of these suggestions appear to be suitable melt sources for the Cassiar plutonic suite. Most Proterozoic basement in the northern Cordillera (or their proxies) have much more evolved radiogenic isotopic compositions (Fig. 3.16a) ± more primitive oxygen isotopic compositions (e.g., Archean to Paleoproterozoic gneisses in southeast British Columbia; Fig. 3.9a) than the Cassiar plutonic suite. Only mafic Archean to Paleoproterozoic orthogneisses and the upper crustal Bonnet Plume River intrusions actually span the neodymium and strontium isotopic range of the Cassiar plutonic suite (Figs. 3.13 and 3.14a). However, even a mafic infracrustal melt source (particularly one of Archean to early Paleoproterozoic age) for the Cassiar plutonic suite is unlikely given the older TDMs that characterize these basement rocks (Fig. 3.16a). The involvement of mafic basement rocks in the genesis of the Cassiar plutonic suite is also chemically improbable for all but the least isotopically evolved samples, in which mafic rocks may be the protolith to a sub-ordinate source of melt. A metasedimentary protolith for the majority of the Cassiar suite is the most consistent with the geochemical interpretation in Chapter 2 and with the isotopic data reported here. The compilation of neodymium isotopic data for basement rocks in northwest Laurentia shows few basement datasets with similar εNd100 and TDM ranges to that of syn- to post-accretionary intrusive rocks in the Cassiar Terrane, none of which fully overlap with the entire compositional range of the Cassiar plutonic suite (Fig. 3.16a). The early Paleozoic and Late Paleozoic (to early Mesozoic) miogeoclinal packages, and the mid-Paleozoic clastic influx that comprise the uppermost crust of the northern Cordillera today (Fig. 3.16b) best approximate the isotopic composition of the intrusive rocks in the Cassiar Terrane. The more isotopically juvenile Paleozoic basinal rocks compared to Precambrian basement (Fig. 3.16a) is thought to be due to the incorporation of abundant lithic detritus eroded from mafic volcanic rocks (e.g., Creaser and Erdmer, 1997; Garzione et al., 1997). It is also apparent that OIB-like Paleozoic volcanic rocks related to several episodes of extension along the western margin of North America (e.g., Goodfellow et al., 1995) could be a reasonable end-member composition to the isotopically juvenile samples (Fig. 3.16b). These neodymium data would suggest a melt source for the Cassiar plutonic suite comprising lithologically variable Paleozoic sedimentary rocks, containing a significant proportion of detritus derived from mafic rocks—that mixed with melts derived from interlayered mafic rocks to produce the least isotopically evolved magmas. The low regional metamorphic grades (greenschist or less; Greenwood et al., 1991) for miogeoclinal sedimentary rocks are incompatible with burial to middle–lower crustal depths of over >25 km, where fluid-absent anatexis is most likely to occur (Gordey and Anderson, 1993; Cook et al., 2004; Sawyer et al., 2011). There is, however, the 203 possibility that Paleozoic miogeoclinal rocks are structurally duplicated at depth, despite interpretations to the contrary (e.g., Cook et al., 2004). In northern British Columbia amphibolite-facies lower Paleozoic (and upper Neoproterozoic) rocks are exposed in structurally exhumed basement (e.g., Evenchick 1988; Plint et al. 1992). Interpretations of seismic data by Evenchick et al. (2005) across the Cassiar Terrane in northern British Columbia argue that intensely thickened Paleozoic and interleaved Paleozoic and Neoproterozoic strata may extend to at least the modern-day middle crust, which would have corresponded to depths of about 22–25 km in the Cretaceous. It may be that Paleozoic rocks were more deeply buried and thickened in northern British Columbia because of more intense deformation that produced steeper (commonly west-verging) fold and thrust structures (Gabrielse, 1991). Although structural styles are somewhat different in Yukon (Gabrielse, 1991), there is no evidence to refute the possibility of a similar degree of structural thickening in Paleozoic rocks. It may be that Paleozoic rocks have been buried to greater depths in the northern Cordillera, perhaps by stacking of décollement thrust faults, than has been recognized by interpretations of regional geophysical surveys. Most structural interpretations are based on the uppermost exposed thrust sheet, which reportedly has a shallow detachment across much of the Selwyn Basin (Gordey and Anderson, 1993; Gordey, in press). Nothing is known about the geometry, depth, or age of the rocks in deeper thrust sheets, which almost certainly exist. If Paleozoic miogeoclinal rocks are an unlikely melt source for the Cassiar plutonic suite in most of the Cassiar Terrane, then older, but isotopically similar metasedimentary rocks with similar or the same sources of detritus are an alternative magma source. The only other supracrustal rocks with neodymium isotopic ranges that overlap with most of the Cassiar plutonic suite are the (<1.2 Ga, possibly 0.88–0.78 Ga) Mackenzie Mountains Supergroup (Fig. 3.2a). Data are scarce or entirely lacking for several other Paleo- to Mesoproterozoic supracrustal packages thought have been deposited as thick successions along the western margin of Laurentia (e.g., Cook et al., 2004; Snyder et al., 2009), although recent and ongoing studies should soon produce isotopic datasets for the Wernecke Supergroup (Furlanetto et al., 2009) and the Pinguicula and Fifteenmile groups (Medig et al., 2010, 2012; Halverson et al., 2012). Even the limited data available for the Mackenzie Mountains Supergroup do not overlap with the full compositional range exhibited by the Cassiar plutonic suite (Figs. 3.15b and 3.16a). Exposures of the Mackenzie Mountains Supergroup are compatible with their deposition on a platform fringing a promontory that extended into an epicratonic basin in northwest Laurentia (e.g., Rainbird et al., 1996). Given their interpreted platformal setting, it is unlikely that these specific rocks comprised an aerially extensive unit that underlies the Cassiar Terrane. Outcropping rocks in the Mackenzie Mountains are thought to delineate the approximate extent of the platformal facies to this Supergroup (e.g., Evenchick et al., 2005). The recent identification of Mackenzie Mountains Supergroup (Turner, 2011) and confirmation of temporal correlations with the Fifteenmile Group (e.g., Halverson et al., 2012) rocks in the Wernecke Mountains suggests that much remains unknown about the distribution of these rocks. However, most studies have shown that this is a platformal package dominated by thick carbonate rocks and containing evaporites (e.g., Rainbird et al., 1996). More distal epicratonic basinal rocks (if they exist west or south of this promontory), however, might comprise mainly fine-grained siliciclastic rocks deposited in a deeper setting. If distal basinal equivalents to the exposed platformal rocks had similar sources of detritus they would also share similar isotopic compositions, but not the sulfur-rich evaporitic and calcareous lithologies. Basinal siliciclastic lithologies would also provide a more fertile (i.e., mica-rich) melt source for fluid-absent partial melting if they were 204 buried to lower–middle crustal depths in the Early Cretaceous. Basinal equivalents to Mackenzie Mountains Supergroup cannot be evaluated as a potential melt source to the Cassiar plutonic suite without additional isotopic data. Another possible supracrustal melt source for the Cassiar plutonic suite is the <1.4–1.0 Ga Pinguicula Group―particularly its lower portions (Unit A), which comprise mainly basinal, laminated mudstones lacking graphite, but containing some diagenetic pyrite (Medig et al., 2010, 2012). Like the Mackenzie Mountains Supergroup (Rainbird et al., 1997) detritus in the Pinguicula Group is also thought to have been originally sourced by a pan-continental river system (Medig et al., 2010, 2012) and, therefore, may exhibit a similar isotopic range to that of the Mackenzie Mountains Supergroup. Exposures of Unit A thicken up to 1400 m towards the south (Medig et al., 2010, 2012), which may be consistent with the late Mesoproterozoic existence of a deep basin to the south (e.g., Thorkelson et al., 2005). If so, this basin would have been located roughly in the region underlain by the Selwyn Basin, although there are no constraints on how far south along the paleocontinental margin it may have extended. Although isotopic and geochemical data for the Pinguicula Group rocks cannot yet be compared to the compositional range of the Cassiar plutonic suite, a detailed study to produce these data is underway (K. Medig, 2012, pers. comm.). 3.5.3.3 Implications Regarding the Age and Composition of Cordilleran Basement None of the Paleoproterozoic proxies for crystalline basement have neodymium isotopic compositions that overlap with the plutonic rocks in the Cassiar Terrane (except for εNd100 for Archean mafic orthogneisses; Fig. 3.16a). These examples of crystalline basement are all much more isotopically evolved that the Cretaceous magmatic event. If older crystalline basement underlies the Cassiar Terrane, the Cassiar suite plutons do not provide any information on the age and composition of this basement as it would not have been involved in the genesis of this syn- to post- accretionary magmatism. This suggests that isotopically highly evolved crystalline basement was either infertile in terms of producing voluminous partial melts (i.e., comprising anhydrous and (or) granulite-facies rocks; Patiño Douce et al., 1990) or that it was not present in the northern Cordillera in the Cretaceous. If anhydrous crystalline basement was present below the Cassiar Terrane during the Cretaceous, it is unlikely to have been very thick (e.g., Cook et al., 2004) because the depths required for crustal anatexis suggest that the supracrustal melt source was buried to at least the lower–middle crust (e.g., >25 km; Sawyer et al., 2011). The isotopic similarities exhibited by Early Cretaceous crustally derived intrusions from east–central Alaska to southeast British Columbia (Section 3.4.3) are probably due to a common source of both isotopically evolved and unevolved detritus from the Laurentian craton to the east. This implies that a large and deep basin may have existed along the western margin of North America at the time that the supracrustal protolith to the Cassiar suite was deposited. Not only are sulfur-rich evaporite (and carbonate) rocks absent or rarer in a deeper depositional setting, but pyritic sedimentary rocks may be limited if the basin is generally oxidized. Furthermore, given that Neoproterozoic rocks are isotopically unlike the Cassiar plutonic suite (and that Paleozoic sedimentary rocks may be an improbable melt source for voluminous magmatism throughout the Cassiar Terrane) it is possible that a deep 205 intracratonic basin or passive margin existed prior to the 0.78 Ga onset of rifting in Rodinia (cf. Thorkelson et al., 2005). A deep marine depositional setting would be compatible with the interpretation of Snyder et al. (2009) that most or all of the westward-tapering wedge of low resistivity rocks in at least the northern Cordillera comprise extensive submarine fan deposits. These thick supracrustal packages were interpreted by Cook et al. (2004) to be distal equivalents to the Mesoproterozoic Muskwa Assemblage or the Mackenzie Mountains Supergroup. Like Neoproterozoic sedimentary rocks, the Muskwa Assemblage is much too isotopically evolved to be the primary source of Cretaceous crustal melts. Distal equivalents to the Mackenzie Mountains Supergroup deposited in an off- shelf setting or a deep intracontinental basin, however, are a potential melt source to the Cassiar plutonic suite. The main supracrustal protolith to the Cassiar suite probably comprised mainly siliciclastic lithologies, ranging from immature quartzo–feldspathic rocks (e.g., greywackes and psammitic shales), to more mature clay-rich rocks (e.g., siltstones and mudstones), with little to no evaporitic, coaly, or carbonate horizons. The supracrustal package may have later been intruded by Paleozoic OIB-like mafic sills, which (together with interlayered supracrustal rocks) provide a reasonable melt source for the isotopically juvenile, alkaline to alkali-calcic Cassiar suite intrusions. Geophysical surveys suggest that aside from the uppermost crust, the basement to the Cassiar Terrane comprises mainly circa 1.7–0.5 Ga metasedimentary rocks that are 16–24 km thick (e.g., Cook et al., 2004; Snyder et al., 2009). A thick package of such supracrustal rocks is consistent with measured stratigraphic thicknesses for exposed Paleo- to Mesoproterozoic packages to the east and north (Fig. 3.2a; e.g., Gabrielse and Yorath, 1991; Thorkelson et al., 2005)―although basin-ward facies could be expected to be even thicker. The results of this study are compatible with a thick package of fine-grained, siliciclastic, deep marine sedimentary rocks, as old as latest Mesoproterozoic age and as young as Paleozoic age, deposited along the length of the Canadian segment of western Laurentia. The presence of Paleozoic basinal rocks structurally buried to mid-crustal depths, however, is the most viable possibility given the isotopic composition of the Cassiar plutonic suite—and coeval crustally derived plutonic rocks from east– central Alaska to southeast British Columbia. 3.5.4 Sources of Isotopic Variation in the Post-Arc A-type Plutonic Suites Isotopic data for the Seagull, Rancheria, and Black River A-type plutonic suites are limited, but there are some compositional features that make it distinct from the Cassiar plutonic suite. The Seagull suite has low initial strontium ratios (Fig. 3.15a) that are not much more isotopically evolved than the enriched mantle compositions (Abraham et al., 2000; Faure, 2001), and 87Sr/86Sr evolves at faster rates in all of the A-type suite samples because of their high Rb/Sr (Figs. 3.6a and 3.14). The Rancheria and Black River suites are form a linear trend with more primitive lead isotopic compositions (i.e., lower 207Pb ± 206Pb and higher 208Pb; Fig. 3.7) than the Cassiar suite. Sulfur contents and δ34S are also generally lower in the A-type granitoids. Like the Cassiar plutonic suite, however, the A-type intrusions exhibit a range of high positive δ18O, and the Rancheria and Black River suites have moderate initial neodymium isotopic compositions, both of which are within the isotopic range of the Cassiar plutonic suite (Fig. 3.15a). The mantle origin that was suggested for these plutonic suites, on the basis of geochemistry in Chapter 2, is 206 consistent with the unevolved radiogenic isotopic compositions of most A-type suites worldwide (Eby, 1990, 1992; Turner et al., 1992; Vander Auwera et al., 2003; Namur et al., 2011). Remarkably low strontium contents in A-type magmas, however, makes them particularly sensitive to contamination by radiogenic, strontium-rich crustal rocks and may be one reason why some these granites have high initial strontium ratios (e.g., 87Sr/86Sri = 0.71–0.72; Turner et al., 1992; Bonin et al., 1998; this study). It is also possible that the undepleted or enriched mantle source of the parental A-type magmas had evolved isotopic compositions not normally associated with mantle rocks (Eby, 1992; Bonin et al., 1998). Subduction-related enrichment in incompatible elements may characterize some old lithospheric mantle (e.g., Mair et al., 2011). Perhaps the deeper mantle also assimilates, or is contaminated, by subducted oceanic crust (comprising hydrothermally altered tholeiitic rocks and pelagic sedimentary rocks; cf. Hofmann, 2003) with higher δ18O and more evolved neodymium, strontium, and lead isotopic compositions. The extremely low strontium and europium concentrations, extremely enriched rubidium concentrations, and low initial 87Sr/86Sr typical of A-type plutonic suites is also compatible with a derivation from biotite dehydration in crustal rocks. While partial melting of sedimentary rocks and magmatic fractionation do not appear to produce the extreme trace element enrichments that typify A-type magmas (at least not with respect to heavy REEs; Chappell et al., 2000; but see example of Breiter, 2012, for a possible exception), arguments have been made for the generation of some A-type magmas from high-temperature partial melting in relatively anhydrous metaigneous rocks of tholeiitic to tonalitic composition (Collins et al., 1982; Creaser et al., 1991; Frost and Frost, 1997; 2011; Yang et al., 2003; Kemp et al., 2005; and “granitic series” example of Yang et al., 2008). Presumably, these melts incorporate any biotite in these relatively anhydrous protoliths, and entrains the contained HFSE- and HREE-rich refractory accessory minerals. Plagioclase or hornblende in the melt source, however, cannot be important reactants given the extremely depleted Sr and Eu/Eu* in A-type magmas. The resulting magmas, therefore, will have peraluminous calc-alkaline compositions (Frost and Frost, 2011). It is not unusual to see an association between metaluminous to peraluminous calc-alkaline intrusions, and (per)alkaline to alkali-calcic A-type intrusions (e.g., Frost and Frost, 2011), as is apparent in the Seagull and Rancheria plutonic suites that are examined here (Chapter 2). However, calc-alkaline and peraluminous compositions are only exhibited by a small proportion of the most siliceous Seagull plutonic suite samples (Chapter 2). The least siliceous Seagull and Rancheria suite samples have peraluminous alkali-calcic to (per)alkaline compositions and more extreme trace element enrichments and depletions (Chapter 2). Therefore, the more isotopically evolved and less geochemically ‘enriched’ examples of the Seagull, Rancheria, and Black River suites may have been partly to completely derived from infracrustal rocks. Their association with crustal discontinuities—and alkaline magmas in the Seagull and Rancheria plutonic suites—is, however, compatible with an origin in the mantle. 3.5.4.1 An Evaluation of Possible A-type Magma Sources The initial 87Sr/86Sr calculated for the Seagull plutonic suite at 100 Ma, and for the heavy REE enriched Allen Stock (Rancheria plutonic suite) at 80 Ma range from 0.706 to 0.708 (Fig. 3.5a), and is compatible with a primitive magma source (Rollinson, 1993; Faure, 2001). More evolved neodymium isotopic compositions and Paleoproterozoic TDMs exhibited by the Rancheria and Black River suites (Figs. 3.6 and 3.16) may appear to be inconsistent with an origin 207 in undepleted mantle or in mantle-derived crust. However, when they are compared with an εNd100 range if –7.8 to – 11.5 (Appendix G.3; Mair et al., 2011) for shoshonitic lamprophyre dykes in the northwest Selwyn Basin at Scheelite Dome, an enriched mantle origin does not seem as improbable. It is possible that the mantle region from which A-type magmas in the Cassiar Terrane may have been derived was similarly isotopically evolved (cf. Bonin et al., 1998). Thus, evolved neodymium compositions in A-type plutonic suites may occur in mantle-derived magmas if the mantle source was enriched in light REEs (i.e., Nd). Similarly, contamination by light REE enriched crust may add non-radiogenic neodymium to the magmas during their ascension through the crust. Sm/Nd might also have been affected by fractionation of these two elements during anatexis or magmatic differentiation. The crystallization of clinopyroxene in a basaltic magma, for example, should produce an early depletion in samarium relative to neodymium and, ultimately, a less radiogenic neodymium isotopic composition if the magma continues to fractionate to a highly siliceous end-product. Instead of a crustal origin, therefore, evolved initial neodymium and strontium ratios for one Rancheria and one Black River suite samples might be due to a combination of contamination by isotopically evolved crust, anatectic or magmatic fractionation from a mafic parent, and (or) an origin in isotopically evolved mantle. The high positive δ18O for the A-type granites in the Cassiar Terrane is, however, difficult to reconcile with a derivation from the mantle or mafic igneous rocks. Unlike the radiogenic (and sulfur) isotopic systems, oxygen isotopes are not strongly affected by contamination processes in magmas (Rollinson, 1993), nor are they fractionated during anatectic or magmatic processes (Taylor and Sheppard, 1986). The least geochemically ‘enriched’ A-type plutons in the Cassiar Terrane not only have the heaviest oxygen and sulfur, but are the most isotopically evolved with respect to strontium and neodymium. While it is possible that an enriched mantle affected by the subduction of oceanic crust could have been contaminated with respect to δ18O, together, the stable and radiogenic isotopic data suggest that the peraluminous calc-alkaline members of the A-type suites (including both samples of the Black River suite) were derived primarily from crustal rocks. Earlier it was stated that mafic–intermediate infracrustal rocks are the most likely source of peraluminous calc-alkaline A-type magmas in the crust (e.g., Collins et al., 1982; Creaser et al., 1991; Frost and Frost, 1997; 2011; Kemp et al., 2005); however, this still does not explain the high positive δ18O that characterizes these plutons. Unless the infracrustal rocks also had ‘heavy’ oxygen isotopic compositions (e.g., hydrothermally altered oceanic crust, or igneous rocks derived from partial melting of supracrustal basement), the high positive δ18O characterizing A-type granites in the Cassiar Terrane remains unexplained. The possible primitive melt sources for the A-type plutonic suites in the Cassiar Terrane will be considered here: (a) subduction enriched mantle (e.g., Eby, 1992), (b) tholeiitic crust (e.g., Frost and Frost, 1997), or (c) isotopically evolved deep mantle (e.g., Eby, 1992; Bonin et al., 1998). The first possibility of mantle-derived magma sourced from a subduction-enriched lithosphere should be enriched in ‘incompatible’ trace element concentrations (e.g., Rb, Th, U, and Nb). This was suggested to be a suitable melt source for the Rancheria plutonic suite (which has more of a shoshonitic, incompatible element enriched character; cf. Bonin et al., 1998) in Chapter 2. This type of enriched mantle, however, is unlikely to be the source of the ‘compatible’ trace element enrichments exhibited by the Seagull and Black River plutonic suites (e.g., Y, Ce, Sm, and Yb; Chapter 2). 208 The second possible source of A-type granites in the Cassiar Terrane could be small-fraction partial melting of tholeiitic rocks in the basement, which will produce strongly reduced, geochemically enriched, alkaline, ferroan, and siliceous melts of A-type affinity (Frost and Frost, 1997). Underplated tholeiitic basalt will lack high positive oxygen isotopic compositions and is an unlikely melt source for the A-type magmas in the Cassiar Terrane. Tholeiitic rocks are typically sourced from depleted mantle—forming the bulk of oceanic crust—and may, therefore, have a relative enrichment in ‘compatible’ trace elements. Partial melting of either pre-existing (old) oceanic lower crust (cf. Snyder et al., 2009), or accreted (Paleozoic) oceanic rocks could produce the less radiogenic strontium and lead isotopic compositions of the A-type plutons. It could also be responsible for the high positive oxygen isotopic compositions if the oceanic crust was altered by seawater (e.g., Taylor, 1986; Rollinson, 1993; Pineau et al., 1999), and (or) if melt was also sourced from overlying sedimentary rocks (Sheppard, 1986; Pineau et al., 1999). The neodymium composition and TDM estimates for the A-type plutonic suites are similar to those of felsic igneous and siliciclastic rocks in the accreted oceanic terranes, although are much more evolved than most of the mafic rocks (Fig. 3.16c). The similarity between the Late Cretaceous Surprise Lake plutonic suite and mafic rocks of the Cache Creek and Slide Mountain terranes (Fig. 3.16c) led Morris and Creaser (2008) to suggest a source in their host rocks (i.e., the Cache Creek terrane) or similar deeper oceanic crust. However, mafic oceanic rocks are markedly more primitive with respect to oxygen and neodymium than the three A-type plutonic suites examined here (Figs. 3.5a,b and 3.13), making a tholeiitic melt source for the Seagull, Rancheria, and Black River plutonic suites unlikely. Furthermore, geophysical surveys have not identified evidence for mafic oceanic crust in the basement to the Cassiar Terrane; instead, the basement appears to be mainly of felsic, or quartzo–feldspathic composition (Cook et al., 2004; Hammer and Clowes, 2004; Ledo et al., 2004; Creaser and Spence, 2005; Evenchick et al., 2005; Fernández-Viejo et al., 2005; Jones et al., 2005). If tholeiitic crust was present at the base of the crust in the Cretaceous, then it must have later delaminated (a process for which evidence in the Cassiar Terrane is currently lacking; see Section 3.5.5.1), or it is geophysically indistinguishable from the mantle (Snyder et al., 2009). There is also the possibility that relatively A-type magmas in the Cassiar Terrane were derived from an isotopically evolved region of the mantle, perhaps one that was affected by subducted oceanic crust. Oceanic crust subducted below the Cassiar Terrane may have stalled in the lower lithospheric or upper asthenospheric mantle, eventually partially to completely assimilated into the mantle. This would alter the stable and radiogenic isotopic composition, and enrich the ‘compatible’ versus ‘incompatible’ trace element chemistry, of this region of the mantle. Subduction of oceanic crust below the western margin of North America was almost certainly occurring in the Early Cretaceous when the Teslin and Whitehorse continental magmatic arcs were formed west-southwest of the Cassiar Terrane (today’s coordinates; Figs. 3.3a and 3.4), however, older subduction that would have preceded accretionary events is also known to have occurred (e.g., Thorkelson et al., 2005; Nelson and Colpron, 2007). A significant proportion of oceanic crust must also have been subducted, for example, during the Late Paleozoic to Mesozoic given that the Slide Mountain and Cache Creek oceanic basins closed during this period. It is likely that the A-type plutonic suites in the Cassiar Terrane were sourced primarily from the mantle sources that were suggested in Chapter 2, on the basis of geochemistry. There are not enough isotopic data available to negate the geochemical interpretations. An undepleted upper asthenospheric to lower lithospheric mantle—that was perhaps 209 isotopically and geochemically contaminated by subducted oceanic crust—is the most probable source of the ‘compatible’ element enriched Seagull, and possibly Black River, plutonic suites. An enriched lithospheric mantle that produces magmas of shoshonitic affinity (enriched in ‘incompatible’ elements) is the most probable primitive source of magma for at least the alkaline members of the Rancheria plutonic suite. Anatectic and (or) magmatic fractionation affecting parental magmas sourced in the mantle, could be partly to entirely responsible for the highly siliceous, peralkaline to peraluminous, and geochemically enriched composition of A-type magmatism in the Cassiar Terrane. Contamination by supracrustal rocks is another possible source of high positive δ18O values exhibited by all three A-type suites, as well as moderately evolved 87Sr/86Sri and εNdi in the Rancheria and Black River suites. Finally, the incorporation of infracrustally derived melt into some of the mantle-derived parental magmas could explain the association between alkaline (mantle-derived) A-type magmas, and more siliceous and isotopically evolved, and less geochemically extreme (enriched or depleted), peraluminous calc-alkaline A-type magmas (Chapter 2; e.g., Kemp et al., 2005; Frost and Frost, 2011). 3.5.5 Petrogenesis of Syn- to Post-accretionary Magmatism in the Cassiar Terrane 3.5.5.1 Early Cretaceous (115–104 Ma) The generation of voluminous Early (to ‘mid’-) Cretaceous magmatism in the Canadian Cordillera inboard of typical calc-alkaline continental magmatic arc has historically been attributed to orogen-related thickening of metasedimentary strata in the mid-Jurassic (Woodsworth et al., 1991; Brandon and Lambert, 1993, 1994; Driver et al., 2000; Mortensen et al., 2000; Hart et al., 2004a). Although locally exhumed Upper Proterozoic rocks in north– central British Columbia were buried to depths of 18–25 km, regional metamorphic grades for these rocks at the surface are suggestive of no more than ~15 km burial in the Cassiar Terrane (e.g., Greenwood et al., 1991). Given current crustal thicknesses of 33–36 km (e.g., Clowes et al., 2005), a total crustal thickness in the Cassiar Terrane of no more than about 50 km during peak (Jurassic) metamorphism (Driver et al., 2000). This is comparable to estimates of about 50 km crustal thickness for most Mesozoic–Cenozoic orogenic belts (Rudnick and Fountain, 1995). One effect of crustal thickening during orogeny is that it depresses the crustal geothermal gradient, at least initially, which leads to lower to mid-crustal temperatures of 500–650°C for 10’s of m.y. (Thompson, 2001). These temperatures will not produce melt in the crust, except perhaps from muscovite-bearing lithologies in the presence of free water (an occurrence that can only be of local importance; e.g., Barbarin, 1996; Kemp and Hawkesworth, 2003). Vapour-absent partial melting of muscovite-bearing assemblages begins at about 750°C (e.g., Thompson, 2001), but the resulting small melt volumes (<15%; Thompson and Connelly, 1995; Clemens and Vielneuf, 1987; Nabelek and Bartlett, 2000; Rushmer, 2001) and higher water contents of these melts will lead to mainly in situ migmatite complexes or small leucocratic plutons (Patiño Douce et al., 1990; Kemp and Hawkesworth, 2003). Voluminous, intermediate–felsic granitoid belts are only produced after the biotite solidus is intersected (~850°C; Clemens and Vielneuf, 1987; London, 1995; Thompson and Connelly, 1995; Patiño Douce and Beard, 1995; Nabelek and Bartlett, 2000; Rushmer, 2001; Thompson, 2001). Temperatures above ~800°C are not characteristic of 210 most lower to middle sections of the crust, however, and can only occur if the rocks are enriched in radioactive elements (an internal source of heat), or if there is an external source of heat (e.g., Bea, 2011). The radioactive decay of uranium, thorium, and potassium-rich rocks containing hydrous minerals can produce enough internal heat to intersect the muscovite ± biotite solidi in lower–middle crustal rocks (e.g., Patiño Douce et al., 1990). This requires a relative enrichment in the radioactive elements (e.g., Clark et al., 2011), not unlike those described in supracrustal rocks from central Spain and the central North China craton (Villaseca et al, 1999; Liu et al, 2001). The most fertile rocks for fluid-absent partial melting are amphibolite-grade (or lower) metasedimentary lithologies containing abundant biotite (and sodic plagioclase) that are present at lower to mid-crustal depths (e.g., Patiño Douce et al., 1990). In order to generate melt in these rocks, a 50–100% increase in crustal thickness must occur during an orogeny and must be maintained for several 10’s of m.y. while the geothermal gradient re- equilibrates to the new crustal architecture (e.g., England and Thompson, 1984; Patiño Douce et al., 1990; Bea et al., 2003; Nabelek and Liu, 2004; Bea, 2011; Clark et al., 2011). Modeling of a 10–15 km-thick fertile lithology in over- thickened Cordilleran crust (of similar composition and thickness to that inferred for the northern Cordillera in the Early Cretaceous) suggests that extensive biotite-dehydration can occur if the above geologic conditions are met (Patiño Douce et al., 1990). This model is compatible with northern Cordilleran geology, and it supports earlier hypotheses that orogen-induced crustal thickening produced plutonic rocks with the S-type characteristics in the northern Cordilleran (Woodsworth et al., 1991; Brandon and Lambert, 1993, 1994; Driver et al., 2000). It provides a satisfactory explanation for genesis of the majority of the Cassiar suite plutons―the calc-alkaline biotite- and muscovite-bearing phases―from a lower to mid-crustal package of metasedimentary rocks that is relatively enriched in LILEs. Despite thermal modeling of the crust that suggests crustal thickening combined with elevated internal heat is adequate for the generation of extensive crustal melting in fertile rocks in the absence of fluid (Patiño Douce et al., 1990; Bea, 2011; Clark et al., 2011), this idea remains contentious. Other studies and reviews of the matter invoke an external source of heat, or the influx from the mantle (England and Thompson, 1984; Dorais and Paige, 2000; Clemens, 2006), to initiate melting reactions at temperatures of ~850°C for biotite dehydration, and 900 ± 50°C for amphibole dehydration (Patiño Douce et al., 1990; Rushmer, 1991; Brandon et al., 1993, 1994; Thompson, 2001). Given the spatial (although temporally unconfirmed) association between hornblende-bearing intrusive phases and biotite- and muscovite-bearing phases of the Cassiar plutonic suite, an external source of heat must is required (e.g., Patiño Douce et al., 1990). Conductive heat transfer from an unusually hot mantle, and (or) advective heat transfer through the injection of mantle-derived magmas (e.g., Petford and Gallagher, 2001), can increase the heat flux from mantle to crust substantially. This high heat flux occurs during hydration and melting of the upper mantle above a subducting oceanic plate, or during upwelling of hot asthenospheric mantle. The former scenario is limited to crust overlying the “mantle wedge” and is important in the generation of magmatic arcs. The latter scenario is more broadly applicable to other tectonic settings, and has been suggested to be important in the generation of post-arc (i.e., post-collisional) to back-arc (i.e., syn-collisional, but inboard and distal to the subduction zone) to within-plate (i.e., anorogenic) igneous rocks. Two mechanisms that would be capable of inducing asthenospheric upwelling in a compressive back-arc—like the Cassiar Terrane in the Early Cretaceous—include: (1) the convective or thermal 211 detachment of lithospheric mantle ± the mechanical removal (delamination) of lithospheric mantle and (or) lower crust (e.g., Schott and Schmeling, 1998); or, (2) high heat-flow related to the deep dehydration of a subduction slab below the back-arc region to an active continental magmatic arc (e.g., Currie and Hyndman, 2006). The probability of each are considered below for back-arc magmatism in the Cassiar Terrane relative to geological constraints. A model invoking the convective removal of over-thickened lower lithosphere (after the attainment of ~60 km crustal thickness in the Jurassic) explains the timing of regional metamorphism and a magmatic lull, followed by exhumation and extensive crustally derived granitoids of the Yukon–Tanana Uplands plutonic suite (Mair et al., 2006b). In east–central Alaska, the widespread Early Cretaceous exhumation of high-grade (upper amphibolite facies) metamorphic basement is evidenced by a large number of pre- to syn-magmatic 40Ar/39Ar cooling ages and U–Pb metamorphic ages, as well as shallow-dipping extensional faults (Dusel–Bacon et al., 2002; Pavlis et al., 1993; Berman et al., 2007). In comparison, late Neoproterozoic and early Paleozoic miogeoclinal rocks in the Cassiar Terrane have only been regionally metamorphosed to greenschist facies (e.g., Greenwood et al., 1991). Amphibolite-grade rocks related to burial metamorphism are only observed in the Cassiar Terrane, where Eocene uplift along steeply dipping dextral strike-slip faults in the Horseranch and Sifton ranges (Fig. 3.2a) has exhumed Neoproterozoic gneisses that experienced Jurassic to Cretaceous peak metamorphism (Greenwood et al., 1991; Plint et al. 1992). In contrast to the plutonic association with shallowly dipping extensional faults in east–central Alaska, steeply dipping transpressional structures dominated the syn-magmatic deformational regime in the Cassiar Terrane (e.g., Gabrielse, 1991; Murphy et al., 1995; Gabrielse et al., 2006; Mair et al., 2006b; Colpron et al., 2011). Thus, the Early Cretaceous setting in the Cassiar Terrane was compressional and produced comparatively little uplift. Any tectonic model for Early Cretaceous magmatism invoking the catastrophic removal of the lower lithosphere is, therefore, highly improbable for the Cassiar Terrane. The observed high heat flow (Lewis et al., 2003) and the setting of Early Cretaceous magmatism in the Cassiar Terrane could be related to anomalously high, back-arc heat flow that may persist up to 300 m.y. after subduction ceases (Hyndman et al., 2005b; Currie and Hyndman, 2006). Elevated heat flow has been attributed to vigorous small-cell convection and reduced viscosity above deep portions of subducting oceanic lithosphere that is releasing fluid to the asthenospheric mantle. The addition of this fluid and the resulting convection produces an efficient mechanism for the conductive and advective transfer of heat from the asthenospheric to lithospheric mantle (e.g., Currie and Hyndman, 2006). This in turn increases the heat flux into the crust where anatexis of a relatively radiogenic supracrustal package at lower-middle crustal depths and temperatures of at least 850°C would produce voluminous high-K calc-alkaline magmas. A 112 Ma “lamprophyre” dyke (Ash, 2001) emplaced near the Cassiar batholith in northern British Columbia is evidence for Early Cretaceous high heat flow in the lithospheric mantle.45 Finally, in addition to high back-arc heat flow, movement along steeply dipping, dextral transpressional faults beginning as early as 140 Ma (Gabrielse et al., 2006) may also have played an important role in the production and 45 High heat flow at the base of the crust may also have contributed to thermal erosion or metamorphism of the lowermost crust, removing geophysical contrasts between pre-existing crystalline basement and the underlying mantle (and resulting in a possible juxtaposition of supracrustal basement and the Moho below the Cassiar Terrane; cf. Snyder et al., 2002). 212 localization of Early Cretaceous crustal melts (e.g., Sylvester, 1989; Glazner, 1991; Gabrielse et al., 2006; Colpron et al., 2011). These faults could have been important in facilitating the ascension and emplacement of the Cassiar suite magmas into the upper crust. Localized decompression related to Early Cretaceous movement along these faults may also have triggered, or contributed to, partial melting in the crust. These structures are geographically associated with intrusive rocks in the Cassiar Terrane and may be responsible for the geometry of Cassiar suite batholiths (elongated to the north-northwest; Figs., 3.3 and 3.4), and the intensely deformed plutonic margins in some of the Cassiar suite intrusions. An association with strike-slip faults has been suggested for other regions where voluminous, post-collisional, high-K calc-alkaline granitoids that are similar to the Cassiar plutonic suite (with respect to chemical composition and inferred tectonic setting; Chapter 2) were emplaced (e.g., the Tuareg Shield; Liégeois et al., 1994, 1998). These steep structures may be intrinsic to the generation and emplacement of large, intermediate to felsic granitoid provinces experiencing transpressional deformation, like the Early Cretaceous syn- to post-accretionary event in the Cassiar Terrane. 3.5.5.2 ‘Mid’ Cretaceous (103–94 Ma) The first intrusions associated with the ‘mid’-Cretaceous Seagull plutonic suite appear to have been generated in a post-arc setting, beginning immediately after the youngest back-arc Cassiar plutonic suite intrusions were emplaced. This post-arc tectonic setting would still have been experiencing high heat flow related to small-scale convection and asthenospheric upwelling (Hyndman et al., 2005b; Currie and Hyndman, 2006), as is evidenced by rare lamprophyric dykes that are spatially associated with this suite (Chapter 2; T. Liverton, pers. comm.). In addition to its younger age, the Seagull suite is also compositionally distinct from the Cassiar suite, and is interpreted to have been derived mainly from isotopically contaminated upper asthenospheric to lower lithospheric mantle. Although vigorous asthenospheric to lithospheric convection below the Cassiar Terrane may explain the widespread and more voluminous Cassiar suite plutonic suite, it does not offer an explanation for the generation and (or) emplacement of the mainly mantle-derived Seagull plutonic suite as a narrow belt sub-paralleling dextral strike-slip structures and the western margin of the Cassiar Terrane (Figs. 3.3b and 3.4). This narrow belt is more compatible with localized anatexis along a deep crust- to lithosphere-scale discontinuity (e.g., Liverton et al., 2005; T. Liverton, 2012, pers. comm.). Possible discontinuities include steeply dipping normal faults related to the inferred Late Proterozoic upper plate rifting geometry (cf. Cecile et al., 1997), the nearby terrane boundary, a Cretaceous strike-slip fault, or a sub- crustal control. Asthenospheric upwelling associated with lithosphere-scale structures is thought to be important in the generation of post-collisional A-type magmas derived from upper asthenospheric to lower lithospheric mantle. Post-collisional intrusions that are compositionally analogous to the Seagull suite, for example, were emplaced during “stress reversals” along shear zones in the transpressional Tuareg Shield of the West African craton (Liégeois et al., 1998). These structures either facilitate the upwelling and decompression-related partial melting of the mantle, or they facilitate the transfer of heat to the lithosphere to melt the crust at high temperatures (e.g., Davies and von Blanckenburg, 1995; Liégeois et al., 1998; Kemp et al., 2005). Brief periods of asthenospheric upwelling are consistent with the short duration of individual magmatic episodes of this composition, both of which may be 213 controlled by sporadic movements along the deep structures (Liégeois et al., 1998). If the Seagull suite emplacement has a structural component to it, localized decompression along the structure may have facilitated the production and ascent of small volume magmas. Syn- to post-magmatic deformation in these plutons, however, is almost entirely absent (the strongly foliated, ~96 Ma ‘Last Peak granite’ emplaced along the d’Abbadie Fault being the sole exception; Fig. 3.4; Gallagher, 1999). It is unlikely that a fault between the Yukon–Tanana terrane and the Cassiar Platform terrane boundary, which is thought to be shallow-dipping (e.g., Cook et al., 2004) localized this plutonic suite—although some geophysical results are more compatible with a steeper contact (e.g., Spence and Creaser, 2005). Nor have steeply dipping terrane boundaries and crustal-scale faults been geophysically detected in the western Cassiar Terrane (e.g., Cook et al., 2004). Thus, an alternate explanation for the emplacement of this narrow belt of magma may be a sub-crustal control on the generation and ascension of upper asthenosphere-derived melts. Termination of subduction, or stalling of the subducting slab, typically slows or ends hydration of the upper mantle wedge above slab, which in turn leads to a cessation of arc magmatism. This typically occurs after the accretion of a block of continental crust, which is too buoyant to allow subduction to continue. Therefore, the common occurrence of post-accretionary magmatism, several to 10’s of m.y. after pre- or syn-collisional magmatism, suggests a role for post-subduction processes in its genesis. The result of down-dip slab drag and up-dip buoyancy on a stalled subducted slab produces tensional stress in the subducted slab. Tearing, or delamination, of this slab at the locus of this stress may result, and should produce linearly focused asthenospheric upwelling and relatively small volumes of magma that are limited to the area of the heat pulse (Davies and von Blanckenburg, 1995; von Blanckenburg and Davies, 1995; Cloos et al., 2005). Terrane accretion and post-collisional tearing in the subducted oceanic crust is one mechanism of producing a narrow belt of mantle-derived igneous rocks in the Cassiar Terrane. In this scenario, the extensional stress would have been focused below the western margin of the Cassiar Terrane, at a relatively deep position in the subducted plate. This deep tearing would allow upwelling of deeper (asthenospheric) parental magmas that appear to have been important in the generation of the Seagull plutonic suite. If these hot magmas ascended through the mantle lithosphere to pool at the base of the crust, they also interact extensively with granulite- grade crustal rocks (possibly generating the associated peraluminous, calc-alkaline infracrustal intrusive phases). Highly fractionated and contaminated derivatives of the parental magma that ascended through the crust, together with sub-ordinate crustally derived magmas, may have produced the Seagull plutonic suite. 3.5.5.3 Late Cretaceous to Eocene (87–75 Ma and 65–47 Ma) The 82–75 Ma Rancheria and 49–47 Ma Black River plutonic suites were both emplaced in proximity to major crustal, possibly lithospheric, structures (Figs. 3.3b and 3.4). The Rancheria suite forms a linear belt that is sub- parallel to the Kechika Fault, a north-trending regional strike-slip structure along which ~300 km of mainly Cenozoic displacement occurred (Evenchick, 1988; Gabrielse et al., 2006). The Black River suite was emplaced immediately southwest of the Tintina Fault (another major transtensional Cenozoic structure) along which ~430 km of dextral offset occurred (e.g., Gabrielse et al., 2006). The Black River plutons are the northernmost examples of several scattered Eocene granites in the Cassiar Terrane that were emplaced southwest of the Tintina Fault, albeit in closer association with the Kechika and other nearby faults (Fig. 3.3b). Both the Rancheria and Black River plutonic 214 suites are geochemically compatible with respective lithospheric and asthenospheric mantle origins; however, isotopic data requires that this mantle was either unusually evolved, or that some or all members of both suites were strongly contaminated by—to entirely sourced from—the crust. Mechanisms for mantle upwelling and generation of these mantle and (or) crustally derived A-type magma in the Cassiar Terrane probably include a combination of regional high heat flow (resulting from several episodes of east to northeast-dipping Early to Late Cretaceous subduction), and localized crustal- to lithospheric-scale extension along shear zones. The Rancheria plutonic suite temporally overlaps the earliest (78–72 Ma) Carmacks Group continental magmatic arc episode in western Yukon, which is generally thought to be related to the onset of north- northeast dipping, highly oblique subduction of the Kula Plate (Mortensen and Hart, 2010). However, the timing and composition of the Rancheria suite intrusions imply that this plutonic suite is not a product of back-arc processes. The Late Cretaceous Rancheria suite magmatism may instead be a response to the appearance of the oceanic Kula Plate and an increased obliquity in plate motions along the continental margin circa 85 Ma (Engebretson et al., 1985; Gabrielse and Yorath, 1991; Kelley, 1993). The change in plate vectors and associated deformation and localized decompression along inboard shear zones was probably more important than subduction along the continental margin in the generation of this A-type suite. Similar arguments may be made for the Black River plutonic suite and changes to plate orientations circa 50–45 Ma, although subduction-related magmatism of the Coast Plutonic Complex appears to have terminated just before the emplacement of the two Black River suite plutons (Fig. 3.3b) with the ~50 Ma subduction of an oceanic spreading centre (e.g., Madsen et al., 2006). Thus, the Black River plutonic suite (and compositionally similar 50–42 Ma plutons in the Cassiar Terrane) may have been controlled by far-field plate motions or re-organizations rather than subduction along the western margin of North America. 3.6 CONCLUSIONS 3.6.1 The Back-Arc Cassiar Plutonic Suite The neodymium and strontium isotopic data imply that the muscovite- and biotite-bearing magmas of the Cassiar suite originated from the same, albeit compositionally heterogeneous, melt source. The typically older age of the muscovite-bearing phases indicate that these phases are not the result of fractionation and (or) differentiation of the biotite-bearing magmas, despite the near-identical average isotopic composition. Instead, the muscovite-bearing magmas are the result of earlier and lower temperature, fluid-absent partial melting involving mainly muscovite dehydration of quartzo–feldspathic to pelitic rocks. The biotite-bearing magmas are the result of later and higher temperature, fluid-absent partial melting involving mainly biotite dehydration, in the same supracrustal protolith. The hornblende-bearing samples appear to have derived from a mixture of melts sourced from both supracrustal rocks and interlayered OIB-like (oxidized and alkaline) hornblende-bearing lithologies at the highest temperatures and latest stages of partial melting. Thus, neither the biotite or muscovite-bearing phases are the result of fractionation in the hornblende-bearing magmas. The isotopic and chemical variation in the Cassiar plutonic suite is most likely due to anatexis of a dominantly 215 siliciclastic package containing lithologically and isotopically heterogeneous rocks. Mineralogical differences, and depositional mixing of isotopically juvenile (e.g., the Coppermine flood ‘basalts’; Creaser and Erdmer, 1997) and isotopically evolved (e.g., Paleoproterozoic arc terranes in northwest Laurentia) detritus in one package of sedimentary rocks, may account for the isotopic variations between the quartzo–feldspathic and pelitic lithological end-members. In this way, most of the systematic isotopic and trace element variations in the Cassiar plutonic suite should be attributable to differing proportions of magma derived from quartzo–feldspathic and pelitic lithological end-members. Although the isotopic range of the Cassiar plutonic suite is very similar to Paleozoic miogeoclinal rocks, geologic constraints require more deeply buried or structurally duplicated equivalents, or an older (but isotopically similar), metasedimentary source of magma. Neoproterozoic and Paleoproterozoic to Archean basement are more isotopically evolved than the Cassiar plutonic suite, which leaves distal equivalents to the Mesoproterozoic rocks (e.g., the Mackenzie Mountains Supergroup or the Pinguicula Group) as the only other possible older magma source(s). There remains the possibility, however, that the Paleozoic miogeoclinal rocks are structurally duplicated at depth. The source of less isotopically evolved, alkalic to alkali-calcic plutonic phases of the Cassiar plutonic suite are oxidized and alkaline (OIB-like) rocks. Mafic volcanic and sub-volcanic rocks, also of Paleozoic age, provide the best isotopic proxy for this melt source. This source of isotopically juvenile magma may have been present in the form of sills and flows interlayered with a Paleozoic supracrustal melt source, and (or) as lithic detritus in the Paleozoic supracrustal rocks. What little sulfur is present in the Cassiar suite magmas was probably imparted by a small proportion of pyritic lithologies in the melt source, together with variable degrees of contamination during the ascension and emplacement of the magmas in the crustal column. The absence of evidence for sulfur from shallow-water carbonate and evaporite rocks makes siliciclastic rocks deposited in a deep epicratonic basin, or as off-shelf turbidite deposits (e.g., Snyder et al., 2009), the most likely protoliths to the Cassiar plutonic suite. Geophysical imaging of a resistive and seismically ‘slow’ westward-tapering “wedge” in the lower–middle crust has been attributed to off-shelf turbidite deposits (e.g., Snyder et al., 2002, 2009). Furthermore, there is also a remarkable isotopic similarity, particularly with respect to neodymium and strontium, between Early Cretaceous igneous rocks along the length of the northern Cordillera (i.e., from east–central Alaska to southeast British Columbia). This implies that these crustally derived magmas were sourced from an extensive package of isotopically and lithologically supracrustal rocks deposited in northwest Laurentia. Supracrustal rocks deposited in epicratonic basin are not likely to be extensive enough to produce compositionally similar intrusive rocks along a strike length of several 1000’s of km, making a continental margin a more appropriate depositional setting for the protolith rocks. Although the suggestion of a passive margin in the latest Mesoproterozoic may be contentious, it is acceptable for much of the Paleozoic. The interpreted lithological variability in the supracrustal melt source to the Cassiar plutonic suite would be consistent with proximal, thick, and heterogeneous turbidite deposits. A proximal, off-shelf protolith is also consistent with the emplacement of this suite near the Neoproterozoic upper plate rifted margin and Paleozoic continental shelf edge (Figs. 3.1, 3.2, and 3.3a). Although thickening and burial of a radioactively hot package of rocks might generate extensive magmas in some orogenic belts, models for the voluminous syn-accretionary magmatism in the Cassiar Terrane require some input of 216 conductive or advective heat from the underlying mantle to produce the associated hornblende-bearing (and possibly voluminous biotite-bearing) intrusions. High heat flow related to effusive asthenospheric upwelling above a subducting oceanic slab, in a transpressive back-arc, is the most likely the source of external heat (e.g., Currie and Hyndman, 2006). In addition to the elevated mantle to crust heat flux, local decompression along pre- to syn- magmatic dextral faults may have been important in triggering and localizing crustal melting (e.g., Sylvester, 1989; Glazner, 1991; Gabrielse et al., 2006; Colpron et al., 2011), as well as facilitating the ascent and upper crustal emplacement of the resulting magmas. 3.6.2 The Post-Arc A-type Plutonic Suites The unique compositions of the A-type plutonic suites emplaced at upper crustal levels in the Cassiar Terrane suggest an origin in the mantle. The Seagull and Black River plutonic suites have trace element compositions that are suggestive of a relatively deep source in undepleted mantle (i.e., upper asthenospheric to lower lithospheric; Eby, 1992; Bonin et al., 1998). The Rancheria plutonic suite has a trace element compositions that are suggestive of a shallower source in continental mantle (i.e., enriched upper lithospheric mantle; Eby, 1992; Bonin et al., 1998). The parental magmas to all three A-type suites appear to have been derived from relatively isotopically evolved mantle reservoirs, although some of the isotopic systems could have been affected by anatectic and magmatic fractionation, as well as by variable degrees of crustal contamination or assimilation (and mixing with crustally derived melts). The most isotopically evolved and least alkaline A-type intrusive phases, however, may have been derived from partial melting of the crust. Both the Seagull and Rancheria plutonic suites may contain both mantle- and crustal components, whereas the Black River plutonic suite may be entirely crustally derived. Although the Cassiar Terrane would have remained an area of high heat flow after arc magmatism to the west terminated (e.g., Currie and Hyndman, 2006), the A-type suites in this region were emplaced as narrow belts (Seagull and Rancheria suites) or as localized bodies (Black River suite) in proximity to crustal discontinuities (e.g., the approximate boundary between Yukon–Tanana terrane and the Cassiar Platform, and the Kechika and Tintina faults). These patterns imply a significant structural control on the generation, ascension, and emplacement of the A- type magmas. Asthenospheric upwelling along crust-penetrating or lithosphere-scale structures is one possible source of mantle-derived magmas. Aside from the Tintina Fault, there are no geophysical expressions of sub-vertical or steeply dipping features in the Cassiar Terrane (e.g., Clowes et al., 2005), although this does not preclude their presence. Another possible source of mantle-derived magmas, in the absence of sub-vertical lithosphere-scale structural discontinuities, is deep tearing in a subducted oceanic plate. Other important factors in the genesis of the A-type plutonic suites appear to be changes to the tectonic or deformation regime in the Cassiar Terrane that can be related to far-field plate vectors and stresses. Specifically, the timing of the Rancheria suite coincides with (or immediately follows) the appearance of the oceanic Kula Plate, and an increase to the obliquity of the angle of plate convergence along the western margin of North America (Engebretson et al., 1985; Gabrielse and Yorath, 1991; Kelley, 1993). Similarly, the emplacement of the Eocene Black River plutons (and other nearly coeval intrusions along the Tintina or Kechika faults) is thought to coincide 217 changes to plate configurations and vectors along the northwest margin of North America (Kelley, 1993; Madsen et al., 2006). 218 Figure 3.1. (next page) Basic elements of northern Cordilleran geology. a) Tectonostratigraphic map; the upper left inset shows the location of the map area relative to Canada. Basinal and platformal rocks of the northern Cordilleran miogeocline have been shaded pale yellow and pale blue, respectively. Shades of grey represent terranes accreted to the western margin of Laurentia in the latest Paleozoic to the Mesozoic. Bolded tectonostratigraphic abbreviations are: SB = Selwyn Basin, YTT = Yukon–Tanana terrane (not distinguished from the Stikinia and Quesnellia terranes), CC = Cache Creek terrane, CP = Cassiar Platform, CT = Cassiar Terrane, MeP = McEvoy Platform, MCE = Misty Creek Embayment, MkP = Mackenzie Platform, MdP = McDonald Platform, and KT = Kechika Trough. Regions of the Selwyn Basin commonly referred to in the text are indicated with the transparent NW, NE, W, and SE labels (YTU = Yukon–Tanana uplands). Geographic abbreviations are: YK = Yukon, AK = Alaska, NWT = Northwest Territories, and BC = British Columbia. Modified from Nelson and Colpron (2007), with contributions from Aitken and Pugh (1984), Cecile et al. (1997), Gordey and Anderson (1993), Gordey and Makepeace (1999), Morrow and Miles (2000), Gabrielse et al. (2006), and Turner and Long (2008). b) Exposed Proterozoic metasedimentary rocks, pre- to syn-magmatic Cretaceous sedimentary rocks, and Neoproterozoic, Paleozoic (Goodfellow et al., 1995), and Quaternary–Recent (Abraham et al., 2001) alkaline igneous rocks. The pale purple shading defines the known sub-surface of the Muskwa Assemblage detected by drilling (e.g., Snyder et al., 2002, 2009). The two red circles highlight areas where Cretaceous sedimentary rocks are preserved in the northern Cordillera. The two dashed red lines illustrate the approximate boundaries of the western edge of the foreland basin during the Albian (112– 99.5 Ma), Cennomanian (99.6–93.5 Ma), and Turonian (93.5–89.3 Ma; Thomson et al., 2011). Neoproterozoic magmatic rocks are: MH = Mount Harper Volcanic Complex, BP = Bonnet Plume intrusions, BR = Bear River dyke, LD = Little Dal ‘basalts’, and PC = Pool Creek syenite. Paleozoic volcanic rocks discussed in the text are: GH = Groundhog, LC = Looncry. Map lines are as in (a) and undivided Insular, Arctic, and Pacific terranes are indicated by the same shade of dark grey as in (a). 219 220 Figure 3.2. (next page) Approximate surface and sub-surface distribution of upper crustal rocks and Precambrian Laurentian basement in the northern Cordillera after the restoration of dextral Cenozoic displacement along major transpressional structures (modified from Nelson and Colpron, 2007). The western continental margin is represented by the heavy dashed line; it is not known what lay west of this line in the Early Cretaceous. Mainly Cenozoic oroclinal bending in Alaska may have begun by the Early Cretaceous (e.g., Gabrielse and Yorath, 1991), but has not been restored here. Map lines and tectonostratigraphic elements are as in Figure 3.1. The Tintina Fault, along which 430 km of strike-slip movement has been restored, is shown with a bolded solid black line. a) 1.8 Ga to present: latest Paleoproterozoic to Neoproterozoic metasedimentary packages (SG = Supergroup, G = Group, and Assem = Assemblage), and Proterozoic, Paleozoic, and Recent alkaline igneous rocks. The sub-surface extent of the Muskwa Assemblage indicated by drill holes is indicated by pale purple shading. The approximate location of the Neoproterozoic Horseranch orthogneiss (“H”) is shown by the bolded yellow polygon in the bottom right of the figure; the Sifton “Si” and Swannell “Sw” orthogneisses plot off the figure to the south–southeast. b) 3.8–1.8 Ga: Archean to Late Paleoproterozoic basement domains. Basement domains are labeled: Na = Nahanni terrane, FN = Fort Nelson “magnetic high” (of unknown origin; shaded light grey), FS = Fort Simpson magmatic arc or “magnetic high”, H = Hottah Terrane, GB = Great Bear magmatic arc, S = Slave craton, N = Nova domain, K = Ksituan domain, C = Chinchaga domain, BH = Buffalo Head domain, T = Taltson domain, and W = Wabamun (a cross-cutting mylonitic domain associated with the Snowbird Tectonic Zone). The approximate locations of the two drill holes that sampled the Fort Simpson magmatic arc are indicated by the circled crosses. The ‘known’ western sub-surface extent of the Nahanni Terrane is indicated by the red dashed line (although it is geophysically inferred to continue west to the edge of the Laurentian miogeocline; e.g., Aspler et al., 2003), and the distribution of Archean to Late Paleoproterozoic basement domains is modified from Ross (1991) and Aspler et al. (2003). 221 222 Figure 3.3. (next page) Distribution of syn- to post-accretionary, Early Cretaceous to Eocene igneous rocks in the northern Cordillera based on the geochronological results of this study, compiled robust age determinations, and the work of Anderson (1988), Murphy (1997), Mortensen et al. (2000), Hart et al. (2004a) and Mortensen and Hart (2010, unpub.). The Cassiar Terrane study area (Fig. 3.4) is outlined in black. Intrusive and volcanic polygons are primarily from Gordey and Makepeace (1999) and Massey et al. (2005), and modified based on the regional maps of Dusel–Bacon et al. (1993), Newberry et al. (1998), Mortensen et al. (2000), Baker and Lang (2001), Murphy et al. (2001), Hart et al. (2004a), Pigage (2004), Pigage et al. (2011), and Gordey (in press). There are some unresolved ages for magmatism in east– central Alaska that have led to the same bodies being labeled Early Cretaceous, Late Cretaceous, and (or) Paleocene. Map lines and tectonostratigraphic units are as in Figure 3.1. a) Early to ‘mid’-Cretaceous magmatic suites. The red belts represent continental arc magmatism; the purple belts represent back-arc magmatism; and the blue, orange, and yellow belts represent post-arc magmatism. The plutonic suites are labeled: A = Anvil, C = Cassiar, DR = Dawson Range, FS = Fairbanks–Salcha, H = Hyland, M = Mayo, S = Seagull, Te = Teslin, To = Tombstone, TR = Tay River, Tu = Tungsten, W = Whitehorse, YTT = Yukon–Tanana Uplands, and TT = Tok–Tetlin. b) Late Cretaceous to Eocene magmatic suites and (or) occurrences. The green belts represent Late Cretaceous magmatism, and the pink belts and stars represent Paleocene–Eocene magmatism. The plutono-volcanic suites are labeled: Ca = Carmacks Group (volcanic > plutonic), RR = Ruby Range (plutonic), NR = Nisling Range (plutonic), CPC = Coast Plutonic Complex (plutonic > volcanic), Sk = Skukum (volcanic), Sl = Sloko (volcanic), WT = Windy Table (volcanic), SL = Surprise Lake (plutonic), and MQ = McQueston (plutonic). Plutono–volcano occurrences mentioned in the text are labeled. 223 224 Figure 3.4. Distribution of syn- to post-accretionary plutonic suites in the Cassiar Terrane study area. Sample locations for this study (white symbols) and plutons referred to in the text are labeled. Sample locations for compiled isotopic data in the Cassiar Terrane are only available for the dataset of Morris and Creaser (2008) and are indicated by the black circles. Intrusions assigned to plutonic suites are coloured, and intrusions and volcanic rocks not assigned to plutonic suites here are shaded as per the legend. The extent of the Cassiar Platform is shaded pale blue. 225 Figure 3.5. (next page) Summary of isotopic data for each plutonic suite in the Cassiar Terrane. Data generated by this study are indicated by the large symbols. Compiled data, indicated by small symbols, are from Cooke and Godwin (1984), Dagenais (1984), Driver et al. (2000), Abraham et al. (2001), Liverton et al. (2001), and Morris and Creaser (2008). The red dotted box represents the compositional range for the Surprise Lake plutonic suite (discussed in the text; Dagenais, 1984; Sinclair, 1995; Mihalnyuk , 1999; Morris and Creaser, 2008). a) Whole rock εNdi and Sri (recalculated to initial ratios at 100, 80, or 50 Ma, as indicated). Average values for muscovite-, biotite-, and hornblende-bearing phases of the Cassiar plutonic suite discussed in the text are illustrated by the unfilled black boxes. b) 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb from magmatic feldspar (measured ratios). c) δ18O from magmatic quartz (unfilled symbols represent δ18O from whole rock analyses, which are ~1 ‰ lower than δ18O from quartz analyses) and whole rock δ34S. 226 227 Figure 3.6. (next page) Neodymium and strontium isotopic compositions of plutonic suites in the Cassiar Terrane. In addition to this study, data for the Cassiar plutonic suite are compiled from Abraham et al. (2001), Driver et al. (2000), and Morris and Creaser (2008). A granitic “migmatite” located near the Big Salmon and Nisutlin batholiths (sampled by Morris and Creaser, 2008) is indicated by the bolded black cross. Data for the arc samples (Teslin, Whitehorse, Carmacks, and the Coast Plutonic Complex or CPC) are compiled from Samson et al. (1991), Selby et al. (1999), Morris and Creaser (2008), and Bineli Betsi and Lentz (2011). The underlying enriched (pre-Jurassic) mantle (EM) is represented by the back- calculated isotopic compositions of Quaternary to Recent ‘basalts’ in the Cassiar Terrane (Abraham et al., 2001). Bulk Earth composition (CHUR) relative to neodymium is represented by the grey line at εNd = 0. Crosses represent xenoliths or rafts from the Cassiar batholith: x = metasedimentary xenolith; x = amphibolitic xenolith (Driver et al., 2000). Cassiar suite samples with anomalously elevated 147Sm/144Nd for their felsic composition are circled with the dashed blue line a) εNdi versus 87Sr/86Sri (at i = 100, 80, and 50 Ma). In addition to this study, data for the Cassiar plutonic suite are compiled from Morris and Creaser (2008), Abraham et al. (2001), and Driver et al. (2000). The possible compositional range of the Seagull plutonic suite are is illustrated by the red boxes (Sri from Liverton et al., 2001, and Morris and Creaser, 2008; εNdi ranges are estimated from Figure 3.9b based on the assumption that the all A-type samples form a linear array when radiogenic isotopic compositions are compared to δ34S). The “upper” and lower” isotopic perturbations exhibited by the Cassiar suite samples and discussed in the text are labeled. The initial strontium isotopic compositions for the Late Cretaceous Cassiar or Storie pluton and the Eocene Mount Haskin pluton are illustrated by the labeled red lines (Sinclair, 1995). b) εNd versus 147Sm/144Nd at 100, 80, and 50 Ma. c) εNd versus TDM (depleted mantle model age). 228 229 Figure 3.7. (next page) Lead isotopic compositions of plutonic suites in the Cassiar Terrane. Symbols are as in Figure 3.6, and compiled data for the Cassiar suite are from Abraham et al. (2001) and Driver et al. (2000). The Buttle Creek pluton discussed in the text is labeled “BC”. Data for the underlying enriched (pre-Jurassic) lithospheric mantle (EM) is represented by isotopic compositions for Quaternary to Recent ‘basalts’ in the Cassiar Terrane (Abraham et al., 2001). Data for igneous feldspars from the Carmacks continental magmatic arc are plotted for comparative purposes (Selby et al., 1999). The upper crustal Shale Curve of Godwin and Sinclair (1982) is represented by the heavy solid line. Common Pb (204Pb) and fractionation error vectors are indicated by thin black lines labeled “c” and “f”, respectively. Grey arrows show trends towards more radiogenic lead compositions (U–Th rich, felsic, unmetamorphosed or late-metamorphosed, and/or older), and the the red dotted line is the possible mixing trend for the A-type plutonic suites described in the text. a) 207Pb/204Pb versus 206Pb/204Pb (“uranogenic” plot). b) 208Pb/204Pb versus 206Pb/204Pb (“thorogenic” plot). c) 207Pb/206Pb versus 208Pb/206Pb. The possible compositional range of the Seagull plutonic suite is illustrated by the red boxes (lead isotopic ranges are estimated from plots of δ34S versus 207Pb/206Pb and 208Pb/206Pb, based on the assumption that the all A-type samples form a linear array when radiogenic isotopic compositions are compared to δ34S; see Fig. 3.9b). 230 231 Figure 3.8. (next page) Stable isotopic compositions of plutonic suites in the Cassiar Terrane. Symbols are as in Figure 3.5, and specific plutons mentioned in the text are labeled: Marker Lake = “ML”, Cabin Creek = “CC”, and Cassiar batholith by the last two digits of the sample number. The range of mantle compositions is indicated by the dark grey shading, using the isotopic ranges: (δ18O = +5.7 ± 0.3‰ from Taylor, 1980; δ34S = 0 ± 3‰ from Chaussidon and Lorand, 1990; 20–50 ppm S; Ionov et al., 1992). a) δ18O versus δ34S. The large grey and black crosses are whole rock data for two felsic orthogneisses of Archean and Neoproterozoic age, respectively (Malton Complex in southeast British Columbia; Vernon– Chamberlain, 1983). The maximum δ18O of a granitic intrusion (except for excessively crustally contaminated phases, such as plutonic margins and small intrusions) is indicated by the dashed line at δ18O = +14.5‰ (Taylor and McLennan, 1986). Samples plotting above the dashed line are from plutonic margins and (or) very small plutons; these heavier isotopic compositions are interpreted in the text to be the result of host rock contamination. b) Calculated minimum sulfur concentrations versus δ34S. Simplified vectors for variations in sulfur concentrations and compositions are indicted by grey arrows and labeled. The maximum concentration of sulfur in granitic magmas at 2–3 kbar is indicated by the dashed line at 500 ppm sulfur (Wallace, 2001, 2005). c) Magnetic susceptibility versus δ34S. 232 233 Figure 3.9. (next page) Stable and radiogenic isotopic comparisons for plutonic suites in the Cassiar Terrane. Symbols for plutonic samples are as in Figure 3.5, and specific plutons discussed in the text are labeled: Marker Lake = ML; Cabin Creek = CC; Buttle Creek = BC; Kuhn Stock = KS. Isotopic ranges for the Seagull suite are indicated by filled red boxes on plots involving 87Sr/86Sr100 (Liverton et al., 2001); the oxygen isotopic range is shaded on plots involving neodymium or lead isotopic compositions (for which there are no available data). The partial to full ranges of mantle isotopic compositions are indicated by the dark grey shading, using the isotopic ranges: δ18O = 5.7 ± 0.3‰ (Taylor, 1980); δ34S = 0 ± 3‰ (Chaussidon and Lorand, 1990); 87Sr/86Sr100 = 0.704–0.705, 208Pb/206Pb = 2.020–2.035, and 207Pb/206Pb = 0.807–0.818 (Quaternary to Recent alkaline ‘basalts’ in the Cassiar Terrane; Abraham et al., 2001). The dotted black boxes (or the area between two dotted lines if only one variable is known) represent the continental magmatic arc rocks west of the Cassiar Terrane (Whitehorse, Prospector Mountain or Carmacks, and Coast Plutonic Complex magmatism; LeCouteur and Tempelman–Kluit, 1976, Dagenais, 1984; Morris and Creaser, 2008). a) 87Sr/86Sr100, εNd100, and 208Pb/206Pb versus δ18O for quartz separates. The Marker Lake and Cabin Creek plutons are circled on plots of δ18O. The possible εNd100 and 208Pb/206Pb compositional range of the Seagull plutonic suite is illustrated by the red boxes (estimated from plots of δ34S versus 207Pb/206Pb and 208Pb/206Pb in (b), based on the assumption that the all A-type samples form a linear array when radiogenic isotopic compositions are compared to δ34S). The isotopic range for the Surprise Lake suite is indicated by the dashed red box labeled “SL”. The grey dashed line at δ18O = 14.5‰ is the approximate upper compositional limit for granitoid rocks globally (Taylor and Sheppard, 1986). The δ18O range of the continental magmatic arc samples are from whole rock data and will be 1–2‰ lower than data from quartz for the same rocks. b) 87Sr/86Sr100, εNd100, and 207Pb/206Pb versus δ34S. Compiled data from the Kuhn Stock (inferred to be a member of the Rancheria plutonic suite on the basis of age and location; Fig. 3.3b; Cooke and Godwin, 1984) are labeled. Red circles highlight trends exhibited by some or all of the samples from the A-type plutonic suites. 234 235 Figure 3.10. (next page) Expanded 87Sr/86Sr100 versus δ18O and δ34S with available basement compositions and ranges for plutonic suites in the Cassiar Terrane. Symbols are as in Figure 3.5, and references for compiled data for arc and enriched mantle are as in Figure 3.9. The blue shaded area is measured 87Sr/86Sr for carbonate rocks in the Windermere Supergroup in the Mackenzie Mountains northeast of the Selwyn Basin (Narbonne et al., 1994). The large grey and black crosses represent two felsic orthogneisses of Archean and Neoproterozoic age, respectively, exposed in southeast British Columbia (whole rock data; the Malton Complex; Vernon–Chamberlain, 1983). a) 87Sr/86Sr100 versus δ18O. Grey arrows highlight trends towards the crustal isotopic reservoirs of Taylor and Sheppard (1986). The field for greywacke rocks is indicated by the shaded box, and there is no set upper 87Sr/86Sr limit to this field (Taylor and Sheppard, 1986; Rollinson, 1993). The compositional range of amphibolite-grade metapelites is indicated by the light grey filled ellipse (Monashee Mountains, southeast British Columbia; Bowman and Ghent, 1986; Kyser et al., 1994). The smaller black crosses are for Neoproterozoic to Archean, mafic to felsic orthogneisses and felsic paragneisses in southeast British Columbia (whole rock data; the Malton Complex; Vernon–Chamberlain, 1983). The range of whole rock oxygen isotopic compositional ranges for siliciclastic and calcareous Neoproterozoic to Triassic basement rocks in the Cassiar Terrane (Dagenais, 1984) and Selwyn Basin (Mair et al., 2006b) are provided. The δ18O range of the continental magmatic arc samples are from whole rock data and will be 1–2‰ lower than data from quartz for the same rocks. b) 87Sr/86Sr100 versus δ34S. The range of sulfur composition for basement rocks is from Mesoproterozoic evaporitic rocks (δ34S = +13.9 to +17.5‰; Gypsum Formation; Little Dal Group; Turner, 2009) and pyrite in Paleozoic siliciclastic rocks (δ34S ≈ –10 to +35‰; Goodfellow, 1987) in the Selwyn Basin, pyrrhotite in Cambrian siliciclastic rocks in the Cassiar Platform (δ34S = +16.7 and +20.9; Cooke and Godwin, 1984), and pyrite in Neoproterozoic siliciclastic rocks in southeast British Columbia (δ34S ≈ –30 to +35‰; Ross et al., 1995) are shown. 236 237 Figure 3.11. (next page) Comparison of lead isotopic comparisons for plutonic suites in the Cassiar Terrane and similar coeval magmatism in the northern and southern Cordillera. Cassiar suite samples are indicated by yellow circles, and A-type suites are indicated by red circles. Symbols and references for compiled data in the Cassiar Terrane are as in Figure 3.6. a) Lead isotopic compositions for the Yukon–Tanana Uplands, east–central Alaska: Paleocene plutons are ~64–59 Ma, Late Cretaceous plutons are ~87–68 Ma, mid-Cretaceous plutons are ~97–91 Ma (94.5 Ma; U–Pb), and Early Cretaceous plutons are 115–101 Ma (109–102 Ma; U–Pb). Mineralogical variations are not differentiated; data and ages are from Aleinikoff et al. (2000) and Dilworth et al. (2007). b) Lead isotopic compositions for muscovite-, biotite-, and hornblende-bearing Early Cretaceous (115– 106 Ma; whole rock Rb–Sr) magmatism, and two metamorphic xenoliths, in the Purcell Mountains, southeast British Columbia (Brandon and Lambert, 1993, 1994). Note the change of scale in (b. 238 239 Figure 3.12. (next page) Comparison of neodymium and strontium isotopic comparisons for plutonic suites in the Cassiar Terrane and similar coeval magmatism in the northern and southern Cordillera. Cassiar suite samples are indicated by yellow circles, and A-type suites are indicated by red circles. Symbols and references for compiled data in the Cassiar Terrane are as in Figure 3.6. The “upper” and lower” isotopic perturbations exhibited by the Cassiar suite samples in Figure 3.6a are labeled. a) Comparison to magmatism Yukon–Tanana Uplands, east–central Alaska: Paleocene plutons are ~64– 59 Ma, Late Cretaceous plutons are ~87–68 Ma, mid-Cretaceous plutons are ~97–91 Ma (94.5 Ma; U– Pb), and Early Cretaceous plutons are 115–101 Ma (109–102 Ma; U–Pb). Mineralogical variations are not differentiated; data and ages are from Newberry et al. (1995), Aleinikoff et al. (2000), and Dilworth et al. (2007). b) Comparison to magmatism in the Purcell Mountains, southeastern British Columbia: muscovite-, biotite-, and hornblende-bearing Early Cretaceous (115–106 Ma; whole rock Rb–Sr) magmatism, and two metamorphic xenoliths (Brandon and Lambert, 1993, 1994). The Fry Creek samples from the Purcell Mountains are circled. 240 241 Figure 3.13. (next 2 pages) Whole rock immobile element ratios (Sm/Nd, La/Sm, Ba/Nb, and Sm/Yb) and isotopic comparisons for plutonic suites in the Cassiar Terrane. Symbols and references for compiled data are as in Figure 3.6. Ratios for average continental crustal reservoirs are indicated by the vertical lines and labels: bulk crust (B) = solid grey; upper crust (U) = dashed light grey; middle crust (M) = dashed grey; lower crust (L) = dashed black (all values are from Rudnick and Gao, 2003); where any or all of the crustal compositions overlap with the bulk composition, only the solid grey line (B) is shown. The composition of the North American Shale Composite (Haskin et al., 1966) is indicated by the red dotted line. a) Immobile element ratios versus ɛNd100 and 87Sr/86Sr100. The isotopic range of the Seagull plutonic suite is indicated by the red boxes (εNd100 ranges are estimated from Figure 3.9b based on the assumption that the all A-type samples form a linear array when lead isotopic compositions are compared to δ34S). The approximate locations of the “upper” and lower” isotopic perturbations exhibited by the Cassiar suite samples are indicated by the grey arrows along the y-axis. The isotopically juvenile to moderately evolved samples with high La/Sm and Ba/Nb, and low Sm/Nd are circled; the isotopically moderately to highly evolved samples with moderate 147Sm/144Nd, Sm/Nd, La/Sm, and Ba/Nb are outlined by the black box; the isotopically very highly evolved samples with anomalously high 147Sm/144Nd and Sm/Nd, and low La/Sm and Ba/Nb are outlined by the dashed blue lines. b) Immobile element ratios versus δ18O and δ34S. All of the stable isotope data apply only to the Cassiar suite samples plotting between the moderately and highly evolved isotopic end-members described in the text, with the exception of a single oxygen value for a isotopically juvenile sample that has been labeled. The isotopically very highly evolved samples with anomalously high 147Sm/144Nd are outlined by the dashed blue lines. 242 243 244 Figure 3.14. Whole rock geochemical parameters and isotopic comparisons for plutonic suites in the Cassiar Terrane. Symbols and references for compiled data are as in Figure 3.6. Cassiar suite samples with anomalously elevated 147Sm/144Nd are outlined with the dashed blue lines. Parameters for average continental crustal reservoirs are indicated by the vertical lines and labels: bulk (B) = solid grey; upper (U) = dashed light grey; middle (M) = dashed grey; lower (L) = dashed black (all values are from Rudnick and Gao, 2003); where any or all of the crustal compositions overlap with the bulk composition, only the solid grey line (B) is shown. The North American Composite Shale average (Haskin et al., 1966) is indicated by the red dotted line. a) Rb/Sr and A/CNK versus 87Sr/86Sr100. The Seagull suite range is indicated by the filled red boxes. Small black arrows indicate general trends imposed by primarily by source rock composition for the Cassiar plutonic suite. Inferred melt sources for the Cassiar plutonic suite, based on the discussion in the text, include: qtz–fsp = quartzo–feldspathic (plagioclase > mica); pelite = metapelite (mica > plagioclase); min. melt = leucogranites formed by minimum melting of the pelitic protolith. b) Rb/Sr and A/CNK versus δ34S. 245 Figure 3.15. (next page) Summary of εNd and 87Sr/86Sr for plutonic suites, proposed magma sources in the Cassiar Terrane, and northwest Laurentian basement rocks. All isotopic compositions are calculated at 100 Ma, except isotopic compositions for the Rancheria and Surprise Lake plutonic suites (80 Ma) and the Black River suite (50 Ma). Back-arc magmatism is indicated by yellow circles (Cassiar plutonic suite), post-arc magmatism is indicated by red circles (Seagull, Rancheria, and Black River plutonic suites), continental arc magmatism is indicated by unfilled circles (Teslin, Whitehorse, and Prospector Mountain plutonic suites, and Coast Plutonic Complex), and enriched mantle magmas are indicated by filled black circles (Quaternary to Recent ‘basalts’ in the Cassiar Terrane). References for compiled plutonic and basaltic data are as in Fig. 3.6, and the data range for the Surprise Lake plutonic suite is from Morris and Creaser (2008). a) Summary of plutonic suite groupings and proposed lithological magma sources. The isotopically juvenile to moderately evolved samples with high La/Sm and Ba/Nb, and low Sm/Nd are circled; the isotopically moderately to highly evolved samples with moderate 147Sm/144Nd, Sm/Nd, La/Sm, and Ba/Nb are outlined by the black box; the isotopically very highly evolved samples with anomalously high 147Sm/144Nd and Sm/Nd, and low La/Sm and Ba/Nb are outlined by the dashed blue lines. Inferred melt sources for the Cassiar plutonic suite, based on the discussion in the text, include: EM = enriched mantle; IJ–qp = isotopically juvenile – quartzo–feldspathic; HE–qp = highly evolved – quartzo–feldspathic; HE–p = highly evolved – pelitic; min. melt = minimum temperature partial melts. b) Comparison of plutonic suite groupings and proposed lithological magma sources with northwest Laurentian basement rocks. All samples for which both 87Sr/86Sr100 and εNd100 data exist are plotted. Data from the entire Canadian Cordillera and westernmost Laurentian basement are considered for Archean to Paleoproterozoic basement, but only data from the northern Cordillera (i.e., northeast British Columbia to east–central Yukon) are considered for Mesoproterozoic to Neoproterozoic basement. Illustrated ranges and averages are based on the combined data from all datasets (see the footnotes to Fig. 3.16).1 1 The “Neoproterozoic (felsic)” strontium isotopic range and average are taken from 40 samples of felsic gneiss in the Malton Complex at the base of the Windermere Supergroup in southeast British Columbia (Vernon– Chamberlain, 1983). Strontium isotopic data for Neoproterozoic sedimentary rocks in the northern Cordillera is limited to two samples, so the 87Sr/86Sr100 range and average for the Windermere Supergroup must be regarded with caution. 246 247 Figure 3.16. (next 2 pages) Comparison of TDM versus εNd100 for northwest Laurentian basement and syn- to post-accretionary magmatism in the northern Cordillera. Symbols are as in Figure 3.15, and references for compiled plutonic and basaltic data are as in Fig. 3.6. All isotopic compositions are calculated at 100 Ma, except isotopic compositions for the Rancheria and Surprise Lake plutonic suites (80 Ma) and the Black River suite (50 Ma)—although there is no substantial change in εNd when calculated at 80 Ma and 50 Ma for the A-type plutonic suites. The inferred εNd100 range for the Seagull plutonic suite is shaded red. The data range for the Surprise Lake plutonic suite is from Morris and Creaser (2008). Isotopic ranges that largely overlap with the compositional range of granitoids in the Cassiar Terrane are illustrated. a) Archean to Neoproterozoic basement.2 b) Paleozoic to Recent basement.3 c) Accreted pericratonic and oceanic terranes.4 2 "Archean paragneiss" from southeast British Columbia to southwest Alberta basement (Frost and O'Nions, 1984); "mafic" and "felsic Archean orthogneisses" from the Slave Craton and the Malton Complex in southeast British Columbia (Vernon–Chamberlain, 1983; Burwash et al, 1985); "Paleoproterozoic paragneiss" from the Malton and Monashee complexes in southeast British Columbia (Vernon–Chamberlain, 1983; Parkinson, 1991); "Paleoproterozoic Wernecke Supergroup" from eastern Yukon (Thorkelson et al, 2005); "mafic Paleoproterozoic orthogneiss" from the Malton Complex in southeast British Columbia (Vernon–Chamberlain, 1983), the Buffalo Head and Nova domains in northeast Alberta (Armstrong et al., 1991), and the Frenchman's Cap core complex in southeast British Columbia (Thériault and Ross, 1991); mafic "Paleoproterozoic Bonnet Plume" intrusions from eastern Yukon (Thorkelson et al., 2001); "felsic Paleoproterozoic orthogneiss" from the Taltson, Buffalo Head, Chinchaga, and Ksituan domains in northeast Alberta (Thériault and Ross, 1991), and the Malton, Monashee, and Frenchman's Cap complexes in southeast British Columbia (McDonough and Parrish, 1991; Parkinson, 1991; Armstrong et al., 1991); "Paleoproterozoic Fort Simpson Terrane" in southeast NT (Villeneuve et al., 1991); "Paleoproterozoic Wernecke Breccia" from eastern Yukon (Thorkelson et al., 2005); "1.27 Ga Coppermine flood basalt" in western Nunuvut (Dupuy et al., 1992); "Mesoproterozoic Muskwa Assemblage" in northeast British Columbia (Ross et al., 2001) and "Mackenzie Mountains Supergroup" in southwest NT (Rainbird et al, 1997); "Neoproterozoic to Early Cambrian Windermere Supergroup" in the northern Cordillera (Abercrombie, 1990; R. Creaser, unpub. 1996 Lithoprobe Snorkle Report; Garzione et al., 1997; Erdmer et al, 2005); "mafic Neoproterozoic sill/volcanic" rocks from the Mackenzie Mountains Supergroup in southeast NT (Dudás and Lustwerk, 1997). 3 "Early Paleozoic miogeocline" from the Selwyn Basin, Mackenzie Platform, and Kechika Trough in eastern Yukon, southwest NT, and northeast British Columbia (R. Creaser, unpub. 1996 Lithoprobe Snorkle Report; Garzione et al., 1997; Erdmer et al, 2005; Beranek, 2009; Hadlari, T., unpub.); "Early Paleozoic mafic" volcanic rocks from a mafic horizon in the MacPass area at the Yukon-NT border (R. Creaser, unpub. 1996 Lithoprobe Snorkle Report) and the Ordovico–Silurian Marmot Formation and Coates Lake Diatreme in southwest NT (Leslie, 2009); "mid-Paleozoic (Devono–Mississippian) clastic influx" from the Selwyn Basin, Mackenzie Platform, and Kechika Trough in eastern Yukon, southwest NT, and northeast British Columbia (R. Creaser, unpub. 1996 Lithoprobe Snorkle Report; Garzione et al., 1997; Hadlari, T., unpub. data); "Late Paleozoic (mid-Mississippian to Triassic) miogeocline" from the Selwyn Basin, Mackenzie Platform, and Kechika Trough in eastern Yukon, southwest NT, and northeast British Columbia (R. Creaser, unpub. 1996 Lithoprobe Snorkle Report; Garzione et al., 1997; Beranek, 2009); "metasedimentary" and "amphibolitic" xenoliths of unknown age from the Cassiar batholith (Driver et al., 2000); "enriched mantle" from Quaternary to Recent ‘basalts’ erupted in the CT near the Yukon– British Columbia border (Abraham et al., 2001). 4 "Slide Mountain (mafic)" and "Slide Mountain (sediments)" from the YTT inlier (Piercey et al., 2004) and from south-central Yukon (Selby et al., 1999; Piercey and Colpron 2009); "Cache Creek (mafic)", "Cache Creek (felsic)", and "Cache Creek (sediments)" from south–central British Columbia (Patchett and Gehrels, 1998); "YTT Snowcap (mafic)" from the YTT inlier (Piercey et al., 2004), the Teslin Tectonic Zone (Creaser et al., 1997), and YTT west of the Teslin Tectonic Zone (Selby et al., 1999; Piercey and Colpron, 2009); "YTT Snowcap (felsic)" from the Teslin Tectonic Zone (Stevens et al., 1996) and YTT west of the Teslin Tectonic Zone (Selby et al., 1999); "YTT Snowcap (carbonate)" from the YTT west of the Teslin Tectonic Zone (Selby et al., 1999; Piercey and Colpron, 2009), "YTT Snowcap Assemblage (sediments)" from the Teslin Tectonic Zone (Creaser et al., 1997) and YTT west of the Teslin Tectonic Zone (Patchett and Gehrels, 1998; Selby et al., 1999; Piercey and Colpron, 2009). 248 249 250 Table 3.1. Basic sample descriptions, crystallization ages, trace element ratios, and geochemical parameters for all samples that underwent radiogenic and (or) stable isotopic analysis. Plutonic Suite Sample Pluton Mag. Susc. Lithology U-Pb age (Ma)a Sm/Ndb La/Sm Ba/Nb Sm/Yb Rb/Sr A/CNK SiO2 Eu/Eu* Cassiar-ms KR-07-01 Dycer Creek 0.13 f-m-gr mc-phyric K=P=Q bt>ms monzogranite 114.8 0.17 8.61 91.60 5.43 0.64 1.17 70.28 0.48 Cassiar-ms KR-07-02 Quiet Lake 0.13 m-gr equigr (mc-phyric) Q>K=P bt>ms monzogranite 114.7 0.18 8.09 60.28 5.36 1.17 1.19 72.00 0.42 Cassiar-ms KR-07-16 Cassiar 1.06 m-(c)-gr mc-phyric Q>K>P bt>>ms (leuco)granite 111.5 0.19 8.12 60.05 2.17 0.51 1.14 73.11 0.53 Cassiar-ms KR-07-17 Cassiar 1.90 f-(m)-gr equigr (mc-phyric) Q=K>P bt>ms monzogranite 111.1 0.19 7.47 48.57 5.39 0.96 1.14 73.00 0.40 Cassiar-ms KR-07-18 Cassiar 0.13 f-m-gr mc-phyric K=P=Q bt>ms monzogranite 114.3 0.18 7.15 70.91 10.57 1.51 1.06 71.87 0.26 Cassiar-ms KR-07-22 Cassiar 0.09 v f-gr mylonitic mc-augen K=Q>P bt-ms monzogranite 113.8 0.21 6.99 74.76 3.30 0.50 1.13 73.02 0.52 Cassiar-bt 07M-152 Buttle Creek 0.19 foliated m-gr equigr Q>P>K bt granodiorite 104.0 0.20 6.72 29.66 3.91 0.46 1.33 68.66 0.72 Cassiar-bt 07M-153 Glenlyon 0.19 f-m-gr equigr P=Q>K bt granodiorite 108.5 0.19 6.96 77.31 3.43 0.39 1.21 68.28 0.71 Cassiar-bt KR-07-07 Young 0.16 f-gr equigr (mc-phyric) K=Q=P bt monzogranite 107.2 0.20 6.48 50.67 4.31 0.67 1.20 73.89 0.51 Cassiar-bt KR-07-08 St. Cyr 12.80 f-(m)-g equigr P=Q>K bt granodiorite 107.7 0.18 6.95 30.57 2.63 1.02 1.10 72.31 0.53 Cassiar-bt KR-07-09 Fox Mountain 0.21 f-gr qtz-fsp-bt (leuco)granitic porphyry 107.8 0.17 8.37 39.11 2.63 0.61 1.14 69.30 0.66 Cassiar-bt KR-07-11 Glenlyon 0.08 f-gr equigr P=K=Q bt leucogranite 107.7 0.18 7.50 58.49 2.30 0.40 1.21 63.61 0.63 Cassiar-bt KR-07-19 Marker Lake 23.30 c-gr mc-crystic K=Q /=P bt monzogranite 106.8 0.16 10.35 196.39 6.24 0.14 1.18 72.17 1.00 Cassiar-bt KR-07-05 Nisutlin 0.08 f-m-gr foliated equigr Q>/=K>P bt leucogranite 113.2 0.17 9.27 94.91 5.52 0.23 1.16 70.26 0.75 Cassiar-hbl KR-07-06 Nisutlin 16.80 c-gr mc-phyric Q>K>P bt>hbl monzogranite 105.1 0.15 10.54 21.12 3.49 0.36 1.08 69.59 0.56 Seagull 07M-154 Little Salmon 0.13 (m)-c-gr equigr P=Q>K bt monzogranite 97.48 0.22 5.90 7.35 2.35 3.43 1.00 73.91 0.18 Seagull KR-07-03 Quiet Lake 0.05 m-gr equigr K>/=Q>P bt>ms (leuco)granite 97.7 0.26 5.20 8.67 0.99 4.64 1.07 76.00 0.24 Seagull? KR-07-04 Nisutlin 0.03 f-gr equigr ms>bt Q>/=P>K leucogranite / 0.22 5.45 0.11 0.94 285.00 1.08 76.64 / Rancheria KR-07-15 Meister Lake 0.09 m-gr equigr K=Q>P ms>bt (leuco)granite 79.2 0.17 10.54 15.32 4.28 0.43 1.07 70.42 0.58 Rancheria KR-07-21 Allen Stock 18.50 c-gr mc-crystic K>Q>P bt monzogranite 77.2 0.19 8.80 8.08 2.55 3.19 1.19 74.10 0.40 Rancheria KR-07-23 Gravel Creek 0.09 c-gr equigr (sub-mc-phyric) Q=K=P bt>ms monzogranite 81.9 0.18 8.34 7.40 6.19 3.39 1.26 73.13 0.25 Black River KR-07-12 Tintina 0.12 f-m-gr mc-crystic qtz-phyric Q>K>P bt (sub-phyric) (leuco)granite sub-porphyry 49.1 0.18 7.49 10.75 2.09 3.39 1.06 73.10 0.31 Black River KR-07-14 Black River 0.13 m-gr mc-phyric bt-(ms) Q=K>>P monzogranite 47.1 0.19 6.90 7.68 2.14 3.90 1.05 73.79 0.22 Notes: a U-Pb ages are reported in Table 2.2 (Chapter 2). b Trace element ratios and geochemical parameters are reported in Appendix D.1. 251 Table 3.2. Radiogenic and stable isotopic data generated by this study for magmatism in the Cassiar Terrane. Neodymium, strontium and sulfur data are from whole rock powders, lead data are from feldspar separates, and oxygen data are from quartz separates. a) Neodymium isotopic compositions and depleted mantle model ages. b) Strontium isotopic compositions. c) Lead isotopic compositions. d) Oxygen and sulfur isotopic compositions, and calculated sulfur concentrations. a) Plutonic Suite Sample 143Nd/ 144Nd ± 2-σ Sm (ppm) Nd (ppm) 147Sm/ 144Nd 143Nd/ 144Ndia εNdib εNd100 εNd80 εNd50 TDM (Ga)c Cassiar-ms KR-07-02 0.51201 0.000007 3.8 21.2 0.1093 0.51193 -10.92 -11.22 -11.72 -12.47 1.66 Cassiar-ms KR-07-18 0.51185 0.000005 9.7 56.5 0.1034 0.51177 -13.97 -14.28 -14.78 -15.53 1.79 Cassiar-bt 07M-152 0.51202 0.000007 5.1 28.2 0.1099 0.51194 -10.96 -11.14 -11.64 -12.39 1.66 Cassiar-bt KR-07-09 0.51190 0.000007 5.2 30.7 0.1023 0.51183 -13.14 -13.36 -13.87 -14.62 1.71 Cassiar-bt KR-07-19 0.51219 0.000008 8.2 49.6 0.0994 0.51212 -7.35 -7.58 -8.08 -8.83 1.28 Cassiar-bt KR-07-20 0.51208 0.000008 3.4 22.7 0.0898 0.51202 -9.32 -9.56 -10.06 -10.82 1.31 Cassiar-bt KR-07-10 0.51198 0.000008 4.7 27.1 0.1058 0.51191 -11.52 -11.73 -12.24 -12.99 1.64 Cassiar-bt KR-07-05 0.51177 0.000007 3.6 17.4 0.1238 0.51168 -15.94 -16.22 -16.72 -17.47 2.33 Rancheria KR-07-15 0.51183 0.000009 3.0 16.2 0.1105 0.51196 -11.33 -10.82 -11.32 -12.07 1.68 Rancheria KR-07-21 0.51183 0.000007 8.0 50.3 0.0962 0.51212 -8.24 -7.70 -8.20 -8.96 1.28 DUPLICATE (Sm–Nd, Rb–Sr) 0.51183 0.000009 7.7 48.8 0.0955 0.51212 -8.11 -7.57 -8.07 -8.82 1.26 Black River KR-07-14 0.51195 0.000007 8.4 44.7 0.1138 0.51191 -12.96 -11.67 -12.18 -12.93 1.83 Notes: a 143Nd/144Ndi (initial) was calculated using a 147Sm decay constant of 6.54x10-12/yr. b εNd values were calculated using 147Sm/144Ndchondrite = 0.512638 and 147Sm/144Ndchondrite = 0.1967 (Patchett et al., 2004). c TDMs (depleted mantle model ages) were calculated using 147Sm/144NdDM = 0.51315 and 147Sm/144NdDM = 0.2137 (Michard et al., 1985). b) Plutonic Suite Sample 87Sr/ 86Sr ± 2-σ Rb (ppm) Sr (ppm) 87Rb/ 86Sr 87Sr/ 86Sria 87Sr/ 86Sr100 87Sr/ 86Sr80 87Sr/ 86Sr50 Cassiar-ms KR-07-02 0.7171 0.00001 242 205 3.422 0.7115 0.7122 0.7132 0.7147 Cassiar-ms KR-07-18 0.7236 0.00001 269 185 4.212 0.7167 0.7176 0.7188 0.7206 Cassiar-bt 07M-152 0.7160 0.00001 134 287 1.353 0.7140 0.7141 0.7145 0.7150 Cassiar-bt KR-07-09 0.7182 0.00001 152 248 1.773 0.7155 0.7157 0.7162 0.7169 Cassiar-bt KR-07-19 0.7104 0.00001 115 459 0.724 0.7093 0.7094 0.7096 0.7099 Cassiar-bt KR-07-20 0.7101 0.00001 109 811 0.390 0.7094 0.7095 0.7096 0.7098 Cassiar-bt KR-07-10 0.7175 0.00001 235 231 2.945 0.7131 0.7134 0.7142 0.7154 Cassiar-bt KR-07-05 0.7239 0.00001 147 246 1.730 0.7211 0.7215 0.7220 0.7227 Rancheria KR-07-15 0.7287 0.00001 330 55 17.561 0.7090 0.7038 0.7088 0.7163 Rancheria KR-07-21 0.7104 0.00001 213 426 1.448 0.7088 0.7083 0.7087 0.7094 DUPLICATE (Sm–Nd, Rb–Sr) 0.7104 0.00001 205 408 1.455 0.7088 0.7083 0.7087 0.7093 Black River KR-07-14 0.7265 0.00001 301 73 11.929 0.7185 0.7095 0.7129 0.7180 Notes: a 87Sr/86Sri (initial) was calculated using a 87Rb decay constant of 1.42x10-11/yr. c) Plutonic Suite Sample 206Pb/ 204Pb ± 2-σ 207Pb/ 204Pb ± 2-σ 208Pb/ 204Pb ± 2-σ 207Pb/ 206Pb 208Pb/ 206Pb Cassiar-ms KR-07-02 19.30 0.001 15.71 0.000 39.22 0.003 0.81 2.03 Cassiar-ms KR-07-18 19.24 0.009 15.74 0.005 39.17 0.021 0.82 2.04 DUPLICATE (Pb–Pb) 19.31 0.001 15.70 0.001 39.36 0.004 0.81 2.04 Cassiar-bt 07M-152 20.17 0.003 15.76 0.002 39.81 0.006 0.78 1.97 Cassiar-bt KR-07-09 19.64 0.009 15.82 0.004 39.82 0.019 0.81 2.03 Cassiar-bt KR-07-19 19.44 0.007 15.76 0.005 39.51 0.015 0.81 2.03 Cassiar-bt KR-07-20 19.32 0.002 15.69 0.001 39.26 0.006 0.81 2.03 Cassiar-bt KR-07-10 19.44 0.004 15.76 0.002 39.55 0.008 0.81 2.03 Cassiar-bt KR-07-05 19.46 0.001 15.73 0.001 39.54 0.003 0.81 2.03 Rancheria KR-07-15 19.60 0.001 15.72 0.000 39.77 0.001 0.80 2.03 Rancheria KR-07-21 19.72 0.001 15.74 0.000 39.86 0.002 0.80 2.02 Rancheria KR-07-23 19.75 0.002 15.74 0.002 39.93 0.005 0.80 2.02 Black River KR-07-14 19.48 0.002 15.70 0.001 39.69 0.004 0.81 2.04 252 d) Plutonic Suite Sample δ 18OVSMOW (‰)a d34SVCDT (‰)b S (ppm)c Cassiar-ms KR-07-01 13.1 7.31 16 Cassiar-ms KR-07-02 14.4 9.04 10 Cassiar-ms KR-07-16 12.2 4.45 5 Cassiar-ms KR-07-17 11.5 8.00 12 Cassiar-ms KR-07-18 12.9 8.35 9 Cassiar-ms KR-07-22 12.6 8.58 35 Cassiar-bt 07M-152 14.3 9.31 66 Cassiar-bt 07M-153 11.8 9.08 36 Cassiar-bt KR-07-07 15.4 8.68 62 Cassiar-bt KR-07-08 12.2 8.02 61 DUPLICATE (S) / 8.02 63 Cassiar-bt KR-07-09 13.5 9.63 13 Cassiar-bt KR-07-11 14.7 7.91 3 Cassiar-bt KR-07-19 13.1 2.06 nd DUPLICATE (S) / 2.03 110 Cassiar-bt KR-07-20 12.5 5.14 8 Cassiar-bt KR-07-24 14.2 3.86 12 Cassiar-bt KR-07-10 11.8 7.92 20 DUPLICATE (O) 12.0 / / Cassiar-bt KR-07-05 15.7 10.12 10 Cassiar-hbl KR-07-06 13.2 8.08 19 Seagull 07M-154 12.5 3.46 8 DUPLICATE (O) 14.0 / / Seagull KR-07-03 13.5 4.89 3 Seagull? KR-07-04 15.4 3.66 3 DUPLICATE (O) 15.3 / / Rancheria KR-07-15 11.6 / 2 Rancheria KR-07-21 12.0 6.00 11 Rancheria KR-07-23 16.5 6.31 4 Black River KR-07-12 14.0 9.00 7 Black River KR-07-14 12.5 10.23 11 Notes: a Reported relative to the Vienna Standard Mean Ocean Water (VSMOW; Coplen, 1994). b Reported relative to the Vienna Cañon Diablo troiolite (VCDT; Ding et al., 2001); underlined values are corrected for low peak heights (<100 nA) as described in the text. c grams of contained sulfur were calculated from the weight of precipitated Ag2S used for sulfur isotopic analysis, converted to a percentage of the total sample weight, and re-calculated in ppm, and reported as minimum estimates of sulfur content. 253 4 CHAPTER: ISOTOPIC CHARACTERIZATION OF SYN- TO POST- ACCRETIONARY, EARLY TO ‘MID’-CRETACEOUS MAGMATISM IN THE SELWYN BASIN The origin of syn- to post-accretionary, Early to ‘mid’-Cretaceous intrusive rocks in the northern Cordilleran miogeocline has already been discussed from a regional perspective by previous workers (e.g., Gordey and Anderson, 1993; Mortensen et al., 2000; Hart et al., 2002, 2004a). Prior to this study, however, characterizations of these igneous rocks rarely included a combination of robust geochronology, extensive geochemical data (including trace element suites), and radiogenic (other than Sr) and stable isotopic information. Chapter 2 provided an updated temporal and geochemical characterization of Early to ‘mid’-Cretaceous plutonic suites in the northern Cordilleran miogeocline, based mainly on the two regions—the Cassiar Terrane and the Selwyn Basin—examined by this study. The geochemical data provided some insight on the composition of possible magma sources to each of the suites discussed in Chapter 2, however, Chapter 3 demonstrated the utility of using radiogenic and stable isotopic compositions as a tool for better characterizing magma sources. This led to a detailed description of the petrogenesis of intermediate–felsic intrusions in the Cassiar Terrane study area. In this chapter, a similarly comprehensive isotopic dataset―including whole rock Sm–Nd, Rb–Sr, and sulfur, as well as lead from feldspar separates, and oxygen from quartz separates―is used to test the geochemically based petrogenetic interpretations from Chapter 2 and to better characterize Early to ‘mid’-Cretaceous magmatism in the Selwyn Basin. In the interest of brevity, and because detailed descriptions of northern Cordillera geology are already presented in Chapters 1 to 3, this chapter provides only: 1) basic background information regarding Early to ‘mid’-Cretaceous igneous rocks in, and southwest of, the Selwyn Basin (Section 4.1), 2) a summary of the isotopic data generated by this study (Section 4.2), and 3) and the most relevant petrogenetic interpretations (integrated with the results of previous isotopic studies; Section 4.3). 4.1 BACKGROUND ON SYN- TO POST-ACCRETIONARY MAGMATISM 4.1.1 Magmatism ‘Outboard’ (southwest) of the Selwyn Basin Early to ‘mid’-Cretaceous magmatism located in an outboard position, or southwest of the Selwyn Basin, is interpreted to be related to northeast-dipping subduction along the western margin of the North America (Fig. 4.1; e.g., Mortensen et al., 2000; Hart et al., 2004a). These magmas were emplaced into continentally-derived, miogeoclinal rocks in east–central Alaska, and into pericratonic terranes (Yukon–Tanana terrane and Stikinia) in western Yukon. The 112–99 Ma Dawson Range plutonic suite (comprising mainly oxidized, hornblende–biotite bearing quartz diorite, diorite, and granodiorite phases) and Casino plutonic suite (comprising leucocratic granite, quartz monzonite, and alaskite) are thought to represent plutonic expressions of a continental magmatic arc located 254 southwest of the Selwyn Basin (Fig. 4.1; Souther, 1991; Woodsworth et al., 1991; Hart, 1995, 1997; Gordey et al., 1998; Mihalnyuk, 1999; Selby et al., 1999; Mortensen et al., 2000; Hart et al., 2004a; Morris and Creaser, 2008; Mortensen and Hart, 2010; Bineli Betsi and Lentz, 2011, 2012). An isotopic study of the Dawson Range and Casino plutonic suites demonstrated that up to 60–80% of the Dawson Range magmas were generated from partial melting of isotopically evolved metasedimentary basement mixed with mantle-derived basalt (Selby et al., 1999). The Casino plutonic suite is considered to comprise late-stage fractionates of this mixed, crust–mantle-derived magma (Selby et al., 1999). This led to an interpretation that the Dawson Range batholith formed by partial melting of over-thickened Laurentian crust (Selby et al., 1999), and probably in association with movement along strike-slip faults (Payne et al. 1987). Recent isotopic work by Bineli Betsi and Lentz (2012) is also compatible with the interpretation that the Dawson Range batholith represents a continental arc that incorporated a large proportion of isotopically evolved basement. Isotopic data for the Dawson Range batholith and Casino plutonic suite (Selby et al., 1999; Bineli Betsi and Lentz, 2012) are presented with the results of this study for comparative purposes. 4.1.2 ‘Inboard’ Magmatism within the Selwyn Basin ‘Inboard’ igneous rocks, located northeast of the interpreted continental magmatic arc in the Selwyn Basin, are of Early to ‘mid’-Cretaceous age, and are thought to have been generated from over-thickened crust in a back-arc setting (e.g., Woodsworth et al., 1991; Mortensen et al., 2000; Hart et al., 2004a). This magmatic event in the Selwyn Basin had been previously subdivided into several plutonic suites (e.g., Mortensen et al., 2000; Hart et al., 2004, 2005; Heffernan, 2004). Using extensive geochronological and geochemical data in Chapter 2, this study has modified the temporal, spatial, and geochemical constraints on these plutonic suites, particularly in the eastern Selwyn Basin. The main features characterizing each of these plutonic suites, and their interpreted source composition(s), are summarized below from Chapter 2. The oldest syn- to post-accretionary igneous rocks in the Selwyn Basin are members of the Anvil plutonic suite. This suite includes several small 114–111 Ma plutons emplaced in the Yukon–Tanana terrane inlier (Figs. 4.1 and 4.2), and the 109–104 Ma Anvil batholith and scattered plutons emplaced in the western Selwyn Basin (Fig. 4.1). The suite comprises strongly reduced, highly siliceous, slightly to moderately peraluminous, biotite ± muscovite- bearing leucogranites of high-K calc-alkaline composition. An aluminous mineralogy (Pigage and Anderson, 1985; Smith and Erdmer, 1990), and the composition of the limited samples available for this plutonic suite, are characteristic of a derivation by low-temperature partial melting of pelitic rocks (e.g., Kemp and Hawkesworth, 2003). This magma source either lacked abundant plagioclase, or plagioclase was not a reactant phase during anatexis. The latter requires that partial melting occurred in the absence of fluid, and at relatively low temperatures. Garnet and staurolite with local fibrolite and cordierite in the contact aureole for the Anvil batholith are suggestive of relatively deep levels of emplacement in the western Selwyn Basin (~10 km depth; Pigage and Anderson, 1985; Smith and Erdmer, 1990). Also, for the Anvil batholith, a shallow-dipping mineralogical foliation and the laccolithic (sill-like) geometry (Pigage and Anderson, 1985), are compatible with emplacement during exhumation and (or) in 255 an extensional setting (Murphy et al., 1995; Gabrielse et al., 2006). Coeval exhumation in the Yukon–Tanana terrane inlier is indicated by ~113–106 Ma 40Ar/39Ar mica ages from exposed amphibolite-grade rocks (Murphy et al., 2001) and Early Cretaceous, low-angle normal faults (Murphy, 2004). Nearby volcanic rocks of the South Fork calderas (preserved nearby in down-dropped blocks) erupted ~10 m.y. after the emplacement of the Anvil batholith (Fig. 4.1) also indicate that exhumation of the Anvil batholith in the western Selwyn Basin was completed by this time. The 106–100 Ma Hyland plutonic suite comprises large (Figs. 4.1 and 4.2), medium-K to high-K calc-alkaline, biotite granodiorite to monzogranite intrusions, and subordinate dioritic plagioclase–biotite porphyry phases and dykes. Like the Anvil batholith in the western Selwyn Basin, the Hyland plutonic suite is of supracrustal affinity and was emplaced as thick, concordant, laccolithic bodies that are strongly foliated near their margins (Murphy et al., 1995; Hart and Lewis, 2006). Schistose to gneissic contact aureoles fringing the largest Hyland suite batholiths are also intensely “migmatitic” (Roots et al., 1966). Thus, as with the Anvil suite to the west, anatexis during regional exhumation in a zone of crustal extension has also been invoked for the Hyland suite magmas (e.g., Murphy et al., 1995; Gabrielse et al., 2006). The Hyland plutonic suite , however, is much more chemically heterogeneous, and appears to lack aluminous minerals (except for muscovite in a few phases; e.g., Kuehnbaum, 1977). Associated, but subordinate, hornblende-bearing intrusive phases and enclaves of intermediate composition (e.g., Brock, 1978; Durgin, 1979) require some input of melt from mafic–intermediate (i.e., hornblende-bearing) rocks. Thus, as with the Cassiar plutonic suite (Chapter 3), a heterogeneous package of mainly quartzo–feldspathic sedimentary rocks and containing lesser interlayered amphibolite (and (or) carbonate) rocks, is an appropriate protolith for the Hyland plutonic suite. Decompression-related anatexis of Neoproterozoic host rocks to the Hyland suite intrusions (a package of relatively oxidized supracrustal quartzo–feldspathic metasedimentary ± metavolcanic rocks) is one possible source of magma for these intrusions (e.g., Hart et al., 2004a). A near-in situ derivation for the Hyland suite batholiths is also compatible with, but not necessitated by, their migmatitic contact aureoles and concordant laccolithic geometries. The 99–96 Ma Tay River plutonic suite was emplaced across the Selwyn Basin (Figs. 4.1 and 4.2). Although small compositional differences across Selwyn Basin are present46, they are relatively minor. This suite—and associated ~97 Ma South Fork volcanic rocks in the western Selwyn Basin (Figs. 4.1 and 4.2)—is of high-K calc-alkaline composition and probable I-type affinity. Based on the geochemical data and the common presence of hornblende, a dominantly infracrustal protolith of intermediate composition (corresponding to 60–65 wt. % SiO2) is thought to be the most reasonable melt source for the Tay River and South Fork magmas. This plutonic suite was derived by high- temperature fluid-absent partial melting of a slightly LILE (Rb, Th, and U) and HFSE (Nb and Ta) enriched, probable metaigneous, lower to middle crustal basement in the Selwyn Basin. Associated mafic dykes in the eastern Selwyn Basin may be evidence for the influx of mantle-derived magma into the crust, which contributed to the heat 46 The Tay River suite in the eastern Selwyn Basin, for example, is almost always moderately to highly oxidized (vs. reduced in the western Selwyn Basin), and lacks associated volcanic rocks. Abundant highly siliceous K-feldspar- megacrystic phases, less siliceous quartz diorite plugs, and quartz monzodiorite dykes are also associated with the Tay River suite in the eastern Selwyn Basin. Shallower levels of emplacement, or lesser amounts of exhumation, appear to characterize the intrusions in the western Selwyn Basin. The Tay River and South Fork suites in the western Selwyn Basin are also comparatively depleted in tantalum and niobium, and have slightly lower SiO2 contents overall. 256 required to produce intermediate members (i.e., hornblende ± biotite-bearing) of the Tay River suite. In the eastern Selwyn Basin, highly siliceous plutons (‘evolved’; Fig. 4.2) may comprise more fractionated and (or) crustally contaminated magmas, or smaller melt fractions from the same infracrustal source (or more siliceous portions thereof). Several of the less siliceous intrusions are chemically distinct and form small isolated plugs in the southeast Selwyn Basin (Fig. 4.2). These intrusions, or the Coal River sub-suite, include equigranular or fine-grained, plagioclase ± hornblende ± biotite ± quartz porphyritic, granodiorite to quartz monzodiorite (and rare monzogranite). Markedly homogeneous trace element compositions for this sub-suite and a relative sodium enrichment compared to other members of the Tay River suite (Chapter 2)—as well as the presence of inherited cores and (or) xenocrystic zircon of Paleozoic age (see Table 2.2, Chapter 2; Appendices B.6 and B.9)—are compatible with a possible partial to full derivation from basalt flows and sills of this age in the eastern Selwyn Basin. The 98–95 Ma Tungsten plutonic suite comprises reduced, slightly peraluminous, high-K calc-alkaline intrusions that are associated with late-stage, silica-undersaturated, alkalic or near-alkalic shoshonitic lamprophyre dykes. The Tungsten suite plutons are typically small, and commonly form rounded plugs (possibly cupolas to buried batholiths; Rasmussen et al., 2011b; Figs. 4.1 and 4.2). Monazite and biotite are ubiquitous, and muscovite (where present) typically forms small ragged sub-solidus grains (although euhedral magmatic muscovite is observed in some intrusions; e.g., Lened; Fig. 4.2; Appendix A.1). Highly evolved, or crustally contaminated, leucocratic phases may also contain garnet or tourmaline and occur as marginal plutonic phases, and cross-cutting aplite and pegmatite dykes. Dioritic enclaves are also observed in some plutons (e.g., van Middelaar, 1988; Gordey and Anderson, 1993). Although minor compositional differences between the Tungsten suite intrusions in the northeast and southeast Selwyn Basin suggest a greater degree of supracrustal input to magmas emplaced in the northeast Selwyn Basin, this suite is interpreted to an intermediate-composition infracrustal protolith. Magmatic fractionation, and (or) small- fraction partial melting are responsible for the relatively felsic composition of this plutonic suite. A shared association with late-stage or cross-cutting shoshonitic lamprophyre dykes was used to tentatively link the Tungsten and Mayo plutonic suites (e.g., Hart et al., 2005), which together form a compositional continuum of magmas derived from the same protolith. The 98–93 Ma Mayo plutonic suite (95–93 Ma in the northwest Selwyn Basin) comprises mainly high-K calc- alkaline, I-type intrusions that are associated with, but not necessarily cogenetic to, late-stage shoshonitic lamprophyre dykes. These plutons are mainly small (Figs. 4.1 and 4.2), reduced to slightly oxidized (titanite ± magnetite bearing), and metaluminous to slightly peraluminous in composition. Intrusive phases range from hornblende–biotite (± trace clinopyroxene) quartz monzonite or monzodiorite, to biotite–hornblende (± trace clinopyroxene) monzogranite to granodiorite. Associated leucogranitic, muscovite ± garnet-bearing phases are rare. The overall chemistry of this plutonic suite is compatible with a derivation from mainly, if not entirely, high- temperature partial melting of intermediate-composition metaigneous basement, similar to that inferred for the Tay River plutonic suite. Associated lamprophyre dykes do not appear to form a compositional end-member to the Mayo suite, although some of the more siliceous intrusive phases may display some geochemical evidence for contamination by the shoshonitic mantle-derived magmas, The lamprophyre dykes, however, mainly provide evidence for the influx of mantle-derived heat into the crust at this time, which would have been required for 257 hornblende ± biotite dehydration reactions in the infracrustal protolith. The 94–89 Ma Tombstone plutonic suite (93–90 Ma in the northwest Selwyn Basin) is found in the most inboard position of the ‘mid’-Cretaceous plutonic suites in the Selwyn Basin (Figs. 4.1 and 4.2). The Tombstone suite intrusions also comprise the youngest phase of this syn- to post-accretionary magmatic event in the Selwyn Basin. Tombstone suite intrusions form clusters of rounded, commonly concentrically zoned plutons (Figs. 4.1 and 4.2). Most intrusions belonging to this plutonic suite are sub-alkaline and metaluminous, however, the Tombstone plutonic suite is characterized by an association with abundant mafic and (or) silica-undersaturated intrusive phases (e.g., Mortensen et al., 2000; Hart et al., 2004a, 2005). In Chapter 2, this plutonic suite is described as comprising aluminous alkali-calcic (high-K calc-alkaline) to alkaline intrusions of mixed I- and A-type affinity, not unlike shoshonitic to high-K calc-alkaline post-collisional magmatism (e.g., Liégeois et al., 1998; Bonin et al., 1998). The Tombstone suite was, therefore, derived from a mixture of magmas from at least two main magma sources. Alkaline (shoshonitic) and LILE-rich magmas thought to have been sourced in an enriched or lithospheric mantle include the low-SiO2 tinguaite and several mafic intrusive phases (although the least siliceous mafic intrusions border on high-K calc-alkaline to shoshonitic compositions). High-K calc-alkaline magmas are thought to have been sourced in infracrustal basement of intermediate composition (similar to that inferred for the Tay River and Mayo plutonic suites). Diverging geochemical trends and an increase in scatter with decreased SiO2 on Harker variation diagrams are compatible with mixing between these two parental magmas in some intrusive phases. The Tombstone, Mayo, and Tungsten plutonic suite are collectively referred to as the Tombstone–Tungsten Belt (TTB: Hart et al., 2004a). 4.2 METHODS AND RESULTS Sample(s) inferred to be representative of the general composition of each intrusion (more than one sample was taken from several batholiths) were chosen for isotopic analysis. These samples were typically a medium- to coarse- grained, K-feldspar-phyric, phase of intermediate to felsic composition. Basic descriptions, as well as U–Pb crystallization ages, trace element ratios, and geochemical parameters presented in Chapter 2 for all samples that underwent isotopic analysis are summarized in Table 4.1, and are located in Figure 4.2. Prior to crushing, hand samples and thin sections were examined and sample material containing quartz ± sulphide bearing veins or fractures, or pervasively altered samples, were avoided. 4.2.1 Radiogenic Isotopes All sample preparation (with the exception of crushing and grinding, done at ALS Chemex, Vancouver, for whole rock isotopic samples), geochemical separation, isotopic analysis, and data correction and reduction for radiogenic isotopic analyses were done at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia (UBC), as detailed in Chapter 3. 258 4.2.1.1 Sm–Nd and Rb–Sr Sm–Nd and Rb–Sr isotopic analysis was done on 38 whole rock powders of intrusive samples. Present-day neodymium and strontium isotopic compositions, and calculated neodymium and strontium isotopic compositions for 100 Ma are reported in Table 4.1a,b and summarized in Figure 4.3a. Compiled isotopic data are also presented and referenced in Figure 4.3a, and tabulated in Appendix G.3. Relatively low εNd100, ranging from –5.8 to –20.9 (Fig. 4.3a), characterizes syn- to post-accretionary intrusions in the Selwyn Basin. Older plutonic suites in the southwest Selwyn Basin (including the Anvil and Hyland plutonic suites, as well as the western portion of the Tay River plutonic suite), and younger members of the TTB in the eastern (particularly northeast) Selwyn Basin are relatively isotopically evolved, with εNd100 typically below –13. The bulk of the Coal River sub-suite (of the Tay River plutonic suite) and the Tombstone plutonic suite in the southeast Selwyn Basin are less isotopically evolved and typically have εNd100 above –13. A group of plutons with ages and geochemical compositions similar to those of the Mayo and Tombstone suite plutons in the northeast Selwyn Basin is distinguished here on the basis of their anomalous isotopic compositions (i.e., εNd100 < –18). These isotopically unique plutons comprise the Hess–MacMillan sub-suite, named here for the proximity of two examples (Pelly River and Mount Wilson plutons; Fig. 4.2) to the Hess–MacMillan fault system along the Yukon–Northwest Territories border. Neodymium isotopic compositions of a few mafic dykes that were also analyzed are notably more radiogenic than plutonic examples with respect to neodymium. These include a high-K calc-alkaline quartz diorite dyke associated with the Tay River plutonic suite in the southeast Selwyn Basin, a shoshonitic lamprophyre (kersantite) dyke and a leucocratic silica-undersaturated dyke associated with the Tungsten plutonic suite in the southeast Selwyn Basin, and an isotopically highly evolved shoshonitic lamprophyre (kersantite) dyke associated with the Hess–MacMillan sub-suite of the Mayo plutonic suite in the northeast Selwyn Basin (Fig. 4.2). Strontium isotopes range from 87Sr/86Sr100 ≈ 0.71–0.73 in the eastern Selwyn Basin. Outliers to this range include samples of the Anvil plutonic suite in the western Selwyn Basin, which has extremely radiogenic 87Sr/86Sr100 (0.738– 0.749), particularly with respect to the sole εNd100 value for this plutonic suite (–16.5). The wide range of 87Sr/86Sr100 for Tay River suite plutonism in the western Selwyn Basin contrasts with a more restricted εNd100 range, although at least half of the 87Sr/86Sr100 values for the Tay River suite in the western Selwyn Basin are within 0.717 ± 0.01. 4.2.1.2 Pb–Pb Lead isotopic analysis was done on feldspar separates from 32 intrusive samples. Lead isotopic data are reported at their present-day values in Table 4.1c and summarized in Figure 4.3b47. Compiled isotopic data are also presented and referenced in Figure 4.3b, and tabulated in Appendix G.3. Lead isotopic compositions for intrusive rocks (206Pb/204Pb: 19.03–19.77; 207Pb/204Pb: 15.64–15.86; 208Pb/204Pb: 38.80–40.02) indicate a dominantly crustal source of lead for intrusions in the eastern Selwyn Basin. There is a 47 The lead isotopic compositions were previously reported and briefly discussed by Rasmussen et al. (2007). 259 distinct gap in the range of 206Pb/204Pb and (to a lesser degree) the range of 208Pb/204Pb compositions. Samples that have 206Pb/204Pb below 19.3 (and 208Pb/204Pb < 39.4) include all members of the Mayo plutonic suite in the northeast Selwyn Basin, most of the Tombstone plutonic suite in the northeast Selwyn Basin, and one member of the Coal River sub-suite of the Tay River suite in the eastern Selwyn Basin (the Park pluton; Figs. 4.2 and 4.3b). The single sample of the Tay River plutonic suite from the western Selwyn Basin also has a similarly non-radiogenic lead isotopic composition (Fig. 4.3b). This bimodal lead isotope distribution appears to be geographically controlled, which suggests that two distinct lead reservoirs or basement domains are present in the Selwyn Basin. Less lead- radiogenic basement may underlie the northeast and eastern Selwyn Basin (and possibly the westernmost Selwyn Basin), despite geochemical arguments for a geochemically similar infracrustal source of magma for the Tay River, Tungsten, Mayo, and Tombstone plutonic suites across the Selwyn Basin. The single sample of a Tay River suite intrusion in the western Selwyn Basin is the least radiogenic with respect to all lead isotopic compositions, and could even suggest a third lead reservoir in this region. It is notable that samples with less radiogenic lead isotopic compositions are associated with more evolved neodymium and strontium isotopic compositions (Fig. 4.3a,b). This unusual coupling might have resulted if relatively early metamorphism of a samarium-poor and rubidium-rich magma protolith removed uranium and thorium (relative to the non-radioactive lead isotope: 204Pb). 4.2.2 Stable Isotopes Most of the plutonic rocks sampled in the Selwyn Basin for this study, as well as 18 samples from S. Heffernan’s collection, were also analyzed for stable isotopes. All sample preparation, analysis, and data reduction followed the analytical methodology detailed in Chapter 3. Oxygen and sulfur isotopic data are reported in Table 4.1d and summarized in Figure 4.3c48. Compiled isotopic data are also presented and referenced in Figure 4.3 and tabulated in Appendix G.3). 4.2.2.1 Oxygen Oxygen isotopic compositions were determined from quartz separates for 53 plutonic samples. All of the intrusions analyzed in the Selwyn Basin study area have a normal distribution of high positive (i.e., isotopically ‘heavy’) δ18O ranging from +8.4 to +16.9‰ (Fig. 4.3c). Many samples have δ18O values above +14.5‰, which is considered to be unusually high, even for granitic rocks (e.g., Taylor and Sheppard, 1986). In Chapter 3, δ18O above +14.5‰ was considered to be the result of incorporation material from pelitic and (or) carbonate rocks. There is very little systematic variation in the δ18O across the study area, indicating that the majority of the rocks in the region were derived from one or more protolith with a similar range of oxygen compositions. Generally lower δ18O (i.e., < +11.9‰), however, is found in samples of the Hyland and Tay River plutonic suites (including the Coal River sub- suite) in the southeast Selwyn Basin. Although there appears to be a decrease in δ18O variation from older (i.e., Tay River) to younger (i.e., Tombstone) plutonic suites in the eastern Selwyn Basin, the plutonic suites with the smallest range in δ18O (i.e., the TTB intrusions) were also sampled from smaller geographic areas. The increase in δ18O 48 The oxygen and sulfur isotopic compositions were previously reported and discussed briefly by Rasmussen et al. (2010). 260 variation from the Tombstone to the Tay River suite samples in Figure 4.3c, therefore, may be related to regional compositional variations within the infracrustal basement, rather than intra-suite variations. Lower δ18O in some samples from the southeast Selwyn Basin is correlated broadly to less evolved neodymium and strontium isotopic compositions (Fig. 4.3a,b). 4.2.2.2 Sulfur Sulfur isotopic compositions were determined from whole rock powders for 60 plutonic and mafic dyke samples. Throughout the Selwyn Basin study area δ34S ranges from +0.8 to +13.6‰, but has a bimodal distribution. The majority of the intrusions have δ34S values ranging from +6 to +10‰. All of the intrusions emplaced in the northeast Selwyn Basin appear to have high positive sulfur isotopic compositions, but only some of the intrusions in the southeast Selwyn Basin have elevated δ34S. The lower δ34S population of +3 to +6‰ is, therefore, typical of plutons in the southeast Selwyn Basin, and (as with δ18O) include several samples of the Hyland and Tay River plutonic suites. Lower δ34S is correlated broadly to less evolved neodymium and strontium isotopic compositions (Fig. 4.3a,c). 4.3 DISCUSSION AND INTERPRETATION OF THE ISOTOPIC DATASETS Due to the large datasets that are presented here (including the new and compiled data), the plutonic suites in the Selwyn Basin are informally subdivided into two groups that will be plotted and discussed separately. Plutonic suites containing a large number of batholiths are examined first, which includes the Early Cretaceous S-type Anvil and Hyland plutonic suites, and the ‘mid’-Cretaceous I-type Tay River plutonic suite (and associated South Fork volcanic rocks; Figs. 4.1 and 4.2). Second, the TTB intrusions (the Tungsten, Mayo, and Tombstone plutonic suites) are collectively plotted and discussed. The TTB plutons are all interpreted to be of mainly I-type affinity and from a magma source that is chemically (if not necessarily isotopically) similar to that inferred for the Tay River plutonic suite (Chapter 2). However, obvious input from enriched lithospheric mantle-derived magmas is apparent in the late- stage lamprophyres associated with the Tungsten and Mayo plutonic suites, and the mafic to alkalic intrusive phases associated with the Tombstone plutonic suite (e.g., Hart et al., 2005). 4.3.1 Anvil, Hyland, and Tay River Plutonic Suites 4.3.1.1 Radiogenic Isotopic Datasets The predominantly S-type Anvil and Hyland plutonic suites, and the mainly I-type Tay River plutonic suite typically have isotopic compositions that are more evolved than the Early Cretaceous Dawson Range and Casino continental magmatic arc rocks (Fig. 4.4a). Neither the Hyland and Tay River plutonic suites approach the very highly evolved isotopic compositions that characterize ‘minimum melt’ magmas in the Cassiar Terrane (represented by the “migmatite” of Morris and Creaser, 2008, and indicated by the black “+”; Fig. 4.4a). The very high 87Sr/86Sr100 range for the Anvil plutonic suite (Fig. 4.4a), however, is compatible with its interpreted petrogenesis by low-temperature, 261 ‘minimum melting’ of a pelitic protolith. Some samples of the Hyland and Tay River plutonic suites (as well as some magmatic arc samples) have anomalously high 147Sm/144Nd (outlined with dashed blue line; Fig. 4.4b), given their intermediate to felsic compositions. Discussions in Chapter 3 have already established that these elevated 147Sm/144Nd values are likely due to Sm/Nd fractionation during partial melting of crustal rocks (e.g., Ayres and Harris, 1997). The calculated depleted mantle model ages (TDMs) in Figure 4.4c for samples with anomalous 147Sm/144Nd, therefore, are not geologically meaningful. More reasonable TDMs exhibited by the Anvil, Hyland, and Tay River plutonic suites range from 1.2 Ga for a mafic Tay River suite dyke, to 2.2 Ga for an Anvil suite sample— although most TDMs are less than 2.0 Ga. These model ages suggest that the magma source(s) to these plutonic suites were, on average (i.e., including crustal assimilants and contaminants incorporated into any mantle-derived magmas), extracted from the mantle between 1.2–2.0 Ga. The most primitive isotopic results were obtained from a mafic dyke and two Coal River sub-suite plutons, both of the Tay River plutonic suite, that are isotopically similar to the most evolved continental magmatic arc samples. The remaining samples of the Anvil, Hyland, and Tay River plutonic suites plot between moderately evolved compositions (ca. εNd100 ≈ –9 and 87Sr/86Sr100 ≈ 0.710—exemplified by the majority of the Tay River plutonic suite in the eastern Selwyn Basin, particularly the Coal River sub-suite), to highly evolved compositions (ca. εNd100 ≈ –17 and 87Sr/86Sr100 ≈ 0.720—exemplified by the Tay River plutonic suite in the western Selwyn Basin and by most samples of the Hyland plutonic suite in the eastern Selwyn Basin). There is a break in the range of neodymium isotopic compositions separating Tay River suite samples in the eastern Selwyn Basin from the western Selwyn Basin (Fig. 4.4a). The Hyland suite overlaps isotopically with the Tay River suite from both regions of the Selwyn Basin, and might also display a similar compositional break (although this could be due to a combination of the limited samples and its greater compositional heterogeneity; Fig. 4.4a). Neodymium isotopic data for the Anvil plutonic suite are limited to one sample, but the range of strontium isotopic compositions shown by the red bar in Figure 4.4a are sufficient to establish that this plutonic suite is the most radiogenic with respect to strontium. ‘Evolved’ members of the Tay River plutonic suite also have anomalously high 87Sr/86Sr100, and form an array extending from the main cluster of Tay River suite samples towards 87Sr/86Sr100 values up to ~0.73 (Fig. 4.4a). Similar trends exhibited by some members of the Cassiar plutonic suite in Chapter 3 were attributed to elevated Rb/Sr in the magma. ‘Evolved’ members of the Tay River plutonic suite have the highest Rb/Sr (e.g., 2–6) of the plutonic suites in the eastern Selwyn Basin, probably as a result of either fluid-absent small-fraction partial melting of biotite-rich rocks (e.g., Farina and Stevens, 2011), or the removal of plagioclase during magmatic fractionation (the former explanation was preferred in Chapter 3). Lead isotopic compositions for the Hyland and Tay River plutonic suites generally cluster immediately above the upper crustal “Shale Curve” of Godwin and Sinclair (1982) in Figure 4.5a. Several outliers are also present. More radiogenic compositions are due to the addition of radiogenic lead, possibly from unmetamorphosed country rocks. The Tay River plutonic suite sample from the western Selwyn Basin is particularly non-radiogenic when compared to plutonic rocks in the eastern Selwyn Basin, but this single sample may not be representative of igneous rocks in the western Selwyn Basin. At present there are no data available for the Anvil plutonic suite. 262 4.3.1.2 Stable Isotopic Datasets Oxygen and sulfur isotopic compositions for the Anvil, Hyland, and Tay River plutonic suites are highly variable and not coupled, however, when considered as a group they are broadly negatively correlated. The Hyland plutonic suite (and possibly the Tay River and South Fork suites in the western Selwyn Basin) generally have lower δ18O and higher δ34S, whereas the Tay River plutonic suite in the eastern Selwyn Basin (and possibly the Anvil suite in the western Selwyn Basin) have the highest δ18O and the lowest δ34S. Sulfur concentrations in most of the intrusions (e.g., typically <200 ppm; Fig. 4.6b) are low for granitic igneous rocks, which can contain up to 500 ppm sulfur at upper crustal pressures (e.g., 2–3 kbar; Wallace, 2001, 2005). This is in contrast to the high concentrations of sulfur found in typical shales (~1 wt.%) and quartzo–feldspathic clastic rocks (~1500 ppm; Ohmoto et al., 1990). If all, or most, of the sulfur in these sulfur-poor plutonic suites was derived from the melt source, then δ34S would not vary significantly with magmatic sulfur content. If the magmas were derived from a sulfur-free (or sulfur-poor) reservoir, and sulfur was added by later contamination of the magmas during ascension and emplacement, a better correlation between sulfur content and isotopic composition would be expected in Figure 4.6b. The gentle overall increase in sulfur content with δ34S (Fig. 4.6b) supports a primary and compositionally variable sulfur source in the magma protolith(s), but does not rule out some contamination by an isotopically ‘heavy’ source. Although sulfur and oxygen isotopes are not correlated for individual plutonic suites in Figure 4.6a, overall it appears that source of isotopically heavy oxygen (i.e., clay- or carbonate-rich rocks) probably contained relatively light sulfur. This correlation eliminates lithologies precipitated from seawater (evaporites or carbonates) as possible sources of sulfur to the Anvil, Hyland, and Tay River plutonic suites. Remaining possible sources of sulfur in these magmas are the sulfur-poor mantle (δ34S = 0 ± 3‰; Chaussidon and Lorand, 1990), or relatively sulfur-rich, but geologically minor, carbonaceous or pyritic siliciclastic rocks (not unlike a pyritic black shale sampled near the Lened pluton in the eastern Selwyn Basin; Fig. 4.7b; this study, unpub. data). A source of sulfur in pyritic or carbonaceous shales is the most consistent with the observed oxygen and radiogenic isotopic compositions, particularly for the dominantly to entirely supracrustally derived Anvil and Hyland plutonic suites. 4.3.1.3 Possible Sources of the Anvil, Hyland, and Tay River Suite Magmas The Anvil plutonic suite―which is highly strontium-radiogenic―has 87Sr/86Sr100 and δ18O values that overlap with those of pelitic rocks in southeast British Columbia (Fig. 4.7a), and δ34S values that overlap with a pyritic black shale of Upper Cambrian to Lower Ordovician-age in the eastern Selwyn Basin (Fig. 4.7b; this study, unpub. data). These observations are compatible with the pelitic protolith that was suggested for this plutonic suite in Chapter 2 on the basis of geochemical and mineralogical information. There are not enough neodymium isotopic data, however, to compare this plutonic suite with the northwest Laurentian basement rocks that are plotted in Figure 4.8, thus, no specific protoliths for the Anvil plutonic suite can be inferred. Hyland plutonic suite samples have scattered oxygen and sulfur isotopic compositions, but generally plot in a trajectory from less radiogenic Proterozoic mafic orthogneisses, to more radiogenic Neoproterozoic sedimentary 263 rocks (with respect to Sr; Fig. 4.7). None of the Hyland suite samples actually overlap with the compositional range exhibited by siliciclastic rocks of the Neoproterozoic Hyland Group in Figure 4.7a, despite field evidence that suggests this plutonic suite might have been derived by near-in situ anatexis of its Neoproterozoic host rocks (see Section 4.1.2). Neodymium isotopic compositions of the Hyland Group also are much more evolved (i.e., typically εNd100 > –19; Fig. 4.8a) than those of the Hyland plutonic suite. Instead, the Hyland plutonic suite is within the neodymium isotopic ranges defined by the Mesoproterozoic Mackenzie Mountains Supergroup (εNd100 = –18 to –9), and Paleozoic to Triassic miogeoclinal rocks in the Selwyn Basin (εNd100 = –19 to –7; Fig. 4.8). It is uncertain if the Mackenzie Mountains Supergroup (exposed northeast of the Selwyn Basin; see Fig. 3.2a, Chapter 3) underlies the southern Selwyn Basin where the Hyland plutonic suite was emplaced, particularly as these rocks are thought to be aerially restricted (e.g., Rainbird et al., 1997; Evenchick et al., 2005). It is also not known whether miogeoclinal rocks are present at depths appropriate to partial melting below the Selwyn Basin. There are not enough plutonic and basement isotopic data available, however, to more conclusively determine which basement rocks are the most likely protolith to the supracrustally derived Hyland plutonic suite. The Tay River (including the Coal River sub-suite) and South Fork suites have oxygen and strontium compositions that overlap with typical greywacke compositions (Taylor and Sheppard, 1986), although low 87Sr/86Sr100 and high δ18O in several samples suggest some input of isotopically ‘heavy’ oxygen was from carbonate rocks (Fig. 4.7a). The infracrustal protolith that was inferred for these magmatic rocks from geochemical and mineralogical observations, in Chapter 2, may itself have been derived by the anatexis of mainly quartzo–feldspathic metasedimentary rocks with subordinate carbonate rocks. Alternatively, these isotopic data may be compatible with a direct derivation of the Tay River plutonic suite (and associated volcanism) from a supracrustal melt source, albeit one with a relatively homogenous geochemical composition and abundant lithic detritus. The presence of common hornblende in the Tay River and South Fork rocks, however, is more likely due to a significant input of magma from mafic to intermediate igneous rocks. Mixing of materials derived from supracrustal and igneous rocks, therefore, might be one suggested source of the geochemical and mineralogical variation exhibited by these magmas. However, the geographically and geochemically defined clusters of neodymium and strontium isotopic data for the Tay River suite rocks in Figures 4.2a and 4.8 (as well as the relatively homogeneous trace element contents; Chapter 2) are not compatible with an origin by magma mixing. It is unlikely that magmas derived from distinctly different lithologies (e.g., mafic–intermediate metaigneous and quartzo–feldspathic metasedimentary) would share the same isotopic compositional range. Two such distinct sources of melt (regardless of whether mixing occurs in the source rocks prior to anatexis—as described in Chapter 3—or by mixing of disparate magma batches following anatexis) should produce a wider range of radiogenic isotopic compositions and trace element contents. Thus, the most likely source of the Tay River and South Fork magmas is an infracrustal basement of intermediate composition. If the infracrustal basement was derived from partial melting of sedimentary rocks, it would have retained some of the isotopic and geochemical features of supracrustal rocks that are also evident in the Tay River suite samples (e.g., high δ18O and K/Na). Some of the oxygen (and possibly sulfur) isotopes in the Tay River plutonic suite may also have been controlled by input from upper crustal pelitic rocks during ascension and emplacement, which might also be responsible for imparting the peraluminous composition to this plutonic suite noted in Chapter 2. 264 ‘Evolved’ members of the Tay River plutonic suite have a similar range of δ18O and other Tay River suite samples, for example, but their elevated Rb/Sr has resulted in a tendency towards more radiogenic 87Sr/86Sr100 compositions. The strontium and oxygen isotopic composition of these intrusions approach that for pelitic rocks in general, and Neoproterozoic quartzo–feldspathic rocks in the Selwyn Basin. Rather than a derivation from these upper crustal rocks, however, these ‘evolved’ Tay River suite samples may still have an I-type affinity if they were highly contaminated by, or that assimilated a large proportion of, upper crustal sedimentary rocks. The ‘evolved’ Tay River suite plutons were frequently observed to contain miarolitic cavities and (or) ribboned aplite–pegmatite zones (e.g., Rasmussen et al., 2007). These samples may have been taken from upper portions of the magma chambers, where local assimilation of isotopically evolved country rocks could have also have had a strong effect on their isotopic composition. It is clear from the data presented in Figure 4.8 that the Coal River sub-suite (of the Tay River plutonic suite) was not derived solely from partial melting of mafic volcanic and intrusive rocks in the Selwyn Basin, as was suggested in Chapter 2. Although the Coal River sub-suite samples appear to trend towards similar neodymium isotopic compositions and TDMs, they do not actually overlap with the Paleozoic igneous rocks (Fig. 4.8). The Coal River sub-suite must also contain magma derived from more isotopically evolved crustal rocks, perhaps the infracrustal basement (suggested for the remainder of the Tay River plutonic suite), or supracrustal rocks (possibly also of Paleozoic age). An argument could be made for the incorporation of magmas derived from pelitic (or other strongly aluminous supracrustal) rocks in this plutonic sub-suite, based on the peraluminous compositions apparent in Chapter 2, and on the zircon morphologies and compositions presented in Chapter 5. The neodymium isotopic range exhibited by the Coal River sub-suite does overlap with the least isotopically evolved Paleozoic miogeoclinal rocks containing interlayered mafic volcanic flows or sills (Fig. 4.8), but does not require that these specific rocks were involved in the genesis of this plutonic sub-suite. 4.3.2 Tungsten, Mayo, and Tombstone (TTB) Plutonic Suites 4.3.2.1 Radiogenic Isotopic Datasets The Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin have been plotted according to their geographic subdivisions in Figure 4.10 to illustrate the isotopic variations exhibited by within the plutonic suites across the Selwyn Basin. These TTB intrusions plot as three main clusters: • a moderately isotopically evolved population (εNd100 ≈ –9 and 87Sr/86Sr100 ≈ 0.711) defined by Tombstone and Mayo suite intrusions in the northwest Selwyn Basin, as well as shoshonitic lamprophyre dykes across the Selwyn Basin, • a highly isotopically evolved population (εNd100 ≈ –15 and 87Sr/86Sr100 ≈ 0.722) similar to that of the Tay River suite intrusions in the western Selwyn Basin, and • an extremely highly isotopically evolved population defined by a several Tombstone-age and Mayo-age intrusions (including one shoshonitic lamprophyre dyke) in the northeast Selwyn Basin (referred to as the Hess–MacMillan sub-suite of the Mayo and Tombstone plutonic suites). 265 The first two isotopic populations broadly correlate with much of the Tay River plutonic suite in the eastern and western Selwyn Basin, respectively (see Fig. 4.4a). The latter isotopic population is unique to the TTB intrusions. When compared to the continental magmatic arc rocks to the southwest, TTB intrusions are always more isotopically evolved (Fig. 4.10a). For the most part, the TTB intrusions overlap with the Tay River suite in the eastern Selwyn Basin, but do not extend to the juvenile compositions exhibited by a mafic dyke and several Coal River sub-suite samples from the Tay River suite (Fig. 4.4a). Nor do any of the TTB rocks approach isotopic compositions similar to the single proxy for an enriched mantle composition, represented by a Recent basalt in the southern Selwyn Basin (see Figs. 1.1b and 1.2a). However, nearly all of the shoshonitic lamprophyre dykes in the Selwyn Basin yield tightly constrained compositions of εNd100 = –8 ± 1 and 87Sr/86Sr100 = 0.712 ± 0.02 from the northern to eastern Selwyn Basin (Fig. 4.10a). These compositions are remarkably evolved for mantle-derived magmas, and they indicate that the lithospheric mantle underlying this region was relatively homogeneous. If the shoshonitic lamprophyres are representative of enriched mantle compositions across most of the Selwyn Basin, then the isotopic data in Figure 4.9a are compatible with a significant enriched mantle component in the Mayo and Tombstone plutonic suites in the northwest Selwyn Basin (although not in the northeast Selwyn Basin) as well as the Tombstone plutonic suite in the southeast Selwyn Basin. The interpretation of an enriched mantle component in the Tombstone plutonic suite is not new, and previous workers have proposed input from mantle- derived melts for this suite in the northern Selwyn Basin (Mortensen et al., 2000; Hart et al., 2004a, 2005). The chemical composition of the Mayo plutonic suite, however, does not necessarily indicate a significant mantle- derived component, despite its association with late shoshonitic lamprophyre dykes (although some input has not been precluded; Hart et al., 2004a, 2005; Chapter 2) and interpretations otherwise (e.g., Mair et al., 2011). The isotopic composition of the Mayo plutonic suite in the northwest Selwyn Basin might instead reflect the isotopic composition of what is inferred here to be an infracrustal source of high-K calc-alkaline magmas. In the northeast Selwyn Basin, the infracrustal protolith to the Mayo plutonic suite (as well as the Tombstone and Tungsten plutonic suites sampled in this region) appears to have been more isotopically evolved than in the northwest Selwyn Basin (Fig. 4.9a). All of the Tungsten plutonic suite has a similar range of highly evolved neodymium and strontium isotopic compositions (if associated mafic dykes are ignored). However, there is barely discernible difference in the isotopic composition of Tungsten suite samples from the northeast and southeast Selwyn Basin. Most samples of the Tungsten suite from the northeast Selwyn Basin are slightly more isotopically evolved49, which may be compatible with the greater degree of crustal contamination that was suggested as one possible explanation for the slight geochemical differences also observed in these intrusions relative to those in the southeast Selwyn Basin (Chapter 2). The most isotopically evolved intrusions in the eastern Selwyn Basin are members of the Hess–MacMillan sub- 49 This small, geographically based difference in neodymium and strontium isotopic compositions within the Tungsten plutonic suite is more evident on plots involving lead and neodymium or strontium isotopes (not shown). 266 suite, which are distinguished on the basis of their extremely highly evolved neodymium and strontium isotopes. This sub-suite has not been previously discussed because these intrusions are not geochemically, mineralogically, or temporally distinct from the Mayo and Tombstone plutonic suites. In addition to its extremely highly evolved isotopic composition, however, this sub-suite is limited to intrusions of both Mayo- and Tombstone-age that are confined to a small region near the Yukon–Northwest Territories border (Fig. 4.2). The Mayo and Tombstone suite samples identified as members of the Hess–MacMillan sub-suite examples by this study are named for their close proximity to the Hess–MacMillan fault system. The most proximal of these intrusions display textural evidence of syn-magmatic deformation. The Pelly River pluton (Fig. 4.2), for example, has a moderately to intensely foliated or lineated contact phase adjacent to mylonitized country rock along the northwest margin of the pluton (Anderson, 1982). Unlike many examples of the isotopically more evolved Cassiar plutonic suite presented in Chapter 3, calculated 147Sm/144Nd is unusually high in only one Tungsten suite sample (Fig. 4.9b). This implies that the ~1.4 Ga to ~2.4 Ga range of TDMs exhibited by most of the TTB plutonic rocks (Fig. 4.9c), therefore, were not affected by REE fractionation during anatectic (or magmatic) processes. These TDMs provide a maximum estimate of the age of melt extraction from the depleted mantle for the mantle and crustal materials that ultimately were melted and incorporated into the TTB magmas. The high TDMs require an old source of magma for the TTB plutonic suites, and (or) a significant input of non-radiogenic neodymium from an isotopically evolved crust. Slightly younger depleted mantle model ages for the mafic dykes located predominantly in the northwest Selwyn Basin at the Scheelite Dome intrusions and mineralization (Fig. 4.1; Mair et al., 2011)―but also found in the southeast Selwyn Basin (Fig. 4.2)―suggest that enrichment of the mantle occurred between 1.0–1.5 Ga (Fig. 4.9c). Older depleted mantle model ages determined for lamprophyre dykes (e.g., 1.8 Ga and 2.1 Ga; Fig. 4.9c) are likely due to more extensive crustal contamination of these specific dykes. Lead isotopic compositions for TTB intrusions in the Selwyn Basin cluster near the upper crustal “Shale Curve” of Godwin and Sinclair (1982; Fig. 4.10). This would be consistent with a derivation entirely from crustal rocks for these plutons; however, even the shoshonitic lamprophyres from the northwest Selwyn Basin overlap with this upper crustal lead evolutionary curve (Fig. 4.10). Contamination of lead-poor, mantle-derived magmas during their ascent and emplacement in lead-rich crust might explain these radiogenic compositions of the lamprophyre dykes in this region. Another explanation, however, is that enrichment of portions of the mantle that melted to form the lamprophyre dykes also added LILEs (including U and Th) to these magmas (e.g., Roden and Murthy, 1985). This could produce lamprophyres that have lead isotopic compositions similar to that of upper crustal growth curves (like the Shale Curve; Fig. 4.10). Mantle enrichment events are commonly attributed to the subduction of pelagic or continentally derived sediments on a down-going oceanic plate (e.g., Zindler and Hart, 1986), and enriched regions are typically found in old continental lithosphere (Roden and Murthy, 1985). It is not improbable, therefore, that enriched mantle-derived mafic magmas have lead isotopic compositions similar to those of upper crustal rocks. The lead isotopic plots in Figure 4.10 do not appear to be capable of distinguishing between enriched mantle and crustal origins for the TTB intrusions. 267 There is also a bimodal distribution of 206Pb/204Pb values (Fig. 4.10). Samples in Figure 4.10 with low 206Pb (i.e., 206Pb/204Pb < 19.3) comprise the Mayo plutonic suite and most of the Tombstone plutonic suite located in the northeast Selwyn Basin only (as well as one sample of the Coal River sub-suite in the eastern, southeast Selwyn Basin; Figs. 4.2 and 4.5a). These samples also have lower 207Pb/204Pb than most of the TTB intrusions in the southeast Selwyn Basin (Fig. 4.10a). The Mayo and Tombstone suite samples from the northeast Selwyn Basin share similar εNd100 and 87Sr/86Sr100 ranges with the Tungsten suite samples from both the northeast and southeast Selwyn Basin (Fig. 4.9), although lead isotopic compositions of the Tungsten plutonic suite are in the higher 206Pb/204Pb population (Fig. 4.10). Despite the distinct neodymium and strontium isotopic compositions of the Hess– MacMillan sub-suite in the northeast Selwyn Basin, its lead isotopic compositions of this sub-suite overlap with both 206Pb/204Pb populations (Fig. 4.10a). Altogether, it seems that the basement involved in the generation most of the intrusions in the northeast Selwyn Basin (excluding some examples of the Tombstone suite and Hess–MacMillan sub-suite), has a different lead isotopic signature than basement in the southeast and northwest Selwyn Basin. The lead source for most of the magmas emplaced into the northeast Selwyn Basin was less radiogenic with respect to 206Pb (and, to a lesser degree, 207Pb) compared to the southeast Selwyn Basin. Isotopically more primitive basement (e.g., U–Th-poor mafic rocks) in the northeast Selwyn Basin, however, is incompatible with the more evolved neodymium and strontium isotopic compositions. Therefore, the lead compositions in this region could instead be due to early, high-grade metamorphism that would have depleted uranium and led to lower 206Pb/204Pb and 207Pb/204Pb. The controls on the lead isotopic compositions in magmas across the Selwyn basin, therefore, may be largely controlled by the metamorphic history of the magma source rocks. The bimodal 206Pb/204Pb distribution does not appear to extend to the Mayo and Tombstone suite samples in the northwest Selwyn Basin, which have a large range of 206Pb/204Pb and relatively non-radiogenic 207Pb/204Pb and 208Pb/204Pb (Fig. 4.10). In the northwest Selwyn Basin, lower 207Pb/204Pb is indicative of younger basement, and the lower 208Pb/204Pb is incompatible with high-grade metamorphism and instead may indicate the presence of rocks of oceanic affinity in the basement (e.g., Tosdal et al., 1999). An interpretation invoking younger, isotopically primitive crust enriched in thorium by seawater (such as island arc volcanic rocks) is also compatible by the lower neodymium and strontium isotopic compositions of the TTB intrusions in the northwest Selwyn Basin (Fig. 4.9a). 4.3.2.2 Stable Isotopic Datasets Oxygen isotopic compositions of the TTB intrusions are high overall, commonly exceeding a δ18O value of +14.5‰ in all but the Tombstone plutonic suite (Fig. 4.11a). The highest oxygen isotopic compositions that were documented by Taylor and Sheppard (1986) for plutonic quartz reach a δ18O value of ~16‰, but are only observed in crustally contaminated plutonic border zones and (or) small volume intrusions (e.g., dykes; Taylor and Sheppard, 1986). However, significant magmatic contamination from clay- or carbonate-rich rocks, or the assimilation of such rocks by magmas, could also cause δ18O to exceed values of about +14.5‰ in the TTB intrusions(G. Arehart, 2012, pers. comm.). Oxygen isotopic compositions are also less variable overall for the TTB plutonic suites (Fig. 4.11a) than they are for the Anvil, Hyland, and Tay River plutonic suites (Fig. 4.6a). 268 Sulfur isotopes display a wider range than the oxygen isotopes for the TTB plutons (Fig. 4.11). Also unlike the Anvil, Hyland, and Tay Rivers suites, the sulfur and oxygen isotopic compositions of the TTB intrusions are positively correlated. TTB samples from the southeast Selwyn Basin generally have the lowest (‘lightest’) stable isotopic compositions, whereas those from the northeast Selwyn Basin generally have the highest (‘heaviest’) stable isotopic compositions. The lightest oxygen and sulfur isotopic compositions (as well as lowest sulfur concentrations; Fig. 4.11b) are typical of Tombstone and Tungsten suite plutons emplaced in the southeast Selwyn Basin (where more juvenile Nd and Sr isotopic compositions also occur; e.g., Fig. 4.3a). Intrusions emplaced in the northeast Selwyn Basin typically have higher positive sulfur compositions (and concentrations), and generally higher positive oxygen compositions (Fig. 4.11a). The broadly positive oxygen–sulfur coupling exhibited by the TTB samples suggests that the sulfur isotopic compositions were inherited from the magma source (with the oxygen isotopes), rather than from a sulfur-rich contaminant. This positive association suggests that plutonic rocks in the northeast Selwyn Basin were derived from rocks containing heavier oxygen and sulfur isotopes, and higher sulfur concentrations (as well as more evolved Nd and Sr isotopes) than in the southeast Selwyn Basin. The gradual change in oxygen and sulfur isotopic compositions from southeast to northeast is supportive of a lithological control on the composition of TTB plutons in the eastern Selwyn Basin. For example, these variations could result from the variable incorporation of sulfur-rich (and isotopically heavy) rocks into the TTB magmas across the eastern Selwyn Basin. Alternatively, the stable isotopic variations could also reflect compositionally distinct infracrustal basement domains in the southeast and northeast Selwyn Basin. This latter interpretation may be more compatible with the geographically control on radiogenic isotopic data in Figures 4.9 and 4.10. 4.3.2.3 Possible Sources of the Tungsten, Mayo, and Tombstone Suite Magmas The stable isotopic compositions of the TTB intrusions are positively coupled with radiogenic isotopic compositions, which is consistent with an increased input of material derived from supracrustal rocks containing high positive sulfur and oxygen (e.g., chemically precipitated sedimentary rocks and (or) clay-rich siliciclastic rocks). For example, there is a general coupling between the strontium and oxygen isotopic systems in the Selwyn Basin intrusions in Figure 4.12a—from isotopically primitive samples from the northwest (the Mayo and Tombstone suites) to isotopically evolved samples from the southeast (the Tombstone and Tungsten suites), and then the northeast Selwyn Basin (Tungsten, Mayo, and Tombstone suites). A similar coupling is also apparent between the strontium and sulphur isotopic systems (although there are no sulphur data for TTB plutons in the northwest Selwyn Basin). Only in the most isotopically evolved samples of the TTB intrusions (the Hess–MacMillan sub-suite) are the radiogenic and stable isotopic systems not as closely coupled (i.e., they do not always have the highest O or S isotopic compositions, despite their extremely evolved Nd and Sr isotopes). When the neodymium isotopic compositions and TDMs of TTB intrusions are compared to those of Laurentian basement, they overlap in composition with the same supracrustal basement rocks as the Anvil, Hyland, and Tay River plutonic suites (Fig. 4.13a): the Mesoproterozoic Mackenzie Mountains Supergroup, and (mainly) Paleozoic 269 miogeoclinal rocks in the Selwyn Basin (Fig. 4.13b). However, the TTB suites have been interpreted on the basis of geochemistry and mineralogy to have been derived from a mainly, or entirely, metaigneous crust of intermediate composition (Chapter 2). The clusters of neodymium, strontium, and oxygen isotopic compositions exhibited by these intrusions (Figs. 4.3a,c, 4.4a, 4.12a, and 4.13a) support this interpreted I-type affinity for the majority of the intrusions. Thus, despite their isotopic similarity to at least metasedimentary packages (Fig. 4.13b), metaigneous sources of magma should instead be considered for the TTB magmas. None of the exposed igneous rocks or proxies for metaigneous basement in the northern Cordilleran overlap with the TDM range of the TTB intrusions (Fig. 4.13b). This suggests that igneous rocks involved in the genesis of the TTB intrusions are not exposed in northwest Laurentia. Although there are many scattered occurrences of relatively small-volume mafic to felsic flows, sills, and intrusions (see Figs. 1.1b and 1.2a, Chapter 1), any widespread and large-volume igneous rocks that might provide a magma source to ‘mid’-Cretaceous magmatic rocks across the Selwyn Basin can only be present in the basement, buried below thick packages supracrustal rocks (see Fig. 3.2a, Chapter 3; e.g., Cook et al., 2004). If the TTB intrusions are indeed largely of I-type affinity, the infracrustal basement below the Selwyn Basin must be differ from exhumed metaigneous basement in southeast British Columbia (e.g., the Malton Complex; Vernon-Chamberlain, 1983) and Archean to Paleoproterozoic basement domains located farther east (Fig. 4.13b, and see Fig. 3.2a, Chapter 3). The Mayo plutonic suite and associated lamprophyre dykes from the northwest Selwyn Basin are the least isotopically evolved with respect to radiogenic isotopic compositions, and also have the lowest oxygen isotopes of all the TTB intrusions (Fig. 4.12a). Even the mantle-derived shoshonitic lamprophyres have high positive δ18O (i.e., +10 to +12‰) and plot with plutonic examples of the Mayo suite in the greywacke field of Taylor and Sheppard (1986)— although they form a trajectory towards enriched mantle (and mafic orthogneiss; Vernon–Chamberlain, 1983) compositions. Lead isotopic compositions, and less evolved neodymium and strontium isotopic compositions, for the Mayo (and Tombstone) intrusions in the northwest Selwyn Basin are compatible with a relatively primitive infracrustal basement, possibly of oceanic island arc affinity. The Mayo suite intrusions, however, do not appear to have mixed with enriched mantle derived magmas, at least with respect to their chemical composition (Chapter 2). The isotopic compositions of these intermediate to felsic plutonic rocks and shoshonitic lamprophyres are remarkably similar in the northwest Selwyn Basin. It is not clear whether the isotopic similarity between the enriched mantle underlying the entire Selwyn Basin and the suggested isotopically primitive infracrustal basement is coincidental, or whether the high-K calc-alkaline intrusive rocks of the Mayo (and Tombstone) plutonic suites originated in the underlying enriched mantle (Mair et al., 2011), as has been suggested for other high-K calc-alkaline suites that have had a more isotopically primitive protolith (e.g., the West African and Variscan orogenies; Liégeois et al., 1998; Barbarin, 1999; Janoušek et al., 2000; Ferré and Leake, 2001; Bonin, 2004). The Tombstone plutonic suite from the southeast Selwyn Basin is less isotopically primitive with respect to oxygen and strontium, but does overlap with the isotopic composition of the Mayo plutonic suite in the northwest Selwyn Basin (Fig. 4.12a). Lead isotopic compositions are also more radiogenic in the southeast Selwyn Basin, which is suggestive of an older magma source overall and (or) a U/Th enrichment (resulting from a more geochemically 270 evolved composition, or a lack of old metamorphism) in this basement. If a more isotopically and geochemically evolved (i.e., old and felsic) basement is present below the southeast Selwyn Basin, the isotopic composition of the Tombstone suite intrusions in this region may be the result of mixing between evolved basement melts and enriched mantle-derived magmas similar to those represented by shoshonitic lamprophyre dykes. This scenario would produce the comparatively lower εNd100, and higher 87Sr/86Sr100 and δ18O that characterize the Tombstone plutonic suite in the southeast Selwyn Basin (compared to the northeast Selwyn Basin). Alternatively, it may be that the basement to this region is just more evolved than in the northwest Selwyn Basin, but less evolved than in the northeast Selwyn Basin. In the southeast Selwyn Basin, however, the Tay River suite and Tombstone suite share similar radiogenic isotopic compositional ranges (for Nd, Sr, S ± Pb; Fig. 4.3), even though the Tay River suite was not modified by mantle-derived magmas (Chapter 2). These overlapping isotopic compositions are compatible with their derivation from the same (or a similar) isotopically intermediate, possibly more felsic, infracrustal basement in the southeast Selwyn Basin. The lack of evidence for mantle-derived magmas in the Tay River plutonic suite indicate that mantle input is not necessary to produce the less evolved isotopic compositions of plutonic rocks in the southeast Selwyn Basin. Strontium (and Nd; Fig. 4.9a) isotopic compositions for the all Tungsten suite plutons are similar to those for the Mayo and Tombstone suite intrusions in the northeast Selwyn Basin, however, lead isotopic signatures for the Tungsten plutonic suite are similar to rocks from the southeast Selwyn Basin. The similar neodymium, strontium, oxygen, and sulphur isotopic compositions of all TTB intrusions in the northeast Selwyn Basin (e.g., Fig. 4.3) argue for the same bulk source of magma (with the exception of the Hess–MacMillan sub-suite, discussed below). Chapter 2 postulated that the Tungsten plutonic suite represents smaller melt fractions or more highly fractionated magmas derived from the same infracrustal source rocks as the Mayo plutonic suite. The Tungsten plutonic suite may have sampled basement that is more characteristic of the north with respect to neodymium and strontium. The evolved lead isotopic compositions that typify the Tungsten plutonic suite, therefore, may be the result of LILE enrichment related to anatectic (i.e., small-fraction partial melting), magmatic (i.e., fractionation), and (or) contamination or assimilation processes. The Hess–MacMillan sub-suite samples are isotopically unique, even though they cannot be distinguished texturally, mineralogically, temporally, or geochemically from most Mayo and Tombstone suite intrusions (Chapter 2). One Mayo-age sample of this sub-suite has oxygen–strontium isotopic compositions that overlap with the range for Neoproterozoic metasedimentary rocks in the Selwyn Basin (Fig. 4.12a). Furthermore, all examples of this sub-suite have low εNd100 and old TDMs that overlap with the least isotopically evolved samples of Neoproterozoic metasedimentary rocks in the Selwyn Basin, as well as the least evolved samples from felsic Paleoproterozoic gneisses and basement domains (Fig. 4.13b). Thus, two possible sources of magma for the Hess–MacMillan sub- suite can be proposed. An orthogneiss of Paleoproterozoic age that was more isotopically evolved composition than the widespread infracrustal basement invoked for other post-arc ‘mid’-Cretaceous plutonic suites across the Selwyn Basin. Second, the Neoproterozoic Hyland Group supracrustal rocks (that comprise detritus derived from uplifted isotopically evolved basement domains farther the east; e.g., Leslie, 2009) are another isotopically appropriate protolith. In either case, however, the localization of these plutons as a narrow belt centred on the Yukon–Northwest 271 Territories border (the continental divide) requires an explanation. One possibility is that a particularly thick thrust sheet buried Neoproterozoic sedimentary rocks to depths compatible with fluid-absent partial melting (e.g., >25 km) in this particular region. Alternatively, decompression-assisted partial melting in this area may be related to syn- magmatic movement along the nearby Hess–MacMillan fault system (Fig. 4.2), particularly if the protolith rocks were already anomalously hot (i.e., enriched in heat-producing elements, or heated advectively by mantle-derived magmas). Additionally, zircon morphologies and compositions discussed in Chapter 5 for this sub-suite favour an origin in supracrustal rocks, although some magmatic zircon morphologies are indicative of high magmatic temperatures. An extremely crustally contaminated, shoshonitic lamprophyre dyke associated with a Mayo-age example of this sub-suite provides evidence for advective heat transfer from the mantle during the genesis of the Hess–MacMillan magmas (Fig. 4.9a). Despite similarly elevated radiogenic isotopic compositions, the Tombstone- age sample from the Hess–MacMillan sub-suite has a significantly different stable isotope composition (lower O and higher S), which may be related to lithological variations within the inferred supracrustal magma source. Thus, of all the TTB plutons examined here, only the Hess–MacMillan sub-suite (of both Mayo and Tombstone suite-age) appears to have been derived solely from partial melting of supracrustal rocks. 4.4 SUMMARY OF PETROGENETIC INTERPRETATIONS FOR SYN- TO POST- ACCRETIONARY MAGMATISM IN THE SELWYN BASIN Although the isotopic data for syn- to post-accretionary plutonism emplaced into the Selwyn Basin (particularly the eastern Selwyn Basin) are not examined as closely as similar datasets in Chapter 3 for the Cassiar Terrane, this simpler treatment of the isotopic data still provides information on the composition of magma protoliths. First, the isotopic data are generally supportive of conclusions drawn from geochemical (and mineralogical) data summarized and presented in Chapter 2. In most cases, however, the geochemical data appear to better discriminate between mantle versus crustal (and infracrustal vs. supracrustal) magma protoliths, although the isotopic data are better at detecting regional variations in the composition, age, and (or) metamorphic history of the proposed magma sources. Overall, the scattered isotopic (and geochemical) compositions of the Anvil and Hyland plutonic suite are compatible with a supracrustal origin, whereas the geographically controlled clusters of radiogenic isotopic compositions (and a lack of chemical variation) for the Tay River and TTB plutonic suites are more compatible with an infracrustal origin. The Early Cretaceous back-arc plutonic suites are interpreted to have been derived entirely (i.e., the Anvil suite) to predominantly (i.e., the Hyland suite) from supracrustal rocks (Chapter 2). The Anvil plutonic suite was mainly derived by low-temperature partial melting of homogeneous pelitic rocks, whereas the Hyland plutonic suite was mainly derived from moderate-temperature partial melting of more heterogeneous quartzo–feldspathic rocks containing calcareous or amphibolitic horizons (Chapter 2). The limited amount of isotopic data available for these two plutonic suites do not allow conclusive statements to be made regarding specific supracrustal protoliths for these plutonic suites. It was noted, however, that the Hyland plutonic suite overlaps isotopically (with respect to εNd100 and TDM) with the Mesoproterozoic Mackenzie Mountains Supergroup and with Paleozoic to Triassic miogeoclinal rocks in the Selwyn Basin—both of which may provide suitable, compositionally heterogeneous melt sources 272 (although basinal equivalents to the Mackenzie Mountains Supergroup would be necessitated). The difference between the pelitic and quartzo–feldspathic supracrustal protoliths inferred for these two plutonic suites may have been controlled by the relative distance from the ancestral margin of North America during deposition. Clay-rich rocks would have been deposited more distally, providing a pelitic source of magma for the Anvil plutonic suite in the western Selwyn Basin. Conversely, more heterogeneous quartzo–feldspathic rocks would have been deposited closer to the northwest margin of North America, would provide a fertile source of magma for the Hyland plutonic suite, which is located east of the Anvil suite. The ‘mid’-Cretaceous post-arc Tay River plutonic suite across the Selwyn Basin appears to have been derived from partial melting of relatively homogeneous rocks, although some variation in the isotopic composition of these source rocks exists between the east and west. In the eastern Selwyn Basin, the Tay River suite magmas are relatively moderately isotopically evolved with respect to neodymium and strontium, but strongly radiogenic with respect to lead. Oxygen and strontium isotopic compositions are more compatible with a supracrustal melt source of average greywacke composition (i.e., high positive δ18O and moderate 87Sr/86Sr100; Fig. 4.7a). While partial melting of relatively fine-grained, aluminous siliciclastic rocks might appear to have been important in the generation of the Tay River plutonic suite, geochemical and mineralogical data presented in Chapter 2 are more compatible with an infracrustal melt source of homogenous, intermediate composition. It is, therefore, inferred that this plutonic suite was derived by high-temperature partial melting of metaigneous basement with an intermediate composition (60–65 wt.% SiO2). Lead isotopic data for this plutonic suite in the eastern Selwyn Basin suggest either a relatively unmetamorphosed and (or) felsic (i.e., U–Th-rich) basement—the latter is also compatible with the moderately evolved neodymium and strontium isotopic compositions in this region. Extremely non-radiogenic lead isotopic compositions for one pluton in the western Selwyn Basin, conversely, suggest either the basement was depleted in U–Th during early high-grade metamorphism, or that the crust below this region is isotopically primitive (i.e., mafic and (or) young)—the former is the most compatible with the highly evolved neodymium and strontium isotopic composition in this region. Additional work is needed to determine whether the one sample analyzed for lead isotopes in the western Selwyn Basin is representative. Regardless of the regional variations, broadly intermediate- composition source rocks probably characterize the infracrustal basement across the Selwyn Basin. The Tay River and Mayo plutonic suites (and high-K calc-alkaline, I-type phases of the Tombstone plutonic suite) have similar major element geochemical compositions, despite a greater range of trace element compositional variation in the Mayo and Tombstone suites (Chapter 2). However, different basement domains appear to have been involved in the genesis of post-arc, I-type magmatism across the Selwyn Basin, despite the similar geochemical composition of all of these high-K calc-alkaline suites. For example, the Mayo and Tombstone plutonic suites in the northwest Selwyn Basin, and the Tay River and Tombstone plutonic suites in the southeast Selwyn Basin are moderately isotopically evolved (with respect to Nd and Sr). Slightly more sodic intrusions in the southeast Selwyn Basin (Chapter 2), and more evolved radiogenic isotopic compositions (Nd, Sr, and Pb) may suggest a more felsic infracrustal basement in the southeast Selwyn Basin (e.g., a continental magmatic arc that incorporated a high proportion of crustal rocks). Conversely, in the northwest Selwyn Basin, less geochemically and isotopically evolved rocks with high U/Th (e.g., oceanic island arc rocks that lack input from isotopically evolved crust) might 273 be a suitable infracrustal protolith to the ‘mid’-Cretaceous plutonic rocks in this region. Finally, the Tay River plutonic suite in the western Selwyn Basin, and the Mayo and Tombstone plutonic suites in the northeast Selwyn Basin, are highly isotopically evolved with respect to neodymium and strontium (but not with respect to lead, especially in the southwest Selwyn Basin). These characteristics are compatible with magma sourced from older, if not more felsic, basement that was metamorphosed (depleted in U–Th) relatively early in its history. Given these regional isotopic variations, it might be more appropriate to sub-divide the Tay River and Mayo plutonic suites on the basis of geography. There is some question as to whether the Tungsten plutonic suite was derived from supracrustal rocks, particularly in the northeast Selwyn Basin (Chapter 2). Its limited isotopic range (particularly with respect to Nd), however, is more compatible with the infracrustal melt source that is similar to that inferred for other TTB plutonic suites. It seems that the chemical composition of the Tungsten plutonic suite reflects anatectic or magmatic processes rather than protolith compositions. This plutonic suite appears to comprises either highly fractionated equivalents of the Mayo plutonic suite, or smaller-fraction partial melts from the same protolith rocks. Either process could also lead to an enrichment in LILEs, which may be why this particular plutonic suite has more radiogenic lead isotopic compositions than the Mayo or Tombstone plutonic suites. There is better evidence for a supracrustal protolith to the Hess–MacMillan sub-suite in the northeast Selwyn Basin. This sub-suite, which contains examples of both Mayo and Tombstone age, overlaps compositionally with the neodymium (including TDMs), strontium, and oxygen isotopic ranges for the Neoproterozoic Hyland Group. These localized S-type plutons, surrounded by regionally extensive I-type igneous rocks, may be the product of partial melting in deeply buried Neoproterozoic metasedimentary rocks along the Yukon–Northwest Territories border. This anatexis of supracrustal rocks may have been aided by decompression associated with the nearby Hess– MacMillan fault system and (or) advective heat influx in the form of lamprophyre dykes. 4.4.1 A Tectonic Model for Syn- to Post-Accretionary Magmatism in the Selwyn Basin A large volume of plutonic rocks was emplaced across the Selwyn Basin (e.g., Figs. 4.1 and 4.2). Early Cretaceous S-type intrusions of the Anvil and Hyland plutonic suites appear to have been emplaced in a back-arc setting that was experiencing relatively high heat flow from the mantle to crust. This scenario is similar to that envisioned for the S-type Cassiar plutonic suite in the Cassiar Terrane, which was located south of the Selwyn Basin in the Cretaceous (Chapter 3). The most widespread magmatism in the Selwyn Basin, however, is of ‘mid’-Cretaceous age, and was emplaced over a period of ~10 m.y. following the cessation of continental arc magmatism in the southwest at ~100 Ma. The termination of the continental arc magmatism might have been an effect of a stalled subducting plate (possibly due to an accretionary event), or a change in the geometry and orientation of the subducting plate (such as a shallowing to the angle of the down-going slab). A model invoking flat slab subduction for magmatism well inboard of the Cretaceous continental margin has already been proposed (e.g., Plafker and Berg, 1994; Newberry, 2000), but discounted several times since (e.g., Heffernan, 274 2004; Mair et al., 2006b; Johnston, 2008). Regardless, the post-arc ‘mid’-Cretaceous magmas of predominantly I- type affinity required high heat flow across a relatively large back-arc region. One way of generating high heat flow well inboard of a continental margin in the absence of a magmatic arc—as well as an inboard progression of the locus of magmatism similar to that observed across the Selwyn Basin (Fig. 4.1)—is effusive mantle convection (cf. Currie and Hyndman, 2006) related to ongoing dehydration above a shallowly subducting oceanic slab. There is other evidence for flat slab subduction at this time. For example, ~10 km of uplift was recorded between ~106–97 Ma in the western Selwyn Basin, following the emplacement of the Anvil batholith and prior to the eruption of the South Fork volcanic rocks (Pigage and Anderson, 1988; Smith and Erdmer, 1990; Murphy et al., 1995; Gordey, in press). Major changes to the source and locus of foreland basin sedimentation east of the northern Cordillera also occurred during this period. Prior to ~100 Ma, east- and north-sourced detritus was being deposited in the ‘mid’- Cretaceous inland sea (e.g., Yorath, 1991; Hadlari et al., 2009). A sea level drop occurring in the foreland basin at ~100 Ma has been attributed to uplift in the northern Cordillera and resulted in an influx of west-derived detritus to the Cretaceous inland sea (Yorath, 1991; Hadlari et al., 2009; Thomson et al., 2011). The foreland basin depocentre continued to shift eastward until about 93 Ma, then returned to a more westerly position by ~89 Ma (Thomson et al., 2011). Not only are the changes to foreland basin deposition related to uplift in the Selwyn Basin in the ‘mid’- Cretaceous, but they also coincide with changes in the locus and composition of syn- to post-accretionary plutonism in the Selwyn Basin. Cordilleran uplift, a shift in the source of detritus and the locus of deposition in the foreland basin, and the progression of high-temperature I-type (crustally derived) magmas across the Selwyn Basin are all compatible with the passage of a flat subducting plate below this region. Compressive stresses related to the flat subduction may have been largely accommodated by dextral transpressional (‘escape’?; Redfield et al., 2007) structures, and reversely activated rift-related transfer and listric detachment faults. Together, these structures in the Selwyn Basin may have minimized the thick-skinned exhumation of basement that typically occurs during flat slab subduction (e.g., Giambiagi and Ramos, 2002). The ~93 Ma shift to the foreland basin depocentre coincides with the emplacement of the most inboard (northerly and easterly) igneous rocks, the Tombstone plutonic suite, which mark the end to syn- to post-accretionary magmatism in the Selwyn Basin. The association between the Tombstone plutonic suite and abundant mafic and alkaline intrusive phases indicates that, at this time, significant melting in the mantle led to an influx of mantle- derived magma into the crust. It is possible that the flat slab subduction event ended with delamination (i.e., break- off or tearing) at the leading edge, which in turn would have triggered decompression-related upwelling and partial melting of the uppermost mantle (e.g., Cloos et al., 2005). Regardless of whether flat slab subduction was the driving force for ‘mid’-Cretaceous plutonism (and localized volcanism) across the Selwyn Basin, high heat flow below the Selwyn Basin was probably the most significant factor in the generation of I-type granitoids. The emplacement of mantle-derived magmas at the end of the ‘mid’- Cretaceous magmatic episode certainly attests to the importance of high heat flow in the lithospheric mantle, which ultimately generated the alkaline intrusive phases associated with the Tombstone plutonic suite. Syn-magmatic, pull- apart regions along dextral transpressional strike-slip faults could also have been important in the generation and localization of both Early and ‘mid’-Cretaceous magmas (e.g., Gabrielse et al., 2006; Colpron et al., 2011). The 275 emplacement of the Hyland plutonic suite, for example, occurs in an area that may have experienced uplift and decompression related to northwest-directed compression in the northwest Selwyn Basin (Murphy et al., 1995). Lithosphere-scale structures may also have been important in localizing this heat and the resulting mantle-derived magmatism, and would have aided the influx and ascension of mantle-derived magmas to upper crustal levels. 276 Figure 4.1. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous plutonic suites in the northern Cordillera based on the geochronological results of this study, compiled robust age determinations, and the work of Mortensen et al. (2000) and Hart et al. (2004a). The Selwyn Basin study area (Fig. 4.2) is outlined in black, and plutons mentioned in the text that are outside the study area are labeled. The red belts represent continental arc magmatism; the purple belts represent back-arc magmatism; and the blue, orange, and yellow belts represent post-arc magmatism. The plutonic suites are labeled: A = Anvil, C = Cassiar, DR = Dawson Range, FS = Fairbanks–Salcha, H = Hyland, M = Mayo, S = Seagull, Te = Teslin, To = Tombstone, TR = Tay River, Tu = Tungsten, W = Whitehorse, YTT = Yukon–Tanana Uplands, and TT = Tok–Tetlin. Intrusive and volcanic polygons are primarily from Gordey and Makepeace (1999) and Massey et al. (2005), and modified based on the regional maps of Dusel-Bacon et al. (1993), Newberry et al. (1998), Mortensen et al. (2000), Baker and Lang (2001), Murphy et al. (2001b), Hart et al. (2004a), Pigage (2004), Pigage et al. (2011), and Gordey (in press). Map lines and tectonostratigraphic elements are as in Figure 1.1 (Chapter 1). 277 Figure 4.2. Distribution of syn- to post-accretionary plutonic suites in the Selwyn Basin study area. Plutons mentioned in the text are labeled. The upper right inset shows samples of the Tungsten plutonic suite (southeast) in the Cantung mine area, and the upper left inset shows the extent of the Hess– MacMillan sub-suite (outlined with orange dashed line). Plutons belonging to the Tungsten plutonic suite are circled in red, and plutons belonging to the Coal River sub-suite (Tay River suite in the southeast Selwyn Basin) are circled in dark blue. Sample locations for this study (white circles) are labeled; sample locations for isotopic data compiled from Morris and Creaser (2008) and Heffernan (2004) in the Selwyn Basin are indicated by the black circles (samples from Heffernan, 2004, analyzed for stable isotopes by this study are also labeled). Map lines and tectonostratigraphic elements are as in Figure 1.1 (Chapter 1) ), and abbreviations are: Ft = fault, FS = Fault System, and H–M = Hess–MacMillan. 278 Figure 4.3. (next page) Summary of isotopic data for each plutonic suite in the Selwyn Basin (mafic dykes are indicated by green symbols); data generated by this study are indicated by the large symbols. a) Whole rock εNd100 and Sr100. Compiled neodymium and strontium data are indicated by small symbols (Wood and Armstrong, 1982; Pigage and Anderson, 1985; Patchett et al, 1998; Lang, 2001; Heffernan, 2004; Morris and Creaser, 2008; Gordey, in press). Strontium isotopic data from Lang (2001) were not recalculated to 100 Ma, but are presented as reported by their ‘initial’ (95–90 Ma) value. b) 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb in plutonic feldspar (measured ratios). Compiled lead data are indicated by small symbols (Heffernan, 2004). c) δ18O in plutonic quartz and whole rock δ34S. Compiled oxygen data (small symbols) for the Tungsten suite in the southeast Selwyn Basin (quartz; Bowman et al., 1985), the Tombstone suite (whole rock; Duncan, 1999), for the Anvil suite in the Yukon–Tanana inlier (quartz; Neufeld, 2004), and the Tay River suite in the western Selwyn Basin (whole rock; Gordey, in press). Whole rock analyses are shown with unfilled symbols with grey outlines. 279 280 Figure 4.4. (next page) Neodymium and strontium isotopic compositions of the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin at 100 Ma. Compiled isotopic data are from Wood and Armstrong (1982), Pigage and Anderson (1985), Patchett et al. (1998), Heffernan (2004), Morris and Creaser (2008), and Gordey (in press). Data for the Dawson Range continental magmatic arc samples are from Selby et al. (1999) and Bineli Betsi and Lentz (2011). The underlying enriched mantle (EM) in the Selwyn Basin is represented by the back-calculated isotopic compositions of a Recent basalt in the southern Selwyn Basin (Abraham et al., 2001). Bulk Earth composition (CHUR) relative to neodymium is represented by the line at εNd = 0. A granitic “migmatite” (Morris and Creaser, 2008) that may be representative of a minimum melt composition in the Cassiar Terrane is indicated by the bolded black plus sign for comparative purposes. Samples with anomalously elevated 147Sm/144Nd are outlined with the dashed blue line. a) εNd100 versus 87Sr/86Sr100. The possible compositional range of the Anvil plutonic suite is illustrated by the red bar based on a range of five 87Sr/86Sr100 values (Pigage and Anderson, 1985) and one εNd100 value (Patchett et al., 1998). b) εNd100 versus 147Sm/144Nd. c) εNd100 versus TDM (depleted mantle model age). 281 282 Figure 4.5. (next page) Lead isotopic compositions of the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. Symbols are as in Figure 4.4. Compiled isotopic data are from Heffernan (2004). Data for the Dawson Range continental magmatic arc samples are from Selby et al. (1999). The underlying enriched mantle (EM) in the Selwyn Basin is represented by the a Recent basalt sample from the southern Selwyn Basin (Abraham et al., 2001). The upper crustal Shale Curve of Godwin and Sinclair (1982) is represented by the heavy solid line, and two basement rocks sampled in proximity to plutons in the Selwyn Basin are represented by black “+” symbols (this study, unpub. data; Heffernan, 2004). Common Pb (204Pb) and fractionation error vectors are indicated on applicable plots by the thin black lines labeled “c” and “f”, respectively. Grey arrows illustrate trends towards more radiogenic lead compositions (U–Th rich, felsic, unmetamorphosed or late-metamorphosed, and (or) older). a) 207Pb/204Pb versus 206Pb/204Pb (“uranogenic” plot). b) 208Pb/204Pb versus 206Pb/204Pb (“thorogenic” plot). c) 207Pb/206Pb versus 208Pb/206Pb. 283 284 Figure 4.6. Stable isotopic compositions of the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. Symbols are as in Figure 4.4. The range of mantle compositions (δ18O = +5.7 ± 0.3‰ from Taylor, 1980; δ34S = 0 ± 3‰ from Chaussidon and Lorand, 1990; 20–50 ppm S; Ionov et al., 1992) is indicated by the grey shading. a) δ18O versus δ34S. The large grey and black crosses are whole rock data for two felsic orthogneisses of Archean and Neoproterozoic age, respectively (Malton Complex in southeast British Columbia; Vernon– Chamberlain, 1983). The maximum δ18O of a granitic intrusion (except for excessively crustally contaminated phases, such as plutonic margins and small intrusions) is indicated by the dashed line at δ18O = +14.5‰ (Taylor and McLennan, 1986). Samples plotting above the dashed line are from plutonic margins and (or) very small plutons; these heavier isotopic compositions are interpreted in the text to be the result of host rock contamination. b) Calculated minimum sulfur concentrations versus δ34S. Simplified vectors for variations in sulfur concentration and composition are indicted by grey arrows. The maximum concentration of sulfur in granitic magma at 2–3 kbar is indicated by the dashed line at 500 ppm sulfur (Wallace, 2001, 2005). 285 Figure 4.7. (next page) Strontium and stable isotopic comparisons and available basement compositions and ranges for the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. Symbols are as in Figure 4.4, and references for compiled data are as in Figure 4.6. Isotopic ranges for the Anvil suite are indicated by red box or red bar. Isotopic ranges for the Tay River suite in the western Selwyn Basin are indicated by the blue bar, and the South Fork point is based on an average of 14 strontium analyses. The ranges of enriched mantle isotopic compositions are indicated by the black bar (EM) using: δ18O = 5.7 ± 0.3‰ (Taylor, 1980); δ34S = 0 ± 3‰ (Chaussidon and Lorand, 1990); Sr100 = 0.703 (Abraham et al., 2001). The blue shaded area is the range of measured 87Sr/86Sr compositions for carbonate rocks in the Windermere Supergroup in the Mackenzie Mountains (northeast of the Selwyn Basin; Narbonne et al., 1994). The area between the two dotted lines represent the 87Sr/86Sr100 for the Dawson Range continental magmatic arc (Selby et al., 1999). Grey arrows highlight trends towards specific isotopic reservoirs. a) δ18O versus 87Sr/86Sr100. δ18O is from quartz for the plutonic rocks. The grey dashed line at δ18O = 14.5‰ is the approximate upper compositional limit for granitoid rocks globally, and the field for greywacke rocks is indicated by the shaded box (after Taylor and Sheppard, 1986; Rollinson, 1993). The range for amphibolite-grade metapelite rocks from the Monashee Mountains, southeast British Columbia, is indicated by the pale grey ellipse (whole rock; Bowman and Ghent, 1986; Kyser et al., 1994). The small black crosses are for Neoproterozoic to Archean, mafic to felsic orthogneisses and felsic paragneisses in southeast British Columbia (whole rock; the Malton Complex; Vernon–Chamberlain, 1983). The range of whole rock oxygen isotopic compositions in siliciclastic and calcareous Neoproterozoic–Triassic basement rocks in the Cassiar Terrane and Selwyn Basin are from Dagenais (1984) and Mair et al. (2006b). b) δ34S versus 87Sr/86Sr100. The range of sulfur compositions in basement rocks are for Mesoproterozoic evaporitic rocks (δ34S = +13.9 to +17.5‰; Gypsum Formation; Little Dal Group; Turner, 2009) and Paleozoic pyrites in siliciclastic rocks (δ34S ≈ –10 to +35‰; Goodfellow, 1987) in the Selwyn Basin, Cambrian pyrrhotites in siliciclastic rocks in the Cassiar Platform (δ34S = +16.7 and +20.9; Cooke and Godwin, 1984), and Neoproterozoic pyrites in siliciclastic rocks in southeast British Columbia (δ34S ≈ – 30 to +35‰; Ross et al., 1995). 286 287 Figure 4.8. Comparison of TDM versus εNd100 for northwest Laurentian basement and the Anvil, Hyland, and Tay River plutonic suites in the Selwyn Basin. A Recent alkaline basalt represents the composition of enriched mantle below the southern Selwyn Basin. References for compiled data are provided in footnotes to Figure 3.16 (Chapter 3). Isotopic ranges that largely overlap with the compositional range of granitoids in the Selwyn Basin are illustrated. a) Archean to Neoproterozoic basement. b) Paleozoic to Recent basement. 288 Figure 4.9. (next page) Neodymium and strontium isotopic compositions of the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin at 100 Ma. Mafic dykes associated with these suites are indicated by green symbols matching those in the legend. Compiled isotopic data are from Abercrombie (1990), Patchett et al. (1998), Lang (2001), Heffernan (2004), Morris and Creaser (2008), Mair et al. (2011), and Gordey (in press); 87Sr/86Sr from Lang (2001) are plotted initial ratios (t = 95–90 Ma). Data for the Dawson Range continental magmatic arc samples are from Selby et al. (1999) and Bineli Betsi and Lentz (2011). The underlying enriched mantle (EM) in the Selwyn Basin is represented by the back- calculated isotopic compositions of a Recent basalt in the southern Selwyn Basin (Abraham et al., 2001). Bulk Earth composition (CHUR) relative to neodymium is represented by the line at εNd = 0. A granitic “migmatite” (Morris and Creaser, 2008) that may be representative of a minimum melt composition in the Cassiar Terrane is indicated by the bolded black plus sign for comparative purposes. One Tungsten suite sample with anomalously elevated 147Sm/144Nd is circled with the dashed blue line. a) εNd100 versus 87Sr/86Sr100. b) εNd100 versus 147Sm/144Nd. c) εNd100 versus TDM (depleted mantle model age). 289 290 Figure 4.10. (next page) Lead isotopic compositions of the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. Symbols are as in Figure 4.9. Compiled isotopic data are from Heffernan (2004), Abercrombie (1990), and Mair et al. (2011). Mafic dykes associated with the these suites are indicated by green equivalents to the symbols in the legend. Data for the Dawson Range continental magmatic arc samples are from Selby et al. (1999). The underlying enriched mantle (EM) in the Selwyn Basin is represented by the a Recent basalt sample from the southern Selwyn Basin (Abraham et al., 2001). The upper crustal Shale Curve of Godwin and Sinclair (1982) is represented by the heavy solid line, and two basement rocks sampled in proximity to plutons in the Selwyn Basin are represented by black “+” symbols (this study, unpub. data; Heffernan, 2004). Common Pb (204Pb) and fractionation error vectors are indicated on applicable plots by the thin black lines labeled “c” and “f”, respectively. Grey arrows illustrate trends towards more radiogenic lead compositions (U–Th rich, felsic, unmetamorphosed or late-metamorphosed, and (or) older). a) 207Pb/204Pb versus 206Pb/204Pb (“uranogenic” plot). b) 208Pb/204Pb versus 206Pb/204Pb (“thorogenic” plot). c) 207Pb/206Pb versus 208Pb/206Pb. 291 292 Figure 4.11. Stable isotopic compositions of the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. Symbols are as in Figure 4.9. The range of mantle compositions (δ18O = +5.7 ± 0.3‰ from Taylor, 1980; δ34S = 0 ± 3‰ from Chaussidon and Lorand, 1990; 20–50 ppm S; Ionov et al., 1992) is indicated by the grey shading. a) δ18O versus δ34S. The large grey and black crosses are whole rock data for two felsic orthogneisses of Archean and Neoproterozoic age, respectively (Malton Complex in southeast British Columbia; Vernon– Chamberlain, 1983). The maximum δ18O of a granitic intrusion (except for excessively crustally contaminated phases, such as plutonic margins and small intrusions) is indicated by the dashed line at δ18O = +14.5‰ (Taylor and Sheppard, 1986). Samples plotting above the dashed line are from plutonic margins and (or) very small plutons; these heavier isotopic compositions are interpreted in the text to be the result of host rock contamination. b) δ34S versus calculated sulfur concentrations. Simplified vectors for variations in sulfur concentration and composition are indicted by grey arrows. The maximum concentration of sulfur in granitic magma at 2-3 kbar is indicated by the dashed line at 500 ppm sulfur (Wallace, 2001, 2005). 293 Figure 4.12. Strontium and stable isotopic comparisons and available basement compositions and ranges for the Tungsten, Mayo, and Tombstone plutonic suites in the Selwyn Basin. Symbols (and references for compiled plutonic data) are as in Figure 4.9. See caption to Figure 4.7 for non-plutonic data sources and a description of the figure. a) δ18O versus 87Sr/86Sr100. Compiled δ18O isotopic data compiled for the Mayo plutonic suite and associated mafic dykes in the northwest Selwyn Basin (Mair et al., 2006b) are whole rock data and will be ~2‰ lower than equivalent quartz data (e.g., Taylor and Sheppard, 1986). b) δ34S versus 87Sr/86Sr100. 294 Figure 4.13. Comparison of TDM versus εNd100 for northwest Laurentian basement and the Tungsten, Mayo, and Tombstone plutonic suites. A Recent alkaline basalt represents the composition of enriched mantle below the southern Selwyn Basin. References for compiled data are provided in footnotes to Figure 3.16 (Chapter 3). Isotopic ranges that largely overlap with the compositional range of granitoids in the Selwyn Basin are illustrated. a) Archean to Neoproterozoic basement. b) Paleozoic to Recent basement. 295 Table 4.1. Basic sample descriptions and crystallization ages for all samples that underwent radiogenic and (or) stable isotopic analysis. Plutonic Suite (sub- suite/region) Sample Pluton Lithology Mag. Susc. Age (Ma)a Anvil (YTT inlier) KR-07-13 Square Lake m-gr qtz-phyric Q>/=P>/=K bt-ms-tourmaline (leuco)granite 0.08 111.7 South Fork (west) 07M-151 volcanic welded dark grey pl-qtz-bt-(hbl) lapilli tuff w/aphanitic to microcrystalline groundmass 0.23 97.2 Tay River (west) 07M-150 Carolyn pl-hbl-qtz-phyric (chl-ep ± ms-cal) diorite-granodiorite porphyry 0.34 95.7 Tay River (east) KR-05-191 Marion (m)-c-gr mc-phyric (sparse) Q>/=K=P bt monzogranite 5.5 96.7 Tay River (east) KR-05-196 Coal River m-gr equigr K=P>Q bt>hbl monzogranite 9.01 98.4 Tay River (east) KR-05-22 Roy c-gr equigr to pl-phyric P>K>Q bt monzogranite 5.5 96.7 Tay River (east) KR-05-26 Fish f-m-gr pl-phyric P>Q=K bt monzogranite 9.01 98.4 Tay River (east) KR-05-157 Roy-mafic dyke v f-gr pl-qtz-phyric quartz diorite dyke 0.68 96.5 Tay River (Coal River) KR-05-43 Park f-m-gr equigr P>>Q>K hbl>bt granodiorite 16.4 96.1 Tay River (Coal River) KR-05-08 Jorgensen f-gr pl-qtz-hbl-phyric qtz monzodiorite porphyry 0.19 95.9 Tay River (Coal River) KR-05-10 Powers f-gr pl-(hbl)-phyric qtz monzodiorite porphyry 8.32 98.9 Tay River ('evolved') KR-05-164 Hole-in-the-Wall f-m-gr mc-crystic K>P>Q bt monzogranite 0.11 95.3 Tay River ('evolved') KR-05-32 Ivo f-gr equigr K>/=Q>P bt>ms grt leucogranite 0.06 96.0 Tungsten (southeast) KR-05-143 Circular Stock m-gr mc-phyric K>P>Q bt monzogranite 0.11 96.9 Tungsten (southeast) KR-05-148 Rifle Range c-gr mc-phyric (sparse) Q>/=K>P bt monzogranite 0.12 96.3 Tungsten (southeast) KR-05-215 Cantung f-m-gr mc-phyric (small and sparse) K=Q>P bt>>>ms monzogranite 0.11 nd Tungsten (southeast) KR-05-207 West Tuna nd 0.11 nd Tungsten (southeast) KR-05-208 Little Hyland strongly foliated (mc-augen) f-m-gr equigr Q=K>P bt>>ms monzogranite 0.11 96.8 Tungsten (southeast) KR-05-210 Nahanni Range m-c-gr equigr Q=K>P bt>>ms monzogranite 0.09 97.2 Tungsten (southeast) KR-05-212 Cantung-QFP dyke leucocratic qtz/feldspathoid-phyric monzonitic dyke 1.22 96.7 Tungsten (southeast) KR-05-213 Cantung-kersantite altered v f-gr bt-phyric kersantite (lamprophyric) dyke 0.62 nd Tungsten (northeast) KR-05-110 Cac m-gr mc-phyric K=P=Q bt monzogranite 0.15 97.4 Tungsten (northeast) KR-05-175 Lened f-m-gr mc-crystic (pl sub-phyric) K>/=Q=P bt monzogranite 0.14 97.6 Mayo (northeast) CL-06-35A Logan f-gr equigr P>K>Q bt>hbl>>cpx qtz monzodiorite 0.31 95.6 Mayo (northeast) CL-06-36A Mt. Christie (f)-m-gr mc-phyric (crowded) K>P>Q hbl>bt>>cpx qtz monzonite 1.6 96.8 Mayo (northeast) CL-06-37A Christie Pass f-m-gr mc-phyric (near-crowded) K>>P=Q hbl>>bt>>cpx qtz monzonite 0.96 96.7 Mayo (northeast) CL-06-39A Keele River f-gr mc-crystic hbl>bt >>cpx K>P>/=Q qtz monzonite 0.27 93.3 Mayo (northeast) CL-06-40A Mile 222 f-m-gr mc-phyric hbl>bt>>cpx K=P>Q qtz monzonite 0.49 96.6 Mayo (northeast) KR-05-113 South Nahanni m-c-gr mc-crystic (pl-qtz sub-phyric) K=Q=P bt>hbl>>cpx monzogranite 0.13 nd Mayo (northeast) KR-05-130 Central Nahanni m-gr equigr bt granodiorite 0.22 97.1 Mayo (northeast) KR-05-136 North Nahanni f-gr equigr P=K>Q bt>hbl monzogranite 0.19 96.9 Mayo (northeast) KR-05-62 Mount Christie m-gr mc-phyric K>P>Q hbl>>bt qtz monzonite 0.34 95.1 Mayo (Hess–MacMillan) KR-05-68 Mount Wilson (foliated) f-gr mc-crystic Q=K=P bt monzogranite 0.12 97.7 Mayo (Hess–MacMillan) KR-05-76 Pelly River foliated f-gr equigr Q=K=P bt qtz diorite 0.1 98.0 Mayo (Hess–MacMillan) KR-05-77 Pelly River- kersantite altered bt-phyric kersantite (lamprophyric) dyke 0.23 95.7 Tombstone (northeast) CL-06-34A Natla f-m-gr mc-crystic K>P>Q bt monzogranite 0.21 92.3 Tombstone (northeast) KR-05-97b O'Grady f-m-gr mc-phyric (crowded) K>>P>>Q hbl>cpx>bt tur-bearing qtz monzonite 4.06 92.8 Tombstone (northeast) CL-06-33A Dechen’La m-gr mc-crystic hbl-phyric K>P>/=Q hbl>bt>>cpx monzogranite 0.29 93.7 Tombstone (southeast) KR-05-194 Coal River m-gr mc-phyric K=P>/=Q bt monzogranite 0.14 92.8 Tombstone (southeast) KR-05-198 Hole-in-the-Wall c-gr mc-phyric (sparse) Q>>P>K bt monzogranite 0.18 91.3 Tombstone (Hess–MacMillan) CL-06-38A Ross River f-m-gr mc-phyric (scattered) Q=P>K bt (leuco)granite 0.15 93.0 Notes: a U–Pb (plutonic samples) and 40Ar/39Ar (mafic dyke samples) ages are reported in Table 2.2 (Chapter 2). 296 Table 4.2. Radiogenic and stable isotopic data generated by this study for magmatism in the Selwyn Basin. Neodymium, strontium and sulfur data are from whole rock powders, lead data are from feldspar separates, and oxygen data are from quartz separates. a) Neodymium isotopic compositions and relevant calculations. b) Strontium isotopic compositions and relevant calculations. c) Lead isotopic compositions. d) Oxygen and sulfur isotopic compositions, and calculated sulfur concentrations. a) Plutonic Suite (sub- suite/region) Sample 143Nd/ 144Nd ± 2-σ Sm (ppm) Nd (ppm) 147Sm/ 144Nd 143Nd/ 144Ndia εNdib εNd100 εNd80 εNd50 TDM (Ga)c Tay River (west) 07M-150 0.51183 6 4.6 24.9 0.1126 -14.78 -14.78 -14.76 -15.26 -16.01 1.99 DUPLICATE (Sm–Nd) 0.51182 8 4.5 24.3 0.1132 -15.00 -15.00 -14.97 -15.47 -16.23 2.02 Tay River (east) KR-05-191 0.51210 7 5.8 34.1 0.1021 -9.35 -9.35 -9.31 -9.81 -10.56 1.43 Tay River (east) KR-05-196 0.51209 7 6.4 42.7 0.0901 -9.35 -9.35 -9.33 -9.83 -10.59 1.31 DUPLICATE (Sm–Nd) 0.51208 6 6.5 45.5 0.0860 -9.50 -9.50 -9.47 -9.98 -10.73 1.28 Tay River (east) KR-05-22 0.51210 9 6.1 37.3 0.0994 -9.33 -9.33 -9.27 -9.78 -10.53 1.40 Tay River (east) KR-05-26 0.51207 11 7.4 36.9 0.1215 -10.11 -10.11 -10.06 -10.57 -11.32 1.78 Tay River (east) KR-05-157 0.51228 7 7.2 42.4 0.1031 -5.85 -5.85 -5.81 -6.31 -7.06 1.20 Tay River (Coal River) KR-05-43 0.51227 7 3.8 22.9 0.0996 -6.05 -6.05 -6.00 -6.50 -7.25 1.18 Tay River (Coal River) KR-05-08 0.51209 6 3.7 21.7 0.1040 -9.65 -9.65 -9.61 -10.11 -10.86 1.48 Tay River (Coal River) KR-05-10 0.51226 12 4.4 23.3 0.1141 -6.33 -6.33 -6.28 -6.78 -7.54 1.36 Tay River ('evolved') KR-05-164 0.51196 7 4.9 25.9 0.1136 -12.24 -12.24 -12.19 -12.69 -13.44 1.81 Tay River ('evolved') KR-05-32 0.51204 12 7.5 39.5 0.1147 -10.64 -10.64 -10.62 -11.12 -11.88 1.71 Tungsten (southeast) KR-05-143 0.51184 6 5.4 29.9 0.1090 -14.55 -14.55 -14.51 -15.01 -15.76 1.91 Tungsten (southeast) KR-05-148 0.51182 7 6.5 39.5 0.0994 -14.74 -14.74 -14.68 -15.18 -15.93 1.77 Tungsten (southeast) KR-05-215 0.51179 6 5.1 27.8 0.1115 -15.48 -15.48 -15.50 -16.00 -16.75 2.02 Tungsten (southeast) KR-05-207 0.51185 7 5.4 28.6 0.1142 -14.42 -14.42 -14.42 -14.92 -15.67 1.99 Tungsten (southeast) KR-05-208 0.51186 7 4.3 20.4 0.1262 -14.27 -14.27 -14.26 -14.76 -15.51 2.24 Tungsten (southeast) KR-05-210 0.51187 7 4.4 20.9 0.1262 -14.11 -14.11 -14.11 -14.63 -15.38 2.22 Tungsten (southeast) KR-05-212 0.51213 6 8.6 49.8 0.1042 -8.69 -8.69 -8.65 -9.15 -9.91 1.41 Tungsten (southeast) KR-05-213 0.51215 5 8.2 46.4 0.1063 -8.45 -8.45 -8.45 -8.95 -9.70 1.42 Tungsten (northeast) KR-05-110 0.51176 6 7.7 46.2 0.1011 -16.02 -16.02 -15.99 -16.49 -17.24 1.88 Tungsten (northeast) KR-05-175 0.51176 5 7.9 48.4 0.0989 -15.91 -15.91 -15.84 -16.35 -17.10 1.84 Mayo (northeast) CL–06-35A 0.51186 9 6.0 32.1 0.1132 -14.29 -14.29 -14.25 -14.76 -15.51 1.96 Mayo (northeast) CL–06-36A 0.51180 9 11.3 63.6 0.1073 -15.35 -15.35 -15.31 -15.81 -16.56 1.94 Mayo (northeast) CL–06-37A 0.51181 9 10.0 57.5 0.1054 -15.05 -15.05 -15.01 -15.51 -16.26 1.88 DUPLICATE (Sm–Nd, Rb–Sr, S) 0.51181 6 10.7 64.3 0.1006 -15.18 -15.18 -15.14 -15.64 -16.40 1.81 Mayo (northeast) CL–06-39A 0.51176 8 9.4 57.4 0.0989 -16.12 -16.12 -16.07 -16.57 -17.33 1.85 Mayo (northeast) CL–06-40A 0.51177 8 10.2 60.9 0.1012 -15.83 -15.83 -15.78 -16.28 -17.03 1.87 Mayo (northeast) KR-05-113 0.51177 8 7.4 44.8 0.1003 -15.73 -15.73 -15.73 -16.23 -16.98 1.85 Mayo (northeast) KR-05-130 0.51177 6 5.3 32.5 0.0984 -15.76 -15.76 -15.74 -16.24 -16.99 1.82 Mayo (northeast) KR-05-136 0.51186 5 5.1 32.0 0.0970 -13.98 -13.98 -13.97 -14.47 -15.22 1.68 Mayo (northeast) KR-05-62 0.51178 7 11.3 66.4 0.1030 -15.62 -15.62 -15.57 -16.07 -16.82 1.88 Mayo (Hess– MacMillan) KR-05-68 0.51150 6 6.6 39.9 0.1007 -20.95 -20.95 -20.92 -21.42 -22.17 2.21 Mayo (Hess– MacMillan) KR-05-76 0.51151 6 5.2 32.1 0.0983 -20.80 -20.80 -20.77 -21.27 -22.03 2.16 Mayo (Hess– MacMillan) KR-05-77 0.51161 7 9.9 57.9 0.1038 -18.96 -18.96 -18.91 -19.41 -20.16 2.13 Tombstone (northeast) CL–06-34A 0.51185 6 7.4 42.0 0.1072 -14.35 -14.35 -14.27 -14.77 -15.52 1.86 Tombstone (northeast) KR-05-97b 0.51189 6 10.3 61.5 0.1009 -13.42 -13.42 -13.33 -13.84 -14.59 1.70 Tombstone (northeast) CL–06-33A 0.51182 7 6.8 38.4 0.1073 -14.99 -14.99 -14.92 -15.42 -16.17 1.91 Tombstone (southeast) KR-05-194 0.51199 6 7.5 44.7 0.1008 -11.50 -11.50 -11.38 -11.88 -12.63 1.56 Tombstone (southeast) KR-05-198 0.51211 6 8.6 46.4 0.1122 -9.32 -9.32 -9.22 -9.73 -10.48 1.56 Tombstone (Hess– MacMillan) CL–06-38A 0.51158 5 6.1 32.8 0.1123 -19.75 -19.75 -19.67 -20.17 -20.92 2.36 Notes: a 143Nd/144Ndi (initial) was calculated using a 147Sm decay constant of 6.54x10-12/yr. b εNd values were calculated using 147Sm/144Ndchondrite = 0.512638 and 147Sm/144Ndchondrite = 0.1967 (Patchett et al., 2004). c TDM (depleted mantle) model ages were calculated using 147Sm/144NdDM = 0.51315 and 147Sm/144NdDM = 0.2137 (Michard et al., 1985). 297 b) Plutonic Suite (sub- suite/region) Sample 87Sr/ 86Sr ± 2-σ Rb (ppm) Sr (ppm) 87Rb/ 86Sr 87Sr/ 86Srid 87Sr/ 86Sr100 87Sr/ 86Sr80 87Sr/ 86Sr50 Tay River (west) 07M-150 0.7191 0.00001 105 299 1.016 0.7177 0.7177 0.7179 0.7184 Tay River (east) KR-05-191 0.7133 0.00001 167 306 1.583 0.7111 0.7111 0.7115 0.7122 Tay River (east) KR-05-196 0.7124 0.00001 158 360 1.271 0.7106 0.7106 0.7109 0.7115 Tay River (east) KR-05-22 0.7124 0.00001 146 337 1.256 0.7107 0.7106 0.7110 0.7115 Tay River (east) KR-05-26 0.7159 0.00001 185 164 3.274 0.7115 0.7112 0.7122 0.7136 Tay River (east) KR-05-157 0.7086 0.00001 107 687 0.451 0.7080 0.7079 0.7081 0.7083 Tay River (Coal River) KR-05-43 0.7093 0.00001 86 484 0.515 0.7086 0.7085 0.7087 0.7089 Tay River (Coal River) KR-05-08 0.7128 0.00001 107 459 0.678 0.7119 0.7119 0.7121 0.7123 Tay River (Coal River) KR-05-10 0.7089 0.00001 70 730 0.278 0.7085 0.7085 0.7086 0.7087 Tay River ('evolved') KR-05-164 0.7340 0.00001 266 112 6.885 0.7247 0.7242 0.7262 0.7291 Tay River ('evolved') KR-05-32 0.7222 0.00001 147 84 5.076 0.7151 0.7150 0.7164 0.7186 Tungsten (southeast) KR-05-143 0.7238 0.00001 164 219 2.180 0.7208 0.7207 0.7213 0.7222 Tungsten (southeast) KR-05-148 0.7210 0.00001 206 349 1.711 0.7187 0.7186 0.7190 0.7198 Tungsten (southeast) KR-05-215 0.7250 0.00001 191 212 2.611 0.7212 0.7213 0.7220 0.7231 Tungsten (southeast) KR-05-207 0.7241 0.00001 180 178 2.924 0.7200 0.7200 0.7208 0.7221 Tungsten (southeast) KR-05-208 0.7266 0.00001 248 128 5.600 0.7188 0.7187 0.7203 0.7227 Tungsten (southeast) KR-05-210 0.7280 0.00001 206 98 6.125 0.7193 0.7193 0.7211 0.7237 Tungsten (southeast) KR-05-212 0.7149 0.00001 97 652 0.431 0.7143 0.7143 0.7144 0.7146 Tungsten (southeast) KR-05-213 0.7128 0.00001 98 607 0.469 0.7122 0.7122 0.7123 0.7125 Tungsten (northeast) KR-05-110 0.7233 0.00001 173 311 1.609 0.7211 0.7211 0.7215 0.7222 Tungsten (northeast) KR-05-175 0.7234 0.00001 194 327 1.718 0.7211 0.7210 0.7215 0.7222 Mayo (northeast) CL–06-35A 0.7239 0.00001 141 329 1.246 0.7222 0.7221 0.7225 0.7230 Mayo (northeast) CL–06-36A 0.7244 0.00001 172 517 0.962 0.7230 0.7230 0.7233 0.7237 Mayo (northeast) CL–06-37A 0.7245 0.00001 304 467 1.886 0.7219 0.7218 0.7224 0.7232 DUPLICATE (Sm–Nd, Rb–Sr, S) 0.7245 0.00001 276 455 1.759 0.7221 0.7220 0.7225 0.7233 Mayo (northeast) CL–06-39A 0.7253 0.00001 245 438 1.621 0.7230 0.7230 0.7234 0.7241 Mayo (northeast) CL–06-40A 0.7249 0.00001 263 488 1.562 0.7228 0.7227 0.7231 0.7238 Mayo (northeast) KR-05-113 0.7232 0.00001 200 365 1.591 0.7210 0.7210 0.7214 0.7221 Mayo (northeast) KR-05-130 0.7227 0.00001 142 398 1.037 0.7212 0.7212 0.7215 0.7220 Mayo (northeast) KR-05-136 0.7249 0.00001 147 300 1.425 0.7229 0.7229 0.7233 0.7239 Mayo (northeast) KR-05-62 0.7240 0.00001 235 527 1.295 0.7223 0.7222 0.7226 0.7231 Mayo (Hess–MacMillan) KR-05-68 0.7350 0.00001 166 259 1.863 0.7324 0.7323 0.7328 0.7336 Mayo (Hess–MacMillan) KR-05-76 0.7348 0.00001 157 298 1.526 0.7326 0.7326 0.7330 0.7337 Mayo (Hess–MacMillan) KR-05-77 0.7233 0.00001 176 411 1.243 0.7216 0.7216 0.7219 0.7224 Tombstone (northeast) CL–06-34A 0.7258 0.00001 82 299 0.794 0.7247 0.7246 0.7249 0.7252 Tombstone (northeast) KR-05-97b 0.7214 0.00001 305 573 1.544 0.7194 0.7192 0.7197 0.7203 Tombstone (northeast) CL–06-33A 0.7236 0.00001 234 464 1.460 0.7217 0.7215 0.7219 0.7226 Tombstone (southeast) KR-05-194 0.7181 0.00001 211 187 3.257 0.7140 0.7135 0.7144 0.7158 Tombstone (southeast) KR-05-198 0.7209 0.00001 206 155 3.863 0.7159 0.7154 0.7165 0.7182 Tombstone (Hess– MacMillan) CL–06-38A 0.7342 0.00001 223 176 3.680 0.7294 0.7290 0.7300 0.7316 Notes (cont.): d 87Sr/86Sri (initial) was calculated using a 87Rb decay constant of 1.42x10-11/yr. 298 c) Plutonic Suite (sub-suite/region) Sample 206Pb/ 204Pb ± 2-σ 207Pb/ 204Pb ± 2-σ 208Pb/ 204Pb ± 2-σ 207Pb/ 206Pb 208Pb/ 206Pb Tay River (west) 07M-150 19.033 0.110 15.643 0.072 38.804 0.242 0.822 2.039 Tay River (east) KR-05-191 19.669 0.011 15.857 0.013 39.952 0.042 0.806 2.031 Tay River (east) KR-05-196 19.509 0.013 15.705 0.014 39.471 0.044 0.805 2.023 Tay River (east) KR-05-22 19.545 0.012 15.746 0.013 39.662 0.044 0.806 2.029 Tay River (east) KR-05-26 19.566 0.010 15.736 0.012 39.623 0.041 0.804 2.025 Tay River (Coal River) KR-05-43 19.220 0.014 15.661 0.014 39.187 0.045 0.815 2.039 Tay River ('evolved') KR-05-164 19.588 0.010 15.773 0.012 39.658 0.041 0.805 2.025 Tay River ('evolved') KR-05-32 19.459 0.010 15.698 0.012 39.458 0.041 0.807 2.028 DUPLICATE (Pb) 19.474 0.010 15.729 0.012 39.560 0.041 0.808 2.031 Tungsten (southeast) KR-05-143 19.461 0.010 15.732 0.012 39.514 0.041 0.808 2.030 Tungsten (southeast) KR-05-148 19.555 0.011 15.752 0.013 39.614 0.042 0.806 2.026 Tungsten (southeast) KR-05-215 19.636 0.038 15.841 0.030 40.025 0.081 0.807 2.038 Tungsten (southeast) KR-05-208 19.486 0.010 15.742 0.012 39.547 0.041 0.808 2.029 Tungsten (southeast) KR-05-210 19.484 0.010 15.745 0.012 39.536 0.041 0.808 2.029 Tungsten (northeast) KR-05-110 19.494 0.010 15.744 0.012 39.539 0.041 0.808 2.028 Tungsten (northeast) KR-05-175 19.463 0.010 15.752 0.012 39.575 0.041 0.809 2.033 Mayo (northeast) CL–06-35A 19.174 0.011 15.707 0.009 39.349 0.023 0.819 2.052 Mayo (northeast) CL–06-36A 19.265 0.019 15.801 0.015 39.431 0.044 0.820 2.047 Mayo (northeast) CL–06-37A 19.157 0.012 15.724 0.010 39.264 0.026 0.821 2.050 Mayo (northeast) CL–06-39A 19.155 0.002 15.688 0.001 39.207 0.003 0.819 2.047 Mayo (northeast) CL–06-40A 19.132 0.012 15.673 0.010 39.138 0.024 0.819 2.046 Mayo (northeast) KR-05-113 19.230 0.010 15.709 0.012 39.311 0.041 0.817 2.044 Mayo (northeast) KR-05-130 19.243 0.010 15.692 0.012 39.264 0.041 0.815 2.040 Mayo (northeast) KR-05-136 19.173 0.011 15.712 0.012 39.408 0.042 0.819 2.055 Mayo (northeast) KR-05-62 19.128 0.010 15.690 0.012 39.178 0.040 0.820 2.048 Mayo (Hess–MacMillan) KR-05-68 19.185 0.010 15.707 0.012 39.285 0.041 0.819 2.048 Mayo (Hess–MacMillan) KR-05-76 19.148 0.010 15.703 0.012 39.292 0.041 0.820 2.052 Mayo (Hess–MacMillan) KR-05-77 nd nd nd nd nd nd nd nd Tombstone (northeast) CL–06-34A 19.258 0.003 15.694 0.002 39.233 0.007 0.815 2.037 Tombstone (northeast) KR-05-97b 19.161 0.010 15.686 0.012 39.166 0.040 0.819 2.044 Tombstone (northeast) CL–06-33A 19.429 0.012 15.727 0.009 39.485 0.024 0.809 2.032 Tombstone (southeast) KR-05-194 19.613 0.021 15.771 0.016 39.746 0.060 0.804 2.026 Tombstone (southeast) KR-05-198 19.395 0.010 15.699 0.012 39.283 0.041 0.809 2.025 Tombstone (Hess–MacMillan) CL–06-38A 19.462 0.001 15.751 0.000 39.597 0.001 0.809 2.035 299 d) Plutonic Suite (sub-suite/region) Sample δ18OVSMOW (‰)e d34SVCDT (‰)f S (ppm)g Anvil (YTT inlier) KR-07-13 13.9 0.82 4 South Fork (west) 07M-151 13.8 10.24 31 Tay River (west) 07M-150 11.2 8.86 10 Tay River (east) KR-05-191 15.7 8.08 47 Tay River (east) KR-05-196 11.7 8.32 20 Tay River (east) KR-05-22 13.8 5.38 50 Tay River (east) KR-05-26 16.9 3.93 39 Tay River (east) KR-05-157 nd 11.13 1500 Tay River (Coal River) KR-05-43 13.2 3.94 51 Tay River (Coal River) KR-05-08 16.3 7.39 920 Tay River (Coal River) KR-05-10 14.6 4.23 120 Tay River ('evolved') KR-05-164 15.7 9.91 49 Tay River ('evolved') KR-05-32 16.0 4.59 150 Tungsten (southeast) KR-05-143 15.2 7.29 35 DUPLICATE (S) / 7.35 33 Tungsten (southeast) KR-05-148 14.3 8.68 110 Tungsten (southeast) KR-05-215 13.6 6.54 24 DUPLICATE (O) 14.1 / / Tungsten (southeast) KR-05-207 nd 9.14 29 Tungsten (southeast) KR-05-208 13.2 9.02 12 Tungsten (southeast) KR-05-210 13.1 9.75 7 Tungsten (southeast) KR-05-212 nd 8.62 5400 Tungsten (southeast) KR-05-213 nd 3.97 1000 DUPLICATE (S) / 4.56 nd Tungsten (northeast) KR-05-110 15.5 10.54 46 Tungsten (northeast) KR-05-175 13.7 9.47 26 DUPLICATE (O) 13.2 / / Mayo (northeast) CL–06-35A 13.6 12.57 28 Mayo (northeast) CL–06-36A 13.5 9.93 48 Mayo (northeast) CL–06-37A 14.3 10.00 240 DUPLICATE (Sm–Nd, Rb–Sr, S) / 9.91 360 Mayo (northeast) CL–06-39A 14.2 8.16 62 DUPLICATE (O) 15.0 / / Mayo (northeast) CL–06-40A 14.5 10.40 28 Mayo (northeast) KR-05-113 14.8 13.30 29 Mayo (northeast) KR-05-130 14.6 11.02 50 Mayo (northeast) KR-05-136 13.9 11.04 200 Mayo (northeast) KR-05-62 13.9 9.27 240 DUPLICATE (O) 14.1 / / Mayo (Hess–MacMillan) KR-05-68 15.1 7.42 nd DUPLICATE (S) / 7.85 240 Mayo (Hess–MacMillan) KR-05-76 nd 6.27 24 Mayo (Hess–MacMillan) KR-05-77 nd 7.77 250 Tombstone (northeast) CL–06-34A 14.5 11.24 26 Tombstone (northeast) KR-05-97b 14.5 8.40 61 Tombstone (northeast) CL–06-33A 13.0 8.79 33 DUPLICATE (O) 14.4 / / Tombstone (southeast) KR-05-194 13.9 6.73 14 Tombstone (southeast) KR-05-198 14.0 2.92 18 Tombstone (Hess–MacMillan) CL–06-38A 12.9 11.00 57 DUPLICATE (O) 12.9 11.00 57 Notes (cont.): e Reported relative to the Vienna Standard Mean Ocean Water (VSMOW; Coplen, 1994). f Reported relative to the Vienna Cañon Diablo troiolite (VCDT; Ding et al., 2001); underlined values are corrected for peak heights of <100 nA. g grams of contained sulfur were calculated from the weight of precipitated Ag2S used for sulfur isotopic analysis, converted to a percentage of the total sample weight, and re-calculated in parts per million; these are minimum estimates of sulfur content. 300 Table 4.3. Stable isotopic data generated by this study for samples from the eastern Selwyn Basin from Heffernan (2004). Oxygen data are from quartz separates and sulfur data are from whole rock powders. Plutonic Suite (sub- suite/region) Sample Pluton Lithology δ18Oqtz (‰)a δ34SVCDT (‰)b S (ppm)c Hyland SH-005 Mount Billings altered m-gr bt granodiorite 8.4 8.51 10 Hyland SH-011E Mount Billings massive f-gr bt granite 11.9 13.57 130 Hyland SH-029 Mount Billings m-gr bt granodiorite 10.3 6.12 14 Hyland SH-070 Logan altered f-gr bt granite 13.1 9.41 240 Hyland SH-99-022 Mount Billings massive m-gr bt granodiorite / 3.17 49 Tay River (east) 98-HAS-03 Faille megacrystic bt granodiorite 13.6 4.46 20 Tay River (east) 98-HAS-06 Mt. Sir James MacBrien megacrystic hbl-bt granodiorite 13.3 7.41 11 Tay River (east) SH-99-006 Coal River massive m-gr hbl-bt granodiorite 11.7 5.08 25 Tay River (east) SH-99-008 Coal River massive m-gr bt granodiorite 10.9 5.78 10 Tay River (east) SH-99-009 Coal River massive m-gr bt granodiorite 14.2 7.50 61 Tay River (east) SH-99-013 Ceasar Lakes massive f-gr bt granite 11.7 / 17 DUPLICATE (O) Ceasar Lakes massive f-gr bt granite 11.7 / 17 Tay River (Coal River) 98-HAS-07 Jorgensen hbl-Pl-bt porphyry (dioritic) 12.2 2.41 8 DUPLICATE (O) Jorgensen hbl-Pl-bt porphyry (dioritic) 11.7 2.41 8 Tay River (Coal River) 98-HAS-12 Patterson massive bt granodiorite 11.3 9.21 58 DUPLICATE (O,S) Patterson massive bt granodiorite 13.0 9.30 110 Tay River (Coal River) 98-Z-12 Powers Pl-hbl-bt porphyry (dioritic) 11.2 12.75 660 Tay River ('evolved') 98-HAS-02 Mount Appler massive m-gr bt granite 13.8 8.25 5 Tay River ('evolved') SH-99-001 Shannon Creek nd 12.4 4.16 7 Tungsten (northeast) 98-Z-C-028 Rudi megacrystic bt granite 13.5 7.27 6 Tungsten (southeast) SH-99-016 West Tuna megacrystic bt granite 12.7 4.46 27 Tungsten (southeast) 98-HAS-14 McLeod megacrystic bt granite 12.0 5.85 8 Tungsten (southeast) SH-99-011 Big Charlie megacrystic bt granite / 4.93 12 Notes: a Reported relative to the Vienna Standard Mean Ocean Water (VSMOW; Coplen, 1994). b Reported relative to the Vienna Cañon Diablo troiolite (VCDT; Ding et al., 2001); underlined values are corrected for peak heights of <100 nA. c grams of contained sulfur were calculated from the weight of precipitated Ag2S used for sulfur isotopic analysis, converted to a percentage of the total sample weight, and re-calculated in parts per million; these are minimum estimates of sulfur content. 301 5 CHAPTER: A RECONNAISSANCE-SCALE STUDY OF ZIRCON TRACE ELEMENT AND HAFNIUM ISOTOPIC COMPOSITIONS, EASTERN SELWYN BASIN 5.1 INTRODUCTION Our understanding of syn- to post-accretionary igneous rocks in the northern Cordillera has improved significantly over the last few decades with the advent of more precise, and increasingly affordable, geochronology. This, together with the integrated use of detailed trace element geochemical datasets and radiogenic and (less commonly) stable isotopic systems, has resulted in better constraints on the timing and composition of magmatism in the northern Cordillera. The use of multiple isotopic systems also provides additional information on magma petrogenesis related to sources of melt, as well as processes such as mixing, assimilation, and (or) contamination. The result has been a better understanding of the tectonic environment of magma genesis, particularly in the Early to ‘mid’- Cretaceous (~115–90 Ma), which was the most metallogenically important period for intrusion-related mineralization in the northern Cordillera (Hart et al., 2004a). This integration of several isotopic systems with geochemical and geochronological data (Chapters 2 and 4) has advanced our knowledge of the type of magmas and magma sources involved in syn- to post-accretionary magmatism in the Selwyn Basin. This in turn provided indirect information on the composition of the basement to the Selwyn Basin (Chapter 4). In this chapter, additional details (and questions) regarding the magma and magma source composition(s) are derived from a pilot study of trace element and hafnium isotopic compositions in zircon. Zircon is a remarkably useful mineral for isotopic dating and tracer studies of felsic to intermediate igneous rocks because it contains detectable quantities of uranium, thorium, and lead (used for U–Pb geochronology), as well as lutetium, hafnium, samarium, neodymium, and oxygen (e.g., Corfu et al., 2003; Hoskin and Schaltegger, 2003; Kinney and Maas, 2003). Other trace elements, including the heavy rare earth elements (REEs) and some high field strength elements (HFSEs; particularly Ti), are also concentrated in zircon (e.g., Kinney and Maas, 2003). Although zircon has been criticized for producing relatively uniform trace element compositions (e.g., Hoskin et al., 2000; Claiborne et al., 2010), the combination of trace element contents and ratios, and isotopic compositions (including the ability to accurately determine the age of different sectors of zircon grains) make this mineral analysis a powerful tool for the study of igneous rocks (e.g., Kinney and Maas, 2003). The implications of detailed zircon studies are many. Isotopic and (or) compositional studies of magmatic rims or antecrystic (magmatic) cores in zircon are commonly used to characterize granitoid composition and petrogenesis (Kemp et al., 2005, 2007; Belousova et al., 2006; Appleby et al., 2008; Peytcheva et al., 2008; Shaw and Flood, 2009; Siebel and Chen, 2009; Barth and Wooden, 2010). These studies are frequently done in tandem with in situ U–Pb geochronological analysis of individual zircon grains. Well planned in situ analysis preserve the grains for additional work, allowing precise U–Pb ages to be integrated with other compositional parameters, including hafnium and oxygen isotopic compositions, and trace element chemistry (Kinney and Maas, 2003; Barth and 302 Wooden, 2010). The ability to determine the age and composition of different sectors in a zircon provides important information on the timing of, and type of processes experienced by, the zircon grains—and the magma from which the zircon precipitated (e.g., Belousova et al., 2006; Claiborne et al., 2006, 2010; Appleby et al., 2008; Peytcheva et al., 2008). In part due to its ability to survive magmatic, metamorphic, and erosional processes, zircons from igneous, metamorphic, and sedimentary rocks can also provide a record of multiple crystallization and resorption episodes (Corfu et al., 2003). Additional information on the composition, provenance, and (or) age of the magma protolith(s) may also be obtained when inherited zircon cores are systematically analyzed (e.g., Pupin, 2000; Belousova et al., 2002; Grimes et al., 2007; Barth and Wooden, 2010). Inherited zircon in magmas derived from a relatively heterogeneous melt source, such as a package of lithologically variable metasedimentary rocks sourced from multiple terranes, will display wide variations in age and composition. Inherited zircon in magmas derived from a single homogeneous melt source should display minimal variation in age and composition (Belousova et al., 2006). Zircon morphologies and the evolution of various crystal forms, visible as internal magmatic zoning, have been linked to: the alkalinity and temperature of the magma (Pupin, 1980), trace element and isotopic variations providing information on the magma petrogenesis (e.g., Hanchar and Miller, 1993), and used as evidence for magmatic mixing and fractionation processes (e.g., Belousova et al., 2006). Finally, the amount of titanium in the zircon structure (which is present in high enough quantities to analyze on the electron-probe microanalyzer, or EPM; Watson and Harrison, 2005), is a promising and simple magmatic thermometer (Watson et al., 2006; Ferry and Watson, 2007), although some fine-tuning of this method is still needed (Fu et al., 2008; Claiborne et al., 2010). Here, reconnaissance-scale trace element and Lu–Hf compositional data for magmatic (and inherited) zircons from representative intrusions in the eastern Selwyn Basin are reported and integrated with isotopic data (for the same samples) that was presented in Chapter 450. Together, these data strengthen the conclusions reached in Chapter 4, but also add some new petrogenetic information. Should the initial results of this study justify additional work, however, these zircon analyses will help to guide future studies of syn- to post-accretionary igneous rocks in the northern Cordillera, particularly in the eastern Selwyn Basin. 5.2 ANALYTICAL METHODS Trace element concentrations and Lu–Hf isotopic compositions in zircon were determined for nine samples representing all of the plutonic suites and sub-suites identified in the eastern Selwyn Basin in Chapter 2 (Fig. 5.1). This includes the Tay River and Tungsten suites, and the Coal River sub-suite of the Tay River suite in the southeast Selwyn Basin (Figs. 5.1 and 5.2). Plutonic suites sampled in the northeast Selwyn Basin include the Tombstone, Mayo, and Tungsten plutonic suites—collectively referred to as the Tombstone–Tungsten Belt (TTB), and the Hess– MacMillan sub-suite comprising samples of both Mayo and Tombstone age (Figs 5.1 and 5.2). Basic sample descriptions, zircon textural observations, and previously reported geochemical and isotopic data are compiled in Table 5.1. Samples of epoxy mounted zircons (initially prepared for Laser Ablation Inductively Coupled Plasma Mass Spectrometry, or LA–ICP–MS, an in situ method of U–Pb geochronological analysis; Appendix B.1; Chapter 50 This study did not determine the oxygen isotopic composition of zircon grains; instead, oxygen compositions are reported from Chapter 4 for quartz separates from the same samples. 303 2) and back-scattered electron (BSE) and cathodoluminescent (CL) images collected from the scanning electron microscope (SEM) were used to characterize zircon morphologies (after Pupin, 1980) and select spots for trace element and Lu–Hf isotopic analysis (Appendix E.1). Hafnium concentrations―used as an internal standard for the trace element analyses―were determined prior to LA–ICP–MS analysis using the EMP at the University of British Columbia (UBC; Table 5.2).51 Selected zircons from the representative samples were then analyzed by LA–ICP–MS in the INCO Innovation Centre Microanalytical Facility at the Memorial University of Newfoundland (MUN). The Finnigan ELEMENT XR―a high resolution double focusing magnetic sector inductively coupled plasma mass spectrometer (HR-ICPMS) coupled to a GEOLAS 193 nm excimer LASER system―was used for the analyses. The LA–ICP–MS method has the advantage of allowing in situ analysis, making it possible to visually characterize and then select specific domains in zircon grains for hafnium isotopic composition. LA–ICP–MS is a highly accurate analytical method, due largely to the volume of material that is ablated (e.g., Kinney and Maas, 2003). One corollary to this, however, is that the relatively large spot sizes of 40–50 μm required for trace element and hafnium laser-ablation analysis is at least an order of magnitude larger than the scale of typical oscillatory zoning in magmatic zircons. However, oscillatory zoning reflects specific enrichments and depletions in trace elements that are controlled by kinetic effects (e.g., Mattinson et al., 1996) rather than magmatic processes (i.e., mixing). When averaged, bright and dark oscillatory zones together may be interpreted as representative of the bulk zircon composition for the entire oscillatory-zoned sector (Clairborne et al., 2006, 2010). While the large spot size during laser ablation analysis is, therefore, an advantage in determining the bulk composition of zircon sectors, it can be larger than many inherited and magmatic grain domains (particularly in grains that are smaller than ~200 μm and (or) grains that have large inherited cores with thin magmatic rims). Another limit to LA–ICP–MS analysis is that follow-up work on the exact same spot will be impossible because of the removal of a large amount of material during ablation. Therefore, if U–Pb and oxygen, and trace element information for the same zircon spot is desired, these must be carried out prior to hafnium analysis using a less-destructive technique (e.g., SHRIMP). In this study, however, budgetary and time constraints did not allow the prior analysis of hafnium spots by less-destructive U–Pb and trace element techniques. Laser ablation U–Pb geochronological was done first using line scans, which form long, narrow, shallow ablation pits that were easily removed with a second polishing of the grain mount. The hafnium and trace element spot analyses were done after the U–Pb work on re-polished zircons; however, no attempt was made to analyze the same portions of the same grains. It was impossible to do so because of the difference in grain area and depth sampled by line and spot scans. 5.2.1 Trace Element Analysis The trace element analytical methods are summarized here from M. Tubbrett (2012, pers. comm.). A helium flow 51 Electron-probe micro-analyses of zircon were done on a fully automated CAMECA SX–50 instrument, operating in the wavelength-dispersion mode with the following operating conditions: excitation voltage, 15 kV; beam current, 80 nA; peak count time, 60 s; background count-time, 30 s; spot diameter, 5 μm. Data reduction was done using the 'PAP' φ(ρZ) method (Pouchou & Pichoir 1985). Hafnium was analyzed in trace-element mode. Nominal amounts of zirconium, silica, and oxygen in zircon were fixed for analysis. For hafnium, the following standard, X-ray line and crystal were used: Hf element, Hf Lα, LIF. The detection limit for hafnium is ~0.1 wt.%, and precision is 6–8% relative at 95% confidence. 304 rate of 1.25 L/min was used to carry ablated material to the inductively coupled plasma, with additional argon make- up gas added after the ablation cell. A spot size of 40 μm was used with energy density of approximately 3 J/cm2 with a laser repetition rate of 10 Hz. Time resolved intensity data were acquired by peak-jumping in a combination of pulse-counting, analog and Faraday modes, depending on signal strength, with one point measured per peak. Concentrations were calibrated with the NIST 612 glasses. Hafnium concentrations (pre-determined by EPM analysis at UBC) were used as an internal standard to deal with differences in ablation yields and matrix effects between the zircon and the calibration material (NIST glass). Approximately 30 s of gas background data were collected prior to each 60 s ablation of both standards and unknowns. The data acquisition methodology employed an analytical sequence of two analyses of the NIST 612 standard glass (the external calibration standard) and one of zircon 91500 (treated as an unknown), followed by analyses of up to fourteen unknowns. The sequence closed with a repetition of the same standards in reverse order. Six replicate analyses were also done to monitor precision. The concentrations of twenty-seven elements are reported in Table 5.2, although not all of these are presented and discussed below. The error for this method, when measuring homogeneous materials, is estimated to be better than 4–7% (relative), based on the reproducibility of results for various reference materials measured from day to day over several months in the MUN laboratory. High relative standard deviation values for some elements in Table 5.2 (see RSDs in Appendix E.2) can be attributed to heterogeneities in the 91500 zircon standard for these elements, especially at low concentrations and for elements prone to fractionation during analysis. Data were reduced using MUN’s in-house “CONVERT” and “LAMTRACE” spreadsheet programs, which employ procedures described by Longerich et al. (1996). “LAMTRACE” allows the selection of representative signal intervals and background subtraction, as well as internal standard correction for ablation yield differences and instrument sensitivity drift during the analytical session. “LAMTRACE” also performs calculations converting count rates into concentrations by reference to the standards. 5.2.2 Lu–Hf Isotopic Analysis and Data Reduction The general method of hafnium isotopic analysis by laser ablation multicollector inductively coupled plasma mass spectrometry (LA-MC-ICP-MS) performed at MUN is described in detail by Fisher et al. (2011) and is summarized here from R. Lam (2012, pers. comm.). A helium flow rate of 1.25 L/min was used to carry ablated material to the inductively coupled plasma, with an additional argon make up gas added via T-piece, after the ablation cell and just before the torch (other analytical conditions included: auxiliary argon = 0.75 L /min, sample argon = 0.75 L /min, helium carrier gas = 1.3 L /min). A spot size of 49 μm was used with energy density of approximately 5 J/cm2 with a laser repetition rate of 10 Hz. Approximately 30 s of gas background data were collected prior to each 60 s ablation of both standards and unknowns; an on-peak integration time of ~1 s led to 50–60 isotopic ratios for each analysis. Each sample set was bracketed by several analyses of the standard zircons Plešovice (PL) and FC-1 (Appendix E.3). A total of eighteen PL analyses produced an average mass bias and interference corrected 176Hf/177Hf value of 0.282477 ± 0.000036 (2-σ), which is within the 95% confidence interval for the accepted value of 0.282482 ± 0.000013 (Slama et al., 2008). A total of eleven FC-1 analyses produced an average mass bias and interference corrected 176Hf/177Hf value of 0.282186 ± 0.000062 (2-σ), which agrees closely to the accepted value of 0.282184 ± 305 0.000016 (Woodhead et al., 2005). Before proceeding with the hafnium isotopic calculations, all raw signals were blank corrected using the mean of ~20–25 pts recorded as the gas blank baseline before ablation. A user-selected interval of about 50 data pts (~50 s of ablation) covering the sample transient peak (~60 s of ablation) was used for calculation of the hafnium ratio with mass bias correction by exponential law. The blank-corrected signal intensities were then also corrected for the isobaric interferences of ytterbium and lutetium on 176Hf, as described in detail by Fisher et al. (2011) using the invariant ratios of ytterbium from Segal et al. (2003) for the mass interference corrections. The mass bias and interference-corrected isotopic ratios were generated using “MCTool-Hf”, which uses macros from “LAMTOOL- Hf” (J. Kosler, University of Bergen), following the protocol described in Appendix E.4. Outlying 176Hf/177Hf values were eliminated based on 2-σ criteria, and the calculation of uncertainties is based on 2-σ error of the mean. The chondritic values of Bouvier et al. (2008) for 176Hf/177Hf0 (0.282785) and 176Lu/177Hf (0.0336), and the 176Lu decay constant of 1.867 x 10-11/yr from Scherer et al. (2001) were adopted for the calculation of initial 176Hf/177Hf and εHf100 compositions (Table 5.3). Depleted mantle model ages (Table 5.3) were calculated using the same 176Hf/177Hf0 (0.28325) and 176Lu/177Hf (0.0384) values also used by Griffin et al. (2000). 5.3 DISCUSSION OF RESULTS 5.3.1 Texture and Morphology of Zircons from Representative Samples Although the external morphology of zircons can readily be documented using a binocular microscope, the complexity of internal textures is best examined using BSE and CL methods (e.g., Corfu et al., 2003; Hoskin and Schaltegger, 2003). BSE images highlight differences in the average atomic number of zircon zones and sectors, with brighter BSE resulting from higher concentrations of ‘heavier’ elements. Hafnium, uranium, then thorium are typically the main controls on BSE brightness in zircon (Hanchar and Miller, 1993; Gagnevin et al., 2010). Zones and sectors that are bright in BSE images are usually dark in CL images, and vice versa (e.g., Hanchar and Miller, 1993; Corfu et al., 2003), even though the CL spectra are controlled by different trace element concentrations. CL brightness mainly reflects the abundance of trivalent REEs. In zircon, bright CL reflectance is thought to be mainly related to concentrations of Dy3+ although Sm3+, Eu2+, Tb3+, and Y3+ may also cause luminescence (Hanchar and Miller, 1993; Corfu et al., 2003). Traditionally, CL images are favoured for zircon studies because of the greater range in intensity of emission spectra (including the presence of colour zoning when zircons are imaged using a CL-equipped petrographic microscope; Corfu et al., 2003). BSE images, however, may be more appropriate for compositional and isotopic studies of zircon, particularly using destructive analytical methods. This is because BSE brightness is the result of trace element variations that are more important with respect to zircon (and magma) petrogenesis (e.g., Hf, U, Th; Gagnevin et al., 2010) than the trivalent REE. It was found in this study that that CL imaging did highlight magmatic zoning, resorption surfaces, and other internal textures (e.g., patchy, sector, and convolute zoning) better than BSE imaging. 306 The CL spectra, however, were unable to image most inclusions, porous textures, fractures, and other physical features that were obvious in the BSE images, and the CL spectra also produced poorer (grainer) images than the BSE spectra. Micron-scale zoning in magmatic zircon cores and rims oscillates between BSE-bright and BSE-dark (e.g., Corfu et al., 2003). Oscillatory zoning was observed in all magmatic grains examined here, and was used to characterize internal zircon morphology (after Pupin, 1980; Appendix E.1). This type of zoning is controlled by the particular stage of zircon growth, the composition of the crystal–liquid interface, the degree of supersaturation of trace elements and zirconium in the melt, the rate of trace element diffusion in the melt, and the oxidation state of the magma (Mattinson et al., 1996; Claiborne et al., 2006, 2010). Changes to magma composition and temperature, however, reduce the saturation of zircon in magmas (e.g., Watson and Harrison, 1983), which may result in smooth or rounded resorption surfaces that cross earlier magmatic zoning (Corfu et al., 2003). These resorption textures typically occur when a hotter and (or) less siliceous magma (cf. Harrison and Watson, 1983) intrudes a cooler magma chamber and temporarily increases zircon solubility. Resorption surfaces can also result from changes to local kinetic phenomena within a relatively stable, closed magmatic system (Corfu et al., 2003). Major resorption surfaces were identified in this study by the termination of oscillatory magmatic zoning or radiating fractures against a smooth rounded surface in magmatic zircons for all of the representative samples examined—with the exception of the Tay River suite sample KR-05-196, which lacked resorption surfaces (Table 5.1). In some cases, more than one resorption surface was observed within the same zircon grain, defining several major boundaries between inner, intermediate, and outer magmatic zircon sectors. Changes in reflectance at resorption surfaces were also commonly observed. Changes to oscillatory zoning patterns (i.e., from coarsely or poorly zoned to finely or well zoned) and zircon morphologies were less commonly observed, suggesting that resorption did not accompany major changes in magma composition or temperature (and instead is likely the result of local kinetic phenomena). The outermost sector of magmatic zircon will be referred to herein as the magmatic rim, and inner magmatic sectors (that may or may not be separated from the magmatic rim by a resorption surface) are referred to as magmatic cores if inherited cores are not present), and as intermediate magmatic ‘rims’ if inherited cores (or two or more resorption surfaces) are present. Magmatic cores observed in this study may exhibit oscillatory zoning, but are more commonly unzoned to poorly zoned and of moderate to low BSE-brightness. Magmatic cores were usually distinguished from inherited cores in that they are almost always ‘pristine’ or lacking textures that are indicative of alteration, dissolution, or metamictization. Inherited zircon cores are commonly BSE-bright (and CL-dark), suggesting a uranium enrichment that could be responsible for expansion and the formation of radial to concentric fractures, from the core outwards (e.g., Corfu et al., 2003). This fracture pattern is particularly common in zircons from the Tungsten suites samples in the northeast Selwyn Basin (Appendices B.3 and E.1). Distinctions between inherited and magmatic cores were occasionally U–Pb confirmed by geochronological analysis, although in most cases magmatic cores were chronologically indistinguishable from magmatic rims (making them ‘autocrystic’ in the sense of Miller et al., 2007), particularly in smaller grains with unzoned magmatic cores of moderate to low reflectance. In some cases, 307 however, some magmatic cores in larger grains were found to be several m.y. older than the magmatic rims for many of the plutonic suites (see AC age; Table 5.1), making them ‘antecrystic’ in the sense of Miller et al. (2007). The Hess–MacMillan sub-suite, however, contains the most old inherited cores of any of the representative samples examined here (about 40 and 70% of zircons grains contain inherited cores in the two samples examined; Appendices B.3 and E.1), which is compatible with its interpreted derivation from supracrustal rocks in Chapter 2. Inherited zircon cores are noticeably less abundant (i.e., observed in <15% of zircon grains examined) in samples of the Tombstone, Tay River (including the Coal River sub-suite), and Mayo plutonic suites. A morphological classification for zircon that has been broadly related to magma alkalinity and temperature (although kinetic or other compositional factors may have a strong control zircon morphology; Vavra, 1993; Corfu et al., 2003) has been developed and widely applied using natural examples of zircons from granitoids with known petrogenetic histories (Pupin, 1980). The typology classification focuses on the relative dominance of the {211} versus {101} crystal pyramids, which are linked to magma alkalinity, and the {110} versus {100} crystal prisms, which are linked to magma temperature (Fig. 5.3a). Zircons from peralkaline to alkaline (i.e., mantle-derived) magmas typically display a range of prismatic forms, depending on the temperature of crystallization, but will be dominated by the {101} pyramid (Fig. 5.3b). Conversely, zircons from peraluminous, cordierite-bearing (and muscovite-bearing) examples of supracrustally derived granites are dominated by the {110} prism and the {211} pyramid (Fig. 5.3b). Zircons from calc-alkaline (mixtures of mantle and crust) and sub-alkaline (I-type) granitoids are typically transitional in morphology (and thus temperature and composition) to the supracrustal and mantle end- member typologies (Fig. 5.3b). The changes in magma composition to higher sodium and potassium contents and a decrease in temperature as crystallization progresses causes zircon morphologies from all magma types to converge (Fig. 5.3b). In this way, late-magmatic zircons ultimately trend towards the dominant {101} pyramid and {110} prism forms, although magma mixing, assimilation, or even contamination may impose a different trend (e.g., Belousova et al., 2006). Weaknesses in this generalized typology have been pointed out (e.g., Hoskin and Schaltegger, 2003), but this classification still provides a simple method of describing zircon morphology (e.g., Belousova et al., 2006) and relating it to real, if overly simple, magmatic parameters. Figure 5.4 shows the distribution of observed internal and external morphologies of zircon (magmatic cores and rims) for each plutonic suite (Fig. 5.3a). Based on the charts in Figure 5.4, inferences can be made regarding the alkalinity and temperature of the magma (melt) from which zircon crystallized for each of the plutonic suites. A significant proportion of the Tay River, Mayo, and Tombstone plutonic suites (excluding sub-suites) have zircons morphologies that are typical of high-temperature (ca. 850 ± 100°C), “sub-alkaline (I-type)” to “calc-alkaline” magmas (Figs. 5.3b and 5.4). All three of these suites are interpreted to have been derived predominantly (if not entirely) from infracrustal source rocks in Chapter 2, which is consistent with the higher temperatures suggested by zircon morphologies in these samples. The continuum of zircon morphologies between the higher temperature population and a smaller lower temperature (ca. 650 ± 100°C) and slightly more alkaline population (Fig. 5.4) in these three infracrustal plutonic suites may be a consequence of normal magmatic cooling. Some scatter towards aluminous compositions in the lower temperature population (particularly in the case of the Tay River suite; Fig. 5.4) is suggestive of contamination by, or the assimilation of supracrustal material. The actual direction of zircon 308 morphological evolution in these samples is uncertain and requires a more detailed evaluation that is beyond the original scope of this study. It can be inferred, however, that the zircons in these three dominantly I-type plutonic suites evolved from morphologies typical of higher temperature conditions to morphologies typical of lower temperature conditions. In comparison, the Tungsten plutonic suite and the Hess–MacMillan sub-suite are characterized by mainly zircon morphologies that are more indicative of lower magmatic temperatures (ca. 650 ± 100°C; Fig 5.4). The Tungsten suite samples from the northeast Selwyn Basin, in particular, define a relatively well-constrained population of lower temperature zircon morphologies that appear to trend towards higher alkalinity as temperature decreases (Fig. 5.4). This pattern is very similar to that exhibited by zircons from peraluminous (cordierite- and muscovite-bearing) granites that were interpreted to be derived from partial melting of sedimentary rocks (“S-type”; Fig. 5.3b; Pupin, 1980). Zircon morphologies from northeast members of the Tungsten plutonic suite, therefore, are consistent with a supracrustal origin, although similar trends could result from lower-temperature partial melting. Conversely, zircons from the Tungsten suite in the southeast Selwyn Basin, and from the Hess–MacMillan sub-suite in the northeast Selwyn Basin, display greater morphological variation in Figure 5.4. A small population of high temperature zircon morphologies are present in both of these suites (ca. 850 ± 100°C; Fig. 5.4), and could suggest a role for higher magmatic temperatures in the genesis of these intrusions (cf. Belousova et al., 2006). Unfortunately a distinction was not made as to whether the morphologies indicative of higher magmatic temperatures are from ~103 Ma antecrystic cores, autocrystic cores, or magmatic rims. There is a disconnect on the typology chart between morphologies that are consistent with hotter (and possibly more alkaline) magma(s) and morphologies that define a larger, lower temperature population for zircon in the Hess–MacMillan sub-suite. Furthermore, both populations (high and low temperature) of zircon morphologies in Hess–MacMillan sub-suite also exhibit a range of pyramidal faces, from those typical of aluminous to alkali-rich magmas (Fig. 5.4). These multiple zircon populations are consistent with the involvement of two or more distinct magmas in the genesis of the Hess–MacMillan sub-suite (e.g., Belousova et al., 2006). Overall, the most of the zircon growth in the Tungsten suite and Hess–MacMillan sub- suite appears to have been relatively low temperatures, which is consistent with an origin by small-fraction partial melting (as was suggested in Chapter 4 for the Tungsten plutonic suite) and (or) by partial melting of micaceous supracrustal rocks (as was suggested in Chapter 4 for the Hess–MacMillan sub-suite). The Coal River sub-suite of the Tay River plutonic suite is an interesting case. The morphological evolutionary pathway (from magmatic core to rim) illustrated by the arrows in Figure 5.4 was apparent in many of the zircon grains. The overall trend from higher temperature zircon morphology to low temperature (late-stage magmatic) morphology with dominant {101} pyramids and {110} prisms is typical of calc-alkaline magmas (Fig. 5.3b). However, magmatic zircon rims (and (or) ~60% of the observed zircon morphologies in this sample) are dominated by {211}-dominant pyramids and {110} prisms (Fig. 5.4). This suggests that the majority of the zircon, and the latest-crystallizing zircon, formed in a magma that was much more aluminous than the earlier magmas (Fig. 5.3b), which is also consistent with the peraluminous whole rock compositions exhibited by all members of this plutonic suite (Chapter 2). Based on these observations, the Coal River suite magmas appear to have been heavily 309 incorporated aluminous material, despite other evidence indicating that this sub-suite incorporated a significant component of magma derived from mafic (probably Paleozoic; Chapter 4) metaigneous rocks.52 Unlike the Coal River sub-suite, most representative samples of the other plutonic suites do not contain zircons that have more than one or two morphological domains, even if several different morphologies are observed in difference zircons within the sample (Appendix E.1). For cases in which more than one morphological domain is present, there is rarely a simple relationship between the observed morphology and the zircon sector (i.e., antecrystic or autocrystic cores, intermediate zones, or magmatic rims), making reconstructions of the morphological evolution from early to late crystallizing zircon for the pluton difficult. The lack of consistent patterns is probably mostly strongly related to different temperatures of zircon saturation and growth in the cooling magmas (i.e., some zircons may be only late-stage magmatic grains, whereas others may have nucleated early and grown continuously or sporadically in the slowly cooling magma chamber). This is also complicated by abundant resorption surfaces in zircons from nearly all of the representative samples examined here (Table 5.1). 5.3.2 Characterization of Zircon Trace Element Compositions When normalized to the chondritic composition of Sun and McDonough (1995), zircons from representative samples of all the plutonic suites in the eastern Selwyn Basin display a gradual enrichment in REEs with decreasing ionic radius, a pronounced positive cerium anomaly, and a small to pronounced negative europium anomaly (Fig. 5.5). These are typical patterns as the range of REE compositions exhibited by terrestrial zircons is limited (although the lowest REE contents and smallest europium anomalies are typical of zircons in mantle-derived rocks; e.g., Hoskin and Schaltegger, 2003; Kinney and Maas, 2003). Despite the lack of major compositional variation in zircon, there are commonly small but significant differences between zircons from igneous rocks that have experienced a range of petrogenetic histories (e.g., Belousova et al., 2002, 2006; Hoskin and Schaltegger, 2003; Barth and Wooden, 2010; Clairborne et al., 2010; Castiñeiras et al., 2011). The Coal River sub-suite of the Tay River plutonic suite probably comprises the most distinct and compositionally homogeneous zircon populations of the samples examined in Figure 5.5a. This sub-suite has the largest chondrite- normalized cerium anomalies (average Ce/Ce* = 83) and the smallest europium anomalies (average Eu/Eu* = 0.51; Figs. 5.5a and 5.6a). These features are consistent with the relatively undifferentiated compositions of most members of this sub-suite in general, and of this sample in particular (Table 5.1).53 The magnitude of these REE anomalies, as well as an overall heavy REE enrichment, also distinguish the Coal River sub-suite from the Tay River plutonic suite (Fig. 5.5a). Two of the magmatic cores analyzed from the Coal River sub-suite have deeper europium anomalies and are more enriched in light to medium REEs than the magmatic rims (Fig. 5.5b). Inherited 52 In Chapter 2, the elevated peraluminicity exhibited by these intrusions was suggested to be related to the unusually low K2O contents in the Coal River plutonic sub-suite, rather than an overabundance of aluminum. 53 The differentiation index (D.I.; after Thornton and Tuttle, 1960) for this sample (KR-05-08) is only 62 (a value of 100 would be assigned to a rock comprised entirely of quartz–albite–K-feldspar); this sample is one of the least evolved plutonic samples from the eastern Selwyn Basin (see Chapter 7 for a comparison of D.I. for other plutonic samples in the eastern Selwyn Basin). 310 zircon cores in the Coal River sub-suite also have smaller cerium anomalies due to their light REE enrichment, but shallow europium anomalies that are similar to those of the magmatic rims (Fig. 5.5b). Like the Coal River sub- suite, zircons from the Tay River suite sample display a remarkable uniformity in chondrite-normalized REE patterns (Fig. 5.5a). This suggests that precipitation of zircon in the Tay River suite (and Coal River sub-suite) samples was from a homogeneous magma that was in turn probably had a relatively homogeneous protolith. The single magmatic zircon core analysis for the Tay River suite has essentially the same REE profile as the rims (Fig. 5.5b). Analyses of inherited zircon cores have similar light REE patterns to those of inherited cores from the Coal River sub-suite (i.e., higher light REE contents and a smaller cerium anomaly; Fig. 5.5b). Inherited zircon in the Tay River suite, however, has slightly deeper europium anomalies than inherited cores in the Coal River sub-suite (Fig. 5.5b). Zircons from the Tungsten plutonic suite in the northeast Selwyn Basin generally have smaller cerium anomalies than in the southeast Selwyn Basin because of higher praseodymium (Fig. 5.5a), although all Tungsten suite zircons have relatively deep europium anomalies (average Eu/Eu* = 0.13– 0.24; Figs. 5.5a and 5.6a) and less pronounced positive cerium anomalies when compared to the Tay River suite zircons. Otherwise the Tungsten suite in the southeast Selwyn Basin is similar to the Tay River suite (including the Coal River sub-suite) in terms of relative REE enrichments and depletions (Figs. 5.5a and 5.6a). Magmatic and inherited cores in the southeast display similar patterns, although in the southeast Selwyn Basin two analyses of inherited cores are anomalously light REE enriched and two analyses of magmatic cores are generally light REE-richer (but still display a pattern similar to that of the magmatic cores). Zircon cores from the northeast Selwyn Basin are relatively lanthanum and heavy REE depleted and have a smaller negative europium anomaly compared to their magmatic rims. These properties are consistent with crystallization in a less differentiated magma. One of the magmatic zircon cores from this suite in the northeast Selwyn Basin is anomalously light REE enriched. Magmatic rims and cores in zircons from the Mayo suite are similar to those from the Tungsten suite in the northeast Selwyn Basin. Magmatic zircon is typically relatively heterogeneous and has smaller cerium anomalies compared to the Tay River suite (Figs. 5.5a and 5.6a; one magmatic core analysis is anomalously light REE-rich). Moderate (Mayo suite) to high (Hess–MacMillan sub-suite) lanthanum and praseodymium enrichments in several analyses are also apparent (Fig. 5.5a). The negative europium anomaly is more pronounced, and heavy REE contents are higher for the Hess–MacMillan sub-suite than for the Mayo suite variety (Figs. 5.5a and 5.6a). Both of these features suggest a higher that the Hess–MacMillan sub-suite experienced a higher degree of magmatic differentiation, although a more incompatible element enriched supracrustal protolith and (or) anatectic processes for this sub-suite could also be responsible for these variations. The presence or absence of plagioclase, hornblende, or certain accessory minerals and their involvement (or lack thereof) in partial melting, for example, will control REE concentrations in the melt and zircon. Cores in the Hess–MacMillan sub-suite also display a similar range of heterogeneous compositions—most notably, one magmatic core is unusually heavy REE depleted, which may be a result of crystallization in the presence of garnet. Chondrite-normalized REE patterns for zircons from the Tombstone suite are similar to those of the Mayo suite in 311 terms of heterogeneity, light REE distribution, and lower heavy REEs when compared to plutonic suites in the southeast Selwyn Basin. Several samples of magmatic rims from the Tombstone suite are, however, anomalously light REE enriched. Magmatic and inherited cores from the Tombstone suite also display a range of compositions, particularly with respect to light REEs, but fall within the range of magmatic rims (with the exception of one particularly REE-rich inherited core). 5.3.2.1 Markers of Magmatic Differentiation in Zircon Hafnium concentrations zircon are unquestionably related to magmatic evolution (e.g., Clairborne et al., 2006). As temperature of a magma decreases, hafnium is increasingly partitioned into zircon, creating an internal zoning from hafnium-poorer cores to hafnium-richer rims (e.g., Hoskin and Schaltegger, 2003; Claiborne et al., 2006). Titanium concentrations are inversely correlated to hafnium, as are the abundances of light and middle REEs, although the heavy REEs do not vary as strongly with hafnium in zircon (Clairborne et al., 2006). Similar correlations are evident in Figure 5.6a, where the plutonic suites have been broadly sub-divided to highlight differences between the southeast and northeast Selwyn Basin. Titanium concentrations in zircons from the southeast Selwyn Basin are low, typically between 3–8 ppm, which is consistent with lower titanium–in–zircon temperature of crystallization for the zircons from these plutonic suites (Fig. 5.6). Titanium concentrations in zircons from the TTB suites in the northeast Selwyn Basin, however, display a decrease from ~24 to 2 ppm over a smaller range of hafnium concentrations (Fig. 5.6). The implied higher magmatic temperatures could be the result of high-temperature partial melting of infracrustal rocks to form the Tombstone and Mayo suite intrusions, as well as the input of mantle-derived magma into the Tombstone plutonic suite (e.g., Hart et al., 2004a). Zircons from the TTB plutons also have lower yttrium and lutetium, and higher lanthanum contents. Thorium and samarium also tend towards higher concentrations in zircons from the TTB intrusions, despite substantial overlap with zircons from the southeast Selwyn Basin. Aside from titanium, all of the trace elements in Figure 5.6 are either relatively static with changing hafnium concentration, or increase with hafnium concentration. Arrays of zircon compositions in Figure 5.6 for the TTB plutonic suites display greater trace elemental variation (i.e., have steeper slopes) than plutonic suites in the southeast Selwyn Basin. Zircons from the Tungsten plutonic suite as a group (barring some analyses from the southeast Selwyn Basin) and the Hess–MacMillan sub-suite have the highest hafnium concentrations of all the plutonic suites examined, as well as some of the highest yttrium, thorium, samarium, and lutetium contents54. This ‘enriched’ trace element signature is typical of zircons from highly evolved plutons. Trace element ratios, such as Th/U and Yb/Gd, also vary with hafnium contents (Fig. 5.7a). Thorium and uranium vary widely in magmatic zircon, although the typical negative coupling between Th/U and hafnium is apparent here (Hoskin and Schaltegger, 2003; Clairborne et al., 2010; Castiñeiras et al., 2011). Yb/Gd, which is one measure of the slope of heavy REEs in zircon (e.g., Fig. 5.4), increases with magmatic evolution. Trace element ratios in zircon may also be useful for identifying populations of grains from magmas that have experienced distinct petrogenetic 54 Interestingly, lanthanum may be lower in zircons from the Tungsten suite than those from less evolved plutonic suites. Several zircons from the Tombstone plutonic suite have the highest lanthanum concentrations, which could be inherited from the light REE- (and LILE-) enriched, lithospheric mantled-derived magmas included in the petrogenesis of this suite (Hart et al., 2004a). 312 histories (e.g., Barth and Wooden, 2010). Ratios such as Th/U, Lu/Hf, Zr/Hf, and Yb/Gd in zircon have been used to distinguish differences between magmatic suites (Barth and Wooden, 2010; Castiñeiras et al., 2011). In Figure 5.7, Yb/Gd and Lu/Hf increase with magmatic evolution (indicated by Hf and Th concentrations, respectively), but are also more elevated overall in zircons from the southeast Selwyn Basin. There are no obvious differences between Th/U for zircons from the southeast and northeast Selwyn Basin as both regions display a well-defined negative correlation with hafnium concentrations (Fig. 5.7a). Just as with the trace element concentrations in Figure 5.6, the Tungsten plutonic suite and the Hess–MacMillan sub-suite have the lowest Th/U and highest Lu/Hf. Zircons from these suites northeast Selwyn Basin also have some of the highest Yb/Gd (Fig. 5.7). When Ce/Ce* and Ce/Sm (a proxy for Ce/Ce* that is not influenced by small variations in light REEs; Wooden et al., 2006) are plotted with hafnium concentrations in zircon, they do not appear to be directly linked to hafnium and, therefore, to the degree of magmatic evolution (Fig. 5.8). Instead, cerium is controlled by the oxygen fugacity of the magma. Ce4+ is more abundant in oxidized magmas and has an ionic charge and radius that makes it more compatible in zircon than the trivalent form (Hoskin and Schaltegger, 2003). The high Ce/Ce* and Ce/Sm in the Tay River suite (especially the Coal River sub-suite), and the high Ce/Sm in the Tombstone and Mayo suites (Fig. 5.8) indicate that these magmas were oxidized magma. Conversely, low Ce/Sm in zircons from the Tungsten plutonic suite and the Hess–MacMillan sub-suite are compatible with reduced magmatic conditions (Fig. 5.8). Europium concentrations are typically better correlated to hafnium concentrations in zircon in Figure 5.8 because it is usually controlled by magmatic processes related to temperature (Hoskin and Schaltegger, 2003). The fractionation of plagioclase and (or) other middle REE-bearing minerals (e.g., monazite; Ayres and Harris, 1997), for example, has been linked to the decrease in Eu/Eu* in whole rock and zircon geochemical datasets (Hoskin and Schaltegger, 2003). The high Eu/Eu* for only zircons from the Coal River sub-suite plutons, however, does not vary with hafnium (and, thus, temperature), which implies that it is controlled by oxidized magmatic conditions (e.g., Castiñeiras et al., 2011). This is consistent with the elevated Ce/Ce* and Ce/Sm, and magnetic susceptibility measurements, for these plutons (Appendix A.1). Conversely, the steep slope exhibited by the Tungsten suite samples when Eu/Eu* and Ce/Sm are compared (circled; Fig. 5.8) is more compatible with the effects of magmatic fractionation on Eu/Eu*. On the basis of zircons from the representative samples, the Tungsten plutonic suite does not appear to comprise magmatic fractionates of the Mayo plutonic suite—if they were, consistently lower Eu/Eu* than that exhibited by the Mayo suite zircons would be expected. Overall, there appears to be a simple geographic control on the composition of zircons from representative samples of the plutonic suites in the eastern Selwyn Basin. Zircons from the southeast are typically more light REE depleted (with the exception of Ce), and slightly heavy REE enriched (i.e., have higher Yb/Gd) compared to TTB zircons from the northeast. This geographic difference even extends to the Tungsten plutonic suite, even though it is interpreted to have been derived from the same infracrustal rocks as the Mayo plutonic suite (Chapter 4). TTB zircons generally have more heterogeneous compositions, particularly the Mayo (including coeval examples of the Hess–MacMillan sub-suite) and Tombstone plutonic suites, and deeper europium anomalies than the Tay River plutonic suite (including the Coal River sub-suite). When Yb/Gd and Lu/Hf are compared to indices of magmatic evolution, they also appear to be effective at discriminating between intrusions from the two regions of the Selwyn 313 Basin. 5.3.2.2 Magmatic and Inherited Zircon Compositional Comparisons A comparison of trace element ratios between magmatic rims, and magmatic and inherited cores (Fig. 5.9) demonstrates a lack of major differences between rim and core compositions. The inherited and magmatic components appear to have similar trace element ratio ranges, and the cores even display similar geographic variations identified in Section 5.3.1.1 with respect to Lu/Hf and Yb/Gd. There are, however, several small differences between the magmatic rim and magmatic core compositions. Th/U typically extends to slightly higher values for magmatic cores, particularly for the Tungsten plutonic suite in the northeast Selwyn Basin (Fig. 5.9). Higher Th/U in the magmatic cores may be due to early zircon crystallization in a less evolved magma, whereas the lower Th/U in magmatic rims (especially for the Tungsten plutonic suite) is consistent with late zircon crystallization in a more evolved magma chamber. Lu/Hf in magmatic cores does not extend to values as high as those found in the magmatic rims (Fig. 5.9), which is another reflection of the lower degree of magmatic differentiation when zircon first saturated in the magma. The similar range of Ce/Sm for magmatic rims and cores is consistent with relatively static oxygen fugacity conditions in the magma chambers while zircon was precipitating (Fig. 5.9). Inherited zircon cores are not substantially different, compositionally, from magmatic cores and rims (Fig. 5.9). Even Th/U, which is typically less than 0.1 in metamorphic zircons (Hoskin and Schaltegger, 2003), does not appear to be markedly depleted in inherited cores from any of the plutonic suites (Fig. 5.9). This suggests that magmatic zircon compositions reflect the zircon compositions of the protolith rocks, although probably in part because of the minimal compositional variation in terrestrial zircons. The dataset for inherited cores in zircons from the representative samples is hardly comprehensive, but if any of these magmas were derived from highly metamorphosed (i.e., lower crustal) rocks, Th/U below 0.1 would be expected in the inherited cores sourced from granulite-facies crystalline basement. The lack of Th/U evidence for a high-grade metamorphic protolith in any of these suites makes shallower (lower–middle crustal) amphibolite-facies rocks the most likely source of magma in the Selwyn Basin. 5.3.2.3 Anomalously Light REE Enriched Zircon When REE concentrations are normalized to chondritic compositions in Figure 5.5, some magmatic rims, cores, and (or) inherited core analyses are unusually enriched in light REEs (Fig. 5.5). The light REE-rich grains are anomalous compared to well-established REE patterns for nearly all zircons (e.g., Belosouva et al., 2002; Kinney and Maas, 2003). Several magmatic rims from the Tombstone, Tay River, and Tungsten (northeast Selwyn Basin) plutonic suites, as well as Hess–MacMillan sub-suite, are anomalously light REE-rich (Fig. 5.5; Table 5.2). Magmatic cores with similar light REE enrichments were only detected in zircons from TTB intrusions in the northeast Selwyn Basin. 314 In sectors of magmatic zircon, areas displaying a lack of zoning and anomalously BSE-bright images caused by heavy element enrichment (U, Th, and REEs; Belousova et al., 2002) suggestive of metamictization were not analyzed. Similarly, magmatic zircon in which hydrothermal alteration may have occurred (indicated by feathery or botryoidal patterns associated with sector boundaries and fractures) was also avoided because of the possibility of hydrothermal or late-magmatic re-mobilization and (or) precipitation of incompatible elements. Some ablation pits, however, were done in highly fractured portions of zircon because of a lack of suitably unfractured material that was large enough to analyze. This was mainly an issue for Tungsten suite zircons from the northeast Selwyn Basin, although only three such analyses in magmatic zircon demonstrate an anomalous light REE enrichment (Fig. 5.5a). Given the general avoidance of zircon sectors affected by sub-solidus processes, the best explanation for elevated light REE contents in analyses of magmatic zircon are miniscule melt or mineral inclusions (e.g., monazite) that are too small to be visible in SEM images (e.g., Whitehouse, 2003; Claiborne et al., 2010; M. Tubrett, pers. comm., 2010). One implication for this is that magmatic zircon cores (i.e., early saturating zircon) from the TTB magmas in the northeast Selwyn Basin, specifically, may have crystallized in a magma that was relatively light REE-rich in its earlier stages, which further implies a light REE enrichment in the parental magmas or protoliths to these suites. Conversely, light REEs do not appear to have been enriched in plutonic suites in the southeast Selwyn Basin until later stages of magmatic fractionation. Light REE enrichment is also present in one inherited zircon core from the Tombstone plutonic suite. Radiating and concentric fractures emanate from this core, which is also speckled with BSE-bright micro-inclusions (Appendix E.1). Metamictization and the addition of uranium to this old core is likely responsible for the expansion fractures (e.g., Corfu et al., 2003; Hoskin, 2005; Pettke et al., 2005) and the higher light REE contents. Although most inherited cores were not analyzed, common radiating and concentric fracture patterns emanating from BSE-bright inherited cores are suggestive of similar processes in all of the plutonic suites (Appendices B.3 and E.1). 5.3.3 εHf100 and TDM As with the radiogenic isotopic systems discussed in Chapters 3 and 4, the Lu–Hf isotopic system operates by the decay of a parent isotope (176Lu) to a daughter isotope (176Hf). The isotopic ratio of interest is the ratio between the daughter produce and then non-radiogenic hafnium isotope (177Hf), which is commonly presented in epsilon notation. Despite similarities to the Sm–Nd isotopic system, one of the unique attributes of the Lu–Hf isotopic system is that 176Hf/177Hf is very low in rocks overall, as well as zircon (<0.0005; Kinney and Maas, 2003). This minimizes the rate of post-crystallization isotopic decay, and essentially preserves initial 176Hf/177Hf, meaning that this ratio does not need to be corrected for decay, especially in zircon. Lutetium is typically found in heavy REE- rich minerals (e.g., garnet, zircon, xenotime) and hafnium is almost always contained within zirconium-bearing minerals (e.g., zircon). Variations in ionic size and charge cause Lu/Hf to fractionate more extensively than Sm/Nd during melt extraction events from the depleted mantle, making hafnium isotopic compositions one of the most finely-tuned isotopic tracers in common use today. In this way, chondritic Lu/Hf has been modified by the extraction of basaltic magmas from the upper mantle, depleting the mantle but enriching the crust in hafnium so that the crust is typically characterized by negative εHf and the mantle by positive εHf. When this sensitive isotopic system is 315 well-preserved preserved robust minerals like zircon, which is able to survive multiple melting and metamorphic events and still preserve the initial hafnium isotopic composition (even in grains where the U–Pb system has been disrupted by lead loss; Kinney and Maas, 2003). The utility of the Lu–Hf isotopic system in petrogenetic studies involving zircon, therefore, is only limited by analytical precision. Hafnium isotopic compositions (εHf100) for zircons from the eastern Selwyn Basin are moderately to strongly negative (Table 5.3, Fig. 5.10a), which is consistent with either a significant crustal input or an entire derivation from the crust (e.g., Kinney and Maas, 2003). In general, suites in the southeast Selwyn Basin have more radiogenic compositions compared to the TTB suites in the northeast Selwyn Basin, which is compatible with a relatively primitive source of magma in the southeast (Fig. 5.10a). One particular feature of note is that εHf100 values from magmatic and inherited zircon cores are remarkably similar to εHf100 values from magmatic rims for zircons from plutonic suites in the southeast Selwyn Basin (Fig. 5.10a). Like geochemical and isotopic date presented previously (Chapters 2 and 4), these similar zircon core and rim compositions are consistent with a homogeneous melt source in the basement to this region composition. Inheritance in the northeast Selwyn Basin was only examined in the Tungsten plutonic suite and the Hess– MacMillan sub-suite. Even though magmatic cores and rims have similar εHf100 in these samples, old inheritance, is markedly more evolved and heterogeneous (for multiple analyses from the Hess–MacMillan sub-suite zircons) than in the southeast (Fig. 5.10a). These observations are compatible with a more isotopically evolved melt source than in the southeast Selwyn Basin (which is consistent with interpretations from Chapter 4). For the Hess–MacMillan sub- suite in particular, a heterogeneous melt source of supracrustal composition (i.e., containing detritus that was sourced from multiple terranes) is implied by the wide range of εHf100 exhibited by inherited cores. Depleted mantle model ages for the plutonic suites display the opposite pattern (Fig. 5.10b), whereby a greater range of model ages is compatible with a more heterogeneous melt source. TDM’s for intrusions in the southeast Selwyn Basin range from 1.1–1.5 Ga and display minimal variation, even for inherited zircon cores (Fig. 5.10b). TDM’s for TTB intrusions range from 1.4-1.8 Ga for magmatic rims and cores, although TDM’s for inherited cores from the Tungsten and Hess–MacMillan sub-suite in the same region are typically greater than 2.0 Ga (Fig. 5.10b). 5.3.3.1 Integration of Lu–Hf isotopic Compositions in Zircon with Other Isotopic Datasets When εHf100 from zircon is plotted with whole rock radiogenic isotopic compositions, it is clear that these isotopic systems are closely coupled (Fig. 5.11). There is a positive correlation between εHf100 and εNd100 (Fig. 5.11a) and a negative correlation between εHf100 and Sr100 (Fig. 5.11b). The Hess–MacMillan sub-suite defines the most evolved isotopic end-member and appears to have incorporated the greatest amount of isotopically evolved crust, and the Tay River suite sample defines the least evolved isotopic end-member. When Hf100 is plotted with stable isotopic compositions, however, there are a few unexpected results. First, the Tungsten plutonic suite in the southeast Selwyn Basin appear to correlate better with the TTB intrusions, despite the chemical and Hf100 similarities in shares with other intrusive rocks in the southeast. These hafnium, sulfur, and oxygen isotopic data provide one of the few links 316 between Tungsten suite plutons in the southeast Selwyn Basin, and TTB plutons in the northeast Selwyn Basin. Another interesting feature is that Hf100 and δ18O appear to be positively coupled in both the southeast and northeast Selwyn Basin (Fig. 5.11c), which presents an isotopic conundrum. The Coal River sub-suite of the Tay River plutonic suite, for example, has the heaviest oxygen isotopic compositions and the least isotopically evolved hafnium compositions (Fig. 5.11c). Typically, hafnium isotopic compositions decrease, or become more negative, as oxygen isotopic compositions increase, as indicated by the arrow in Figure 5.11c. This decoupling between the hafnium and oxygen isotopic systems—in zircon and quartz, respectively—requires that the more isotopically juvenile component to these magmas was sourced from rocks that experienced near-surface processes, such as weathering or low-temperature hydrothermal alteration (e.g., Rollinson, 1993). The influence of isotopically juvenile, altered oceanic crust and (or) young sedimentary rocks eroded from isotopically juvenile crust, on the granitoids in the eastern Selwyn Basin is the most likely explanation for this unusual isotopic correlation. When εHf100 and δ34S are compared, however, they display a positive correlation for all of the samples (with the exception of the Hess–MacMillan sub-suite; Fig. 5.11d). This implies that a crustal reservoir is responsible for imparting the heavier sulfur to these plutonic suites, and that as the magmas become more isotopically evolved (i.e., appear to have incorporated a greater proportion of evolved crust), isotopically heavy sulfur derived from crustal rocks was added. A comparison of TDMs determined from zircon hafnium isotopic composition and whole rock neodymium isotopic composition, the array of data has a slope of less than one, and suggests a discrepancy between model ages estimated from the two isotopic systems (Fig. 5.12). TDMs determined from magmatic sectors of zircon could be younger from those determined by whole rock analyses as the latter will include all inherited components in the magma. This would make TDMs determined from whole rock isotopic datasets older if inheritance is present, which it is (to varying degrees) in all of the plutonic suites examined here (Table 5.1). This may be why all of the plutonic suites have older neodymium-based model ages when compared to the hafnium-based model ages determined from magmatic zircon. A similar line of reasoning explains why the inherited cores from the Tungsten suite and the Hess– MacMillan sub-suite (both in the northeast Selwyn Basin) plot above the line with a slope of 1.0 (Fig. 5.12). There is also the possibility that these two isotopic systems will inherently produce different estimates of model ages in some situations. Schimitz et al. (2004) have documented significant decoupling between the whole rock Lu–Hf and Sm– Nd isotopic systems in lower crustal, granulite-facies rocks of mafic composition (which have high Lu/Hf and low Sm/Nd) and metasedimentary composition (which have low Lu/Hf and widely variable Sm/Nd). This isotopic decoupling is thought to occur during anatexis in rocks containing residual garnet or zircon (that preferentially incorporate Lu), and residual rutile (that preferentially incorporates Hf; Schmidtz et al., 2004). Therefore, it may be that the TDMs determined from the Lu–Hf isotopic system (whether from zircon or whole rock datasets) often underestimate crustal residence ages, just as TDMs determined from the Sm–Nd isotopic system may often overestimate crustal residence ages (particularly if Sm/Nd was fractionated during partial melting). 317 5.4 PETROGENETIC IMPLICATIONS 5.4.1 Crustal Protolith Compositions All of the zircons in Figure 5.11a plot within a terrestrial array defined by lower crustal xenoliths, although the Tungsten suite and Coal River sub-suite samples appear to have slightly more radiogenic hafnium when compared to zircons with similar εNd100 from other samples. The tendency towards more radiogenic hafnium could be related to higher Lu/Hf in these suites, implying that their parental magmas were depleted in lutetium relative to hafnium. If so, garnet or zircon would have comprised residual minerals in the protolith rocks (cf. Schmidtz et al., 2004) during the extraction of magma. If, however, zircon and garnet were not residual, the plutonic samples should inherit the Lu/Hf signature of the protolith rocks. In the case of the Coal River sub-suite, for example, the slightly more radiogenic εHf100 could have been influenced by magma derived from mafic orthogneissic rocks (Schmidtz et al., 2004), which may be consistent with geochemical and isotopic interpretations in Chapters 2 and 4 that are compatible with some magma input from mafic rocks. Lu/Hf fractionation is also controlled by depositional sorting, which can be an important control on this ration in magmas derived from supracrustal rocks. Zircon—the primary host of hafnium in the crust—is concentrated in sandy siliciclastic rocks, such as those deposited on continental shelves (Kinney and Maas, 2003). Conversely, lutetium (along with other REEs, including Sm and Nd) is concentrated in clay-rich siliciclastic rocks, which are deposited in large quantities in deeper marine settings (Kinney and Maas, 2003). This lithological fractionation of Lu/Hf during erosion, combined with the degree of hafnium isotopic evolution, makes this isotopic system useful for inferences regarding the protolith lithologies for S-type granitic rocks. Lu/Hf is consistently highest in the Tungsten plutonic suite zircons, and to a much lesser degree in the Coal River sub-suite zircons (Fig. 5.7b). Although high Lu/Hf could be linked to magmatic fractionation for the Tungsten suite (and may have been inherited from a mafic orthogneissic protolith in the case of the Coal River sub-suite), it could also have been derived during partial melting, assimilation, and (or) contamination involving pelitic rocks. This pelitic influence on the Tungsten plutonic suite in the southeast Selwyn Basin would have been in the form of contamination or assimilation if previous interpretations that this is a dominantly I-type suite are correct. Conversely, the peraluminous Coal River sub-suite is interpreted to be derived from both infracrustal (mafic) and supracrustal rocks (Chapter 4), which suggest the supracrustal component was mainly metapelitic. For the one suite thought to be entirely derived from supracrustal rocks— the Hess–MacMillan sub-suite—Lu/Hf in zircon is not particularly elevated (Fig. 5.7b). The range of Lu/Hf in these zircons is compatible with a derivation from a mixture of pelitic and quartzo–feldspathic sedimentary rocks, which is consistent with the interpreted origin in Neoproterozoic ‘grits’ for this sub-suite (Chapter 4). 5.4.2 Magma Mixing Some authors have used hafnium ± oxygen isotopic compositions of zircon in combination with other isotopic systems (e.g., Sm–Nd, Rb–Sr) to argue for magma mixing between two compositionally distinct melt sources in the genesis of some I- and S-type magmas (e.g., Griffin et al., 2002; Kemp et al., 2007; Appleby et al., 2008, 2009; 318 Peytcheva et al., 2008; Shaw and Flood, 2009). Typically only an isotopically juvenile mantle-derived magma can be distinguished from a more isotopically evolved crustally derived magma in plutons that have formed by mixing of two or more magmas; mixing of isotopically similar magmas would be virtually impossible to detect by isotopic comparison. When the former occurs, a wide range of hafnium isotopic compositions can result, even within zircons from the same sample (Griffin et al., 2002; Shaw and Flood, 2009). Despite the presence of significant resorption surfaces in magmatic zircons for all but the Tay River suite (Table 5.1), none of the other plutonic suites have large hafnium isotopic variations in magmatic zircon. Thus, mixing between two or more magmas is unlikely to have occurred in any of the representative intrusive phases selected from the eastern Selwyn Basin, unless the magmas were of relatively similar isotopic composition. This presents a problem for the Coal River sub-suite, which is isotopically homogeneous, but also interpreted to have been derived from mixing of magmas derived from mafic infracrustal rocks and pelitic supracrustal rocks. One alternative, therefore, might be that this sub-suite was derived from a mechanically well-mixed and fine-grained supracrustal rock containing a significant proportion of detritus derived from mafic volcanic rocks. 5.5 A COMPARISON OF ZIRCON CRYSTALLIZATION TEMPERATURES Results for three methods of estimating zircon crystallization temperatures are presented in Figure 5.13. Estimates of temperature based on zircon morphology (after Pupin, 1980) provide a very broad approximation of the temperature of zircon crystallization. The morphology-based temperature estimates here are limited to ranges of either 650 ± 100°C or 850 ± 100°C in Figure 5.13 as greater accuracy is not realistic given the continuous range of morphological variations in Figure 5.4. For samples in which both high- and low-temperature zircon morphologies are present, the dominant population is outlined with the black box, and the sub-ordinate population is outlined with the grey box in Figure 5.13. A detailed examination of magmatic cores and rim morphologies might allow the evolution of the temperature and composition of a magma saturated in zircon to be better determined (e.g., Belousova et al., 2006). This more detailed work, however, has not been done here, and linkages between magmatic core versus rim temperatures (and the assumption that these cores represent the earliest crystallizing zircon in the magma) are not possible here. Zircon saturation temperatures (ZSTs; Watson and Harrison, 1983) are calculated from whole rock geochemical analyses using the concentrations of alkali elements, aluminum, silica, and zirconium. Lower temperatures determined by this method are typical of rocks that have higher SiO2 and zirconium contents. ZSTs have been used extensively in the literature to estimate the temperatures of magmas (see examples in Hanchar and Watson, 2003) and in some cases to determine the petrogenetic history of rocks (e.g., Chappell et al., 2000). The accuracy of these estimates of magma temperature, however, can be reduced by the presence of inherited zircon, which tends to be significant in crustal melts (including all of the representative samples examined here; Table 5.1). Inherited zircon is found in intrusions that were oversaturated in zirconium when the melt was extracted from the source rock and, in such cases, ZSTs will only provide an estimate of the maximum temperature of melt extraction (Miller et al., 2003). In intrusions lacking inherited zircon, however, ZSTs do provide a reasonable estimate of the temperature of magma emplacement (i.e., the onset of extensive crystallization, typically beginning with zircon saturation in intermediate to 319 felsic magmas), as well as a minimum estimate for the temperature of partial melting (Miller et al., 2003). More recently, zircon compositional studies have employed a titanium–in–zircon thermometer to estimate the temperatures of magmas and petrogenetic studies (e.g., Watson and Harrison, 2005). The titanium–in–zircon thermometer uses the observed negative correlation between titanium concentration in zircon and temperature of the melt, based on experimental results and natural samples (Watson and Harrison, 2005; Watson et al., 2006; Ferry and Watson, 2007). Assuming all of the zirconium in a rock is contained within zircon, this method should theoretically reproduce the ZST estimates; however, the calculated titanium–in–zircon temperatures are almost always lower in natural samples than ZSTs (e.g., Fu et al., 2008). This discrepancy may be due to overestimating the activity of titanium in the melt (i.e., assuming rutile is in equilibrium with zircon), underestimating the activity of silica (typically assumed to be 1.0 when quartz and zircon are in equilibrium), and (or) other factors that may influence the partitioning of titanium into zircon (e.g., pressure55, sub-solidus Ti exchange, non-equilibrium behaviour; Ferry and Watson, 2007; Ferris et al., 2008; Fu et al., 2008). Claiborne et al. (2010) used an aSiO2 of 1.0 and an aTiO2 of 0.7 (based on the presence of titanite and possibly ilmenite in intrusive rocks) to calculate titanium–in–zircon temperature estimates. The temperatures presented in Figure 5.13 have been calculated using the same activities; if the activity of TiO2 is actually lower in the plutonic suites investigated (all of which have crystallized ilmenite, and several of which have also crystallized titanite; Appendix A.1), then these temperature estimates may be too low. The titanium–in–zircon temperatures plotted in Figure 5.13 vary widely within samples, largely because these temperatures will differ depending on which zone of zircon is analyzed. If the analysis is in a magmatic core, for example, this method should provide a higher temperature estimate than if the analysis is in a magmatic rim. However, analyses from the same zircon sections (e.g., rims, intermediate magmatic zones, or cores) cannot always be directly compared as there is also significant variation in terms of when specific cores and rims may have crystallized. If a zircon grain is separated from the melt early in the crystallization of an intrusion (i.e., as an inclusion in an early formed mafic mineral) then the magmatic rim of that zircon should yield a higher temperature estimate than the magmatic rim of a late-crystallizing zircon that remained in equilibrium with the melt until sub- solidus conditions. With this logic in mind, the highest titanium–in–zircon temperature estimate from a magmatic core provides a minimum estimate for the temperature of the parental (i.e., pre-emplacement) magma (in cases where the magma was undersaturated in zirconium in the melt source; Miller et al., 2003). The lowest titanium in zircon temperature estimate from a magmatic rim, thus, provides a maximum temperature estimate for the solidification of the magma body. Titanium–in–zircon temperatures for the samples examined by this study are mostly within the range of ~820°C to 660°C, although some high temperature outliers range up to ~920°C and some low temperature outliers range down to ~630°C in the TTB intrusions. Temperature estimates from magmatic cores vary widely, although at least one magmatic core from each of the TTB suites (except the Tungsten suite in the southeast Selwyn Basin, in which magmatic cores were not analyzed) have the highest titanium–in–zircon temperature estimates. 55 The titanium–in–zircon thermometer is calibrated at 10 kbar, but there is a change of 5–10°C/kbar in the resulting temperature estimate (Ferry and Watson, 2007; Ferris et al., 2008); however, in upper crustal magmas this actually leads to an overestimate of temperature so this cannot be the cause of the different ZST and titanium–in–zircon temperature estimates. 320 The ZSTs typically plot within the titanium–in–zircon temperature ranges, although generally towards higher temperature limit of this range. This bias towards higher ZST temperatures may be because magmatic rims were targeted for most of the trace element analyses and, therefore, a greater abundance of lower titanium–in–zircon temperatures would have been determined. That said, titanium–in–zircon temperatures calculated from a smaller number of magmatic core analyses, which might be expected to be closer to ZSTs, have an equally wide range of temperature estimates. It may be that some of these magmatic cores are actually from smaller, late-nucleating and crystallizing grains, which may also be why these particular grains have magmatic (and not inherited) cores. It was commonly observed in BSE and CL images that the smallest zircons did not normally contain resorbed inherited cores; these grains typically have euhedral, unzoned magmatic cores with moderate BSE-brightness. The small size of these grains, and the similarity between ages produced by U–Pb analyses of magmatic rims from larger grains and magmatic cores (and rims) from smaller grains suggests that these magmatic cores are from autocrystic grains (e.g., Miller et al., 2007) that nucleated at lower magmatic temperatures. The higher ZST estimates for the Mayo and Tay River suites are consistent with their zircon morphologies, amphibole-bearing mineral assemblages, and interpreted I-type characters (Chapter 2), but are anomalously high compared to estimates of magmatic temperature from the titanium–in–zircon method. It could be that the assumed aTiO2 = 0.7 for the titanium–in–zircon calculations may be incorrect for these particular samples. Accounting for the emplacement pressures for magmas in the eastern Selwyn Basin (~2 kbar; Gordey and Anderson, 1993) would raise the titanium–in–zircon temperature estimates by 40–80°C. Even if these measurements are corrected for pressure and different aTiO2, the slightly higher titanium–in–zircon temperature range would still barely overlap with the ZSTs, if at all, in Figure 5.13. While the notably elevated ZST for the Mayo suite sample in Figure 5.13 may be due to its K-feldspar-crowded texture (Appendix A.1) and overall elevated whole rock potassium content (Chapter 2), the high ZST for the Tay River suite sample requires an alternate explanation. One possible explanation may be that the Tay River suite was derived from zircon-poor rocks; therefore, zirconium was so undersaturated in these magmas that it did not begin to precipitate until the magma cooled to temperatures of ~800°C (Fig. 5.13). Given that the titanium–in–zircon thermometer generally underestimates the temperature of zircon crystallization for all but the Tombstone plutonic suite in Figure 5.13, it appears that uncertainties regarding the controlling factors on the substitution of titanium into zircon are far from resolved (e.g., Fu et al., 2008). Conversely, the titanium–in–zircon thermometer appears to significantly overestimate the temperature of zircon crystallization for some analyses (although this is not unusual; e.g., Harrison et al., 2007). Higher temperatures estimates for several Tungsten suite and Hess–MacMillan sub-suite analyses in magmatic zircon are similar to the ZSTs, but anomalous compared to most of the titanium–in–zircon temperatures (Fig. 5.13). These high titanium–in– zircon temperatures are in contrast to the dominant population of lower-temperature zircon morphologies that characterize these samples, particularly in zircons from the northeast Selwyn Basin where no high temperature morphologies are recorded (Fig. 5.4). Titanium–in–zircon temperature estimates for most inherited cores are unremarkable and plot within the range of magmatic rim temperatures. Exceptions to this are two analyses of inherited cores in the Tungsten plutonic suite in 321 the southeast Selwyn Basin (T = 1150 and 1630°C; Table 5.2b) and one analysis of a inherited core in the Tombstone plutonic suite in the northeast Selwyn Basin (T = 1040°C; Table 5.2b). The titanium contents in these cores are anomalously high, and also associated with a light REE enrichment that has already been attributed to metamictization for the Tombstone suite sample (Section 5.3.2.3). Radiating fractures extending from zircon cores in two samples (the Tombstone suite sample, and one of the Tungsten suite samples) is diagnostic of this process, thus the high titanium contents (and the resulting titanium–in–zircon temperature estimates) in these inherited core analyses cannot be regarded as primary. These comparisons between different methods of estimating zircon crystallization temperatures highlight the complexity of magmatic systems, and some of the issues associated with these three thermometers (e.g., Watson and Harrison, 1983; Miller et al., 2003; Fu et al., 2008). 5.6 SUMMARY OF PRELIMINARY OBSERVATIONS AND POTENTIAL DIRECTIONS FOR FUTURE ZIRCON COMPOSITION WORK Zircon trace element and hafnium compositions, coupled with BSE and CL images from the same grains, are a useful petrogenetic tool in studies of felsic to intermediate plutonic rocks, particularly when in situ U–Pb ages are also available. The zircon compositional data reported here are generally supportive of previous petrogenetic interpretations for the plutonic suites in the eastern Selwyn Basin, despite the reconnaissance scale of this study, and the relatively limited number of zircon analyses (particularly for Hf isotopic compositions). Most of the zircons from the representative samples for the plutonic suites emplaced into the eastern Selwyn Basin are textually complex. They typically contain multiple zircon sectors or domains, including: magmatic rims, magmatic (both antecrystic and autocrystic; cf. Miller et al., 2007) or inherited cores, and commonly more than one magmatically zoned sector. Zircon morphologies, which provide information on the temperature and alkalinity of the magma(s) from which the zircons precipitated, are best observed from polished sections through zircon grains using the BSE and CL modes on the SEM. Overall, the morphologies documented for the representative samples are consistent with earlier interpretations related to magma temperatures. Zircons from the Mayo, Tay River, and Tombstone plutonic suites (and sub-suites) are dominated by morphologies that are compatible with higher magmatic temperatures and their interpreted I-type petrogenesis in Chapter 2. The Tungsten plutonic suite (southeast Selwyn Basin) and the Hess–MacMillan sub-suite have mainly low-temperature zircon morphologies that are compatible with the small-fraction partial melting conditions and (or) supracrustal protoliths interpreted for these suites in Chapters 2 and 4. The Coal River sub-suite of the Tay River plutonic suite is unusual in that it is dominated by low temperature zircon morphologies that are indicative of equilibration with an unusually aluminous, low- temperature magma. Trace element concentrations and ratios in magmatic zircons were not only reflective of the whole rock trace elemental variation (e.g., the Tay River suite and Coal River sub-suite have homogeneous trace element concentrations in both zircons and whole rock samples), they also elucidated: 322 • differing degrees of magmatic differentiation between the representative samples, highlighting the Tungsten suite and the Hess–MacMillan sub-suite as the most differentiated of those examined; • magmatic conditions, such as the unusually oxidized magma conditions from which zircons from the Coal River sub-suite were precipitated compared to the other representative samples; and • at least two distinct magmatic series, defined mainly by relatively high Yb/Gd for the Tay River and Tungsten suite zircons from the southeast Selwyn Basin, and relatively low Yb/Gd for the Tombstone, Mayo, and Tungsten suite zircons from the northeast Selwyn Basin (other elemental differences include higher Ti and La, and lower Y and Lu in the TTB zircons). Additionally, TTB zircons (especially those from the Tombstone suite sample) are generally more enriched in light REEs and Th, even in zircon analyses displaying evidence for low degrees of magmatic fractionation (e.g., lower Hf, higher Eu/Eu*), which is attributed to the presence of melt or mineral micro-inclusions inherited from an incompatible element enriched parental magma. In general, most magmatic rims and cores share similar trace element concentrations and ratios, and commonly inherited cores (although a smaller dataset) appear to do the same, especially for the I-type plutonic suites. Similarly, the hafnium isotopic composition of magmatic cores typically mirrors that of the magmatic rims in all of the plutonic suites where cores were examined. Old inherited zircon cores from the southeast Selwyn Basin were also isotopically similar to their magmatic rims, which is consistent with the interpreted homogeneous I-type melt source for plutonism in this region. However, in the northeast Selwyn Basin (the Tungsten and Hess-MacMillan samples in this region) inherited zircon cores have markedly more evolved hafnium compositions (and older TDMs) that would be more compatible with a heterogeneous (i.e., supracrustal) melt source for these samples. Like the distinct Yb/Gd in zircons from plutons in the southeast and northeast Selwyn Basin, hafnium isotopic compositions are also regionally divided. Zircons from the Tay River plutonic suite (including the Coal River sub- suite) and to a lesser degree the Tungsten plutonic suite in the southeast Selwyn Basin are typically slightly more radiogenic than zircons from the TTB plutons in the northeast Selwyn Basin. This broad change in zircon isotopic composition from southeast to northeast could reflect the presence of a more isotopically juvenile basement below the southeast Selwyn Basin (which is also consistent with other plutonic isotopic datasets for this region, including Nd, Sr, and O; Chapter 4). None of the plutonic suites in the eastern Selwyn Basin display hafnium isotopic evidence for magma mixing between two or more compositionally distinct magmas, despite interpretation in Chapter 4 that the Coal River sub-suite was sourced from a mixture of mafic igneous and supracrustal rocks. Hafnium in zircon, and whole rock strontium, neodymium, and sulfur isotopic compositions are well coupled according to the isotopic variations that reflect the input of crustal materials. Stable isotopic comparisons with hafnium isotopic compositions also link the Tungsten plutonic suite in the southeast Selwyn Basin to the TTB plutonic suites in the northeast Selwyn Basin. Oxygen from quartz separates, however, is positively coupled with hafnium in zircon, which is at odds with the trend expected if isotopically heavy supracrustal rocks are responsible for more evolved (i.e., crustal) hafnium isotopic compositions in the zircons. This implies that a supracrustal (or low-temperature hydrothermally altered) material is partly responsible for imparting the more primitive (or 323 radiogenic) hafnium isotopic compositions observed in zircons from plutonic suites in the eastern Selwyn Basin. The heavy oxygen, radiogenic hafnium (as well as neodymium, and non-radiogenic strontium) zircons from the Coal River sub-suite provide the best example of this isotopic conundrum. In addition to these isotopic characteristics, this sub-suite also has zircon morphologies and whole rock chemistries (Chapter 2) that argue for an important aluminous component in the late-stage magma. Several questions remain regarding the petrogenesis of the Coal River sub-suite, as well as the positive coupling between hafnium and oxygen isotopic compositions Overall, zircon trace element and hafnium isotopic compositions appear to support the petrogenetic interpretations based on whole rock datasets for igneous rocks in the eastern Selwyn Basin. The reconnaissance scale of this particular study, however, leaves room for additional work to confirm or disprove the petrogenetic interpretations, and has resulted in additional questions concerning the petrogenesis of plutonic suites in this region (e.g., the Coal River sub-suite). 5.6.1 Recommended Future Zircon Compositional Work Not only do intrusive rocks provide some of the only information on the age and composition of the basement to the northern Cordilleran miogeocline, but inherited zircon within these mainly crustally derived rocks are the most direct source of information on this basement. The presence of extensive crystalline rocks below the Selwyn Basin, particularly in the southeast, may be inferred based on the whole rock geochemical (Chapter 2) and isotopic (Chapter 4) composition of plutonic suites in this region. Future zircon compositional work in the Selwyn Basin, if done in a systematic manner, will also help to answer questions regarding the age and composition of the interpreted metaigneous basement to this region. Extensive work on inherited zircon cores, which are abundant in most of the plutonic suites in the Selwyn Basin, would provide statistically valid information on the age and composition of plutonic melt sources, regardless of their affinity. Additional work on both magmatic and inherited zircon should include in situ U–Pb ages, hafnium and oxygen isotopic compositions (in order to determine if δ18O from quartz is accurate), and trace element chemistry. The characterization of zircon morphologies, and particular attention to their relative location (as well as the morphological sector analyzed for trace elements and isotopic compositions) would also provide information on the compositional evolution of the magmas. Smaller analytical spot sizes and (or) higher precision for U–Pb, trace element, and oxygen isotopic measurements are several advantages that might help to resolve magmatic evolutionary pathways for selected representative samples, if the SHRIMP is employed for data collection, although LA–ICP–MS analytical methods could be used for more regional characterizations where high precision and small spot sizes are not as important. 324 Figure 5.1. Distribution of Early to ‘mid’-Cretaceous, syn- to post-accretionary plutonic suites in the northern Cordillera (AK = Alaska; YK = Yukon; NWT = Northwest Territories; BC = British Columbia). The study area (Fig. 5.2) is outlined in black. Plutonic suites discussed in the text are labeled: To = Tombstone, M = Mayo, Tu = Tungsten, TR = Tay River. The Tombstone, Mayo, and Tungsten plutonic suites for the Tombstone–Tungsten Belt (TTB) referred to in the text. Tectonostratigraphic units and map lines are as in Figure 1.1 (Chapter 1). 325 Figure 5.2. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous magmatism in the eastern Selwyn Basin study area. Representative samples used for reconnaissance-scale zircon compositional work are indicated by zircon symbols and labeled with the sample number. The approximate southeast– northeast boundary between the Tungsten plutonic suite in the southeast Selwyn Basin is at a latitude of 62°N. Map lines and tectonostratigraphic units as in Figure 1.1 (Chapter 1), and abbreviations are: Ft = fault, FS = Fault System, and H–M = Hess–MacMillan. 326 Figure 5.3. The link between zircon morphology, composition, and temperature of crystallization (modified from Pupin, 1980, and Belousova et al., 2006). a) Morphological classification chart for zircon. The crystal faces on the y-axis describe zircon prisms, and the crystal faces on the x-axis describe zircon pyramids. Alkalinity is relative Al/(Na+K). b) Summarized morphological evolution for magmas of various compositions. 327 Figure 5.4. (next 2 pages) Classification of internal zircon morphologies (after Pupin, 1980). Percentages are given by the large numbers in each morphological type, and the grey shading corresponds directly to 2x the percentage (where 0% is represented by no shading, and 50% would be represented by black). The bold outline shows the morphological range of most zircon (Pupin, 1980). 328 329 330 Figure 5.5. (next page) Chondrite-normalized multi-rare earth element (REE) plots of zircon compositions (Sun and McDonough, 1989). a) Magmatic zircon rim compositions. Average europium and cerium anomalies are reported (excluding values from grains that are anomalously light REE-rich). b) Magmatic and xenocrystic zircon core compositions. Magmatic core analyses are indicated by solid lines and inherited core analyses are indicated by dashed lines. 331 332 Figure 5.6. (next page) Trace elemental variation related to magmatic evolution. Blue (southeast Selwyn Basin) and red (northeast Selwyn Basin) shading is used to make regional compositional differences more apparent. a) Hafnium versus titanium, yttrium, and thorium. The approximate titanium–in–zircon temperature is indicated by the grey dashed lines (calculated assuming aTiO2 = 0.7 and aSiO2 = 1). b) Hafnium versus lanthanum (a light REE), samarium (a middle REE), and lutetium (a heavy REE). 333 334 Figure 5.7. (next page) Trace element ratios related to magmatic evolution. Symbols are as in Figure 5.6. Blue (southeast Selwyn Basin) and red (northeast Selwyn Basin) shading is used to make regional compositional differences more apparent. a) Hafnium versus Th/U and Yb/Gd. b) Thorium versus Lu/Hf. The Tungsten suite as a group (both northeast and southeast members) is outlined by the dashed black lines. 335 336 Figure 5.8. Magnitude of chondrite-normalized europium and cerium anomalies expressed as Ce/Ce*, Ce/Sm (as a preferred proxy for Ce/Ce*; Wooden et al., 2006), and Eu/Eu*. Symbols are as in Figure 5.6. Blue (southeast Selwyn Basin) and red (northeast Selwyn Basin) shading is used to make regional compositional differences more apparent. The Tungsten suite as a group (both northeast and southeast members) is outlined by the dashed black lines in some of the plots. 337 Figure 5.9. (next page) Comparison of trace element ratios in zircon rims and cores (magmatic and inherited). Symbols are as in Figure 5.6; however, for the plots showing zircon inheritance, magmatic core analyses are indicated by the large symbols, and inherited core analyses are indicated by the small symbols. Blue (southeast Selwyn Basin) and red (northeast Selwyn Basin) shading is used to make regional compositional differences more apparent. The Tungsten suite as a group (both northeast and southeast members) is outlined by the dashed black lines in some of the plots. 338 339 Figure 5.10. Hafnium isotopic composition and model ages in zircon. Symbols are as in Figure 5.6; however, magmatic rims are represented by the filled symbols, magmatic cores are represented by unfilled black symbols, and inherited cores are represented by unfilled grey symbols. Blue (southeast Selwyn Basin) and red (northeast Selwyn Basin) shading is used to make regional compositional differences more apparent Note that samples are not arranged in any particular order relative to the x-axis. a) Range of εHf100. b) Range of TDM model ages calculated from hafnium isotopic compositions. 340 Figure 5.11. Radiogenic and stable isotopic compositions versus hafnium isotopic compositions in zircon. Symbols are as in Figure 5.6; Blue (southeast Selwyn Basin) and red (northeast Selwyn Basin) shading is used to make regional compositional differences more apparent. a) Whole rock εNd100 versus εHf100 in zircon. The dashed grey line shows an approximate “terrestrial” array of neodymium and hafnium isotopic compositions for lower crustal xenoliths and felsic orthogneisses; illustrated vectors show the general direction of divergence from the terrestrial trend if Lu/Hf are fractionated during the anatexis of mafic orthogneiss and paragneiss protoliths (after Schmidtz et al., 2004). Large arrows show the general direction of typical isotopic coupling that is indicative of an increasing degree of crustal input. b) Whole rock Sr100 versus εHf100 in zircon. c) δ18O from quartz separates versus εHf100 in zircon. d) Whole rock δ34S versus εHf100 in zircon. 341 Figure 5.12. Comparison of depleted mantle model ages estimated from whole rock neodymium isotopic compositions (x-axis) and zircon hafnium isotopic compositions (y-axis). Symbols are as in Figure 5.6; however, magmatic rim analyses are represented by the filled symbols, magmatic core analyses are represented by unfilled black symbols, and inherited core analyses are represented by unfilled grey symbols. The dashed grey line has a slope of 1. 342 Figure 5.13. Comparison of zircon crystallization temperatures estimated from zircon morphology (after Pupin, 1980), whole rock chemistry (zircon saturation temperature, or ZST, after Harrison and Watson, 1983), magmatic zircon composition (Ti–in–zircon after Claiborne et al., 2010). Symbols are as in Figure 5.6; however, magmatic rim analyses are represented by the filled symbols, magmatic core analyses are represented by unfilled blue outlines, and inherited core analyses are represented by small unfilled blue outlines. The grey and black boxes show the dominant and sub-ordinate, respectively, broad temperature ranges estimated from zircon morphology (after Pupin, 1980). The ZST for each sample is indicated by the red symbol (note that zircons from two samples were examined, so both ZST’s are shown for some plutonic suites). Titanium–in–zircon temperature estimates are plotted for magmatic rims using the symbols above, for magmatic cores using the large unfilled blue symbols, and for xenocrystic cores using the small unfilled blue symbols. 343 Table 5.1. Summary of relevant features of representative samples chosen for zircon compositional analysis. U–Pb ages were previously reported in Chapter 2, and radiogenic and isotopic data were previously reported in Chapter 4. Note all isotopic data from Chapter 4 are for whole rock analyses, except for δ18O (quartz separates). Sub-suites are denoted by: HM = Hess–MacMillan, and CR = Coal River. Plutonic Suite (sub- suite or region) Pluton Sample Sample Descriptiona Summary of Zircon Textures (modified from Appendix B.3)b D.I.c ZST (°C)d MR (Ma) AC (Ma) IC (Ma) δ18O (‰) δ34S (‰) 87Sr/86Sr100 εNd100 TDM (Ga)e Tombstone O'Grady KR-05-97b hbl>cpx>bt tur- bearing qtz monzonite thick w-z MR with at least 1 resorption event, and older p- resorbed poorly/thickly zoned MC; IC present in ~14% of grains 71 790 92.8 ~96.5 nd 14.5 8.4 0.721 -13.33 1.70 Mayo Mt. Christie CL-06-36 hbl>bt>>cpx qtz monzonite thick w-z MR and p-resorbed w-z MC; at least 1 significant resorption surface in rim; IC present in ~2% of grains 68 985 96.8 nd nd 13.5 9.9 0.724 -15.31 1.94 Mayo (Hess– MacMillan) Mt. Wilson KR-05-68 foliated bt monzogranite thick w-z MR; typically 2 major resorption surfaces in MR; p- resorbed to pristine MC; resorbed IC present in ~41% of grains (primarily u-z and/or porous + highly fractured) 71 851 97.7 104 1591, 1766 15.1 7.4 0.734 -20.92 2.21 Mayo (Hess– MacMillan) Pelly River KR-05-76 foliated bt qtz diorite w-z MR; typically 1+ significant resorption surfaces in MR; some small p-resorbed w-z MC; resorbed IC present in ~68% of grains (large and porous) 54 786 98.0 103 1792 nd 6.3 0.734 -20.77 2.16 Tungsten (NE) Cac KR-05-110 bt monzogranite thick w-z MR; typically 1-2 major resorption surfaces in MR; very highly fractured MR ± MC; p-resorbed w-z MC; resorbed IC present in ~10% of grains 77 833 97.4 nd nd 15.5 10.5 0.722 -15.99 1.88 Tungsten (NE) Lened KR-05-175 bt monzogranite thick w-z MR; typically 3+ significant resorption surfaces; highly fractured MR + cores; w-z p-resorbed MC; un- to p-z resorbed IC present in ~17% of grains 79 843 97.6 nd 1794 13.7 9.5 0.722 -15.84 1.84 Tungsten (SE) Rifle Range KR-05-143 bt monzogranite thick w-z MR; highly fractured MR; typically 2 significant resorption surfaces; un- to w-z p-resorbed MC (CL-bright); resorbed IC (u-z ± porous; present in ~23% of grains) 80 832 96.9 103 nd 15.2 7.3 0.722 -14.51 1.91 Tay River (SE) Coal River KR-05-196 bt>hbl monzogranite thick p-z non-resorbed MR; pristine to p-resorbed poorly- to u-z MC; IC present in ~8% of grains 75 827 98.4 nd nd 11.7 8.3 0.711 -9.33 1.31 Tay River (Coal River) Jorgensen KR-05-08 pl-qtz-hbl-phyric qtz monzodiorite porphyry thick w-z MR; typically 1+ resorption surface in larger grains; small p-resorbed w-z MC; non-resorbed IC present in ~13% of grains (u-z ± porous) 62 781 95.9 nd 340, 370 16.3 7.4 0.712 -9.65 1.48 Notes: a mineral abbreviations are after Kretz (1983); b abbreviations are: MR = magmatic rims; MC = magmatic cores (autocrystic ± antecrystic); IC = inherited cores; w-z = well-zoned; u-z = unzoned; p-z = poorly-zoned; w = well; p- resorbed = partially resorbed; WR = whole rock; qtz = quartz; nd = no data. c D.I. = Differentiation Index (after Thornton and Tuttle, 1960); d ZST = Zircon Saturation Temperature (after Watson and Harrison, 1983); e calculated from Sm–Nd isotopic compositions. 344 Table 5.2. (next 6 pages) Results of zircon geochemical analysis. All values given in ppm, unless otherwise noted. Pale grey shading indicates analyses with anomalously high light REE that have been attributed to miniscule inclusions of REE-rich accessory minerals in otherwise pristine magmatic rims and cores, and (or) to metamictization in xenocrystic cores. Sub-suites are denoted by: HM = Hess– MacMillan, and CR = Coal River. a) Magmatic rims. b) Magmatic cores (MC; including autocrystic and antecrystic) and xenocrystic (or inherited) cores (IC). 345 a) Analysis Spota Plutonic Suite (sub- suite/region) Descriptionb Morphologyc SiO2d Ti MnO Ga Sr Y Nb 97b-TE-1 Tombstone p-z L1 44.63 14.6 <0.00002 0.33 0.33 414 2.44 97b-TE-10 Tombstone p-z interm-outer L3-S3 41.98 6.24 0.00027 0.89 1.94 960 5.82 97b-TE-11 Tombstone p-z interm-inner S5 42.28 8.87 <0.00002 0.49 0.37 994 4.30 97b-TE-13 Tombstone w-z interm S23 35.98 13.79 <0.000 0.66 <0.26 665 4.22 97b-TE-14 Tombstone w-z interm S23 41.28 7.43 0.00029 0.47 0.41 676 5.67 97b-TE-18 Tombstone w-z J2 45.51 3.30 0.00100 3.39 6.91 812 6.37 97b-TE-19 Tombstone w-z interm S17 46.89 3.06 0.00002 0.67 0.50 1075 4.32 97b-TE-2 Tombstone p-z interm S19-S24 42.18 24.1 <0.00002 0.32 0.42 493 1.95 97b-TE-20 Tombstone w-z P1-P2 37.22 7.00 0.00166 0.35 0.51 1027 3.78 97b-TE-4 Tombstone w-z S19-S24 42.19 7.19 0.00010 0.76 0.63 768 5.93 97b-TE-6 Tombstone w-z S18 38.61 6.06 0.00232 6.11 13.33 723 5.67 97b-TE-7 Tombstone w-z S24-J4 41.14 7.91 0.00132 2.21 5.96 763 6.87 97b-TE-8 Tombstone w-z S25-J5 38.08 4.56 <0.00002 0.47 0.36 696 5.65 36-TE-1 Mayo w-z S18 36.85 4.69 0.00018 0.38 0.50 511 2.32 36-TE-4 Mayo w-z S23-S24 31.88 12.8 0.00001 0.56 <0.42 655 2.97 36-TE-5 Mayo w-z interm S17-S22 35.20 7.95 0.00002 0.67 0.48 1804 2.57 76_TE-1 Mayo (HM) w-z interm S4-S5 57.39 5.88 0.00150 0.42 2.92 3009 14.70 76_TE-2 Mayo (HM) w-z S21-S22 75.71 7.49 0.00659 1.51 10.66 3596 20.88 68-TE-1 Mayo (HM) w-z interm S16-S17 51.79 5.70 0.00005 0.28 0.51 1879 10.56 68-TE-8 Mayo (HM) w-z S1 36.92 17.6 0.00001 0.45 0.31 923 1.68 110-TE-2 Tungsten (NE) w-z interm L2-S2 47.39 2.88 0.00011 0.25 <0.67 2266 11.28 110-TE-2 Tungsten (NE) w-z interm L2-S3 38.78 4.33 0.00005 0.28 0.61 2281 11.23 110-TE-5 Tungsten (NE) w-z interm S4-S5 36.48 8.81 0.00100 0.23 1.19 3481 10.90 175-TE-1 Tungsten (NE) w-z S3-S4 60.86 6.41 0.00044 1.20 1.16 3497 20.39 175-TE-3 Tungsten (NE) w-z S3 56.67 14.8 0.00521 2.72 3.38 3410 19.08 175-TE-6 Tungsten (NE) p-z interm S13 47.75 2.10 <0.00002 0.27 0.29 995 9.96 143-TE-2 Tungsten (SE) p-z interm Q2-Q3 38.62 4.08 0.00002 0.34 0.41 2208 3.16 143-TE-3 Tungsten (SE) w-z interm L3-S3 58.76 4.38 0.00002 0.22 0.64 2956 16.59 143-TE-4 Tungsten (SE) w-z interm S11-S12 41.91 4.38 0.00002 0.33 0.53 2710 10.90 143-TE-5 Tungsten (SE) w-z P2-P3 44.30 27.5 0.00150 <0.20 0.72 2030 8.62 143-TE-8 Tungsten (SE) w-z S9-S10 39.74 6.00 0.00004 0.21 0.48 2052 6.19 143-TE-9 Tungsten (SE) w-z interm S9-S14 39.03 8.15 0.00019 0.22 0.37 1751 5.25 143-TE-11 Tungsten (SE) w-z S4-S5-L5 46.36 25.9 0.00020 0.37 0.82 4984 22.37 143-TE-12 Tungsten (SE) w-z S6-S2-L3 44.54 3.00 0.00055 0.22 0.80 5240 23.05 196-TE-1 Tay River p-z S3 37.23 3.60 0.00065 0.28 0.37 775 3.88 196-TE-2 Tay River w-z interm S23-S18 35.49 4.20 0.00068 0.17 0.46 861 3.48 196-TE-4 Tay River w-z S22 34.88 4.20 0.00073 0.41 0.24 829 4.50 196-TE-6 Tay River w-z (interm) S11-S12 39.28 4.80 0.00078 0.30 0.26 908 4.85 196-TE-7 Tay River w-z S19-S24 38.91 4.80 0.00200 1.37 5.65 845 3.56 196-TE-8 Tay River p-z S3 40.70 5.24 0.00103 0.27 0.39 843 2.93 196-TE-10 Tay River p-z S2-S3 42.69 6.59 0.00010 0.41 0.44 844 4.44 08-TE-10 Tay River (CR) w-z interm S14-S19 40.79 4.68 <0.00002 0.34 <0.64 1489 11.56 08-TE-11 Tay River (CR) w-z interm S6-S7 45.15 8.09 <0.00002 0.57 0.51 1590 12.29 08-TE-2 Tay River (CR) w-z interm S1-S6 44.83 6.47 <0.00002 0.40 <1.10 1193 9.72 08-TE-2 Tay River (CR) w-z interm S1-S6 45.31 8.63 0.00002 0.56 0.43 1364 9.99 08-TE-3 Tay River (CR) w-z S1-S6, S2-S7 41.93 8.33 <0.00002 0.60 <0.73 1327 11.42 08-TE-3 Tay River (CR) w-z S1-S6, S2-S7 37.32 11.6 0.00005 0.72 0.54 1449 10.71 08-TE-5 Tay River (CR) w-z interm S3-S8 39.93 5.52 <0.00002 0.44 <0.64 1617 16.34 08-TE-5 Tay River (CR) w-z interm S3-S8 41.00 4.71 0.00006 0.39 0.54 2006 24.31 08-TE-7 Tay River (CR) w-z S1-S6, S2-S7 41.84 5.88 <0.00002 0.40 <0.84 1345 15.27 08-TE-8 Tay River (CR) u-z P1-P2 33.68 7.01 0.00027 0.57 <0.85 610 3.00 Notes: a sample label in Appendix E.1 (first digits of the label are the last digits of the sample number); b description of zoning; abbreviations are: u-z = unzoned, p-z = poorly zoned, w-z = well-zoned, interm = intermediate zone (between inner and outer zones); c morphological classification after Pupin (1980); d in wt.%; 346 Analysis Spota Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 97b-TE-1 <0.10 0.01 19.1 0.08 1.32 2.41 0.35 10.59 3.21 35.41 12.52 57.9 13.66 135.5 28.5 97b-TE-10 <0.10 7.35 46.1 2.25 10.84 5.62 0.72 22.15 6.72 81.43 29.32 141.8 33.28 345.4 69.9 97b-TE-11 <0.10 0.04 33.1 0.15 2.37 4.80 0.80 23.65 7.49 85.22 29.92 149.4 34.38 354.1 71.7 97b-TE-13 <0.10 0.61 34.0 0.35 3.06 5.35 0.54 20.04 5.50 62.96 21.55 97.0 20.58 204.4 41.4 97b-TE-14 <0.10 0.02 27.5 0.04 1.29 2.65 0.42 14.40 4.63 56.45 21.39 101.1 24.05 268.3 54.2 97b-TE-18 <0.10 43.64 133 13.36 57.34 12.85 0.85 22.71 5.92 69.55 24.55 121.3 28.72 284.3 58.8 97b-TE-19 <0.10 2.68 38.4 0.98 6.36 6.76 0.97 28.55 8.38 98.22 35.58 172.7 40.41 411.5 81.7 97b-TE-2 <0.10 0.83 19.2 0.35 2.73 3.41 0.61 13.29 3.96 43.11 14.88 67.2 14.87 150.5 31.5 97b-TE-20 <0.10 0.04 19.3 0.12 1.96 4.67 0.66 23.62 7.42 92.98 34.55 161.5 36.76 354.4 68.9 97b-TE-4 <0.10 2.80 37.5 0.79 4.17 3.84 0.45 16.63 5.44 63.35 23.85 115.1 26.99 279.9 57.0 97b-TE-6 <0.10 99.22 267 28.42 120.0 23.96 1.51 28.45 5.85 60.74 22.82 110.1 25.47 255.0 51.6 97b-TE-7 <0.10 37.24 105 8.27 31.85 7.72 0.98 19.83 5.34 63.58 23.86 115.0 26.53 265.5 55.4 97b-TE-8 <0.10 0.64 27.5 0.27 1.92 2.79 0.32 13.80 4.16 56.29 21.55 101.7 23.41 260.6 54.2 36-TE-1 <0.31 1.37 23.5 0.60 3.35 2.06 0.26 9.83 3.09 39.93 14.71 72.1 17.61 179.1 39.1 36-TE-4 <0.47 0.04 29.3 0.05 1.21 2.85 0.55 14.54 4.62 55.01 21.75 102.3 22.32 225.5 47.5 36-TE-5 <0.29 0.66 31.4 0.54 8.00 14.06 4.23 57.80 15.62 178.54 58.36 252.8 50.61 463.3 93.4 76_TE-1 <0.10 2.59 16.6 1.55 9.38 7.24 0.67 40.79 17.88 259.80 89.34 416.4 92.29 814.2 153.9 76_TE-2 <0.10 10.64 46.5 5.62 33.74 18.22 1.66 73.30 25.51 347.27 135.1 627.5 128.5 1138 209.7 68-TE-1 <0.10 1.02 15.5 0.53 4.53 6.60 0.99 38.83 13.45 168.74 62.53 299.2 64.26 555.4 102.9 68-TE-8 <0.14 0.02 22.6 0.12 1.86 4.16 0.79 21.33 6.57 81.71 30.34 140.8 30.72 286.8 54.2 110-TE-2 <0.20 0.45 14.4 0.42 3.37 5.38 0.55 34.61 13.17 192.22 77.02 365.8 83.51 819.4 172.5 110-TE-2 <0.20 0.25 13.9 0.25 2.15 5.23 0.62 35.70 13.81 196.46 78.96 396.2 89.00 861.6 164.4 110-TE-5 <0.10 0.63 6.59 0.33 2.51 4.27 0.78 31.20 13.29 214.03 94.62 522.5 134.8 1264 276.8 175-TE-1 <0.10 11.95 42.1 4.49 23.40 11.21 2.69 45.23 18.44 272.17 110.8 583.1 135.1 1367 269.9 175-TE-3 <0.10 24.76 79.3 13.22 67.90 26.58 5.66 61.74 20.76 289.64 111.4 563.1 131.1 1302 253.2 175-TE-6 <0.10 0.23 10.5 0.16 1.97 7.43 1.44 46.51 12.79 122.41 28.52 93.3 17.05 153.6 27.9 143-TE-2 <0.10 0.03 15.0 0.29 5.51 9.78 1.61 49.59 15.21 205.73 80.38 376.7 80.31 742.8 148.1 143-TE-3 <0.10 0.01 12.6 0.04 0.73 3.89 0.45 34.53 14.90 236.70 99.19 491.1 115.5 1138 228.2 143-TE-4 <0.10 0.01 18.0 0.11 1.92 5.87 0.80 42.52 15.47 226.96 93.03 473.5 100.8 989.0 198.7 143-TE-5 <0.10 0.01 6.04 0.03 0.59 4.00 0.33 25.18 11.87 189.36 82.91 419.6 95.65 978.1 188.6 143-TE-8 <0.10 0.02 12.9 0.11 1.94 4.81 0.69 27.88 10.83 164.76 73.26 379.6 84.36 794.3 155.6 143-TE-9 <0.10 0.01 12.8 0.12 2.17 4.66 0.71 28.37 9.96 144.54 58.77 296.2 68.92 663.2 127.5 143-TE-11 <0.10 0.48 9.77 0.24 2.24 7.87 0.48 61.24 28.17 424.67 170.3 837.9 183.7 1747 318.1 143-TE-12 <0.10 0.01 11.8 0.07 1.63 7.57 0.45 67.75 30.69 459.33 180.0 870.3 188.7 1801 320.7 196-TE-1 <0.10 0.00 22.2 0.03 0.76 1.97 0.46 11.89 4.40 58.35 24.13 126.1 30.31 317.5 67.8 196-TE-2 <0.10 0.03 13.9 0.05 0.90 2.22 0.59 14.89 4.81 70.40 28.31 141.6 34.51 368.5 76.8 196-TE-4 <0.10 0.01 23.1 0.05 0.88 2.15 0.53 13.54 4.62 66.01 28.11 147.7 33.65 356.4 76.1 196-TE-6 <0.10 0.05 18.3 0.06 1.08 2.47 0.54 13.90 4.82 70.18 28.31 151.4 35.92 369.8 78.0 196-TE-7 <0.10 20.34 62.5 4.93 19.50 5.39 0.87 17.26 5.21 68.53 27.81 144.1 33.81 347.3 72.4 196-TE-8 <0.10 0.17 14.7 0.09 1.05 2.12 0.57 13.77 4.62 62.54 25.52 133.5 33.09 333.1 67.8 196-TE-10 <0.10 0.01 21.7 0.05 1.05 2.56 0.53 13.32 4.71 64.58 26.57 139.4 33.46 342.9 72.7 08-TE-10 <0.10 0.03 27.1 0.06 1.10 2.80 1.32 21.11 7.60 105.86 46.98 259.6 62.33 624.4 139.8 08-TE-11 <0.10 0.07 28.4 0.06 0.96 2.99 1.37 21.45 7.89 115.14 47.01 250.8 60.25 637.7 142.1 08-TE-2 <0.10 0.07 28.8 0.07 1.27 2.50 1.32 17.32 6.00 85.72 35.97 188.3 47.58 467.7 113.2 08-TE-2 <0.20 0.05 29.3 0.07 1.41 2.92 1.30 18.83 6.95 98.23 39.19 212.5 52.60 567.4 130.5 08-TE-3 <0.10 0.05 40.9 0.06 1.87 3.42 1.54 22.14 7.38 105.39 43.75 221.1 52.94 528.1 122.5 08-TE-3 <0.22 0.09 34.0 0.11 1.42 3.45 1.29 19.61 7.82 111.94 44.60 240.5 58.80 608.4 135.1 08-TE-5 <0.10 0.01 23.2 0.04 0.83 2.46 1.32 21.04 7.72 118.66 50.48 271.3 67.05 676.0 153.8 08-TE-5 <0.14 0.02 21.1 0.04 0.68 2.70 1.40 22.33 9.89 148.69 62.92 348.5 83.53 883.3 183.2 08-TE-7 <0.10 0.07 31.9 0.04 1.16 3.15 1.26 21.53 7.40 96.26 41.92 217.5 54.16 514.6 115.7 08-TE-8 <0.10 0.23 15.3 0.17 1.33 2.10 0.56 9.78 2.64 43.12 19.31 103.8 28.55 272.3 67.7 347 Analysis Spota Hf Ta Pb Th U Th/U Yb/Gd Lu/Hf Ce/Sm SmCNe EuCNe GdCNe Eu/Eu*e LaCNe 97b-TE-1 11826 1.36 3.14 305 743 0.41 12.80 0.0024 7.93 15.27 5.29 52.20 0.19 -0.96 97b-TE-10 13100 3.86 6.79 524 2044 0.26 15.60 0.0053 8.20 36.94 11.77 110.28 0.18 30.00 97b-TE-11 13260 2.52 6.05 639 1847 0.35 14.97 0.0054 6.89 31.41 13.17 117.83 0.22 -0.83 97b-TE-13 12533 2.13 6.69 637 1247 0.51 10.20 0.0033 6.36 35.13 8.50 99.69 0.14 1.58 97b-TE-14 12751 3.54 5.59 475 1973 0.24 18.64 0.0042 10.39 16.89 6.46 71.34 0.19 -0.92 97b-TE-18 13705 4.02 6.19 504 2108 0.24 12.52 0.0043 10.35 85.83 14.01 113.14 0.14 183.1 97b-TE-19 14844 2.46 9.23 794 2446 0.32 14.41 0.0055 5.69 44.64 16.30 142.49 0.20 10.29 97b-TE-2 11346 1.07 2.88 286 555 0.52 11.33 0.0028 5.63 22.07 9.80 65.76 0.26 2.50 97b-TE-20 13015 2.49 4.97 475 1148 0.41 15.00 0.0053 4.13 30.55 10.63 117.69 0.18 -0.83 97b-TE-4 13138 3.48 5.20 543 2101 0.26 16.83 0.0043 9.77 24.97 6.98 82.55 0.15 10.82 97b-TE-6 12999 3.61 4.63 435 1820 0.24 8.97 0.0040 11.16 160.91 25.75 141.94 0.17 417.7 97b-TE-7 13561 3.96 30.66 520 2112 0.25 13.39 0.0041 13.66 51.15 16.32 98.62 0.23 156.1 97b-TE-8 14087 3.78 5.53 460 2273 0.20 18.89 0.0038 9.86 17.84 4.63 68.33 0.13 1.68 36-TE-1 12813 1.46 3.58 304 1123 0.27 18.23 0.0030 11.42 12.91 3.55 48.38 0.14 4.79 36-TE-4 12497 1.81 4.33 444 980 0.45 15.51 0.0038 10.29 18.26 8.73 72.04 0.24 -0.83 36-TE-5 10364 1.08 4.20 502 615 0.82 8.01 0.0090 2.23 93.99 74.10 289.45 0.45 1.77 76_TE-1 14285 6.88 5.14 465 3168 0.15 19.96 0.0108 2.30 47.92 10.97 203.98 0.11 9.92 76_TE-2 14014 10.04 8.01 800 5294 0.15 15.53 0.0150 2.55 122.10 28.48 367.36 0.13 43.90 68-TE-1 14597 3.77 4.30 410 1867 0.22 14.30 0.0070 2.34 43.57 16.66 194.15 0.18 3.30 68-TE-8 11452 0.68 21.57 120 159 0.75 13.45 0.0047 5.43 27.13 13.00 106.20 0.24 -0.91 110-TE-2 14024 6.51 6.17 544 2520 0.22 23.68 0.0123 2.68 35.33 8.84 172.92 0.11 0.91 110-TE-2 14024 5.91 6.15 622 2749 0.23 24.13 0.0117 2.66 34.32 9.99 178.41 0.13 0.04 110-TE-5 12723 7.85 5.16 275 1850 0.15 40.50 0.0218 1.54 27.86 12.80 155.80 0.19 1.64 175-TE-1 14943 15.31 9.71 904 7755 0.12 30.22 0.0181 3.75 74.74 46.76 226.28 0.36 49.44 175-TE-3 15965 14.64 14.88 778 7631 0.10 21.09 0.0159 2.98 178.61 99.50 309.23 0.42 103.5 175-TE-6 13387 3.60 3.40 321 1954 0.16 3.30 0.0021 1.41 49.17 24.56 232.73 0.23 -0.03 143-TE-2 10490 1.39 3.82 380 704 0.54 14.98 0.0141 1.53 65.05 27.63 248.18 0.22 -0.87 143-TE-3 15013 13.71 5.77 449 3668 0.12 32.95 0.0152 3.23 25.26 7.03 172.51 0.11 -0.97 143-TE-4 12297 4.14 4.86 461 1573 0.29 23.26 0.0162 3.07 38.66 13.17 212.67 0.15 -0.95 143-TE-5 15658 6.42 3.05 176 1990 0.09 38.84 0.0120 1.51 26.02 4.79 125.55 0.08 -0.95 143-TE-8 11364 2.59 2.47 228 776 0.29 28.49 0.0137 2.68 31.47 11.29 139.10 0.17 -0.92 143-TE-9 11204 2.34 2.76 252 843 0.30 23.38 0.0114 2.75 30.47 11.56 141.56 0.18 -0.95 143-TE-11 15345 11.93 6.30 489 6155 0.08 28.53 0.0207 1.24 52.16 7.47 306.74 0.06 1.04 143-TE-12 14859 12.98 6.08 642 6546 0.10 26.58 0.0216 1.55 50.18 6.94 339.46 0.05 -0.95 196-TE-1 11821 2.36 3.06 322 632 0.51 26.70 0.0057 11.25 12.32 7.15 58.76 0.27 -0.98 196-TE-2 12459 1.98 2.48 224 568 0.40 24.75 0.0062 6.25 13.97 9.52 73.82 0.30 -0.86 196-TE-4 11764 2.42 3.00 302 589 0.51 26.32 0.0065 10.75 13.53 8.33 67.05 0.28 -0.97 196-TE-6 11469 2.29 2.57 245 547 0.45 26.61 0.0068 7.42 15.67 8.61 68.85 0.26 -0.77 196-TE-7 11477 1.99 1.98 200 397 0.50 20.13 0.0063 11.60 35.41 14.38 85.72 0.26 84.84 196-TE-8 11994 1.56 2.01 184 372 0.50 24.19 0.0057 6.92 13.30 9.07 68.19 0.30 -0.30 196-TE-10 11531 2.24 2.12 209 442 0.47 25.75 0.0063 8.49 16.29 8.38 65.91 0.26 -0.97 08-TE-10 12612 3.64 2.60 238 786 0.30 29.58 0.0111 9.69 17.93 22.39 105.07 0.52 -0.89 08-TE-11 12215 3.70 2.60 255 895 0.28 29.73 0.0116 9.49 19.20 23.26 106.81 0.51 -0.70 08-TE-2 12934 3.68 3.70 426 897 0.48 27.01 0.0088 11.55 15.86 22.48 86.02 0.61 -0.72 08-TE-2 12934 3.44 2.72 296 785 0.38 30.13 0.0101 10.05 18.72 22.09 93.63 0.53 -0.81 08-TE-3 12496 3.97 4.62 545 985 0.55 23.85 0.0098 11.98 22.09 26.42 110.27 0.54 -0.80 08-TE-3 12496 3.39 3.17 350 749 0.47 31.02 0.0108 9.87 22.30 21.88 97.55 0.47 -0.62 08-TE-5 13195 5.33 2.53 267 967 0.28 32.13 0.0117 9.44 15.64 22.48 104.71 0.56 -0.97 08-TE-5 13195 7.68 2.46 250 1428 0.18 39.55 0.0139 7.81 17.23 23.78 111.23 0.54 -0.93 08-TE-7 13526 4.65 5.55 598 1310 0.46 23.90 0.0086 10.11 20.29 21.36 107.19 0.46 -0.70 08-TE-8 10248 1.10 3.87 153 330 0.46 27.84 0.0066 7.27 13.18 8.93 48.15 0.35 -0.04 Notes (cont.): e normalized to the chondritic composition of Sun and McDonough (1989); 348 Analysis Spota CeCNe PrCNe Ce/Ce*e LaCN/YbCNe T (°C)f 97b-TE-1 30.15 -0.09 100.05 0.00005 820 97b-TE-10 74.12 23.26 2.81 0.01526 735 97b-TE-11 52.93 0.62 / 0.00008 769 97b-TE-13 54.44 2.79 25.90 0.00215 814 97b-TE-14 43.87 -0.57 60.78 0.00005 751 97b-TE-18 216.0 143.0 1.33 0.11009 680 97b-TE-19 61.67 9.56 6.22 0.00466 673 97b-TE-2 30.36 2.76 11.56 0.00395 877 97b-TE-20 30.44 0.29 / 0.00008 746 97b-TE-4 60.24 7.53 6.67 0.00718 748 97b-TE-6 435.2 305.3 1.22 0.27906 732 97b-TE-7 171.0 88.12 1.46 0.10060 757 97b-TE-8 43.85 1.86 24.81 0.00175 707 36-TE-1 37.33 5.47 7.29 0.00550 709 36-TE-4 46.82 -0.52 71.74 0.00013 806 36-TE-5 50.15 4.84 17.14 0.00102 758 76_TE-1 26.16 15.72 2.09 0.00228 730 76_TE-2 74.88 59.60 1.46 0.00670 752 68-TE-1 24.22 4.73 6.13 0.00131 727 68-TE-8 35.87 0.33 / 0.00005 841 110-TE-2 22.54 3.57 12.53 0.00040 669 110-TE-2 21.66 1.74 80.01 0.00021 703 110-TE-5 9.74 2.55 4.77 0.00035 768 175-TE-1 67.60 47.41 1.40 0.00627 738 175-TE-3 128.3 141.4 1.06 0.01364 821 175-TE-6 16.12 0.71 / 0.00107 644 143-TE-2 23.42 2.11 / 0.00003 697 143-TE-3 19.50 -0.59 25.76 0.00000 703 143-TE-4 28.42 0.15 / 0.00001 703 143-TE-5 8.86 -0.70 10.90 0.00001 893 143-TE-8 20.03 0.22 / 0.00002 731 143-TE-9 19.90 0.27 / 0.00001 760 143-TE-11 14.94 1.54 11.80 0.00020 885 143-TE-12 18.20 -0.27 36.21 0.00001 672 196-TE-1 35.20 -0.64 44.22 0.00001 687 196-TE-2 21.60 -0.43 35.63 0.00007 700 196-TE-4 36.70 -0.43 57.01 0.00002 700 196-TE-6 28.88 -0.32 58.00 0.00010 711 196-TE-7 101.0 52.16 1.52 0.04201 711 196-TE-8 22.90 -0.01 / 0.00036 719 196-TE-10 34.42 -0.43 53.36 0.00001 740 08-TE-10 43.28 -0.34 78.54 0.00003 709 08-TE-11 45.28 -0.36 89.64 0.00008 760 08-TE-2 46.00 -0.26 107.25 0.00010 738 08-TE-2 46.84 -0.29 96.85 0.00006 766 08-TE-3 65.75 -0.39 118.23 0.00006 762 08-TE-3 54.54 0.16 / 0.00010 796 08-TE-5 36.92 -0.53 51.57 0.00001 724 08-TE-5 33.37 -0.56 46.36 0.00001 710 08-TE-7 50.98 -0.61 77.86 0.00010 730 08-TE-8 23.89 0.79 / 0.00060 746 Notes (cont.): f titanium–in–zircon temperature after Claiborne et al. (2010). 349 b) Analysis Spota Core Plutonic Suite (sub-suite/region) Descriptionb Morphologyc SiO2d Ti MnO Ga Sr Y Nb Ba La Ce Pr Nd Sm 97b-TE-12 MC Tombstone u-z S25-J5 35.43 8.99 0.00178 4.23 11.02 966 3.42 <0.10 64.57 180.5 18.82 80.41 20.19 97b-TE-15 MC Tombstone w-z S17-S22 37.64 13.01 0.00008 0.34 0.26 613 2.88 <0.10 0.06 21.67 0.11 1.93 3.68 97b-TE-16 MC Tombstone u-z S13 32.60 5.88 0.00002 0.54 0.24 640 3.82 <0.10 0.00 25.68 0.08 1.77 3.44 97b-TE-5 MC Tombstone u-z S20-S25 38.04 39.75 0.00097 1.90 4.47 1316 3.45 <0.10 17.39 65.39 4.54 21.98 14.59 97b-TE-9 MC Tombstone u-z S10 36.41 5.88 0.00241 5.50 10.64 1161 3.88 <0.10 76.33 221.8 23.43 103.6 25.81 36-TE-2 MC Mayo p-z S23-S24 27.16 14.19 0.00006 0.58 <0.35 1495 2.93 <0.41 0.11 33.21 0.23 3.77 8.01 36-TE-3 MC Mayo w-z interm S22-S23 37.65 11.89 0.00200 4.62 13.93 555 2.35 <0.17 100.3 215.0 20.16 78.39 15.95 76_TE-4 MC Mayo (HM) u-z S3-S8 41.71 28.60 0.00173 0.21 0.33 321 1.54 <0.10 0.08 5.60 0.46 7.43 11.68 68-TE-3 MC Mayo (HM) resorbed P1 33.63 9.65 0.00002 0.37 0.30 2063 4.75 <0.10 0.02 14.44 0.46 7.45 13.84 68-TE-6 MC Mayo (HM) w-z S24 40.38 2.88 0.00017 <0.17 0.47 1308 3.90 <0.10 0.20 3.35 0.14 1.44 3.19 68-TE-7 MC-R Mayo (HM) w-z P1 (G1-P2) 37.09 7.31 0.00022 0.20 0.91 1647 3.98 0.29 0.71 8.37 0.46 4.30 6.74 110-TE-1 MC Tungsten (NE) p-z L4 41.62 13.61 <0.00002 0.56 <1.02 2161 2.93 <0.17 0.05 30.61 0.56 8.42 14.66 110-TE-1 MC Tungsten (NE) p-z L4 38.91 14.16 0.00004 0.51 0.47 1757 2.28 <0.17 0.05 22.40 0.44 7.30 10.89 110-TE-4 MC Tungsten (NE) p-z L4-L5 32.96 1.68 0.00037 0.15 6.90 2350 5.10 <0.10 0.05 5.23 0.04 0.63 2.63 175-TE-2 MC? Tungsten (NE) p-z P2-P1 54.05 35.49 0.00243 5.18 3.67 3442 28.89 <0.10 78.68 163.6 27.13 135.2 44.38 175-TE-4 MC Tungsten (NE) u-z S2? 32.63 6.29 0.00005 0.29 0.31 1017 1.84 <0.10 0.02 13.01 0.26 3.90 6.21 196-TE-3 MC Tay River w-z S19-S20-S24-S25 38.81 4.20 0.00073 0.25 0.30 900 4.24 <0.10 0.06 18.46 0.07 1.15 2.31 08-TE-1 MC Tay River (CR) p-z S14-S19 41.29 4.14 0.00019 0.66 <0.67 2709 21.94 <0.10 0.17 46.19 0.39 6.20 11.34 08-TE-1 MC Tay River (CR) p-z S14-S19 35.70 4.75 0.00009 0.92 0.41 1701 22.09 <0.20 0.05 45.63 0.21 3.63 6.80 08-TE-12 MC Tay River (CR) p-z ? S21-S22 48.06 8.15 <0.00002 0.66 0.46 1310 6.04 <0.10 0.00 31.77 0.09 1.91 3.97 97b-TE-17 IC Tombstone resorbed n/a 43.24 79.52 0.00700 16.61 50.17 2372 16.08 <0.10 205.3 701.4 65.67 293.2 83.16 97b-TE-3 IC Tombstone resorbed L2? 36.27 30.76 0.00030 0.93 1.12 1075 4.53 <0.10 0.68 39.66 0.41 4.72 6.65 76_TE-3 IC Mayo (HM) resorbed n/a 36.07 13.13 0.00049 0.53 3.48 1104 5.28 <0.10 0.93 23.15 0.82 6.26 7.71 68-TE-2 IC Mayo (HM) resorbed n/a 44.01 5.76 0.00002 0.16 0.32 492 2.61 <0.10 0.02 9.92 0.10 1.55 2.88 68-TE-9 IC Mayo (HM) resorbed n/a 40.73 7.96 0.00021 0.93 1.80 2655 6.93 0.29 3.15 34.49 1.65 12.92 15.29 68-TE-10 IC? Mayo (HM) p-z L2-S2 38.72 10.50 0.00003 0.15 0.44 1709 11.23 <0.16 0.23 8.26 0.19 1.77 4.28 175-TE-5 IC Tungsten (NE) u-z L1-L2, S1-S2 40.16 10.19 0.00011 0.40 1.12 1040 3.31 <0.10 0.03 20.42 0.14 2.08 4.07 143-TE-1 IC? Tungsten (SE) p-z L4-S4? 34.34 3.90 0.00200 0.18 29.57 926 3.64 <0.10 0.77 2.83 0.23 1.10 1.14 143-TE-7 IC Tungsten (SE) p-z L3-S3-S8 39.19 14.35 0.00198 0.12 0.52 1452 3.07 <0.10 0.00 7.95 0.09 1.75 3.71 143-TE-10 IC Tungsten (SE) resorbed n/a 48.00 1089 0.04800 6.94 3.96 1865 15.96 <0.10 2.10 13.00 1.37 10.99 9.99 143-TE-13 IC Tungsten (SE) u-z Q2-Q3 29.46 3.00 0.00051 0.06 0.32 2091 8.06 <0.10 0.00 4.86 0.03 0.78 3.32 143-TE-14 IC Tungsten (SE) u-z Q2-Q3 41.03 152.3 0.00170 0.64 2.11 1832 2.93 <0.10 5.27 18.83 1.65 10.29 9.47 196-TE-5 IC? Tay River resorbed J3? 33.04 27.58 0.00150 0.83 2.12 977 3.48 <0.10 4.10 26.73 1.05 6.21 4.74 196-TE-9 IC Tay River resorbed n/a 39.39 8.39 0.00020 0.12 0.58 816 2.84 <0.10 1.23 9.04 0.41 2.34 2.42 08-TE-4 IC Tay River (CR) resorbed n/a 45.57 54.38 0.00364 0.77 8.99 1702 16.17 <0.10 2.01 17.44 0.62 3.18 2.97 08-TE-6 IC Tay River (CR) resorbed n/a 34.74 3.66 0.00013 0.31 <0.61 1092 19.75 <0.10 0.03 23.98 0.05 0.94 1.96 08-TE-9 MR Tay River (CR) u-z P1-P2 27.89 4.14 <0.00002 0.26 <0.77 706 3.02 <0.10 0.62 13.64 0.16 1.17 1.30 Notes: a sample label in Appendix A.4 (first two digits of label are from the sample number); b description of zoning; abbreviations are: u-z = unzoned, p-z = poorly zoned, w-z = well-zoned, interm = intermediate zone (between inner and outer zones); c morphological classification after Pupin (1980); d in wt.%; 350 Analysis Spota Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U Th/U Yb/Gd Lu/Hf Ce/Sm SmCNe EuCNe GdCNe 97b-TE-12 1.77 36.39 8.40 97.3 32.03 149.3 34.15 304.3 61.1 12798 2.33 6.81 713 1717 0.42 8.36 0.0048 8.94 135.43 30.37 181.88 97b-TE-15 0.57 15.37 4.51 53.0 18.48 88.3 20.18 197.1 39.9 10980 1.60 4.41 395 817 0.48 12.83 0.0036 5.89 23.84 9.16 76.22 97b-TE-16 0.70 15.08 4.57 56.3 20.95 95.2 21.43 216.1 44.0 10062 2.11 4.65 585 1315 0.44 14.34 0.0044 7.48 22.21 11.47 74.75 97b-TE-5 2.10 46.95 11.80 127.5 40.44 190.8 39.64 357.1 72.5 10830 1.77 11.18 1260 1609 0.78 7.61 0.0067 4.48 97.60 36.37 234.94 97b-TE-9 2.41 41.44 9.93 111.8 37.88 172.4 37.62 381.4 77.1 11585 2.37 10.31 652 2022 0.32 9.20 0.0067 8.60 173.37 41.84 207.24 36-TE-2 1.12 34.73 10.47 134.4 52.43 228.6 51.03 468.2 99.9 12974 1.93 7.50 1002 2178 0.46 13.48 0.0077 4.15 53.13 18.93 173.52 36-TE-3 1.95 21.88 4.79 51.3 17.05 82.2 18.01 177.2 36.8 13015 1.33 3.38 349 660 0.53 8.10 0.0028 13.48 106.74 33.69 108.95 76_TE-4 1.17 40.27 7.72 53.1 9.04 21.8 2.98 21.8 3.6 11667 0.49 1.99 151 497 0.30 0.54 0.0003 0.48 77.93 19.78 201.34 68-TE-3 1.32 59.37 17.43 216.3 72.93 316.7 64.66 584.2 107.2 9427 1.87 4.52 507 835 0.61 9.84 0.0114 1.04 92.51 22.52 297.34 68-TE-6 0.54 23.80 8.45 122.2 47.98 229.1 49.87 463.1 87.6 12526 1.85 2.51 184 1380 0.13 19.46 0.0070 1.05 20.58 8.63 118.59 68-TE-7 0.98 34.00 11.05 147.6 53.40 250.2 54.42 499.4 94.1 11766 1.94 3.00 280 1304 0.22 14.69 0.0080 1.24 44.53 16.39 169.85 110-TE-1 4.93 58.78 16.58 201.6 75.84 364.8 71.24 735.9 145.2 10050 1.13 6.22 823 784 1.05 12.52 0.0144 2.09 98.05 86.55 294.36 110-TE-1 3.75 47.21 13.60 157.5 53.68 255.1 58.13 559.3 116.4 10050 0.82 4.97 542 609 0.89 11.85 0.0116 2.06 72.55 65.68 236.24 110-TE-4 0.52 26.07 10.68 165.2 73.14 374.2 81.97 842.4 180.4 11882 3.46 3.90 327 2356 0.14 32.31 0.0152 1.99 16.80 8.31 130.02 175-TE-2 11.98 88.21 24.79 308.9 108.4 512.3 113.58 1106.9 214.2 14441 11.12 71.48 2397 5706 0.42 12.55 0.0148 3.69 298.87 211.84 442.25 175-TE-4 1.96 26.76 7.86 89.7 31.38 146.9 32.56 310.6 64.6 9093 0.75 2.98 276 481 0.57 11.61 0.0071 2.10 40.95 33.87 133.49 196-TE-3 0.59 13.12 4.82 65.7 27.32 138.8 32.42 337.0 69.2 11798 1.81 2.46 256 488 0.53 25.69 0.0059 8.00 14.58 9.46 64.92 08-TE-1 1.16 62.08 20.67 260.9 98.61 429.6 85.78 677.3 118.9 7016 6.12 9.19 223 260 0.86 10.91 0.0169 4.07 75.64 19.59 310.96 08-TE-1 0.67 37.89 13.56 180.5 66.04 303.3 59.74 516.1 87.1 7016 8.06 9.65 222 293 0.76 13.62 0.0124 6.71 44.92 10.92 189.39 08-TE-12 1.47 20.61 7.15 100.2 40.71 209.6 51.25 550.0 123.5 12752 2.23 3.03 287 716 0.40 26.69 0.0097 8.01 25.81 25.06 102.55 97b-TE-17 7.59 128.5 25.51 249.9 76.35 316.0 63.29 592.1 122.4 9193 4.28 220.5 855 4809 0.18 4.61 0.0133 8.43 560.86 133.81 644.59 97b-TE-3 1.24 27.28 7.78 92.4 33.39 155.2 33.36 320.9 64.3 10214 2.26 13.50 1165 1157 1.01 11.77 0.0063 5.96 43.96 21.04 136.07 76_TE-3 2.81 31.80 9.77 109.3 36.30 168.3 36.19 346.1 64.4 9517 3.22 17.35 273 1111 0.25 10.88 0.0068 3.00 51.11 48.88 158.81 68-TE-2 0.18 12.74 3.89 49.5 17.54 83.1 17.56 162.7 31.5 11404 0.74 8.42 34 67 0.51 12.78 0.0028 3.45 18.43 2.14 62.99 68-TE-9 2.01 65.85 20.95 257.2 89.22 410.5 82.90 731.4 140.6 12985 2.76 11.87 1166 1858 0.63 11.11 0.0108 2.26 102.34 34.75 329.91 68-TE-10 0.69 32.30 11.98 160.3 60.96 283.0 59.97 556.1 104.0 12608 4.53 5.71 604 2698 0.22 17.22 0.0082 1.93 27.93 11.31 161.30 175-TE-5 0.87 21.05 6.88 89.7 33.78 161.7 34.65 347.5 66.3 11374 1.72 12.52 145 494 0.29 16.51 0.0058 5.02 26.50 14.47 104.80 143-TE-1 0.17 9.01 4.36 73.2 30.79 167.6 40.09 394.2 80.0 11557 3.23 4.36 81 916 0.09 43.75 0.0069 2.48 6.70 1.93 44.28 143-TE-7 0.63 24.06 8.72 124.3 52.78 266.5 57.37 571.0 112.9 11120 1.48 1.86 170 558 0.31 23.73 0.0102 2.14 24.05 10.10 119.92 143-TE-10 4.55 43.70 13.38 165.3 59.05 286.8 67.52 644.4 131.6 14958 22.79 10.90 412 1349 0.31 14.75 0.0088 1.30 66.52 79.76 218.58 143-TE-13 0.32 25.04 10.81 161.0 66.60 337.6 73.47 683.0 126.4 10287 4.30 1.92 188 1905 0.10 27.28 0.0123 1.46 21.46 4.65 124.83 143-TE-14 2.22 40.23 12.83 168.6 63.79 301.4 66.00 618.2 114.3 11207 1.22 2.87 208 495 0.42 15.37 0.0102 1.99 62.99 38.43 201.18 196-TE-5 1.01 20.43 6.20 80.5 30.77 154.0 35.18 349.9 72.6 8697 1.24 3.46 297 399 0.74 17.13 0.0083 5.64 31.01 16.99 101.65 196-TE-9 0.73 12.34 5.07 69.0 26.62 137.4 33.76 352.5 69.6 12818 2.00 8.65 146 682 0.21 28.58 0.0054 3.73 15.36 11.97 60.98 08-TE-4 1.57 19.85 8.31 114.6 50.71 302.5 80.18 865.1 204.3 13986 10.48 8.68 362 1297 0.28 43.58 0.0146 5.87 19.07 26.89 98.75 08-TE-6 0.88 14.85 5.66 86.1 34.30 182.6 43.97 427.2 88.9 7787 16.88 10.69 317 697 0.46 28.78 0.0114 12.25 12.22 14.67 73.61 08-TE-9 0.67 10.00 3.38 46.0 21.78 122.0 30.80 293.5 70.8 10007 1.09 1.27 125 367 0.34 29.37 0.0071 10.47 7.80 10.92 49.23 Notes (cont.): e normalized to the chondritic composition of Sun and McDonough (1989); 351 Analysis Spota Eu/Eu*e LaCNe CeCNe PrCNe Ce/Ce*e LaCN/YbCNe T (°C)f 97b-TE-12 0.19 271.4 293.5 201.8 1.25 0.15218 770 97b-TE-15 0.21 -0.77 34.36 0.20 / 0.00020 807 97b-TE-16 0.28 -0.99 40.89 -0.19 93.97 0.00001 730 97b-TE-5 0.24 72.38 105.7 47.94 1.79 0.03493 940 97b-TE-9 0.22 321.1 360.9 251.5 1.27 0.14354 730 36-TE-2 0.20 -0.52 53.18 1.51 / 0.00017 817 36-TE-3 0.31 422.4 349.7 216.3 1.16 0.40618 798 76_TE-4 0.16 -0.67 8.14 3.90 / 0.00260 898 68-TE-3 0.14 -0.90 22.56 3.97 / 0.00003 777 68-TE-6 0.17 -0.15 4.47 0.45 / 0.00031 669 68-TE-7 0.19 1.98 12.66 3.95 4.52 0.00102 750 110-TE-1 0.51 -0.77 48.93 5.08 / 0.00005 812 110-TE-1 0.50 -0.78 35.53 3.78 / 0.00007 817 110-TE-4 0.18 -0.81 7.53 -0.57 11.13 0.00004 627 175-TE-2 0.58 331.0 265.9 291.3 0.86 0.05099 925 175-TE-4 0.46 -0.92 20.22 1.84 0.00005 736 196-TE-3 0.31 -0.74 29.11 -0.26 66.70 0.00013 700 08-TE-1 0.13 -0.30 74.35 3.17 / 0.00017 699 08-TE-1 0.12 -0.78 73.43 1.25 / 0.00007 711 08-TE-12 0.49 -0.99 50.82 -0.03 293.82 0.00000 760 97b-TE-17 0.22 865.2 1143 706.6 1.46 0.24871 1040 97b-TE-3 0.27 1.87 63.69 3.39 25.32 0.00152 907 76_TE-3 0.54 2.92 36.76 7.87 7.66 0.00193 808 68-TE-2 0.06 -0.92 15.18 0.08 / 0.00008 728 68-TE-9 0.19 12.28 55.27 16.77 3.85 0.00309 758 68-TE-10 0.17 -0.04 12.47 1.04 / 0.00029 785 175-TE-5 0.27 -0.86 32.31 0.47 / 0.00007 782 143-TE-1 0.11 2.24 3.61 1.46 2.00 0.00140 694 143-TE-7 0.19 -0.99 11.96 -0.04 59.37 0.00000 818 143-TE-10 0.66 7.87 20.21 13.81 1.94 0.00234 1632 143-TE-13 0.09 -0.98 6.93 -0.63 8.78 0.00000 672 143-TE-14 0.34 21.25 29.72 16.78 1.57 0.00612 1150 196-TE-5 0.30 16.32 42.60 10.26 3.29 0.00841 893 196-TE-9 0.39 4.18 13.75 3.41 3.65 0.00250 763 08-TE-4 0.62 7.46 27.44 5.63 4.24 0.00166 983 08-TE-6 0.49 -0.88 38.12 -0.49 57.78 0.00005 688 08-TE-9 0.56 1.61 21.25 0.76 19.27 0.00151 699 Notes (cont.): f titanium–in–zircon temperature after Claiborne et al. (2010). 352 Table 5.3. Results of zircon hafnium isotopic analysis. Sub-suites are denoted by: HM = Hess–MacMillan, and CR = Coal River. Analysis Spota Zoneb Plutonic Suite (sub- suite/region) Descriptionc Morphologyd Hf (ppm)e Lu (ppm)f 176Hf/ 177Hfg ± 2-σ 176Lu/ 177Hfg ± 2-σ 176Yb/ 177Hfg ± 2-σ 178Hf/ 177Hfg ± 2-σ εHf0 εHfih εHf100h TDM (Ga)i 97b-Hf-1 MR Tombstone p-z L1 11611 23.3 0.282185 0.00005 0.00028 0.00000 0.00928 0.00023 1.46719 0.00005 -21.2 -19.0 -19.0 1.48 97b-Hf-2 MR Tombstone p-z interm S19-S24 11651 16.1 0.282196 0.00005 0.00020 0.00001 0.00626 0.00029 1.46719 0.00006 -20.8 -18.6 -18.6 1.46 97b-Hf-3 MR Tombstone w-z S19-S24 14193 45.3 0.282197 0.00004 0.00045 0.00000 0.01473 0.00054 1.46725 0.00006 -20.8 -18.6 -18.6 1.47 97b-Hf-4 MR Tombstone w-z inner S23 12855 46.1 0.282206 0.00006 0.00051 0.00005 0.01702 0.00153 1.46725 0.00005 -20.5 -18.3 -18.3 1.46 97b-Hf-6 MR Tombstone w-z interm S18-S23 13040 48.4 0.282145 0.00007 0.00053 0.00004 0.01851 0.00097 1.46723 0.00005 -22.6 -20.5 -20.5 1.54 97b-Hf-9 MR? Tombstone p-z interm S13 10564 43.5 0.282234 0.00005 0.00058 0.00001 0.02189 0.00086 1.46724 0.00006 -19.5 -17.3 -17.3 1.42 36-1-Hf-1 MR Mayo w-z S24-J4 12646 37.6 0.282142 0.00009 0.00042 0.00001 0.01569 0.00067 1.46723 0.00008 -22.7 -20.5 -20.5 1.54 36-1-Hf-2 MR Mayo w-z interm S17-S22 12529 39.0 0.282110 0.00007 0.00044 0.00001 0.01606 0.00074 1.46723 0.00006 -23.9 -21.7 -21.7 1.58 36-1-Hf-3 MR Mayo w-z interm S17-S22, S18-S23 14233 43.2 0.282145 0.00007 0.00043 0.00002 0.01570 0.00051 1.46725 0.00007 -22.6 -20.5 -20.5 1.54 68-Hf-4 MR Mayo (HM) w-z interm S1-S6 12741 44.7 0.281937 0.00007 0.00050 0.00005 0.01896 0.00132 1.46726 0.00007 -37.3 -27.8 -27.8 1.82 110-Hf-1 MR Tungsten (NE) w-z P1-G1 17994 427.1 0.282182 0.00005 0.00446 0.00007 0.15750 0.00525 1.46724 0.00008 -54.3 -19.4 -19.4 1.66 143-Hf-3 MR Tungsten (SE) w-z interm L3 13394 160.5 0.282273 0.00004 0.00170 0.00012 0.06169 0.00321 1.46724 0.00007 -30.1 -16.0 -16.0 1.41 143-Hf-4 MR Tungsten (SE) p-z interm S11-S12 11422 131.9 0.282184 0.00006 0.00164 0.00013 0.06167 0.00655 1.46724 0.00007 -30.3 -19.1 -19.1 1.53 143-Hf-6 MR Tungsten (SE) p-z interm S9-S14 10983 124.4 0.282183 0.00006 0.00161 0.00013 0.06141 0.00267 1.46717 0.00008 -30.0 -19.2 -19.2 1.53 196-Hf-1 MR Tay River p-z S3 9000 45.4 0.282386 0.00005 0.00072 0.00011 0.02333 0.00415 1.46725 0.00006 -47.5 -12.0 -12.0 1.21 196-Hf-2 MR Tay River p-z interm S3 12172 54.6 0.282374 0.00003 0.00064 0.00001 0.02017 0.00028 1.46725 0.00005 -21.3 -12.4 -12.4 1.23 196-Hf-4 MR Tay River w-z interm S23-S18 12259 67.2 0.282385 0.00005 0.00078 0.00003 0.02503 0.00077 1.46723 0.00007 -18.1 -12.0 -12.0 1.22 196-Hf-6 MR Tay River p-z S11-S12 11661 72.9 0.282334 0.00006 0.00089 0.00001 0.03078 0.00057 1.46723 0.00008 -21.1 -13.8 -13.8 1.29 196-Hf-7 MR Tay River w-z S1-S2 12166 62.5 0.282389 0.00005 0.00073 0.00004 0.02460 0.00195 1.46725 0.00006 -18.1 -11.8 -11.8 1.21 08-Hf-1 MR Tay River (CR) w-z interm S1-S6, S2- S7 13004 141.0 0.282467 0.00007 0.00154 0.00010 0.04653 0.00127 1.46723 0.00007 -21.3 -9.1 -9.1 1.13 08-Hf-3 MR Tay River (CR) w-z S1-S6 12141 119.9 0.282450 0.00004 0.00140 0.00005 0.04304 0.00234 1.46724 0.00007 -21.9 -9.7 -9.7 1.15 08-Hf-5 MR Tay River (CR) w-z interm S14-S19 12196 119.2 0.282462 0.00007 0.00139 0.00009 0.04227 0.00065 1.46724 0.00009 -21.3 -9.3 -9.3 1.13 08-Hf-6 MR Tay River (CR) w-z S17-S22 11824 99.5 0.282423 0.00005 0.00119 0.00010 0.03887 0.00422 1.46727 0.00007 -22.5 -10.7 -10.7 1.18 68-Hf-2 MC Mayo (HM) resorbed P1 10042 27.1 0.281934 0.00008 0.00038 0.00009 0.01321 0.00268 1.46730 0.00007 -22.3 -27.9 -27.9 1.82 68-Hf-3 MC Mayo (HM) w-z S3-S8 12747 61.8 0.281928 0.00007 0.00069 0.00009 0.02395 0.00268 1.46729 0.00007 -34.6 -28.1 -28.1 1.84 175-Hf-1 MC Tungsten (NE) u-z S2? 9395 53.6 0.282154 0.00005 0.00081 0.00005 0.02724 0.00245 1.46724 0.00008 -14.1 -20.2 -20.2 1.54 143-Hf-5 MC Tungsten (SE) p-z S19-S24 12397 79.2 0.282167 0.00007 0.00091 0.00006 0.03475 0.00396 1.46725 0.00011 -14.5 -19.7 -19.7 1.53 143-Hf-2 MC? Tungsten (SE) p-z Q2-Q3 10395 125.6 0.282187 0.00004 0.00172 0.00004 0.05966 0.00109 1.46724 0.00006 -16.3 -19.0 -19.0 1.53 196-Hf-5 MC Tay River w-z S19-S20- S24-S25 11137 87.9 0.282330 0.00005 0.00112 0.00007 0.03751 0.00217 1.46724 0.00005 -14.2 -14.0 -14.0 1.31 08-Hf-2 MC Tay River (CR) w-z S19-S14 8522 86.4 0.282432 0.00006 0.00144 0.00012 0.04875 0.00553 1.46727 0.00006 -16.1 -10.4 -10.4 1.17 76-Hf-1 IC Mayo (HM) resorbed n/a 10815 102.1 0.281730 0.00005 0.00134 0.00001 0.04991 0.00107 1.46724 0.00005 -15.9 -35.2 -35.2 2.15 68-Hf-1 IC Mayo (HM) resorbed n/a 9771 29.5 0.281248 0.00006 0.00043 0.00003 0.01382 0.00078 1.46718 0.00006 -14.0 -52.2 -52.2 2.75 68-Hf-6 IC Mayo (HM) resorbed n/a 11753 52.6 0.281441 0.00008 0.00064 0.00000 0.02253 0.00041 1.46728 0.00007 -11.3 -45.4 -45.4 2.51 353 Analysis Spota Zoneb Plutonic Suite (sub- suite/region) Descriptionc Morphologyd Hf (ppm)e Lu (ppm)f 176Hf/ 177Hfg ± 2-σ 176Lu/ 177Hfg ± 2-σ 176Yb/ 177Hfg ± 2-σ 178Hf/ 177Hfg ± 2-σ εHf0 εHfih εHf100h TDM (Ga)i 175-Hf-2 IC Tungsten (NE) u-z L1-L2, S1-S2 11874 57.9 0.281806 0.00006 0.00069 0.00010 0.02498 0.00344 1.46724 0.00007 -12.5 -32.4 -32.4 2.01 143-Hf-7 IC Tungsten (SE) u-z Q2-Q3 10455 73.4 0.282150 0.00006 0.00100 0.00005 0.03755 0.00224 1.46719 0.00007 -11.8 -20.3 -20.3 1.55 143-Hf-1 IC? Tungsten (SE) p-z L4-S4? 17240 176.5 0.282274 0.00004 0.00145 0.00011 0.04988 0.00383 1.46725 0.00005 -11.4 -16.0 -16.0 1.40 196-Hf-3 IC Tay River resorbed n/a 15384 142.1 0.282323 0.00004 0.00131 0.00007 0.04492 0.00308 1.46722 0.00006 -12.8 -14.2 -14.2 1.32 Notes: a sample label in Appendix E.1 (first digits of the label are the last digits of the sample number); b MR = magmatic rim, MC = magmatic core (autocrystic and antecrystic), and IC = inherited core. c description of zoning; abbreviations are: u-z = unzoned, p-z = poorly zoned, w-z = well-zoned, interm = intermediate zone (between inner and outer zones); d morphological classification after Pupin (1980); e Hf concentrations from EPM analysis; f Lu concentrations calculated from Hf concentration and 176Lu/177Hf; g mass bias and interference corrected (where required) as detailed in Appendix E.4; h chondritic values of 176Hf/177Hf0 = 0.282785 and 176Lu/177Hf = 0.0336 (Bouvier et al., 2008), and the 176Lu decay constant of 1.867 x 10-11/yr (Scherer et al., 2001) were adopted for the calculation of εHfi and εHf100; i 176Hf/177Hf0 = 0.28325 and 176Lu/177Hf (0.0384) values used by Griffin et al. (2000) were adopted for the calculation of TDM. 354 6 CHAPTER: APPROXIMATING MAGMATIC FLUID COMPOSITIONS OF BARREN AND TUNGSTEN SKARN-ASSOCIATED PLUTONS WITH APATITE AND BIOTITE: CASE STUDIES FROM THE NORTHERN CANADIAN CORDILLERA 6.1 INTRODUCTION The study of the volatile elements fluorine and chlorine in crystallizing granitoid intrusions is of broad interest to the field of economic geology for two primary reasons (cf. Pirajno, 2009). First, increasing the activity or concentration of volatiles in a magma—through processes such as fractional crystallization, magma mixing, or country rock contamination—has a strong effect on the behaviour of melt, minerals, and ionic species. This is mainly manifested by reducing the solidus temperature of the magma and de-polymerizing the melt to release highly charged cations. Together these effects may increase the availability of some metals for partitioning into a fluid phase. Second, many volatile phases also act as neutral to negatively charged species, “ligands”, most commonly NH3, H2O, Cl–, OH–, and HS–. Ligands form soluble complexes with cations, particularly those with higher ionization energies (e.g., Ni, Cu, Zn, Pt, Au, Co, Cr, Mo, and W). When metals behave incompatibly in a magma the formation of charged metal– ligand complexes allows for the sequestration, concentration, and transport of metal complexes in fluid—aqueous and (or) vaporous solution(s)—exsolved from the melt. This would imply that the ability of an exsolved magmatic fluid to extract metals from a magma will be largely controlled by the availability of the ligand species (e.g., Seward and Barnes, 1997), for both their ability to de-polymerize melts and form complexes with metals, regardless of the total content of metals in the magmatic system. That said, the concentration of metals in a magma and their availability for partitioning into a magmatic fluid, combined with the timing of fluid exsolution must also be of central importance to the formation of magmatic–hydrothermal ore deposits (e.g., Candela, 1997). Numerous reviews detail our understanding of the processes leading to magmatic volatile exsolution during the ascension, emplacement, and crystallization of a magma—particularly through the relatively high temperature (ca. 850–650°C) transformation from orthomagmatic to magmatic–hydrothermal conditions (e.g., Burnham, 1979; Crerar et al. 1985; Hedenquist, 1995; Candela 1997; Shinohara and Kazahaya, 1995; Seward and Barnes, 1997; Webster, 1997a,b; Williams–Jones and Heinrich, 2005; Audétat et al., 2008). In short, volatile species become less soluble as magmas evolve to felsic compositions and (or) as they cool through this temperature transition. Upon saturation in a magma, if the dissolved volatiles exceed the confining pressure of the magma they will unmix from the silicate melt and crystals to form an immiscible magmatic fluid, commonly referred to as a magmatic volatile phase (MVP). Carbonate, chlorine, and alkali–chloride complexes are particularly insoluble in granitic magmas, even in small quantities (e.g., Holloway, 1976; Webster, 1997b). Most felsic magmas, therefore, likely exsolve and possibly off-gas (periodically or continuously) small amounts of carbonate-rich or highly saline MVP, or contain an immiscible molten salt phase throughout most of their ascension, emplacement, and crystallization history (Christiansen and Keith, 1996; Downey and Lentz, 2006). Although these fluids may scavenge some incompatible elements from the melt, metals that partition preferentially into a fluid do not concentrate in the MVP unless 355 significant amounts of H2O are also exsolved (Christiansen and Keith, 1996). Unmixing of a H2O-rich fluid is thought to occur mainly by two processes that are strongly tied to pressure– temperature conditions (Burnham, 1979). It may comprise one or more violent pulses of early-stage MVP exsolution if thermodynamic conditions change suddenly, as in adiabatic depressurization (“first boiling”). Or, it may be an ongoing late-stage process if conditions change gradually, as in isobaric cooling and crystallization (“second boiling”). At shallow or subcritical levels of emplacement (e.g., <1.5 kbar; 800°C) the MVP will comprise two phases: a dense and hypersaline brine containing over 26 wt.% NaCl-equivalent, and a H2O ± CO2 enriched vapour. At deeper or supercritical conditions (e.g., >1.5 kbar; 800°C) the MVP will consist of a either a single-phase, volatile-rich, low to moderately saline aqueous solution containing under 20 wt.% NaCl-equivalent, or a vapour dominated by H2O. This single-phase MVP may further unmix to a two-phase daughter fluid if subcritical conditions occur or are encountered as the MVP ascends. Following the exsolution of a saline hydrous MVP, incompatible elements and ligands that partition preferentially into the fluid phase are effectively removed from (i.e., depleted in) the melt and added to (i.e., enriched in) the MVP (e.g., Candela and Piccoli, 1995; Christensen and Keith, 1996). Density contrasts between the MVP and melt allow the MVP to ascend through the magma— potentially scavenging metals along the way to accumulate at the top of the magma chamber. If pre-existing or newly created (from fluid/vapour overpressuring; Burnham, 1979, 1985) permeability or porosity channels exist in the overlying plutonic carapace and country rocks, the metal-rich MVP will continue its ascension. Changes to thermodynamic and compositional conditions of the magmatic fluid—such as those that occur during depressurization and interactions with non-equilibrated country rock and fluids—invariably reduce the stability of the metal–ligand complexes (e.g., Seward and Barnes, 1997). In this way the metals will precipitate from solution and, if locally concentrated, may form a magmatic–hydrothermal ore deposit. There are several applications for the halogen-specific compositional data obtained from hydrous minerals in studies of magmatic, hydrothermal, and metamorphic fluids. Perhaps of greatest interest to the economic geologist is the potential to determine the composition of mineralizing fluids with respect to relative halogen and H2O content. In magmatic settings, an early loss of a saline and hydrous MVP is inferred to lower the concentration of magmatic volatiles that partition preferentially into the fluid (e.g., Cl and H2O; e.g., Piccoli and Candela, 2002) in the crystallizing magma and in subsequent MVP. Hydrous magmatic minerals crystallizing in a volatile depleted magma should be relatively depleted in the exsolved volatile species, and relatively enriched in volatile species that preferentially partition into minerals over fluid (e.g., F; e.g., Piccoli and Candela, 2002). If volatiles that would normally partition into a fluid phase are allowed to accumulate in the magma, the F:Cl:OH composition of the magmatic minerals will reflect that of the dissolved magmatic volatiles, as well as any later-exsolved, saline and hydrous MVP. Thus, the timing of MVP exsolution becomes an important consideration in interpretations of magmatic mineral compositions with respect to F:Cl:OH content for studies of magmatic fluid composition. Detecting the relative composition and timing of an MVP using halogen-hydroxyl variations in magmatic minerals may allow us link to known mineralization to magmatic events (e.g., Candela, 1986; Sheets et al., 1996; Lang, 1998; Selby and Nesbitt, 2000; Coulson et al., 2001; Piccoli and Candela, 2002), and to identify intrusive rocks that are prospective for related mineralization. 356 This study of magmatic fluid compositions, which are calculated for multiple intrusive phases from several composite plutons in the northern Canadian Cordillera, demonstrates how magmatic fluids evolve with respect to F– Cl–OH with magma differentiation. Variations to typical magmatic fluid compositional trend of increasing HF relative to HCl and H2O are the result of alteration of the magmatic minerals, and (or) precipitation of the magmatic minerals in equilibrium with a compositionally anomalous magmatic fluid. First, however, some background on similar studies of halogen concentrations in minerals (and calculated magmatic fluid compositions) and the issues limiting the scope of these studies is necessitated. 6.1.1 Issues Inherent to Magmatic Fluid Compositional Studies There now exists a plethora of observational and experimental data on the partitioning of ligands (e.g., Ludington and Munoz, 1975; Zhu and Sverjensky, 1991; Brenan, 1993b; Icenhower and London, 1997; Villemant and Boudon, 1999; Mathez and Webster, 2005) and metals (e.g., Tingle and Fenn, 1984; Candela and Piccoli, 1995; Wood and Samson, 2000; Hu et al., 2008) between minerals, melt, and fluid in magmatic systems. Several thermodynamic schemes based on the experimental data and examples of natural systems have been developed in order to calculate estimates of relative and actual concentrations of fluorine, chlorine, and H2O in fluids (e.g., Korzhinsky, 1981; Munoz and Swenson, 1981; Munoz, 1984, 1992; Speer, 1984; Candela, 1986; Zhu and Sverjensky, 1991, 1992; Piccoli and Candela, 1994; Piccoli et al., 1999; Aksyuk, 2000; Sallet, 2000). However, interpretations of the resulting compositional data have been disappointing in terms of identifying significant or widely applicable markers that distinguish mineralizing from barren intrusions. In several cases where compositional features associated with mineralized or mineralizing intrusions are found, comparisons with other datasets may fail to confirm the observations or may even produce contradictory results (e.g., Lang, 1998; Piccoli and Candela, 2002). Although there are undoubtedly innumerable factors that together determine which intrusions are capable of producing economic mineralization, several issues that are specific to fluid compositional studies are discussed as follows. Biotite, which accommodates fluorine and chlorine in its hydroxyl site and is found in a wide variety of rocks, is most frequently used to estimate fluid compositions in the form of F:Cl:OH fugacity or activity ratios. The partitioning of chlorine into biotite is dependent upon the temperature and pressure of solid–solid or fluid–solid exchange, and on the composition of the melt when fluid is involved (e.g., pH, aluminum saturation index, and oxygen fugacity; Icenhower and London, 1997; Kesler et al., 1975; Munoz and Ludington, 1974; Stormer and Carmichael, 1971; Zhu and Sverjensky, 1991). However, the most significant control on the uptake of halogens into biotite are cations that have cross-site compositional influences on the hydroxyl site (e.g., Fe2+, Fe3+, Mg2+, AlIV, Ti4+, and Li+; e.g., Zhu and Sverjensky, 1992). Due to their concentrations in biotite, iron and magnesium contents are the most significant controls on the partitioning of fluorine and chlorine into biotite (e.g., Zhu and Sverjensky, 1991). Inverse correlations known as the “Mg–Cl and Fe–F avoidance rules” are the result. The mutual Mg–Cl and Fe–F avoidances are also controlled externally by fluid composition, pressure, and temperature, so these ‘rules’ are not always obeyed (e.g., Zhu and Sverjensky, 1992; Lang, 1998; Selby and Nesbitt, 2000). Biotite-based studies of fluid composition have been problematic due to the difficulty of accounting for cationic and external controls on the 357 substitution of halogens into the hydroxyl site (Zhu and Sverjensky, 1992). Magmatic biotite compositions are also susceptible to partial or full re-equilibration with late-stage magmatic or hydrothermal fluids. Many studies have documented instances where these sub-solidus processes must have occurred in biotite, even when petrographic evidence for these processes is lacking (e.g., Zaw and Clark, 1978; Munoz and Swenson, 1981; van Middelaar and Keith, 1990; Loferski and Ayuso, 1995). Like magmatic minerals crystallizing in a volatile depleted environment, minerals precipitating from a hydrothermal or metamorphic fluid will also be enriched in ligands that have a preference for partitioning into solids, even if the ligands are present in very low concentrations in the fluid (e.g., F; Zhu and Sverjensky, 1991; Zhang et al., 2001). In some cases it may be difficult to ascertain whether halogen concentrations in igneous biotite are inherited from magmatic conditions, or are a result of later fluid incursion and re-equilibration. Although experimental evidence indicate that halogen exchange between biotite and fluid is very slow—even at temperatures greater than 700oC—and that biotite is essentially closed to sub-solidus halogen exchange (Aksyuk, 2000), observations that biotite is more likely to re- equilibrate with later-stage (cooler) fluids than apatite indicate otherwise (Stormer and Carmichael, 1971; Munoz and Ludington, 1974; Dilles, 1987; van Middelaar, 1988; Piccoli and Candela, 1994; Icenhower and London, 1997; Sallet, 2000). Therefore, it has been suggested that halogen exchange in biotite may occur at lower temperatures (i.e., 300–350oC; van Middelaar and Keith, 1990) that are more similar to closure temperatures for other volatiles in biotite (e.g., Ar). Some more recent studies on magmatic fluid compositions have employed compositional data from apatite instead of, or in conjunction with biotite datasets. Like biotite, apatite also accommodates fluorine and chlorine substitution in its hydroxyl site. Other halogen species, as well as other monovalent and divalent anions (e.g., most notably CO32– and S2–) substitute in trace to small quantities for hydroxyl in apatite, but it is generally assumed that these anions and any vacancies are insignificant (Pan and Fleet, 2002; Piccoli and Candela, 2002). Unlike biotite, however, apatite is considered to be relatively immune to re-equilibration with lower temperature or later-stage fluids (e.g., Stormer and Carmichael, 1971; Roegge et al., 1974; Candela, 1986; Munoz, 1990; Piccoli et al., 1999). Also unlike biotite, major element compositional controls on fluorine and chlorine partitioning into apatite either do not exist or have little effect on what appears to be ideal F–Cl–OH mixing in the hydroxyl site (Tacker and Stormer, 1989; Zhu and Sverjensky, 1992; Lang, 1998). Relative proportions of fluorine, chlorine, and hydroxyl (and concentrations of sulphur) in apatite, thus, indirectly reflect the abundances of these species in coexisting fluid at the time of apatite crystallization (e.g., Piccoli and Candela, 1994; Parat and Holtz, 2005). That said, some drawbacks to using apatite to estimate halogen contents in fluids include natural compositional variations between apatite grains within the same sample (e.g., Sha and Chappell, 1999) and the possibility of halogen diffusion between apatite and other hydrous minerals (Parry et al., 1978; Brenan, 1993b). It should also be noted that apatite is always enriched in fluorine compared to the composition of the fluid or melt with which it equilibrated. The fluorine-rich apatite end- member, “fluorapatite”, may precipitate from even a fluorine-poor melt (e.g., containing <0.3 wt.% F; Xiaolin et al., 1998; Kifer and Wolf, 2000) or a fluorine-poor hydrothermal fluid (containing as little as 10–6 molal F–; Zhu and Sverjensky, 1991). When the preferential partitioning of fluorine into apatite over melt and fluid results in fluorapatite compositions, thermodynamic calculations involving chlorine or hydroxyl are impossible (Zhu and 358 Sverjensky, 1992). Apatite-based studies have also been limited by significant analytical difficulties. The primary issue is that falsely inflated fluorine measurements is caused by the accelerated diffusion of fluorine during electron-microprobe (EMP) analysis. This diffusion becomes increasingly pronounced as the angle decreases between the c-axis in apatite and the electron beam; the best measurements are obtained with the long grain axis is perpendicular to the electron beam. Elevated fluorine concentrations may also occur when significant carbonate concentrations are present in the apatite; although a 1:1 coupled substitution of F– and CO32– for PO43– acts to increase fluorine concentrations in apatite, although it is not certain where this additional fluorine resides in the apatite structure (McArthur, 1990). Accelerated fluorine diffusion and (or) elevated carbonate contents may lead studies of apatite composition to report fluorine concentrations well in excess of what may be stoichiometrically accommodated in the hydroxyl site of apatite (3.77 wt.% F). Advances in analytical and data reduction methods for apatite have reduced, but not entirely eliminated errors resulting from accelerated fluorine diffusion along the c-axis in apatite (Stormer, et al., 1993; Raudsepp, 1995). It is also not always obvious in published apatite compositional datasets whether or not these analytical precautions were taken; therefore, even if stoichiometrically possible fluorine concentrations in apatite are reported, in some cases these may be still artificially inflated. The EMP analysis of apatite is also complicated by the absorption of halogens under high accelerating voltage (Piccoli and Candela, 1994), peak overlaps (e.g., F and P; Piccoli and Candela, 1994; Raudsepp, 1995), and the possibility of compositional variations or zoning within apatite standards for some elements (e.g., F). A major issue with most thermodynamic calculations is the strong effect on F–Cl–OH partitioning between minerals and fluid that result from variations in temperature, pressure, and fluid composition (Brenan, 1993a; Zhu and Sverjensky, 1991, 1992). If assumptions of these conditions are incorrect, very large errors may be imparted to the calculated fluid compositions. The treatment of Zhu and Sverjensky (1992) is the only formulation that uses apatite and biotite compositions together to minimize the effects of F–Cl–OH partitioning on the calculated fluid compositions. It also incorporates one of the most extensive accountings of the reciprocal effects of octahedral and hydroxyl site occupancy controlling fluorine and chlorine substitution into biotite (Zhu and Sverjensky, 1991, 1992). Unlike other models that consider mineral-fluid partitioning, the study of F–Cl–OH exchange between two mineral (solid) phases significantly reduces the effect of temperature on mineral composition, and removes most pressure and all fluid composition controls (Zhu and Sverjensky, 1992). Assuming biotite and apatite are not in disequilibrium, both minerals will also provide independent estimates of fluid compositions relative to F:Cl:OH. Comparison of the apatite- and biotite-based estimates may, therefore, be used to check for late-stage re- equilibration of the biotites (Coulson et al., 2001)—assuming apatite is not also affected by late fluids. Alternately, biotite will provide an estimate of fluid compositions in systems that contain only the fluorine-rich apatite end- member (Ekström, 1972). 359 6.2 GEOLOGICAL BACKGROUND 6.2.1 Previous Mineral Chemistry Investigations of Intrusion-related Systems Several halogen-based metallogenic studies have been completed on intrusions and alteration minerals associated with calc-alkaline copper porphyry deposits (e.g., Mason, 1978; Parry et al., 1978; Munoz and Swenson, 1981; Lang, 1998; Selby and Nesbitt, 2000; Idrus et al., 2007), although a handful of studies on layered ultramafic intrusions (e.g., Willmore et al., 2000), molybdenum porphyries (Gunow et al., 1980), base metal deposits (e.g., W, Pb–Zn, and Sn; Zaw and Clark, 1978; Nedachi, 1980; van Middelaar and Keith, 1990; Aksyuk, 2000;), and reduced intrusion-related gold mineralization (Coulson et al., 2001; Yang and Lentz, 2005) also exist. Many of these studies have attempted to classify the compositions of apatite, biotite, amphibole, and other halogen-bearing minerals relative to mineralizing and non-mineralizing intrusions, and (or) have linked changes to halogen content in minerals to compositions of mineralizing or non-mineralizing fluids. Most of this work has focused on biotite, despite the problem of internal and external controls on halogen substitution into biotite. The effects on halogen partitioning combined with the overall composition of the magmatic systems and variations in the timing of hydrous fluid exsolution, may be largely why efforts to link biotite compositions to mineralizing intrusions or fluids have produced variable results (Lang, 1998). These attempts frequently resulted in small or poorly defined correlations that—although relevant to the specific study—have been difficult to extrapolate to other studies of mineralizing systems. The vast majority of work to date has focused on mineral deposits associated with porphyry systems that are typically emplaced at subcritical conditions (e.g., <1.5 kbar). At these relatively shallow levels of emplacement boiling typically occurs by depressurization, a relatively early event occurring during or immediately after emplacement of the magma. These studies have, therefore, focused on detecting an early loss of volatile species from the parent magma. Comparatively few fluid compositional studies on supercritical magmatic–hydrothermal mineralizing systems, including tungsten skarn and reduced intrusion-related gold, have been completed. Fluid inclusion studies of tungsten skarn mineralization indicate that the mineralizing fluids are a late-stage, non-boiling hydrous fluid of low to moderate salinity—typically <10 wt.% NaCl-equivalent—exsolved at supercritical pressures (1.5–2.5 kbar; Newberry and Swanson, 1986; Newberry, 1998; Meinert et al., 2005). At deeper levels of emplacement the isobaric saturation of magmatic volatiles through slow cooling and crystallization is thought to contribute to later fluid exsolution (e.g., Newberry and Swanson, 1986; Webster, 1997a,b; Newberry, 1998), as well as low H2O initial concentrations in these magmas (e.g., Candela, 1997; Strong, 2001). However, unmixing of a saline hydrous fluid related to progressive crystallization of the magma may actually be an ongoing process, occurring from relatively early through to the latest magmatic conditions. Progressive crystallization also increases the concentration of tungsten in the melt of reduced, ilmenite-poor magmas (Candela and Bouton, 1990; Newberry, 1998), and the exsolution of a saline hydrous brine will cause tungsten to partition preferentially into the MVP (e.g., Candela and Piccoli, 1995; Audétat et al., 2008). If MVP exsolution continues until near-solidus magma conditions, the latest fluids exsolved from the residual melt have the potential to be enriched in tungsten by orders of magnitude (e.g., Newberry and Swanson, 1986). This greatly increases the possibility of significant associated tungsten 360 intrusions, with respect to chlorine versus Mg#. No information on fluorine in biotite was provided, and estimates of fluid composition were not made from the apatite and biotite compositions. More detailed work by Coulson et al. (2001) on the 94 Ma Emerald Lake batholith (Fig. 6.1)—a composite intrusion of the Tombstone plutonic suite (e.g., Coulson et al., 2002)—reported magmatic apatite, biotite, and hornblende compositions for several intrusive phases. From the apatite and biotite data, logarithmic activity ratios for HCl/HF (which represent the composition of a magmatic fluid in equilibrium with apatite and biotite) were calculated by Coulson et al. (2001) using the thermodynamic treatment of Zhu and Sverjensky (1991, 1992). Although apatite and biotite were assumed to have equilibrated with the same fluid at similar temperatures, a notable increase in estimated HCl/HF was evident between apatite- and biotite-based estimates of fluid composition. This was attributed to analytical issues associated with apatite analysis and (or) an actual change in magmatic fluid composition related to cooling. Coulson et al. (2001) also demonstrated that there was a gradual decrease in both apatite- and biotite-based estimates of HCl/HF activity, from the earliest and least evolved intrusive phases to the latest and most evolved intrusive phases. This relationship between chlorine relative to fluorine in the magmatic fluid and the increasing degree of magmatic differentiation was attributed to the greater propensity for chlorine to partition into fluid—in this case one that was continually off-gassed. The ongoing off-gassing of chlorine depleted the subsequent magmas, minerals, and fluids in HCl relative to HF. The disseminated and low-grade nature of gold mineralization throughout the Emerald Lake pluton was thought to be one a result of ongoing exsolution of an unfocused, homogenous gold- poor fluid throughout the intrusive body. Two detailed halogen studies were also completed at the Cantung tungsten skarn deposit (Fig. 6.1): one focused on the fluorine contents in mica and amphibole in the underground orebody (Zaw, 1976; Zaw and Clark, 1978), and the other focused on the fluorine and chlorine contents in mica from the underground orebody, as well as several phases of the associated intrusion (van Middelaar, 1988; van Middelaar and Keith, 1990). The first study determined that skarn micas are primarily annite-phlogopite, but trend towards the siderophyllite end-member. Relatively depleted fluorine concentrations in biotite from the associated unaltered granite intrusion (0.30–0.42 wt.% F) in comparison to enriched fluorine in the skarn micas (0.82–2.98 wt.% F) were also reported (Zaw, 1976; Zaw and Clark, 1978). These results were similar to those of van Middelaar and Keith (1990). Chlorine was below the detection limit of 0.2 wt.% in all mica analyses by Zaw (1976), and was less than 0.1 wt.% in mica from all of the intrusive phases (van Middelaar, 1988). Both studies found that fluorine concentrations were highest in the more magnesian micas— consistent with the Fe–F avoidance rule—although apparently not as enriched as expected given the degree of Mg- enrichment (Zaw and Clark, 1978). The observed low ferric iron content in the micas is consistent with the extremely low oxygen fugacity interpreted for the Cantung system, which further implied that aluminum and oxygen had no notable effect on F–Cl–OH exchange in biotite (Ludington and Munoz, 1975; Zaw and Clark, 1978; van Middelaar and Keith, 1990). The H2O/HF fugacity ratio (after Ludington and Munoz; 1975) demonstrated that the skarn mica at Cantung equilibrated with fluid(s) of relatively high, but variable H2O/HF (Zaw and Clark, 1978). A relative measure of HF and HCl fugacities in the mineralizing fluid—the F/Cl intercept ratio, in which higher values indicate lower fluorine 361 mineralization (Newberry and Swanson, 1986; Newberry, 1998; Keith et al., 1989; Webster, 1990; London, 1995; Aksyuk, 2000). The incursion of fluids into overlying country rocks occurs when overpressuring of the more voluminous MVP creates or re-opens fractures (e.g., Candela, 1997). If the relatively deep magmatic conditions are combined with slow fluid ascent rate, the MVP may never intersect its solvus (i.e., produce a two-phase daughter fluid) and will remain as the single-phase mineralizing fluid that characterizes tungsten skarn deposits (Rasmussen et al., 2011b). 6.2.2 The Northern Cordillera Syn- to post-accretionary, Early to ‘mid’-Cretaceous intrusions (~124–90 Ma) were emplaced throughout the northern Cordillera in response to the final stages of terrane accretion and the re-initiation of northwesterly subducting oceanic crust along the western margin of North America (Hart et al., 2004a; Nelson and Colpron, 2007; Fig. 6.1). Much of these magmatic rocks was emplaced into the Selwyn Basin (Fig. 6.1), a lower-plate rifted margin characterized by Early Cambrian to Middle Devonian basinal sedimentary rocks deposited on a package of Neoproterozoic rift-related sedimentary rocks (and presumably underlain by radiogenic crystalline basement; Cook et al., 2004; Gordey and Anderson, 1993). The most northerly and easterly plutonic rocks in the Selwyn Basin occurred well inboard of any subduction zone and related magmatic arc activity in the ‘mid’-Cretaceous (ca. 98-90 Ma). This belt of felsic, sub-alkalic, predominantly reduced intrusions is also known as the Tombstone–Tungsten Belt (TTB; Fig. 6.1). Based on isotopic and geochemical datasets, the parental magmas of the TTB intrusions appear to have been derived from crustal sources, although a contribution of magma from an enriched or lithospheric mantle is evident in associated lamprophyre dykes and in the composition of the youngest plutons (Mortensen et al. 2000; Heffernan, 2004; Hart et al. 2005; Mair et al., 2006a, 2011; Rasmussen et al. 2007, 2011a). Fluid inclusions studies and hornfels mineralogy suggest that most of the TTB intrusions were emplaced at pressures ca. 1.5–3 kbar (Thompson et al., 1999; G.R. Thompson, unpub. data, 2000; Marshall et al., 2003; Baker and Lang, 2001; Mair et al., 2006a; Yuvan et al., 2007; Rasmussen et al., 2011b). A preliminary regional study of apatite and biotite compositions for plutons of the northwestern TTB across central Yukon found that variations in fluorine and chlorine abundances in apatite were most strongly related to the overall aluminicity versus alkalinity of the intrusion, rather than its plutonic suite or age (Lang, 2001). In apatite, fluorine concentrations increased and chlorine concentrations decreased with increasing whole rock silica, from metaluminous (~2.5–4 wt.% F; 1.7 to <0.1 wt.% Cl) to peraluminous (~3.5–5.4 wt.% F; 0.4 to <0.1 wt.% Cl) plutons. In comparison, apatite in alkalic intrusions contained unusually high fluorine and low chlorine given the whole rock SiO2 contents (~3.2–4.6 wt.% F; <0.4 wt.% Cl), but the SiO2 range was too small to illustrate compositional trends with respect to chlorine and fluorine. Fluorine in apatite was negatively correlated with P2O5 for the aluminous plutons—suggestive of some carbonate in the apatite structure (Lang, 2001)—whereas alkalic plutons showed no such correlation. For biotite from metaluminous samples, conformance with the Mg–Cl avoidance rule was evident, although metaluminous samples associated with Cu–Au mineralization had anomalously high chlorine content (e.g., ~1.2–1.8 wt.%). Chlorine concentrations in biotite from alkalic rocks were either moderately high (~1 wt.%) or very low (<0.3 wt.%) and overlapped with biotite from peraluminous 362 fugacity (Munoz, 1984; Munoz and Swenson, 1981)—had a very restricted range of moderate values (van Middelaar and Keith, 1990). The limited range of fugacity ratios for intrusive and skarn samples, despite major element compositional variations between igneous and skarn biotites, led both study authors to infer that all micas were re- equilibrated (with respect to halogen concentrations) with one fluid. The sole exception to this was a leucocratic dyke with an unusually high fluorine fugacity; its position distal to the path of mineralizing fluids suggested that it may be representative of an earlier magmatic fluid composition (van Middelaar and Keith, 1990). The main mineralizing fluid at Cantung was described as a generally homogenous, H2O-rich, late-stage hydrothermal fluid (e.g., 300–350°C) exsolved from the main plutonic body and responsible for the re-equilibrated mica compositions (van Middelaar and Keith, 1990). The observed lack of variation in the F/Cl intercept ratio at Cantung is consistent with a study of Cordilleran granitoids that also found a small range F/Cl intercept ratios overall, regardless of whether the intrusion was related to tungsten mineralization or was barren (Keith et al., 1989). 6.2.3 Justification for this Study Most studies that relate magmatic and hydrothermal mineral compositions to the composition of mineralizing fluids (e.g., Loferski and Ayuso, 1995; Streck and Dilles, 1998; Selby and Nesbitt, 2000) and the compilation of Piccoli and Candela (2002) are based on chlorine-rich mineralizing systems (e.g., Cu–Mo–Au porphyries and layered PGE deposits), or are examples of extremely fluorine-rich systems (e.g., Climax-type Mo-porphyries). The mineralizing fluids and skarn micas of large tungsten deposits are relatively intermediate—with respect to chlorine and fluorine concentrations—to the chlorine-rich copper and fluorine-rich molybdenum porphyry systems ( Zaw and Clark; 1978; van Middelaar and Keith, 1990). Since the earlier biotite compositional studies on the Cantung deposit, however, advancements have been made with respect to greater analytical precision and more robust corrections for coupled substitutions in biotite from experimental and natural systems are now available (Zhu and Sverjensky, 1991, 1992). Another limitation of the earlier studies at Cantung is that without apatite-based estimates of magmatic fluid composition, it is uncertain whether or not the biotites at Cantung were actually re-equilibrated with one late hydrothermal fluid, or if in fact they have retained primary magmatic fluid compositions. Finally, the pressure of skarn mineralization at Cantung has been revisited and determined to be 2–3.8 kbar (Marshall et al., 2003), significantly higher than the 1 kbar estimate used for two earlier fluid compositional studies. While the new pressure estimate for mineralization at Cantung may not dramatically affect the earlier fluid compositional work, these deeper magmatic and mineralizing conditions are important with the development of a model for reduced tungsten skarn systems (given that Cantung is a world-class example of this deposit type based on its high ore grades; e.g., Menzie and Jones, 1986). With the above issues and the limited information provided from previous fluid fugacity studies for reduced tungsten skarn deposits, a re-assessment of the composition of the magmatic fluids in tungsten mineralizing systems, using the particularly robust thermodynamic treatment of Zhu and Sverjensky (1992), is warranted. This study reports relative F:Cl:OH concentrations in fluids that equilibrated with apatite and biotite, characterizes the composition of magmatic fluids in granitic intrusions that are prospective for tungsten skarn mineralization, and identifies specific intrusive phases that are more likely to have exsolved a mineralizing fluid. 363 6.2.4 Description of the Study Plutons The study area contains numerous intrusion-related precious and base-metal mineral occurrences—W-(Cu) skarn occurrences are particularly abundant (Fig. 6.2)—but this region is most notable for the presence of two world-class tungsten skarn deposits in association with two plutons of the TTB (Rasmussen et al., 2007). Many of the tungsten occurrences throughout the study area (Fig. 6.2) have also undergone advanced exploration and drilling (e.g., Rudi, Cac, Lened, and Ivo; NORMIN, 2013). Five multi-phase plutons in the Cantung mine area with varying likelihoods of associated tungsten mineralization were selected for detailed halogen analysis. These are summarized in Table 6.1 and include: the barren Central Nahanni and Rifle Range plutons, and the tungsten-associated Rudi, Lened, and “Mine Stock” (i.e., Cantung) plutons (Fig. 6.2). The Central Nahanni pluton is an example of a multi-phase, sub-alkalic, metaluminous intrusive body that is composed primarily of intermediate to felsic quartz diorite, granodiorite, and granite. This pluton is not associated with any documented mineral occurrences, nor was any evidence for mineralization observed in the field. Therefore, it is used as an example for the evolution of magmatic fluid compositions in a moderately evolved, but apparently unmineralized intrusion. The Rifle Range pluton is an example of a very small, multi-phase, sub-alkalic, peraluminous intrusive body that is composed primarily of felsic granite, leucogranite, porphyry, and aplite. At least one phase of the Rifle Range pluton is cut by scheelite-bearing quartz veins, and it also has a ‘down-glacial till’ geochemical anomaly with concentrations of very high arsenic (3,000 ppm), bismuth (100 ppm), and tungsten (up to 800 ppm; D. Tenney, 2010, pers. comm.). Despite these localized cross-cutting veins, to date large accumulations of tungsten skarn mineralization have not been identified in the Rifle Range pluton. This pluton is, therefore, used as an example a highly evolved intrusion affected by late-stage magmatic or hydrothermal fluids, but not associated with an identified body of tungsten skarn mineralization. The Rudi pluton is another example of a very small, multi-phase, sub-alkalic, peraluminous intrusive body that is composed primarily of only highly differentiated granite, leucogranite, and aplite. However, unlike the Rifle Range pluton, the Rudi pluton is associated with three very small tungsten skarn occurrences along its margins. The Lened pluton is another small, multi-phase, sub-alkalic, peraluminous intrusive body that is composed primarily of felsic granite, leucogranite, and aplite. Mineralization at Lened is more significant (indicated reserve of 0.75 Mt at 1.17% WO3 and 0.15% Cu; Walton, 1998), but sub-economic because of its location. Similarly, Cantung is an example of a relatively small, multi-phase, sub-alkalic, peraluminous intrusive system; however, Cantung is associated with world-class tungsten skarn mineralization (according to the classification by grade; Menzie and Jones, 1986). At Cantung, several phases of intrusion have been simply classified as: early plutonic phases (the biotite-bearing Mine Stock granite), early dykes associated with plutonic phases, and late leucocratic dykes (Rasmussen et al., 2011b). Endoskarn development in the marginal granite in contact with skarn, and in steeply-dipping late aplitic dykes is manifested as pervasive calcic alteration leading to either a “tonalitic” mineralogy (plagioclase–quartz) or a “dioritic” mineralogy (plagioclase–epidote-group minerals ± calcite–sericite; Rasmussen et al., 2011b). “Dioritic” aplite dykes are frequently mineralized with scheelite in fractures, veins, and as disseminations, and are more proximal to (and commonly in contact with) tungsten skarn mineralization. Aplite dykes exhibiting the tonalitic 364 mineral assemblage are unmineralized, and are more distal to tungsten skarn mineralization. The five plutons studied are composed of multiple intrusive phases that can be distinguished on the basis of texture, mineralogy, and cross-cutting relationships. All but the Mine Stock at Cantung (of which very little is exposed) are known to be reversely zoned from less evolved cores to more felsic and highly evolved margins. Between 5 and 12 samples of distinct intrusive phases were sampled from each pluton. Relative degrees of differentiation were determined for each sample by sample mineralogy, texture, and cross-cutting relationships, as well as by a combination of apatite and biotite compositions (detailed in Section 6.5.1). Petrographic textures range from relatively pristine examples that record equilibrium magmatic conditions to variably altered textures, including: • allotriomorphic and equigranular groundmass; • pervasive sericite/saussurite alteration of feldspar; • coarse-grained apatite associated with secondary minerals (e.g., calcite, sericite, Mg-rich biotite, scheelite, and pyrrhotite); and • small ragged mica along grain margins or in veins and clots. Therefore, it was possible to texturally distinguish and then analyze both primary and secondary apatite and biotite from intrusive phases that experienced differing magmatic ± hydrothermal conditions. 6.3 ANALYTICAL METHODS All sample preparation and micro-beam analysis were completed at the University of British Columbia with the assistance of Dr. M. Raudsepp. Data reduction and activity ratio calculations for apatite and biotite were done with the assistance of Dr. G. Dipple and Dr. L. Groat following the methods and calculations described by Dr. I. Coulson (2006, pers. comm.) and Zhu and Sverjensky (1991, 1992). 6.3.1 Apatite Apatite was identified under transmitted light in polished thin sections and each thin section was examined petrographically in transmitted light and imaged in back-scattered electron (BSE) mode on the scanning electron microscope (SEM). Apatite occurs as inclusions in both anhydrous (quartz, feldspar) and hydrous (amphibole, biotite) mineral phases, as well as throughout the groundmass. Igneous apatites from all samples are small (5-200 μm), stubby prismatic grains with well-formed crystal faces. Hydrothermal examples of apatite—typically associated with late-stage veins and scheelite mineralization (at Cantung)—are also stubby and prismatic, but these grains typically measure more than 200 μm along the long (c-) axis and have less well-developed crystal faces. The larger size and changed morphology of the hydrothermal apatites is consistent with precipitation of these grains from an aqueous fluid (Wylie et al., 1962). An attempt was made to select grains from the polished thin sections that were contained predominantly within anhydrous phases in order to avoid the possible effects of grain-boundary halogen diffusion between apatite and other hydrous minerals (Lofurski and Ayuso, 1995; Piccoli et al., 1999; Piccoli and Candela, 2002). Given that the 365 vast majority of apatite was associated with biotite and amphibole in all samples, however, 50-90% of analyses in any given sample were actually performed on apatite grains included within or adjacent to these minerals. No discernible compositional differences were observed between apatites petrographically associated with hydrous or anhydrous mineral hosts. An attempt was also made to select apatite grains in the polished thin sections with appropriate orientations for analysis. Grains with their c- or long axis parallel to the thin section surface experience less fluorine diffusion— occurring mainly along the c-axis—during election beam exposure (Stormer et al., 1993). Perfectly oriented grains are, however, rare in thin sections. Therefore, individual apatite grains from selected intrusive samples, mounted in epoxy, and imaged in BSE mode on the SEM, as described in Chapter 7 (Section 7.3.2.1). Apatite from both polished thin sections and grain mounts was analyzed on a fully automated CAMECA SX-50 EMP operating in the wavelength-dispersion mode and following modified methods of Coulson et al. (2001), Raudsepp (1995), and Stormer et al. (1993). Five repeated 100-second analyses were done in the same location of each grain to record increases in measured fluorine concentrations with time. Between one and fifteen apatite grains from each sample were analyzed using the following operating conditions (based on our test analyses and the recommendations of Piccoli and Candela, 1994): an excitation voltage of 15 kV; a beam current of 10 nA; and spot diameter of 10 μm. Peak and background times, respectively, were: 10 and 5 seconds for magnesium, aluminum, phosphorus, chlorine, calcium, manganese, iron, and strontium; 20 and 10 seconds for sulfur; and 50 and 25 seconds for fluorine. For the elements considered, the following standards, X-ray lines, and crystals were used: topaz, FKα, PC1 (W/Si); scapolite, ClKα, PET albite, NaKα, TAP; grossular, AlKα, TAP; diopside, MgKα, TAP; apatite, PKα, PET; apatite, CaKα, PET; barite, SKα, PET; synthetic rhodonite, MnKα, LIF; synthetic fayalite, FeKα, LIF; and SrTiO3, SrLα, TAP. Data reduction was done using the 'PAP' φ(ρZ) (phi-rho-Z) method (Pouchou and Pichoir 1985). Detection limits for fluorine, chlorine, and sulphur were 0.09, 0.06, and 0.05 oxide wt.%, respectively (all other detection limits are listed in Table 6.2). ‘Measured’ fluorine concentrations were transmuted into ‘initial’ fluorine concentrations, inferred to be the actual fluorine concentration in each grain prior to the start of analysis and the onset of diffusion. This was done using a linear regression of the five spot analyses to a time of zero seconds, as recommended by Raudsepp (1995) and Stormer et al. (1993); only the initial fluorine concentrations are reported in Table 6.2. In most cases, however, it was found that fluorine concentrations increased minimally, if at all, with time. This was particularly the case for the ideally oriented grains mounted in epoxy, although grains with non-ideal orientations in polished thin sections also demonstrated very little change in measured fluorine concentrations with time. It seems that an average of all measurements (as opposed to a linear regression) would suffice for determining the actual fluorine concentrations in most samples presented here. It was found, however, that analyses from polished thin sections were more likely to produce anomalous results. For example, the five spot analyses for some samples produced peaks and troughs in the fluorine measurements that were not consistent with steadily increasing fluorine diffusion with time. In other cases a steady decrease in fluorine with time was observed. These unusual results may have been due to non-ideal grain orientations in polished thin sections; analyses with inconsistent linear regressions were not used in any of the 366 calculations. Isolated analyses with anomalously high or low fluorine concentrations are tentatively interpreted to represent either xenocrystic apatite grains, or apatites that re-equilibrated with local or fracture-controlled fluid(s) of unknown origin; these analyses were also discarded. Atomic proportions and mole fractions of fluorine, chlorine, and hydroxyl (by stoichiometry) were normalized to 12 anions (excluding H2O, F, and Cl). Analyses with totals between 97.0 and 100.9 wt.% were used for plotting and interpretation purposes, although analyses with totals of 97.0–98.0 wt.% were monitored closely and only used if these analyses were indistinguishable from the analyses with higher totals. Averages and 1-sigma standard deviations for all analyses are presented as representative of each sample in Table 6.2; the actual compositional ranges of all the data used for interpretations and figures are provided in the text and in digital Appendix F.1. Procedures were employed to minimize complications associated with apatite EMP analysis, including: analyzing appropriately oriented apatite grains whenever possible to minimize fluorine diffusion along the c-axis, a lower accelerating voltage to reduce the absorption of halogens, relatively long count times to reduce the F-P peak overlap, and a topaz fluorine standard to remove errors associated with using apatite as a standard for fluorine measurements. Despite this care, many of the analyses produced fluorine concentrations higher than what can be stoichiometrically accommodated in apatite (>3.77 wt.% F; Table 6.2). An issue with determining compositions by EMP methods analysis is that several anions of direct interest to the study magmatic fluid compositions cannot be measured, specifically hydroxyl and carbonate. Hydroxyl was calculated by stoichiometry from fluorine and chlorine assuming that it—along with measured fluorine and chlorine—fills the halogen site and that there are no vacancies or other significant anion substitutions. However, the invariable presence of minor amounts of carbonate in the halogen site of apatites from felsic to intermediate intrusions may have contributed to the inflated (>3.77 wt.%) fluorine concentrations in some samples (e.g., Binder and Troll, 1989; Piccoli and Candela, 2002). We suspect, therefore, otherwise ‘good’ analyses that returned fluorine concentrations higher than what can be stoichiometrically accommodated in apatite are due to the combined effects of minor carbonate and errors inherent to fluorine analysis in apatite using the EMP. Another issue with apatite datasets generally is the existence of compositional variations between grains from the same sample (e.g., Sha and Chappell, 1999). This is certainly observed here for most elements present in quantities above their detection limits, which is demonstrated by some of the larger 1-sigma standard deviations reported in Table 6.2. The differences are greater than what can be attributed to analytical uncertainty, given the detection limits listed in Table 6.2, which suggests these compositional variations are a real feature of apatite (Sha and Chappell, 1999). As there were no systematic differences in apatite compositions that could be attributed to hydrous vs. anhydrous host minerals, the most likely sources of this intra-sample variation are: (1) analysis of early versus late apatite grains, whereby protracted magmatic crystallization may record broad changes in magma composition with time (resulting from magma differentiation, mixing, and (or) crystallization); and (or), (2) analysis of unrecognized restitic, xenocrystic, or altered apatite grains. 367 6.3.2 Biotite Biotite was identified in polished thin sections for selected intrusive phases from the five multi-phase bodies. Each thin section was examined petrographically in transmitted light and imaged in BSE mode on the SEM. Euhedral to subhedral, magmatic biotite grains (in accordance with the criteria set by Miller et al., 1981) and ragged hydrothermal biotite in greisens and endoskarn were analyzed on the EMP operating in the wavelength-dispersion mode, following the methods of Coulson et al. (2001). Biotite analyses of 120 seconds each were done on 2-5 grains from each sample, under the following operating conditions: an excitation voltage of 15 kV; a beam current of 10 nA; and, a spot diameter of 10 μm. Respective peak and background count times were: 40 and 20 seconds for both fluorine and chlorine, and 20 and 10 seconds for all other elements. Data reduction was done using the 'PAP' φ(ρZ) method (Pouchou and Pichoir 1985). For the elements considered, the following standards, X-ray lines, and crystals were used: synthetic phlogopite, FKα, PC1 (W/Si); albite, NaKα, TAP; kyanite, AlKα, TAP; synthetic phlogopite, MgKα, TAP; synthetic phlogopite, SiKα, TAP; scapolite, ClKα, PET; synthetic phlogopite, KKα, PET; diopside, CaKα, PET; rutile, TiKα, PET; synthetic magnesiochromite, CrKα, LIF; synthetic rhodonite, MnKα, LIF; synthetic fayalite, FeKα, LIF; and barite, BaLα, PET. Detection limits for fluorine and chlorine, respectively, were 0.08 and 0.01 oxide wt.% (all other detection limits are listed in Table 6.3). Atomic proportions were normalized to 11 anions for biotite (excluding H2O, F, and Cl), so that the atomic proportions of fluorine, chlorine, and hydroxyl would be presented as molar quantities for the halogen site. Only analyses with totals of 98.0–100.9 wt.%, were used for plotting and interpretation purposes, unless there was no significant compositional variation for analyses with totals between 97.0 and 98.0 wt.%. Textural classifications of primary vs. secondary micas were re-evaluated after analysis with the assumption that notably depleted TiO2 and FeO in biotite are indicative of sub-solidus alteration and/or re-equilibration with later fluid(s) (e.g., Robert, 1976; Selby and Nesbitt, 2000; Nachit et al., 2005). The compositional information confirmed all of the primary (magmatic) and secondary (altered) textural classifications in Table 6.1, with the exception of sample KR-05-170 (see below). Averages and 1-sigma standard deviations for all analyses are presented as representative of each sample in Table 6.3; the actual compositional ranges of all the data used for interpretations and figures are provided in the text and in Appendix F.2. 6.4 RESULTS 6.4.1 Apatite Chemistry The two main constituents of apatite, CaO and P2O5, range from 51.7–56.0 wt.% (average 54.2 wt.%) and 39.6–43.6 wt.% (average 41. 7 wt.%), respectively, in primary apatite (Table 6.2; Fig. 6.3a,b). Within apatite grains from single hand samples, CaO and P2O5 content may range significantly (up to as much as 4.53 wt.% and 3.40 wt.%, respectively). Apatite associated with mineralization is very similar with respect to CaO (average 54.6 wt.%) and P2O5 (average 41.4 wt.%) content. In both primary apatite and apatite associated with mineralization, MnO (average: 368 0.51 and 0.22 wt.%, respectively) and FeO (average: 0.30 and 0.18 wt.%, respectively) are present in relatively high quantities (Table 6.2; Fig. 6.3a,b), typically above their minimum detection limits (0.12 wt.% for both oxides). MgO, Al2O3, and SrO are very low, usually below their detection limits (0.04, 0.03, and 0.11 wt.%, respectively), particularly in apatite associated with mineralization (Table 6.2). Sulfur in all but a few rare analyses and for all averaged sample values is below the detection limit of 0.05 oxide wt.%. Fluorine content ranges widely (1.04–4.53 wt.%, typically >3.5 wt.%) and correlates negatively with the predominantly low to very low (i.e., below detection) chlorine concentrations (0.02–1.06 wt.%, typically < 0.12 wt.%; Table 6.2; Fig. 6.3c). The majority of the samples analyzed contain near end-member fluorapatite, although fluorine is typically more enriched in apatite associated with mineralization (3.19–4.52 wt.%; average 3.86 wt.%), whereas chlorine displays the opposite trend (0.02–0.16 wt.%; average 0.07 wt.%; Table 6.2; Fig. 6.3c). Several samples have anomalous chlorine and (or) fluorine concentrations (Fig. 6.3c, circled and (or) labelled). A negative correlation between fluorine and P2O5 in apatite (not shown) was not observed here; therefore, significant quantities of carbonate are not likely to be present in the apatites (McArthur, 1990). 6.4.2 Biotite Chemistry Tri-octahedral mica (Mg + Mn + Cr + Fe2+ + Ti + VIAl > 2.5) or “biotite” compositions were plotted following the International Mineralogical Association (IMA) classification (Rieder et al., 1999) whereby atoms per formula unit (apfu) of total aluminum are plotted vs. Fe# (Fig. 6.4a). In general, the Fe# in fresh biotite (0.34–0.86) increases with increasing ∑Al (2.48–3.75) for all of the plutons (Table 6.3; Fig. 6.4a). Biotite compositions in the plutonic bodies form a well-defined linear trend—typical of calc-alkaline rocks (e.g., Mason, 1978)—from low Fe/Mg annite biotite near the transition to phlogopite towards the siderophyllite biotite end-member (Fig. 6.4a). The biotites examined here typically have unchanged to slightly lower Fe# with increasing AlVI (not shown) on an analysis-by- analysis basis within single samples. Some outlying compositional populations in Figure 6.4a deserve mention. Secondary mica inferred to be related to mineralizing fluids from a dyke associated with the Rudi pluton is phlogopite (Fe# <0.3; Fig. 6.4a, circled with dashed line), consistent with Fe+AlIV/(Fe+AlIV+Mg) less than 0.45 for altered or secondary biotite (Selby and Nesbitt, 2000). Three clusters of analyses that plot below the general linear biotite trend fall closer to the biotite- phlogopite transition (Fig. 6.4a, circled with dotted lines). At Cantung these relatively magnesian biotites are also from pervasively altered dykes affected by mineralizing fluids (Rasmussen et al., 2011b) and are similar in composition to biotite analyzed from the biotite skarn facies at Cantung (Zaw, 1976). Slightly more magnesian biotite was also found in a sulphide-bearing porphyry dyke with primary igneous textures at the Central Nahanni pluton (KR-05-133), and in one ‘fresh’ biotite grain with variable Fe# from a plutonic phase at Lened (KR-05-170). Several analyses of unusually iron-rich biotites from the most highly evolved or fractionated samples (KR-05-124, 173; 03M-203) plot above the general compositional trend (Fig. 6.4a, circled with solid line) and are discussed in Section 6.5.1.1. 369 Primary magmatic biotite is typically enriched in TiO2 (1.18–4.82 wt.%; average 3.30 wt.%) and MnO (0.12–1.44 wt.%; average 0.47 wt.%; Table 6.3). In some samples TiO2 is generally lower, but still within the range of magmatic TiO2 content observed here, including: biotite from several of the more highly evolved samples (e.g., sample KR-05-124: average 2.63 wt.%); biotite associated with mineralization at Rudi (average 2.28 wt.%); and, biotite from the unmineralized Mine Stock at Cantung (average 2.73 wt.%). Na2O (average 0.09 wt.%), Cr2O3 (<0.15 wt.%), and CaO (<0.07 wt.%) are low in all samples (Table 6.3). Fluorine and chlorine are also present in low abundances (<1.73 and <0.32 wt.%) in all samples—but above their respective detection limits of 0.08 and 0.01 wt.%—and generally correlate inversely (Fig. 6.4b). 6.5 DISCUSSION 6.5.1 Magmatic Differentiation from Apatite and Biotite Compositions The concentration of several major element oxides in apatite and biotite may be tied to an increasing degree of magmatic evolution, or magmatic differentiation, by assimilation/contamination and fractional crystallization (AFC) processes. Increasing MnO and FeO, and decreasing CaO and SrO in apatite have been related to compositional features indicative of differentiation, such as elevated aluminicity and reduced oxygen fugacity (Sha and Chappell, 1999; Hoskin et al., 2000; Belousova et al., 2001; Piccoli and Candela, 2002). In the plutons studied here, MnO and FeO concentrations in apatite do not display any discernible pattern or trend within samples and between intrusive phases from the same plutonic body (Fig. 6.3a,b; Sha and Chappell, 1999; Hoskin et al., 2000). However, when the entire apatite dataset is plotted it is clear that MnO—and perhaps, to a lesser extent, FeO—do correlate negatively with CaO (Fig. 6.3a,b). The abundance of MnO (and possibly FeO) vs. CaO in apatite, therefore, appears to allow us to infer the relative degree of differentiation undergone by the plutonic bodies, if not the internal degree of differentiation. Based on apatite compositions, therefore, the Central Nahanni pluton appears to be the least evolved plutonic body sampled here, followed by Cantung and Rudi, Rifle Range, and finally Lened as the most evolved plutonic body. On a plot of MnO vs. XF/XCl in primary apatite—both of which should increase with magmatic differentiation—two clusters of data arguably display positive correlations (Fig. 6.3d): a tightly constrained group of samples from Cantung and Rudi with high XF/XCl, and a very scattered group of samples from Cantung, Lened, and Rifle Range. Apatites from the Central Nahanni pluton, which is the one intrusion sampled here with a large range of internal compositional variation, do not display a correlation between MnO and XF/XCl. The poor or absent correlation between F:Cl and magmatic evolution is not unique to this study (e.g., Hoskin et al., 2000). Although other studies have successfully linked XF/XCl in apatite, or variations thereof, to magmatic differentiation (Tsusue et al., 1981; Candela, 1986; Lang, 1998; Sha and Chappell, 1999; Sallet, 2000; Belousova et al., 2001; Rasmussen et al., 2012), it does not appear to be particularly reliable or well-defined. Biotite compositions also display enough variation to be useful as indicators of differentiation. Here, an overall increase in total aluminum and Fe# in biotite for samples from the Central Nahanni pluton through to the Lened pluton (Fig. 6.4a) mirrors the MnO vs. CaO trend in apatite. The biotite compositional trend, however, provides better constrained information on inter- and intra-intrusion differentiation than do the apatite compositional trends, 370 allowing the relative degree of differentiation to be determined on nearly a sample-by-sample basis. Differentiation indices calculated from the whole rock chemistry of one representative sample of the Central Nahanni, Rifle Range, and Lened plutons (Rasmussen et al., 2007, 2012), however, suggest that the Rifle Range pluton is significantly more evolved than the Lened pluton (or that the Rifle Range melt was closer to a pure SiO2– NaAlSiO4–KAlSiO4 system; Thornton and Tuttle, 1960). This differentiation index—calculated by summing the normative percentage of the minerals quartz, K-feldspar, and albite—has been used successfully to delineate inter- and intra-plutonic differentiation (e.g., Coulson et al., 2001; Rasmussen et al., 2007, 2012). One possible explanation for this discrepancy could be that the Lened pluton was more highly contaminated by (or derived from a greater proportion of) argillaceous supracrustal rocks than the Rifle Range pluton56. This could produce the apatite and biotite geochemical variations (i.e., depleted Ca; elevated Mn, Fe, and Al), and the lower proportion of non- normative minerals (feldspars and quartz) relative to minor minerals (micas) observed in the Lened pluton. Therefore, these suggested major and trace element components in magmatic apatite and biotite are still valid monitors of magmatic differentiation, but will vary depending on the specificities of the associated AFC process(es) and, in some cases, the composition of the magma protolith. 6.5.1.1 Unusually Iron-Rich Biotite In magmatic biotites the Fe# correlates positively with total aluminum content (see Fig. 6.4a), which has been related to an increasing degree of plutonic differentiation (Clarke, 1981). Biotites from highly evolved leucocratic intrusive phases in plutons associated with tungsten skarns (KR-05-173, 124 and 03M-203) approach near end- member siderophyllite compositions. However, these biotites are anomalously iron-rich in comparison to other biotites from highly differentiated intrusive phases with similar total aluminum contents. The unusual iron enrichment in these highly fractionated leucocratic intrusions is difficult to explain, but it is a real feature tied to the measured FeO concentrations (Table 6.3). While fluid/rock interactions with calcareous country rocks may affect the composition of biotite, there is no obvious reason for an iron versus magnesium enrichment in apparently magmatic biotite. In fact, relatively magnesian micas are found in the dyke at Rudi in which a pervasive calcic endoskarn mineral assemblage is suggestive of chemical exchange with carbonate rocks (KR-05-117). It is instead possible that the magmas containing the unusually iron-rich biotite may have been strongly contaminated by argillaceous wall rocks with a high Fe#. The lack of compositional data for nearby country rocks does not allow these possible sedimentary sources of iron to be tested. 6.5.2 Approximations of Magmatic Fluid Compositions Due to the sensitivity of fluorine and chlorine partitioning between fluids and minerals to internal and external factors, abundances of these halogens in magmatic minerals cannot be directly correlated to their abundances in a 56 The presence of coarse–grained magmatic muscovite throughout this intrusion is atypical of most magmatism in this study area, which―if muscovite-bearing—typically contain small, ragged sub-solvus grains (Rasmussen et al., 2007, 2011b). Magmatic muscovite is additional evidence that the Lened magma contains, or was strongly contaminated by a high proportion of argillaceous source material (Barbarin, 1996). 371 fluid or magma. Perhaps the most pronounced example of this are the compositional trends or negative correlations between fluorine and iron, and chlorine and magnesium (also known as the “Fe–F and Mg–Cl avoidance rules”) that commonly occur in sheet and chain silicate minerals. While the resulting trends may appear to indicate a change in fluid composition with respect to fluorine and chlorine, they are actually tied to a coupled substitution mechanism that is more strongly influenced by mineral composition. Changes to temperature and pressure conditions in the magma will affect the partitioning of halogens into minerals, which will likewise lead to compositional variations in the magmatic minerals that do not reflect the halogen contents of the magmatic volatiles and (or) fluid. Halogen concentrations in hydrous magmatic minerals may, therefore, only provide a proxy for magmatic fluid composition if the effects of internal (i.e., mineral composition) and external (e.g., temperature, pressure, and pH) factors are either removed, or are accounted and corrected for. 6.5.2.1 The Reciprocal Solid Solution Model Zhu and Sverjensky (1991) have developed an internally consistent database of natural and experimental data that agree with the thermodynamic principles of aqueous species. This was done in order to allow the estimation of activities for HF, HCl, and H2O species in solution. Relative fluoride and chloride concentrations in a magma or a fluid can be determined from the fluorine and chlorine concentrations in hydrous minerals—apatite and biotite were chosen due to the abundance of experimental and natural data—that have equilibrated with the magma or fluid. Calculated HCl activities only account for a fraction of the chloride species in a hydrothermal fluid (e.g., Cl–, NaCl0, KCl0 ± Mg– and Ca–chlorides; Zhu and Sverjensky, 1991); therefore, the activity of HCl is only used as a first-order approximation for the evolution of chloride species in the fluid (Coulson et al., 2001). Speciation calculations suggest that at HF comprises essentially all of the fluorine species in a hydrothermal fluid (Zhu and Sverjensky, 1991; Piccoli et al., 1999). Building on derived standard state thermodynamic properties for fluor- and chlor- end-member apatite and mica compositions (Zhu and Sverjensky, 1991), the Reciprocal Solid Solution Model (RSSM) was developed in order to account for the inherent solid solution between end-member mineral compositions in natural systems (Zhu and Sverjensky, 1992). Unlike most thermodynamic calculations that consider partitioning of F–Cl–OH between mineral and fluid, which is strongly controlled by pressure, temperature, pH, and fluid composition (e.g., Zhu and Sverjensky, 1991, 1992; Brenan, 1993a; Piccoli and Candela, 1994), the RSSM is based on the partitioning of F–Cl– OH between two mineral phases: apatite and biotite. Halogen exchange between solids essentially removes the variables of fluid composition and pressure, and significantly reduces the influence of temperature on the partitioning of F–Cl–OH. When the effects of fluid composition (e.g., pH) are removed, the variation in fluorine and chlorine concentrations in these minerals will only be related to crystal chemical effects (for biotite) and the actual halogen variations in the magma. The RSSM, thus, also accounts for cationic (primarily Fe2+ and Mg, and secondarily AlVI) temperature-dependent controls on halogen substitution in biotite. Ferric iron has been ignored in the RSSM as it is not a significant factor on fluorine substitution into mica (e.g., Icenhower and London, 1997), nor is it important in reduced tungsten skarn systems (Zaw and Clark, 1978; van Middelaar and Keith, 1990). The effects of other cations (e.g., Ti4+ and Mn2+) are considered to be inconsequential due to their relatively low 372 concentrations in biotite. Details on the thermodynamic calculations employed here are described by Coulson et al. (2001). Aqueous compositions for HF, HCl, and H2O are thus independently calculated from apatite and biotite compositions at a specified temperature and pressure. As calculated activities of HF, HCl, and H2O are only valid in comparison to each other (Zhu and Sverjensky, 1992), activity ratios are used to approximate the relative composition of the volatile material in equilibrium with apatite and biotite. 6.5.2.2 Pressure–Temperature Estimates Although the effects of pressure are negligible on F:Cl partitioning (e.g., Zhu and Sverjensky, 1992), the RSSM calculations do require a pressure estimate for the crystallization of apatite and biotite. A general pressure of 2 kbar based on the presence of andalusite in contact aureole rocks (e.g., Gordey and Anderson, 1993; Rasmussen et al., 2007) was used for the RSSM calculations. The partitioning of fluorine and chlorine between two solid phases is more dependent on temperature rather than pressure conditions (Zhu and Sverjensky, 1992). Apatite and biotite typically crystallize relatively close in time and space in felsic magmas, as is evidenced by their common petrographic association. In the samples examined here, apatite mainly occurs as euhedral inclusions in biotite and amphibole, or is found along the grain boundaries of these minerals. Rarely is apatite observed as inclusions in feldspars and quartz, or in the groundmass, but these grains tend to be very small—in most cases too small to analyze (e.g., <10 μm). In some cases, such as when apatite is included in early-forming minerals like plagioclase, it may be argued that these apatite grains saturated at slightly higher magmatic temperatures (Piccoli and Candela, 1994; Yang et al., 2004). Ubiquitous inclusions of apatite in unaltered biotite and the similar temperature ranges (detailed below) for apatite and biotite formation, however, suggest that in most cases both minerals crystallized under very similar magmatic conditions and in equilibrium with essentially the same magmatic volatiles. Although there is a wide range of apatite saturation temperatures (ASTs) reported in the literature, the temperature of apatite saturation is mainly controlled by the silica content of the magma (Harrison and Watson, 1984; Bea et al., 1992). Over 80% of apatite in felsic to intermediate magmas crystallizes within a temperature range of 30–60°C of the calculated AST, respectively (Piccoli and Candela, 1994); therefore, variation in the range of temperatures over which apatite has crystallized is unlikely to have a large effect on the partitioning of F–Cl–OH into apatite. ASTs were calculated using whole rock geochemical data for representative samples for each of the plutons examined here, and have a range of approximately 100°C: from 786°C for a peraluminous leucogranite dyke, to 870°C for metaluminous granodiorite (Table 6.4). The average AST for the samples in Table 6.4 is 831°C. For biotite at the Cantung tungsten skarn deposit, ferrous/total iron in skarn and magmatic micas are consistent with quartz–magnetite–fayalite to graphite–methane buffer conditions (Zaw and Clark, 1978; van Middelaar, 1988), which is extrapolated to the other intrusions sampled in the study area based on similar pluton compositions and 373 skarn mineralogies. Using the experimentally and theoretically determined QFM buffer equilibria curves of Wones and Eugster (1965) and the Fe/(Fe+Mg) composition of the biotites, the temperature of formation of magmatic biotite in the study area was approximated to range from 750 to 900°C, with an average temperature of about 800– 850°C. This 800–850°C range is also consistent with the upper stability limit for biotite in a synthetic, H2O - undersaturated granitic magma with low H2O contents (e.g., 2–4 wt.%; Naney, 1983) typical of granitoids associated with reduced tungsten skarns (Newberry and Swanson, 1986). For the thermodynamic calculations an approximate temperature of formation of 850°C was assumed for both apatite and biotite; this temperature is assumed to approximate saturation of these minerals species for most of the samples examined. The use of the same temperature conditions for both minerals in all of the samples examined is acceptable given the minimal variation within calculated activity ratios associated with temperature deviations (e.g., Zhu and Sverjensky, 1992; Piccoli and Candela, 1994; Kifer and Wolf, 2000). For example, this study and Coulson et al. (2001) found less than 0.1 log units of variation in the HCl/HF activity ratio (and <0.2 log units of variation in the H2O/HF activity ratio) for biotite formed at 650°C and at 850°C, and less than 0.1 log units of variation in the HCl/HF activity ratio for apatite formed at 700°C and at 900°C. 6.5.2.3 The Application of F:Cl:OH Variations and Activity Ratios One intuitive feature of magmatic–hydrothermal systems is that ratios of F:Cl:OH in a magma should not fluctuate if the conditions of a steady-state and closed-system magma are met (Candela 1986). The existence of a such a magmatic system for any period of time is improbable, and F:Cl:OH ratios will invariably fluctuate if either or both of these circumstances change. A change to steady-stage conditions manifested as a decrease in temperature will, for example, result in a decrease in Cl/F in apatite (Cl/OH should not be affected; Piccoli and Candela, 1994) and possibly in mica (cf. Figs. 17 and 19–21 in Zhu and Sverjensky, 1991). Perhaps one of the most important examples of open-system behaviour with respect to mineralizing systems is the off-gassing of an MVP. Although the partitioning of chlorine and H2O will depend to some degree on the conditions of MVP exsolution, as a general rule chlorine is strongly partitioned into fluid over melt when an hydrous MVP exsolves; fluorine displays the opposite behavior and will consistently partition into minerals in preference of melt (e.g., Zhu and Sverjensky, 1991). The loss of an MVP, therefore, should decrease Cl/F (and Cl/OH at >0.5 kbar) and OH/F in the magmatic system (Candela, 1986; Christiansen and Keith, 1996; Sheets et al., 1996; Webster, 1997a,b; Lang, 1998; Piccoli et al., 1999; Villemant and Boudon, 1999; Piccoli and Candela, 2002). This scenario is particularly applicable to the porphyry environment where decompression-related MVP saturation may occur early, prior to the crystallization of most magmatic minerals, including apatite and (or) biotite. In magmas that were not substantially affected by depressurization, the removal of early-crystallizing anhydrous minerals (e.g., pyroxenes, plagioclase) should enrich the residual melt in all volatile phases, theoretically without producing any variation in F:Cl:OH and behaving as a closed system with respect to these components. If simultaneous off-gassing of chlorine and H2O from the magma is occurring, however, relative fluorine concentrations in the magma should increase. In this way, higher F/Cl and F/OH will frequently characterize more 374 differentiated or felsic plutons (Tsusue et al., 1981; Candela, 1986; Lang, 1998; Sha and Chappell, 1999; Sallet, 2000; Belousova et al., 2001; Coulson et al., 2001; Rasmussen et al., 2007, 2012). If hydrous minerals (e.g., apatite and micas) are also being removed by crystal fractionation, the residual melt should become depleted in fluorine (Sallet, 2000). If off-gassing is not simultaneously occurring with hydrous mineral fractionation, an increase in the concentration of chlorine (Figs. 19 and 23 in Zhu and Sverjensky, 1991) and H2O in the system will compound the effects of hydrous mineral fractionation on the magmatic volatile composition and produce elevated Cl/F and OH/F in the magma. It is conceivable then that a magma evolving along this fractionation path will become depleted in fluorine and enriched in chlorine and H2O, at least until H2O saturates and large quantities of chlorine and H2O are exsolved as hydrous MVP. A similar scenario where crystal fractionation causes volatile saturation and fluid exsolution may be applicable to deep tungsten skarn systems if it occurs at supercritical conditions. To summarize, an early-stage MVP that is lost to a magma will be enriched in the volatile species that are depleted in the melt and post-exsolution magmatic minerals; this will deplete subsequent MVP in the same volatile components (e.g., Cl). A late-stage MVP that is lost to a magma will be enriched in the volatile species that are enriched in the melt and magmatic minerals. The former scenario typically occurs in subcritical (shallower) H2O- rich magmas—such as those associated with copper porphyries and skarns—where the early saturation of fluids is evidenced by the prevalence of porphyry textures in near-by intrusions (e.g., Candela, 1997). The latter scenario would be expected to occur in deeper (supercritical) or H2O-poorer magmas—such as those associated with tungsten skarns—where the late saturation of fluids is evidenced by coarse-grained plutonic textures in associated plutons (Newberry and Swanson, 1986; Newberry, 1998; Candela, 1997). In this way, approximations of F:Cl:OH contents in a magma may be used to evaluate intrusions in terms of their potential for associated mineralization, as well as the possible metal tenor and type of the mineralization (e.g., Candela, 1986; Lang, 1998; Selby and Nesbitt, 2000; Coulson et al., 2001; Rasmussen et al., 2007, 2012). Based on the above analysis we hypothesize that magmas associated with reduced tungsten skarn mineralization in the study area exsolved a late-stage mineralizing MVP enriched in HCl and H2O. Magmatic apatite and biotite in these intrusions, therefore, crystallized in a chlorine and H2O-rich environment and should provide relatively elevated HCl/HF and H2O/HF activities—expressed as the logarithmic functions log[aHCl/aHF] and log[aHOH/aHF]—as estimates of MVP composition. Although log[aHCl/aHF]AP,BT and log[aHOH/aHF]AP do not rely on any assumptions involving the activity of H2O, log[aHOH/aHF]BT was calculated assuming a pure or dilute aqueous fluid at unity (aHOH = 1; Zhu and Sverjensky, 1991). The H2O/HF activity ratios calculated from biotite were then compared to those calculated from apatite, and corrected for the initial assumption of a pure aqueous fluid by the reasoning that log[aHOH/aHF]AP and log[aHOH/aHF]BT should be roughly equivalent if both minerals are closely associated in time and space in a magma, and if neither mineral has re-equilibrated with later fluids. Using the linear relationship between log[aHOH/aHF]AP and log[aHOH/aHF]BT for samples where apatite and biotite data are not clearly decoupled (Fig. 6.5), log[aHOH/aHF]BT was re-calculated using the equation: log[aHOH/aHF]BT-corr. = (log[aHOH/aHF]BT-uncorr. + 1.4023) / 0.3667. Averages of the calculated log[aHCl/aHF] and log[aHOH/aHF] values and respective 1-sigma standard deviations for each sample are tabulated in Tables 6.2 and 6.3. All of the calculated activity ratios, however, are plotted in Figure 375 6.6, and the complete datasets are available in Appendices F.1 and F.2. Apatite with compositions consistent with true end-member fluorapatite could not be used for interpretations due to the impossibility of calculating HF/HCl and H2O/HF activity ratios; the activity ratios for these analyses are denoted by “n/a” in the data tables. In most samples with end-member fluorapatite, primary biotite data provide the only estimate of magmatic volatile compositions. Log[aHCl/aHF] and log[aHOH/aHF] are plotted versus relative degrees of differentiation in Figure 6.6a, and versus known or approximated distances from tungsten skarn mineralization for plutons associated with mineralization in Figure 6.6b. 6.5.2.4 Unmineralized Case Studies Overall log[aHCl/aHF] and log[aHOH/aHF] for the intermediate to felsic, metaluminous Central Nahanni pluton comprise a restricted range of generally positive values (Fig. 6.6a). This dominance of HCl and H2O over HF in the magma is typical of other metaluminous intrusions in the study area (Coulson et al., 2001; Lang, 2001; Rasmussen et al., 2007, 2012). The observed gradual decrease in HCl/HF and H2O/HF with an increasing degree of differentiation in the Central Nahanni pluton has also been observed in other intrusions in the region (Coulson et al., 2001; Lang, 2001; Rasmussen et al., 2007, 2012) and elsewhere (e.g., Tsusue et al., 1981; Lang, 1998; Sha and Chappell, 1999; Belousova et al., 2001). Log[aHCl/aHF] determined by magmatic apatite for the felsic and peraluminous Rifle Range pluton, conversely, has a restricted range of mainly moderately to slightly negative values, whereas log[aHOH/aHF] is typically moderately positive (although less so than in the Central Nahanni pluton; Fig. 6.6a). While the magmatic volatiles in the Rifle Range pluton were also dominated—to a lesser degree—by H2O, unlike in the Central Nahanni pluton HF was generally present in greater abundances than HCl (Fig. 6.6a). The increase in HF relative to HCl for more differentiated (i.e., felsic and aluminous) magmas in the Rifle Range pluton also characterizes other peraluminous intrusions in the study area (Coulson et al., 2001; Rasmussen et al., 2007, 2012). As with the Central Nahanni pluton, HCl/HF (and possibly H2O/HF) activity ratios determined from apatite data decrease with increasing degree of differentiation for the unaltered Rifle Range samples. However, the HCl/HF and H2O/HF activity ratios determined from biotite data for samples KR-05-150b and KR-05-151 actually increase with differentiation. These decoupled estimates activity ratios from apatite and biotite are discussed below. Due to the relative simplicity of the Central Nahanni and Rifle Range plutonic systems and their lack of association with known mineralization, several distinctions between apatite and biotite estimates of magmatic fluid composition that are evident in Figure 6.6a may be addressed. These include small variations in activity ratios, large variations in activity ratios or decoupled trends, and a greater range in calculated activity ratios for most of the apatite datasets. (a) Small Variations in Apatite and Biotite Estimates of Fluid Composition: Many samples display consistent small variations in the activity ratios determined from apatite and biotite data. For the Central Nahanni samples there is a slight decrease in log[aHCl/aHF] and a slight increase in log[aHOH/aHF] values 376 from apatite to biotite (Fig. 6.6a). Taken at face value this implies that magmatic fluids were slightly HCl-poorer and H2O-richer—relative to HF—when biotite equilibrated with the magma. The opposite trend was observed in the Emerald Lake pluton, another mostly metaluminous composite intrusion (Coulson et al., 2001), implying that biotite at Emerald Lake crystallized in an HCl-richer environment. Unlike pressure and fluid composition, temperature will have a small effect on the partitioning of halogens (particularly F) into apatite and biotite (Zhu and Sverjensky, 1992). Changes in temperature, or errors in the temperatures used for thermodynamic calculations could also be responsible for these small variations in volatile compositions between the apatite and biotite datasets. Given the ubiquity of euhedral apatite inclusions in biotite from all of the samples examined here, and the slightly lower upper stability limits for biotite in H2O-undersaturated magmas (e.g., ca. 800°C for < 2 wt.% H2O; Naney, 1983), apatite is likely to have saturated in the magma at somewhat higher temperatures than biotite (e.g., Coulson et al., 2001). Compared to apatite, biotite may also have a lower closure temperature to halogen exchange, which is consistent with the greater susceptibility of biotite to re- equilibration with later or cooler fluids (e.g.,van Middelaar and Keith, 1990; Piccoli et al., 1999). Another factor could be the choice of actual temperatures used in the calculations. Although an average temperature of 850°C was approximated for both apatite and biotite in Section 6.5.2.2 as upper temperature limit for the stability of both minerals in felsic magmas, apatite and biotite actually crystallize over a range of temperatures as the magma cools (e.g., Whitney, 1988; Piccoli and Candela, 1994). Furthermore, the minerals should have ultimately equilibrated with the magmatic fluid at the temperature at which halogen exchange with the melt ceased; this ‘closure temperature’ is likely much lower than the saturation temperature used here for calculations. Unfortunately, the lack of constraints on halogen closure temperatures in apatite and biotite requires that we use the upper temperature limits for both minerals, with the expectation that the overall trends documented by the activity ratios are still valid. The very small early HCl enrichment suggested by the apatite data for the Central Nahanni pluton is certainly possible within the magmatic environment, particularly if apatite equilibrated with higher temperature (earlier) magmatic fluids that were slightly HCl-richer and H2O-poorer. Small quantities of chlorine and alkali–chloride complexes are very insoluble in felsic magmas (e.g., Webster, 1997b) and the off-gassing of some highly saline fluid will deplete subsequent MVP in chlorine without notably depleting H2O in the magma (Christiansen and Keith, 1996). The elevated H2O content suggested by biotite-based estimates of fluid composition for Central Nahanni is consistent with a decrease in the salinity of later MVP given the relatively high (2 kbar) emplacement pressure inferred for plutons in this study (e.g., Audétat et al., 2008). Thus, any changes in magmatic volatile composition— whether real (i.e., due to off-gassing or fluid evolution) or apparent (i.e., due to inappropriate temperature estimates)—may be manifested as these small differences in the activity ratios determined from apatite and biotite. When this reasoning is considered, along with the analytical difficulties associated with EMP analysis of halogens in apatite and biotite (summarized in Zhu and Sverjensky, 1992) and limits on our knowledge on the cationic controls for biotite (Zhu and Sverjensky, 1991), it seems that small variations in calculated activity ratios for apatite and biotite are unavoidable. The combined effects of these issues, however, should not significantly affect the 377 interpretation of F:Cl:OH activity ratios or the general trends exhibited by these ratios. (b) Decoupling Between Apatite and Biotite Estimates of Fluid Composition: Opposing or decoupled trends are present in the activity ratios calculated from biotite ± apatite for at least two, apparently unaltered, Rifle Range samples (KR-05-150b and 151; Fig. 6.6a). In these samples the apatite compositions imply that HCl/HF decreased in the magmas with increasing degree of differentiation, which is consistent with the off-gassing of increasingly HCl depleted magmatic fluids. The biotite compositions imply that HCl/HF and H2O/HF increased in the magmas with increasing degree of differentiation, which is consistent with crystallization in the presence of an HCl- and H2O-rich magmatic fluid. These diverging estimates of magmatic volatile composition suggest a disequilibrium between apatite and biotite compositions. The biotites do not appear to be altered petrographically, and they have magmatic morphologies and compositions (e.g., elevated TiO2 and Fe#; e.g., Moore and Czamanske, 1973; Selby and Nesbitt, 2000; Natchit et al., 2005). Furthermore, re-equilibration with a late hydrothermal fluid would be more likely to increase, rather than decrease, the relative fluorine contents in these micas. This leaves the option that biotite in these unaltered samples equilibrated with a magma that became more enriched in chlorine and H2O as it cooled. These decoupled activity ratios might, therefore, be attributed to an early off-gassing for chlorine and H2O-rich fluid prior to the closure temperature for halogen exchange in apatite, followed by a later-stage build-up of chlorine and H2O in the magma prior to the closure temperature for halogen exchange in biotite. (c) Elevated Variation in Apatite Estimates of Fluid Composition: There is also a higher degree of scatter in log[aHCl/aHF] and log[aHOH/aHF] values for apatite from the Central Nahanni and Rifle Range plutons (as well the other intrusions in Fig. 6.6). Real or natural intra-sample compositional variation in apatite compositions has been reported by other studies (Sha and Chappell, 1999). This could translate to a larger range in calculated activity ratios from apatite. The elevated variation in estimates of magmatic fluid composition from apatite could also be attributed to analytical difficulties and larger calculation errors tied to higher fluorine contents in apatite (Zhu and Sverjensky, 1992; Coulson et al., 2001). More scattered apatite-based data certainly characterizes many samples with near end-member fluorapatite, where chlorine contents are at or below the detection limit or calculated hydroxyl contents are very low to nil. Despite the greater range of variability in activity ratios imparted by the apatite datasets, log[aHCl/aHF] values determined from the apatite and biotite datasets are generally in good agreement. (The observed agreement between log[aHOH/aHF] values from apatite and biotite is inherent here because of the reliance of corrected biotite-based activity ratios on the apatite dataset; see Section 6.5.2.3). 378 6.5.2.5 Mineralized Case Studies (a) Rudi: Small-Volume Tungsten Skarn: At Rudi, slightly negative log[aHCl/aHF] and moderately positive log[aHOH/aHF] for biotite are relatively constant for the two least differentiated intrusive phases (KR-05-121a and 122) and for greisenized samples adjacent to exoskarn mineralization (KR-05-125b and 123; Fig. 6.6a). Exceptions are a marginal leucogranite that equilibrated with a magma depleted in HCl (and H2O) relative to HF (KR-05-124), and a late dyke altered to a “tonalitic” endoskarn mineralogy that equilibrated with a magma strongly enriched in HCl and H2O relative to HF (KR-05-117). The similar activity ratios for pervasively greisenized samples and the least differentiated magmatic phases allow, but do not require, that the greisen minerals and unaltered magmatic minerals equilibrated with a fluid of the same composition. Preferential partitioning of fluorine into minerals, particularly apatite, from both magmatic and hydrothermal fluids will also lead to the precipitation of relatively fluorine-rich (or chlorine-poor) minerals, even if the fluid is very poor in fluorine (e.g., Zhu and Sverjensky, 1991). Rather than partition from fluid into mineral(s), chlorine and other halogens will preferentially remain in the fluid and (or) be dispersed distally into the country rocks (e.g., Zhang et al., 2001). Therefore, the composition of skarn or greisen minerals will almost invariably be fluorine-rich, although the degree of fluorine-enrichment can vary between different magmatic–hydrothermal systems (e.g., Climax-type Mo-porphyries vs. W-skarns vs. Cu-porphyries; van Middelaar and Keith, 1990; Selby and Nesbitt, 2000). The activity ratios determined for the unaltered leucogranitic phase (KR-05-124; Fig. 6.6b) are consistent with strongly HCl-(H2O) depleted magma and fluid, or one that exsolved an early MVP prior to apatite and biotite crystallization. The most proximal phase to mineralization, however, is a late dyke (KR-05-117; Fig. 6.6b) that was metasomatized to a “tonalitic” endoskarn mineral assemblage similar to that exhibited by the late aplitic dykes at Cantung (Rasmussen et al., 2011b). The dyke contains apatite that equilibrated with an HCl-H2O-rich fluid and secondary phlogopite that equilibrated with an H2O-richer and (or) halogen-poorer fluid (given that apatite and biotite estimates of log[aHOH/aHF] vary, but log[aHCl/aHF] do not; Fig. 6.6). This may be due to an increase in H2O relative to HF and HCl—or a loss of HF and HCl at constant HCl/HF from the magma—at temperatures below the closure temperature to halogen exchange in apatite and above the closure temperature to halogen exchange in biotite. Alternatively, exchange with a very halogen-poor hydrothermal fluid could produce a similar compositional change in the phlogopite. Unlike the leucogranite, it appears that there was either a build-up of volatiles—relative to HF—in the dyke magma and late-stage magmatic fluids, or that the minerals in the dyke were re-equilibrated with a late, strongly HF depleted, hydrothermal fluid. (b) Lened: Moderate-Volume Tungsten Skarn: For intrusive phases at Lened, log[aHCl/aHF] is typically negative and log[aHOH/aHF] is moderately positive; both activity ratios decrease gradually with increasing degree of differentiation (Fig. 6.6a). Exceptions to this trend are 379 the two most evolved leucogranitic phases (KR-05-176 and 173), which become increasingly enriched in HCl and H2O relative to HF. As in the Rudi pluton, apatite and biotite from one sample at Lened (KR-05-173) are clearly decoupled with respect to fluid composition: apatite appears to have equilibrated with a HCl–H2O enriched fluid, whereas biotite has equilibrated with an even more HCl–H2O enriched (or HF depleted) fluid. When plotted against the distance of each sample to tungsten skarn mineralization, there is an increase in the HCl/HF activity ratios with increasing proximity to mineralization at Lened (Fig. 6.6b). Apatite and biotite in the two leucocratic samples closest to mineralization crystallized from the most HCl–H2O enriched magma in equilibrium with a similarly enriched fluid. This is particularly the case for biotites from sample KR-05-173, which texturally and geochemically appear to be of magmatic origin. It seems that the magmatic volatiles in this leucogranite became progressively more HCl- and H2O enriched—or HF depleted—prior to the closure of halogen exchange in biotite, despite an overall enrichment in HCl and H2O in comparison to other intrusive phases at Lened. While apatite compositions indicate unusually high HCl and H2O in the early magma and early fluids, the biotite compositions in the most proximal leucogranite appear to be reflecting a more anomalous build-up of HCl–H2O-rich volatiles in the late magma and late fluids. (c) Cantung: Large-Volume Tungsten Skarn: At Cantung the complexity of multiple dyke phases, several of which are pervasively altered and contain hydrothermal biotite ± apatite, complicates the interpretation of the activity ratios in Figure 6.6. The activity ratios (determined from biotite since mainly the fluorine-rich apatite end-member was found at Cantung) arguably decrease with increasing degree of differentiation for several unaltered samples (03M-207; KR-03_07-18 and 09-5). Activity ratios determined for a series of steep-dipping late aplitic dykes altered to a calcic-endoskarn mineral assemblage (03M-204 and 212; KR-03_31-7, 07-22, and 31-7) are essentially indistinguishable when compared to unaltered plutonic phases at Cantung (Fig. 6.6a). As with most samples in the previous examples of peraluminous intrusions, these activity ratios occupy a relatively restricted range of moderately negative values for log[aHCl/aHF] and moderately positive values for log[aHOH/aHF], with the exception of sample 03M-203—a late biotite-bearing leucocratic dyke. The restricted range of activity ratios in Figure 6.6 is not unique to this study. Similarly constrained ranges of fluid composition at Cantung were also determined for the micas in several intrusive and skarn phases; this was interpreted to be the result of a system-wide re-equilibration of micas with a late magmatic–hydrothermal fluid (Zaw and Clark, 1978; van Middelaar and Keith, 1990). Unlike most of the samples from Cantung, however, the dyke sample 03M-203 does not appear to have been affected by the same widespread re-equilibration. This dyke was anomalously enriched with respect to HCl/HF and H2O/HF, although a later enrichment in H2O (or a depletion in halogens) is also apparent in the log[aHOH/aHF] values calculated from biotite. The biotite-bearing leucocratic dyke at Cantung may be representative of a highly evolved residual melt at depth that coexisted and equilibrated with a volatile-rich magma and fluid (Rasmussen et al., 2011b). However, the lack of extremely anomalous activity ratios at Cantung similar to those observed at Rudi or Lened may be in part because the extremely fractionated residual 380 melt from which the bulk of the mineralizing fluids are interpreted to have been exsolved is not exposed, at least not as a pristine melt (Rasmussen et al., 2011b). 6.6 IMPLICATIONS 6.6.1 Magmatic Fluid Compositions of Tungsten-Skarn Systems in the Northern Cordillera The composition of magmatic fluids determined from apatite and biotite compositions in intrusive phases that are not directly of closely associated with tungsten skarn mineralization display reasonably predictable trends and compositions that are similar to those exhibited by other plutonic rocks in the northern Cordillera (van Middelaar and Keith, 1990; Coulson et al., 2001; Lang, 2001; Rasmussen et al., 2007, 2012), and elsewhere (Tsusue et al., 1981; Christiansen and Keith, 1996; Lang, 1998; Sha and Chappell, 1999; Belousova et al., 2001). For example, HCl/HF and H2O/HF decrease with increasing degree of magmatic differentiation, metaluminous intrusions produced fluids with slightly positive HCl/HF and moderately positive H2O/HF, and peraluminous intrusions produced fluids with slightly negative HCl/HF and slightly positive H2O/HF. The plutons associated with tungsten skarn mineralization also have MVP compositions and trends that are consistent with their peraluminicity and relatively high degree of differentiation. However, the intrusive phases that were emplaced closest to—or that are otherwise associated with—tungsten skarn mineralization at Rudi, Lened, and Cantung appear to have produced a compositionally distinct magmatic fluid. This MVP was unusually HCl/HF- and H2O/HF-rich when it equilibrated with magmatic apatites, and it evolved to become even more H2O/HF-rich (and, at Lened, HCl/HF-rich) by the time it equilibrated with magmatic biotites. These results are consistent with our hypothesis that apatite and biotite from intrusive phases that exsolved a mineralizing fluid in a tungsten skarn-forming system will provide elevated HCl/HF and H2O/HF activity ratios that are indicative of built-up of chlorine and H2O in the magma. The H2O/HF enrichment in the magmatic volatiles determined from biotite in the magmas most closely associated with mineralization would suggest that biotite in these samples either was open to F–Cl–OH exchange longer than apatite, or was re-equilibrated with sub-solidus magmatic–hydrothermal fluids. In the latter scenario, an increase or at least stagnation in the ratio of fluorine to H2O in the biotites would be expected (unless the fluid was severely deficient in fluorine; cf. Zhu and Sverjensky, 1991), which is clearly not the case. Instead, it may be argued that the biotite equilibrated with a fluid that had evolved to become particularly H2O-rich (relative to HF), possibly at lower magmatic temperatures. The lack of variation in HCl/HF activity ratios from both minerals implies that the increase in estimated H2O/HF activity ratios, from apatite to biotite, reflects a relative enrichment in H2O compared to both HF and HCl. This inferred H2O/HCl enrichment is consistent with observed and predicted decreases in the salinity of magmatic fluids exsolved at higher degrees of magma crystallinity, in systems at high confining pressures (>1.3 kbar; e.g., Audétat et al., 2008). In this instance then, the activity ratios determined from apatite compositions would provide a record of earlier (i.e., higher temperature and lower magmatic crystallinity) fluid compositions, and the activity ratios determined from biotite compositions would provide a record of later (i.e., lower temperature and higher magmatic crystallinity) fluid compositions. Significantly decoupled estimates of H2O/HF are not observed in 381 any of the more distal intrusive phases in the plutons associated with mineralization, although some decoupling is also observed in the most differentiated phases of the barren intrusions (KR-05-131 and 156). However, when combined with unusually elevated HCl/HF, this late decrease in salinity appears to be a characteristic of magmatic fluids from intrusions that are closely associated with tungsten skarn mineralization. 6.6.1.1 HCl/HF and H2O/HF Activity Ratios as Tungsten Mineralization Vectors Anomalously high HCl/HF and H2O/HF activity ratios characterize the intrusive phases most proximal to mineralization in two of the three mineralized case studies examined (Rudi and Lened). These estimates of magmatic fluid composition do not appear to a product of the degree of differentiation experienced by the proximal intrusive phases (at Lened and Cantung these intrusions are highly evolved leucogranites that would instead be more HF enriched if the degree of differentiation was a factor). In other peraluminous magmatic systems, therefore, unusually elevated HCl/HF and H2O/HF activity ratios calculated from apatite and (or) biotite compositions may also indicate a build-up of magmatic volatiles within a crystallizing magma. The identification of intrusive phases that are anomalously enriched in chlorine and H2O relative to fluorine—particularly when they would normally be enriched in fluorine relative to chlorine given their degree of differentiation—could help to target intrusions from which a late-stage, tungsten-carrying MVP may have exsolved. Identifying these intrusions may be one method of focusing exploration efforts for mineralization in which the late-stage loss of an MVP is an important factor in the formation of an ore deposit, as it is for tungsten skarn systems (e.g., Newberry, 1998; Newberry and Swanson, 1986). 6.6.2 Apatite- versus Biotite-Based Estimates of Magmatic Fluid Composition Although apatite is less vulnerable to sub-solidus halogen exchange than biotite (e.g., Stormer and Carmichael, 1971; Roegge et al., 1974; Candela, 1986; Munoz, 1990; Piccoli et al., 1999), in most of the samples presented here the calculated activity ratios in biotite mirror those of apatite very closely. This implies that biotite in most of the intrusive phases was not re-equilibrated with later fluids, unless apatite was also re-equilibrated with the same fluid. Some exceptions do, however, exist. Large differences in the activity ratios determined from biotite and apatite are limited to three intrusive phases (KR-05-117 and 173; 03M-203). In these three examples, apatite and biotite compositions must have been strongly decoupled by a significant late-stage build-up of H2O in the magmas. Other smaller, but still significant increases in HCl/HF ± H2O/HF activities from apatite to biotite were also observed in several samples not related to mineralization (KR-05-123, 125b, 131, 152, 155, and 156). These smaller variations are likely the result of small build-ups of magmatic volatiles in these intrusive phases over the temperature window defined by apatite and biotite closure temperatures to halogen exchange. Aside from the above exceptions, it appears that the biotite activity ratios provide a good approximation of the composition of volatiles in magmatic systems where apatite either did not saturate in the melt, or where apatite has end-member fluorapatite compositions (Ekström, 1972; Zhu and Sverjensky, 1991; Kifer and Wolf, 2000). Even in cases where end-member fluorapatite is not present, however, we would agree with Ekström (1972) that the 382 collection of compositional data for approximating magmatic fluid conditions is preferably obtained from biotite rather than apatite. There are several reasons for this: (a) the time-consuming and error-prone method of determining reasonable fluorine concentrations in apatite; (b) the limits associated with magmatic fluid composition calculations for apatite when end-member fluorapatite is present; (c) the larger errors associated with fluorine-enrichment in apatite (Zhu and Sverjensky, 1992); and (d) the ability to analyze in situ biotite with relative ease and to test whether apparently primary biotite has been affected by late-stage fluids using biotite compositions (e.g., TiO2 concentration and log[XF/XOH] vs. Fe# in biotite). Unless a study is focused on hydrothermal biotite compositions and associated fluids, however, appropriately oriented coexisting apatite should be analysed whenever possible to test for biotite re- equilibration. 6.7 CONCLUSIONS This study provides information on apatite and biotite compositions from crustally derived plutons in an area of the northern Cordillera that is particularly prospective for reduced tungsten skarn mineralization. The examination of apatite and biotite compositions in intrusive phases from five composite plutons provided additional, although not definitive, information on the petrogenesis of magmatism in the eastern Selwyn Basin. The focus of this study, however, was to: (a) examine the evolution of magmatic volatiles with respect to F:Cl:OH and magmatic differentiation, (b) generally characterize magmatic fluid compositions in representative plutons, and (c) identify specific intrusive phases that are likely to be important in the genesis of tungsten skarn mineralization. 6.7.1 Magmatic Petrogenesis These geochemical markers in apatite may also be used as relative indicators of inter-plutonic differentiation. MnO (± FeO, XF:XCl) very broadly increase with decreasing CaO, implying a differentiation order of (from least to most evolved): Central Nahanni, Cantung and (or) Rudi, Rifle Range, and Lened. Increasing Fe# and total aluminum in biotite mirrors this differentiation trend from the Central Nahanni to Lened plutons, but is a much better constrained indicator of both the inter- and intra-plutonic degree of differentiation. However, the Lened pluton provides an example by which these geochemical indicators of magmatic differentiation may not compare with some other differentiation indices (e.g., Thornton and Tuttle, 1960) due to the specific processes by which differentiation occurred ± the composition of the magma protolith. The unusually iron-rich biotite detected in three leucocratic, highly fractionated and (or) wallrock-contaminated samples from the plutons associated with tungsten mineralization also have anomalous activity ratios that are suggestive of equilibration with either an unusually chlorine and H2O-rich magma compared to other samples from the same intrusive body (KR-05-173; 03M-203), or a relatively chlorine- and H2O-poor magma that may have exsolved an early MVP (KR-05-124). It may be that the iron enrichment is somehow related to magmatic or contamination processes that also led to a late-stage build-up of HCl and H2O in at least of two magmas and the strong depletion in HCl and H2O (due to the early loss of an MVP?) in the third magma. Regardless of the cause of the iron enrichment, it appears that the near end-member siderophyllitic biotites with anomalously high iron contents 383 are found in highly evolved leucogranitic intrusive phases that also have anomalous HCl/HF and H2O/HF activity ratios. Two such intrusive phases are spatially associated with tungsten mineralization and one appears to have exsolved an early saline and hydrous MVP. 6.7.2 Magmatic Fluid Composition Activity ratios—calculated here as log[aHCl/aHF] and log[aHOH/aHF] from apatite and biotite compositions—may be used to predict aqueous fluid compositions, and make inferences regarding the evolution of magmatic fluids in intrusions associated with mineralization. The composition of magmatic fluid in equilibrium with apatite and biotite in the intrusions examined as a part of this study generally evolves to less HCl/HF and H2O/HF-rich compositions with increasing decree of differentiation (or SiO2 content). HCl/HF is typically positive in metaluminous plutons, whereas HCl/HF is typically negative in peraluminous plutons; H2O/HF is generally positive overall, but particularly in the metaluminous plutons. Despite small variations, the activity ratios indicate that apatite and biotite equilibrated with comparatively similar magmatic fluids in most samples. Each of the three plutons associated with tungsten skarn mineralization contain one intrusive phase that equilibrated with an MVP that was anomalously enriched in HCl and H2O relative to HF (despite the relatively high degrees of differentiation exhibited by two of the three intrusive phases). Furthermore, these three samples are also anomalous in that the apatite- and biotite-based estimates of H2O/HF are strongly decoupled, implying an enrichment in overall H2O content (i.e., a reduction in salinity) in the fluid at the time of biotite crystallization (i.e., at greater magmatic crystallinity). In summary, apatite appears to have equilibrated with an earlier HCl- and H2O-rich magmatic fluid; and (b) biotite appears to have equilibrated with a later similarly HCl- rich, but particularly H2O-rich magmatic fluid. The combination of these fluid compositions is unique to the three individual intrusive phases that have been related to the tungsten skarn mineralization at Rudi, Lened, and Cantung. While significant decoupling of apatite and biotite estimates of fluid composition do occur in several other intrusive phases from both the barren and mineralizing plutons, they do not appear to have produced a significant H2O-rich MVP of moderate to low salinity into which tungsten would preferentially partition. 6.7.2.1 A Model for Magmatic Tungsten Skarn Mineralization We suggest that the three specific intrusive phases associated with tungsten mineralization exsolved a supercritical hydrous brine of moderate to low salinity as a result of slow crystallization of the magma at moderate depths (i.e., supercritical conditions). As the magma cooled, apatite and biotite crystallized, and as fluid exsolution occurred on an ongoing basis, the magmatic fluid evolved to less saline compositions. Incompatible elements concentrated in the residual melt—such as tungsten, in reduced and ilmenite-poor magmas—are increasingly partitioned into the fluid with ongoing magma solidification. This would lead to significant tungsten enrichments in the latest-stage (near- solidus) fluids. Ascension of the tungsten-bearing fluid through the cooling magma may also have aided in scavenging tungsten from the melt. Periodic or a continuous trapping of the ascending fluid at the upper crystallization front of the magma chamber (i.e., the plutonic carapace) would have led to an increase in fluid 384 pressure in this region. Eventual overpressuring of the fluid was responsible for creating (or re-opening fractures) and injecting the metal-rich MVP through the carapace into the overlying country rocks, transporting metals to sites of skarn mineralization in calcareous hosts. This may have been a single late-stage process, but it more likely occurred multiple times before the magma was fully solidified in intrusions associated with significant tungsten mineralization (e.g., see evidence for multiple episodes of sub-vertical fluid flow at Cantung described in Rasmussen et al., 2011b). This model agrees with shared characteristics and current interpretations of reduced tungsten skarn mineralizing systems (e.g., Newberry and Swanson, 1986; Newberry, 1998; Meinert et al., 2005; Rasmussen et al., 2011b). 6.7.2.2 Final Remarks Attempts to use concentrations of fluorine, chlorine, and hydroxyl in apatite and biotite to estimate the compositions of magmatic or hydrothermal fluids thought to be associated with many intrusion-related mineral deposits have met with varying degrees of success. Actual concentrations of these species in a fluid are not directly reflected in apatite and biotite compositions due to preferential partitioning of fluorine and chlorine into these respective minerals, as well as internal and external controls on the uptake of halogens into these minerals. Relative concentrations of these volatile species in fluids, therefore, are calculated as fugacity or activity ratios, which attempt to correct for the controls on halogen partitioning. Variations in these ratios may thus be attributed to evolving fluid composition. Activity ratios (from Zhu and Sverjensky, 1991, 1992) are determined based on the partitioning behaviour of halogens between apatite and biotite compositions; this significantly reduces or removes the effects of external controls (e.g., temperature, pressure) on F–Cl–OH exchange. In this way, apatite and biotite are also used as independent monitors of aqueous, rather than gaseous, fluid composition. The calculation of activity ratios from two different magmatic minerals is particularly useful for two reasons: (1) the apatite and biotite activity ratios may be compared to check for sub-solidus re-equilibration of either mineral, particularly biotite; and (2) the biotite activity ratios provide an estimate for fluid compositions in magmatic–hydrothermal systems that crystallized end-member fluorapatite (from which activity ratios cannot be calculated). However, it is much more expedient and less error- prone to analyze and calculate activity ratios from biotite alone for fluid compositional studies, unless there is the possibility that biotite has re-equilibrated with later fluids that are not of interest to the study. 385 Figure 6.1. Distribution of Early to ‘mid’-Cretaceous, syn- to post-accretionary plutonic suites in the northern Cordillera (AK = Alaska; YK = Yukon; NWT = Northwest Territories; BC = British Columbia). The study area (Fig. 7.2) is outlined in black. Plutonic suites discussed in the text are labeled: To = Tombstone, M = Mayo, Tu = Tungsten, TR = Tay River. The Tombstone, Mayo, and Tungsten plutonic suites for the Tombstone–Tungsten Belt (TTB) referred to in the text. The largest tungsten skarn deposits or occurrences, as well as the Dawson Range batholith and the Casino deposit in western Yukon, are labelled. Tectonostratigraphic units and map lines are as in Figure 1.1 (Chapter 1). 386 Figure 6.2. Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous magmatism (grey polygons), and tungsten deposits and occurrences (circles) in NTS map sheets 105I, 105H, 095L, and 095E. Plutons studied here are labeled. The Yukon–Northwest Territories border is indicated by the stippled line, mineral occurrence locations are from the NORMIN (2013) and Yukon MINFILE (2013) databases, and intrusive polygons are modified from Gordey and Makepeace (1999). 387 Figure 6.3. (next page) Apatite chemistry. a) MnO versus CaO (wt.%). b) FeO versus CaO (wt.%). c) Chlorine versus fluorine (wt.%); chlorine and fluorine detection limits of 0.06 wt.% and 0.09 wt.%, respectively, are indicated by the dashed lines. Anomalous samples discussed in the text are circled in red. d) XF/XCl versus MnO; XF/XCl is less than ca. 140 (indicated by the dashed line) for samples with chlorine below the detection limit. 388 389 Figure 6.4. Biotite chemistry. a) Fe/(Fe+Mg), or Fe#, versus aluminum (apfu) biotite classification after Rieder et al. (1998) where biotite end-members are: Ann = annite, Sid = siderophyllite, Phl = phlogopite, and Eas = eastonite. Anomalous samples discussed in the text are circled in red. b) Chlorine versus fluorine (wt.%); detection limits are indicated by dashed lines. 390 Figure 6.5. Linear relationship between log[aHOH/aHF] values from apatite and (uncorrected) biotite datasets. Samples containing biotites that appear to have equilibrated with a distinct fluid―KR-05-117 (decoupled with respect to H2O), KR-05-173 (decoupled with respect to HF ± H2O), and 03M-203 (minor decoupling with respect to H2O)―were not used in determining the correction factor for log[aHOH/aHF]BT. These anomalous samples are labelled and indicated by red symbols and text. 391 Figure 6.6. (next 2 pages) Log(aHCl/aHF) and log(aHF/aHOH) for apatite (black) and biotite (grey). Data for primary magmatic samples are indicated by filled symbols; data for mineralized samples (e.g., greisens, endoskarn) are indicated by unfilled symbols and are plotted nearest to the y-axis (regardless of their initial degree of magmatic differentiation prior to alteration). a) Relative degree of differentiation increasing towards the x-axis. Italicized sample labels (last set of digits for the samples listed in Table 1) are given for each set of symbols. b) Proximity to mineralization increasing towards the x-axis. Sample labels are as in part (a), and anomalous samples that may be associated with mineralization are circled in red. 392 393 394 Table 6.1. Abbreviated descriptions for samples analyzed by this study. Pluton Sample Abbreviated Sample Descriptiona,b,c Relative Degree of Differentiation Distance to mineralization (m) Degree and Type of Alterationd,e,f,g Aph Bth C. Nahanni KR-05-128 vfrg. PHBQ-x quartz diorite porphyry 4 n/a fresh P P C. Nahanni KR-05-129 fgr bt>hbl granodiorite (interm) 3 n/a fresh P P C. Nahanni KR-05-130 mgr bt>>hbl granodiorite (felsic) 2 n/a fresh P P C. Nahanni KR-05-131 mgr bt granite 1 n/a fresh P P C. Nahanni KR-05-133 fgr HPQ-x sulphide-bearing quartz diorite dyke 5 n/a fresh x P Rifle Range KR-05-148 cgr K-x bt>ms granite 4 n/a fresh P P Rifle Range KR-05-149 fgr (K-x) bt granite 5 n/a fresh P P Rifle Range KR-05-150b fgr (K-x) bt>ms granite 3 n/a fresh P P Rifle Range KR-05-151 fgr (K-x) ms-bt granite 2 n/a fresh P P Rifle Range KR-05-152 vfgr QPB-x granitic porphyry 1 n/a altered P P Rifle Range KR-05-154 fgr BP-x sulphide-bearing granitic porphyry 1 n/a (v.) altered P S Rifle Range KR-05-155 fgr bt>ms granite w/min'd fract 3 n/a greisen M M Rifle Range KR-05-156 vfgr (bt) aplite 1 n/a altered P P Rudi KR-05-117 vfgr bt>>ms sulphide-bearing dyke 2 25 v. altered (calc) P P/M Rudi KR-05-121a cgr bt>>ms granite 5 225 altered P P+S Rudi KR-05-122 mgr K-x bt>>ms 3 175 altered P P Rudi KR-05-123 cgr bt>ms granite 4 5 greisen (pot) M M Rudi KR-05-124 fgr ms>trm>>grt>bt leucogranite 1 150 fresh x P Rudi KR-05-125b ms>>trm>bt leucogranite 1 0 greisen (mica) M M Lened KR-05-170 mgr bt>>ms granitic apophysy 3 1350 n/a P P Lened KR-05-173 fgr ms>>bt leucogranite 1 1200 n/a P P Lened KR-05-175 cgr K-x bt>ms granite (same sample as KR-05- 177) 4 1800 n/a P P Lened KR-05-176 mgr ms>>bt leucogranite 1 1900 n/a P P Lened KR-05-178 f-mgr bt>ms leucogranite 2 2400 n/a P P Cantung 03M-203 fgr late ms>bt aplite 1 5-25 fresh P P Cantung 03M-204 vfgr late aplite 1 5-25 calc-ser M M Cantung 03M-207 f-mgr bt>ms apical granite 5 25-100 fresh P P Cantung 03M-212 vfgr early sulphide-bearing porphyry 3 5-25 (calc-ser) M M Cantung KR-05-215 mgr (K-x) bt>>ms granite 5 >100 fresh P P Cantung KR-03_09-5 f-mgr bt>ms apical granite 4 5-25 fresh + calc-ser margin P+M P Cantung KR-03_31-7 f-mgr early sulphide-bearing dyke 4 0-5 calc-ser + min'd M M Cantung KR-03_01-1 vfgr late sulphide-bearing aplite 1 0-5 calc-ser + min'd M x Cantung KR-03_07-18 fgr bt granitic dyke 3 5-25 fresh + calc-ser margin ? P+M Cantung KR-03_07-22 vfgr late aplite 1 5-25 calc-ser M M Cantung KR-05-205 vcgr late trm>ms pegmatite 2 25-100 fresh P x Notes: a vfgr, fgr, mgr, cgr, vcgr = very fine, fine, medium, coarse, and very coarse-grained; b K-, P-, Q-, B-, H-x = K-feldspar, plagioclase, quartz, biotite, hornblende-phyric; c mineral abbreviations are after Kretz (1983); d fresh = minor sericite(locally coarse)/saussurite/clay in feldspar; minor chlorite in biotite/muscovite; e altered = sericitized/saussuritized/clay-speckled feldspar, chlorite ± fine-grained secondary muscovite; f greisen = coarse-grained secondary mica; pot = unaltered K-feldspar and plagioclase intensely sericitized; mica = veins and clots of secondary biotite/muscovite; g v. altered = intensely sericitized/saussuritized/clay-speckled feldspar, chlorite or fine-grained secondary biotite/muscovite; calc = secondary plagioclase ± apatite and epidote–group minerals; ser = late calcite–sericite; h halogen-bearing mineral that was analyzed: Ap = apatite, Bt = biotite; P = primary, S = secondary, M = mineralizing, x = not present. 395 Table 6.2. Representative apatite compositions and calculated magmatic fluid activity ratios. Pluton: Central Nahanni Rifle Range d et ec tio n lim it (o xi de w t.% ) K R -0 5- 12 8 K R -0 5- 12 9 K R -0 5- 13 0 K R -0 5- 13 1 K R -0 5- 14 8 K R -0 5- 14 9 K R -0 5- 15 0b K R -0 5- 15 1 K R -0 5- 15 2 K R -0 5- 15 4 K R -0 5- 15 5 K R -0 5- 15 6 na 8 7 15 8 11 9 3 3 5 7 2 1 Stageb P P P P P P P P P P M M P2O5c 0.113 41.98 41.35 41.77 41.68 41.61 41.45 41.03 40.73 41.39 41.15 41.37 41.69 ± 2-σ 0.92 0.91 0.84 0.57 0.81 0.97 0.47 1.01 0.79 0.66 1.18 n/a CaO 0.067 55.09 54.19 54.90 54.51 53.91 53.82 53.15 53.42 53.69 55.06 53.00 54.00 ± 2-σ 0.16 0.57 0.45 0.41 0.41 0.34 0.38 1.30 0.38 0.56 3.20 n/a SrO 0.116 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 ± 2-σ 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.00 0.00 0.00 0.00 n/a Al2O3 0.032 0.01 0.11 0.03 0.04 0.05 0.01 0.01 0.02 0.01 0.01 0.00 0.05 ± 2-σ 0.01 0.19 0.03 0.06 0.06 0.01 0.01 0.03 0.01 0.02 0.00 n/a FeOT 0.122 0.35 0.20 0.14 0.20 0.37 0.19 0.64 0.65 0.38 0.16 0.56 0.61 ± 2-σ 0.18 0.24 0.15 0.13 0.13 0.13 0.16 0.38 0.16 0.09 0.12 n/a MnO 0.118 0.08 0.03 0.11 0.10 0.83 0.57 0.50 0.30 0.52 0.15 0.33 0.79 ± 2-σ 0.06 0.04 0.06 0.07 0.19 0.17 0.02 0.12 0.08 0.06 0.13 n/a MgO 0.035 0.03 0.04 0.06 0.02 0.05 0.03 0.03 0.03 0.02 0.00 0.02 0.06 ± 2-σ 0.03 0.08 0.09 0.02 0.05 0.03 0.03 0.04 0.02 0.01 0.00 n/a Cl 0.060 0.78 0.58 0.67 0.12 0.08 0.12 0.04 0.03 0.07 0.12 0.03 0.08 ± 2-σ 0.17 0.11 0.21 0.06 0.03 0.07 0.02 0.01 0.01 0.04 0.01 n/a F 0.091 2.56 2.70 2.65 3.46 3.74 3.72 4.04 3.94 4.08 3.50 4.48 4.53 ± 2-σ 0.21 0.06 0.11 0.07 0.11 0.21 0.10 0.09 0.20 0.18 0.06 n/a S 0.046 0.03 0.03 0.02 0.04 0.02 0.03 0.03 0.04 0.03 0.04 0.02 0.02 Total 100.02 98.29 99.41 98.78 99.07 98.37 97.76 97.49 98.45 98.77 97.89 99.90 Se 0.002 0.002 0.001 0.003 0.001 0.002 0.002 0.002 0.002 0.003 0.001 0.001 Fe 0.025 0.014 0.010 0.015 0.026 0.014 0.046 0.046 0.027 0.011 0.040 0.043 Mn 0.006 0.002 0.008 0.007 0.059 0.041 0.036 0.022 0.038 0.011 0.024 0.056 Mg 0.004 0.006 0.007 0.002 0.007 0.004 0.004 0.003 0.002 0.001 0.003 0.008 Al 0.001 0.011 0.003 0.004 0.005 0.001 0.001 0.002 0.001 0.001 0.000 0.005 Ca 4.974 4.968 4.982 4.965 4.907 4.926 4.893 4.938 4.903 5.034 4.845 4.859 P 2.994 2.995 2.995 3.000 2.992 2.997 2.985 2.975 2.986 2.972 2.990 2.964 Sr 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 XF 0.683 0.730 0.710 0.929 1.004 1.004 1.097 1.075 1.099 0.945 1.211 1.203 XCl 0.111 0.084 0.097 0.017 0.012 0.017 0.006 0.005 0.010 0.017 0.004 0.011 XF/XCl 6.2 8.7 7.3 53.5 81.8 59.1 173.2 230.4 105.7 55.2 345.9 109.4 XOHf 0.207 0.186 0.193 0.054 -0.016 -0.021 -0.103 -0.080 -0.109 0.037 -0.214 -0.214 nHCl/HFg 8 7 15 8 11 9 3 3 5 7 1 1 log aHCl/aHFh,i 0.75 0.64 0.67 –0.21 –0.38 –0.30 –0.70 –0.81 –0.47 –0.21 –0.94 –0.48 ± 2-σ 0.10 0.11 0.15 0.20 0.19 0.33 0.19 0.10 0.06 0.18 n/a n/a nHF/HOHg 8 7 15 8 3 3 0 0 0 0 5 0 log aHOH/aHFi 2.16 2.13 2.13 1.42 0.87 1.21 n/a n/a n/a 1.62 n/a n/a ± 2-σ 0.18 0.08 0.10 0.22 0.24 0.13 n/a n/a n/a 0.32 n/a n/a Notes: a number of analyses averaged; b P = primary, M = related to mineralization; c in oxide wt.%; e atomic proportions; normalized to 13 anhydrous O; f determined from stoichiometry; g number of analyses accepted for calculations; h loga(HCl) = –ΔG0R,2 / (2.302585 * RT) – log a(Hap) + log a(ClAp), where ΔG0R,2 = –5860 (850°C, 2 kbar) for the reaction: HCl + HAp = H2O + ClAp; i log a(HF) = –ΔG0R,1 / (2.302585 * RT) – log a(Hap) + log a(FAp), where ΔG0R,1 = – 13874 (850°C, 2 kbar) for the reaction: HF + HAp = H2O + Fap; 396 Pluton: Rudi Lened K R -0 5- 11 7 K R -0 5- 12 1 K R -0 5- 12 2 K R -0 5- 12 3 K R -0 5- 12 5b K R -0 5- 17 0 K R -0 5- 17 3 K R -0 5- 17 5 K R -0 5- 17 6 K R -0 5- 17 7 K R -0 5- 17 8 na 4 3 3 2 3 8 3 8 14 9 6 Stageb P P P P P P P P P P P P2O5c 42.31 41.02 41.34 42.14 41.25 42.28 41.70 42.00 41.91 41.63 41.48 ± 2-σ 0.12 0.17 0.16 0.72 0.60 0.44 0.74 0.44 0.66 0.71 0.72 CaO 55.05 54.02 54.61 54.45 54.93 54.61 54.92 53.34 53.66 53.51 53.03 ± 2-σ 0.62 0.68 0.86 1.15 1.12 0.50 0.27 0.48 0.60 0.34 0.64 SrO 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00 0.00 0.00 ± 2-σ 0.00 0.00 0.00 0.00 0.00 0.03 0.00 0.01 0.00 0.00 0.00 Al2O3 0.03 0.02 0.04 0.01 0.22 0.02 0.01 0.01 0.01 0.02 0.02 ± 2-σ 0.02 0.02 0.06 0.01 0.38 0.02 0.01 0.01 0.01 0.03 0.01 FeOT 0.25 0.16 0.32 0.45 0.08 0.18 0.16 0.25 0.37 0.52 0.64 ± 2-σ 0.07 0.15 0.10 0.12 0.07 0.15 0.17 0.08 0.20 0.15 0.33 MnO 0.03 0.36 0.41 0.19 0.09 0.45 0.42 0.68 1.14 1.02 1.07 ± 2-σ 0.02 0.08 0.12 0.00 0.02 0.12 0.07 0.08 0.35 0.14 0.18 MgO 0.03 0.02 0.02 0.03 0.02 0.03 0.02 0.05 0.03 0.06 0.05 ± 2-σ 0.03 0.02 0.02 0.02 0.04 0.04 0.02 0.03 0.02 0.04 0.02 Cl 0.61 0.02 0.02 0.03 0.02 0.09 0.08 0.04 0.03 0.05 0.03 ± 2-σ 0.10 0.00 0.00 0.01 0.01 0.03 0.04 0.01 0.01 0.02 0.01 F 2.36 4.02 3.97 3.95 3.92 3.77 1.16 3.68 3.60 3.71 3.95 ± 2-σ 0.12 0.13 0.09 0.10 0.09 0.07 0.11 0.07 0.17 0.10 0.15 S 0.03 0.02 0.03 0.06 0.05 0.02 0.01 0.01 0.02 0.02 0.06 Total 100.09 97.97 99.08 99.62 98.93 99.85 99.17 98.52 99.31 98.98 98.66 Se 0.002 0.002 0.002 0.004 0.004 0.001 0.001 0.001 0.001 0.001 0.004 Fe 0.018 0.011 0.023 0.032 0.006 0.012 0.011 0.018 0.026 0.037 0.046 Mn 0.002 0.026 0.029 0.014 0.007 0.032 0.030 0.049 0.082 0.074 0.078 Mg 0.004 0.003 0.003 0.003 0.003 0.004 0.003 0.006 0.003 0.008 0.007 Al 0.003 0.002 0.004 0.001 0.022 0.002 0.001 0.001 0.001 0.002 0.002 Ca 4.951 4.961 4.969 4.908 4.998 4.917 4.982 4.861 4.869 4.877 4.841 P 3.006 2.977 2.972 3.002 2.966 3.008 2.988 3.024 3.004 2.997 2.992 Sr 0.000 0.000 0.000 0.000 0.000 0.001 0.000 0.000 0.000 0.000 0.000 XF 0.625 1.091 1.065 1.051 1.054 1.001 0.310 0.989 0.964 0.997 1.065 XCl 0.087 0.003 0.003 0.004 0.003 0.013 0.011 0.006 0.005 0.008 0.004 XF/XCl 7.2 363.6 355.0 300.1 316.1 76.2 27.4 171.9 198.9 132.9 304.3 XOHf 0.288 -0.094 -0.068 -0.054 -0.057 -0.014 0.678 0.006 0.031 -0.004 -0.069 nHCl/HFg 4 0 0 1 1 8 3 8 13 9 6 log aHCl/aHFh,i 0.70 n/a n/a –0.86 –0.85 –0.34 0.09 –0.75 –0.77 –0.60 –0.93 ± 2-σ 0.09 n/a n/a n/a n/a 0.13 0.20 0.18 0.16 0.19 0.08 nHF/HOHg 2 0 0 0 0 2 3 8 9 3 0 log aHOH/aHFi 1.66 n/a n/a n/a n/a 1.34 2.95 1.61 2.18 1.57 n/a ± 2-σ 0.07 n/a n/a n/a n/a 0.68 0.25 0.34 0.25 0.32 n/a 397 Pluton: Cantung 03 M -2 03 03 M -2 04 03 M -2 07 03 M -2 12 K R -0 5- 21 5 K R -0 3_ 09 -5 K R -0 3_ 31 -7 K R -0 3_ 01 -1 K R -0 3_ 07 -1 8j K R -0 5- 20 5 na 4 5 7 7 6 2 2 1 1 3 Stageb P M P M P P M M M P P2O5c 41.08 41.94 41.54 40.70 42.00 41.99 42.00 41.04 41.87 41.88 ± 2-σ 0.56 1.40 0.56 0.75 0.83 0.46 1.27 n/a n/a 0.42 CaO 54.65 53.97 54.38 54.60 54.35 55.44 55.20 55.12 55.77 54.58 ± 2-σ 0.20 0.50 0.50 0.35 0.37 0.17 0.54 n/a n/a 0.33 SrO 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.02 ± 2-σ 0.00 0.00 0.01 0.02 0.00 0.00 0.00 n/a n/a 0.03 Al2O3 0.02 0.03 0.06 0.00 0.00 0.00 0.02 0.01 0.00 0.03 ± 2-σ 0.02 0.06 0.06 0.00 0.00 0.00 0.03 n/a n/a 0.04 FeOT 0.21 0.10 0.29 0.15 0.45 0.08 0.08 0.00 0.00 0.08 ± 2-σ 0.04 0.04 0.07 0.11 0.17 0.10 0.10 n/a n/a 0.05 MnO 0.06 0.30 0.67 0.29 0.68 0.06 0.14 0.08 0.11 0.21 ± 2-σ 0.04 0.16 0.07 0.09 0.18 0.08 0.06 n/a n/a 0.13 MgO 0.21 0.02 0.13 0.01 0.04 0.01 0.01 0.00 0.01 0.02 ± 2-σ 0.09 0.02 0.09 0.02 0.02 0.01 0.01 n/a n/a 0.02 Cl 0.26 0.02 0.08 0.09 0.08 0.02 0.02 0.02 0.03 0.02 ± 2-σ 0.05 0.01 0.02 0.05 0.02 0.00 0.00 n/a n/a 0.00 F 3.52 4.08 3.60 3.72 4.07 4.09 4.18 3.98 4.04 3.99 ± 2-σ 0.02 0.12 0.06 0.05 0.16 0.24 0.03 n/a n/a 0.18 S 0.03 0.03 0.02 0.01 0.00 0.00 0.00 0.00 0.03 0.01 Total 98.52 98.77 99.28 98.02 99.95 99.96 99.89 98.59 100.15 99.15 Se 0.002 0.002 0.002 0.003 0.001 0.001 0.001 0.001 0.002 0.001 Fe 0.015 0.007 0.020 0.011 0.032 0.006 0.006 0.000 0.000 0.006 Mn 0.004 0.022 0.048 0.021 0.049 0.005 0.010 0.006 0.008 0.015 Mg 0.027 0.002 0.016 0.001 0.005 0.001 0.001 0.000 0.001 0.002 Al 0.002 0.003 0.006 0.000 0.000 0.000 0.002 0.001 0.000 0.003 Ca 5.009 4.897 4.943 5.034 4.895 4.985 4.963 5.038 5.012 4.942 P 2.975 3.005 2.983 2.964 2.988 2.983 2.983 2.964 2.973 2.996 Sr 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.001 XF 0.952 1.093 0.965 1.012 1.081 1.086 1.110 1.074 1.072 1.066 XCl 0.038 0.003 0.011 0.013 0.011 0.003 0.003 0.003 0.004 0.003 XF/XCl 25.2 321.5 85.5 75.4 98.3 362.0 369.8 358.0 268.0 355.2 XOHf 0.010 -0.096 0.024 -0.026 -0.092 -0.089 -0.113 -0.077 -0.076 -0.069 nHCl/HFg 4 5 7 7 6 0 0 0 0 0 log aHCl/aHFh,i 0.15 –0.95 –0.38 –0.38 –0.45 –0.87 n/a n/a n/a n/a ± 2-σ 0.08 0.09 0.11 0.27 0.14 n/a n/a n/a n/a n/a nHF/HOHg 5 0 7 0 0 0 0 0 0 0 log aHOH/aHFi 1.06 n/a 1.05 n/a n/a n/a n/a n/a n/a n/a ± 2-σ 0.26 n/a 0.23 n/a n/a n/a n/a n/a n/a n/a Notes (cont.): j altered margin of intrusive phase. 398 Table 6.3. Representative biotite compositions and calculated Fe#, ∑ Al, and magmatic fluid activity ratios. Pluton: Central Nahanni Rifle Range Rudi de te ct io n lim it (w t.% o xi de ) K R -0 5- 12 8 K R -0 5- 12 9 K R -0 5- 13 0 K R -0 5- 13 1 K R -0 5- 13 3 K R -0 5- 14 8 K R -0 5- 14 9 K R -0 5- 15 0b K R -0 5- 15 1 K R -0 5- 15 2 K R -0 5- 15 4 K R -0 5- 15 5 K R -0 5- 15 6 K R -0 5- 11 7 K R -0 5- 12 1 na 7 5 3 4 2 2 5 5 5 5 5 5 1 4 5 Stageb P P P P P P P P P P P P P P P SiO2c 0.017 36.19 36.32 35.94 36.54 38.51 34.18 35.17 34.43 34.30 34.40 35.78 34.61 34.79 38.35 35.33 SD SiO2d 0.37 0.15 0.31 0.22 0.21 0.42 0.08 0.21 0.26 0.25 0.28 0.24 n/a 0.28 0.09 TiO2 0.021 3.82 3.68 3.65 3.29 1.67 3.03 3.35 3.79 4.03 3.07 3.63 3.97 2.64 2.28 3.66 SD TiO2 0.46 0.16 0.15 0.36 0.65 0.63 0.15 0.15 0.10 0.03 0.08 0.21 n/a 0.12 0.05 Al2O3 0.013 13.82 13.65 14.06 14.36 14.99 16.88 16.11 18.65 18.24 18.39 14.49 18.39 16.09 16.42 16.24 SD Al2O3 0.13 0.10 0.15 0.09 0.26 1.08 0.33 0.19 0.16 0.13 0.03 0.32 n/a 0.52 0.28 Cr2O3 0.034 0.04 0.02 0.04 0.02 0.12 0.03 0.01 0.01 0.01 0.01 0.03 0.03 0.00 0.05 0.01 SD Cr2O3 0.03 0.02 0.03 0.01 0.04 0.02 0.00 0.01 0.02 0.01 0.03 0.02 n/a 0.05 0.01 FeOT 0.041 19.51 19.88 20.17 20.55 14.51 24.60 22.82 23.09 22.47 22.05 19.88 21.96 24.40 9.18 21.72 SD FeOT 0.68 0.50 0.55 0.30 0.42 0.13 0.27 0.18 0.39 0.25 0.32 0.60 n/a 0.19 0.44 MnO 0.036 0.17 0.26 0.23 0.32 0.11 0.37 0.54 0.26 0.27 0.44 0.43 0.22 0.44 0.10 0.46 SD MnO 0.01 0.02 0.04 0.02 0.00 0.08 0.03 0.03 0.07 0.01 0.03 0.02 n/a 0.02 0.05 MgO 0.012 11.84 11.49 11.44 10.65 14.98 6.72 7.93 5.75 5.33 6.98 10.39 5.50 7.00 17.39 8.05 SD MgO 0.47 0.19 0.25 0.31 0.63 0.03 0.17 0.09 0.17 0.12 0.19 0.16 n/a 0.35 0.14 CaO 0.018 0.01 0.01 0.03 0.02 0.03 0.00 0.01 0.02 0.02 0.01 0.02 0.00 0.00 0.04 0.02 SD CaO 0.02 0.01 0.03 0.02 0.02 0.00 0.01 0.01 0.02 0.01 0.01 0.00 n/a 0.02 0.01 Na2O 0.017 0.09 0.06 0.06 0.06 0.05 0.07 0.13 0.15 0.11 0.13 0.10 0.12 0.06 0.08 0.07 SD Na2O 0.02 0.01 0.02 0.01 0.02 0.04 0.02 0.02 0.02 0.02 0.01 0.02 n/a 0.01 0.02 K2O 0.021 9.62 9.49 9.42 9.42 9.57 9.01 9.57 9.54 9.57 9.73 9.45 9.62 9.55 9.89 9.70 SD K2O 0.11 0.07 0.19 0.08 0.16 0.46 0.05 0.09 0.07 0.08 0.12 0.09 n/a 0.12 0.02 Cl 0.014 0.20 0.22 0.19 0.13 0.24 0.15 0.15 0.14 0.15 0.14 0.18 0.12 0.20 0.08 0.06 SD Cl 0.02 0.06 0.02 0.02 0.06 0.05 0.01 0.07 0.02 0.01 0.01 0.02 n/a 0.02 0.00 F 0.078 0.29 0.30 0.32 0.41 0.21 0.89 0.48 0.64 0.48 0.41 0.52 0.61 0.85 0.17 0.74 SD F 0.02 0.03 0.03 0.01 0.02 0.19 0.05 0.04 0.06 0.08 0.03 0.13 n/a 0.03 0.07 Total 99.11 98.87 99.03 99.23 98.66 98.83 99.60 99.66 98.27 99.17 98.16 98.35 98.94 97.90 99.20 Ke 0.948 0.939 0.932 0.930 0.917 0.917 0.956 0.951 0.964 0.970 0.945 0.966 0.972 0.931 0.969 Na 0.013 0.009 0.010 0.010 0.007 0.011 0.020 0.022 0.017 0.020 0.015 0.018 0.009 0.011 0.011 Fe 1.261 1.290 1.308 1.330 0.912 1.640 1.494 1.509 1.485 1.441 1.303 1.446 1.628 0.567 1.422 Mg 1.364 1.329 1.322 1.229 1.679 0.798 0.925 0.669 0.628 0.813 1.214 0.645 0.833 1.914 0.940 Mn 0.011 0.017 0.015 0.021 0.007 0.025 0.036 0.017 0.018 0.029 0.029 0.015 0.030 0.006 0.030 Ti 0.222 0.215 0.213 0.192 0.095 0.182 0.197 0.223 0.239 0.180 0.214 0.235 0.159 0.127 0.216 Ca 0.001 0.001 0.002 0.001 0.002 0.000 0.001 0.001 0.002 0.001 0.002 0.000 0.000 0.003 0.002 VIAl 0.056 0.067 0.073 0.139 0.221 0.311 0.237 0.408 0.409 0.383 0.144 0.432 0.289 0.259 0.264 IVAl 1.203 1.182 1.212 1.172 1.106 1.275 1.248 1.309 1.290 1.312 1.195 1.274 1.224 1.169 1.234 Si 2.797 2.818 2.788 2.828 2.894 2.725 2.752 2.691 2.710 2.688 2.805 2.726 2.776 2.831 2.766 XF 0.071 0.074 0.078 0.102 0.050 0.225 0.120 0.157 0.120 0.101 0.128 0.151 0.214 0.039 0.182 XCl 0.026 0.028 0.025 0.017 0.030 0.021 0.020 0.019 0.019 0.019 0.024 0.015 0.028 0.010 0.008 XOHf 0.903 0.898 0.897 0.882 0.920 0.755 0.860 0.824 0.860 0.881 0.848 0.833 0.758 0.951 0.810 Fe# = Fe/(Fe+Mg) 0.48 0.49 0.50 0.52 0.35 0.67 0.62 0.69 0.70 0.64 0.52 0.69 0.66 0.23 0.60 ∑ Al = 2*(VIAl+IVAl) 2.52 2.50 2.57 2.62 2.65 3.17 2.97 3.43 3.40 3.39 2.68 3.41 3.03 2.86 3.00 log aHCl/aHFg,h,i 0.52 0.52 0.44 0.09 0.86 –0.40 –0.05 –0.32 –0.21 –0.06 0.16 –0.39 –0.23 0.63 –0.60 SD log aHCl/aHF 0.05 0.11 0.07 0.06 0.14 0.06 0.03 0.04 0.02 0.08 0.04 0.09 n/a 0.16 0.08 log aHOH/aHFi 2.26 2.26 2.23 2.08 2.54 1.62 1.90 1.69 1.80 1.95 1.97 1.71 1.58 2.76 1.72 log aHOH/aHF(corr.)j 2.43 2.35 2.27 1.85 3.09 0.40 1.36 0.77 1.09 1.48 1.56 0.84 0.48 3.69 0.88 SD log aHOH/aHF 0.06 0.09 0.10 0.03 0.08 0.27 0.12 0.09 0.15 0.25 0.08 0.28 n/a 0.25 0.13 Notes: a number of analyses averaged; b P = primary, M = related to mineralization; c in oxide wt.%; d SD = standard deviation; e atomic proportions; normalized to 11 anhydrous O (excluding H2O, Cl, and F); f determined from stoichiometry; g log a(HCl) = –( 0.5(ΔG0R,c / 2.303RT) * XMg + 0.5logKa – log (XCl/XOH)bt), where ΔG0R,c = –4862 (850°C, 2 kbar) for the reaction: Chphl + Ann = Phl + Clann; h log a(HF) = –( 0.5(ΔG0R,b / 2.303RT) * XMg + 0.5logKa – log (XF/XOH)bt), where ΔG0R,b = –10099 (850°C, 2 kbar) for the reaction: Phl + Fann = Fphl + Ann; i log aHOH/aHF assuming pure H2O fluid where aHOH = 1 (Zhu and Sverjensky, 1992); j log aHOH/aHF corrected for assumption of a pure fluid by reasoning that log aHOH/aHF(AP) = log aHOH/aHF(BT) and using the linear relationship (R2 = 0.71): log aHOH/aHF(corr.) = (log aHOH/aHF(BT)+1.4023) / 0.3667; 399 Pluton: Rudi Lened K R -0 5- 12 1 K R -0 5- 12 2 K R -0 5- 12 3 K R -0 5- 12 4 K R -0 5- 12 5b K R -0 5- 17 0 K R -0 5- 17 3 K R -0 5- 17 5 K R -0 5- 17 6 K R -0 5- 17 7 K R -0 5- 17 8 na 2 7 5 5 3 6 2 1 5 6 7 Stageb S P S P S P P P P P P SiO2c 35.02 34.96 35.22 35.01 36.30 35.15 32.20 35.02 34.27 34.72 34.57 SD SiO2d 0.12 0.44 0.29 0.53 0.83 0.22 0.66 n/a 0.21 0.22 0.43 TiO2 3.79 3.30 3.39 2.63 3.63 3.44 2.91 3.30 2.98 3.38 2.90 SD TiO2 0.07 0.25 0.14 0.29 0.09 0.11 0.82 n/a 0.15 0.26 0.47 Al2O3 16.42 17.92 16.85 19.77 15.05 18.42 19.46 17.26 19.03 18.74 19.37 SD Al2O3 0.05 0.42 0.26 0.47 0.25 0.17 0.40 n/a 0.21 0.21 0.27 Cr2O3 0.01 0.01 0.01 0.01 0.00 0.02 0.01 0.03 0.01 0.01 0.01 SD Cr2O3 0.02 0.01 0.01 0.01 0.01 0.01 0.02 n/a 0.02 0.01 0.01 FeOT 21.93 22.28 22.01 24.05 22.55 20.35 26.60 21.88 22.90 21.79 22.58 SD FeOT 0.04 0.57 0.17 0.65 0.10 0.26 0.50 n/a 0.30 0.42 0.31 MnO 0.41 0.41 0.40 1.38 0.30 0.34 1.34 0.50 0.74 0.51 0.58 SD MnO 0.00 0.02 0.02 0.05 0.02 0.01 0.06 n/a 0.10 0.04 0.11 MgO 7.47 6.59 7.64 2.42 8.02 7.28 3.53 7.18 4.63 5.78 4.85 SD MgO 0.06 0.35 0.07 0.16 0.38 0.17 0.21 n/a 0.09 0.07 0.32 CaO 0.02 0.01 0.00 0.01 0.05 0.02 0.02 0.03 0.01 0.01 0.00 SD CaO 0.01 0.01 0.00 0.01 0.00 0.01 0.02 n/a 0.01 0.01 0.00 Na2O 0.05 0.08 0.07 0.07 0.07 0.13 0.02 0.10 0.06 0.08 0.08 SD Na2O 0.00 0.03 0.03 0.02 0.01 0.02 0.00 n/a 0.02 0.01 0.02 K2O 9.49 9.73 9.71 9.68 9.16 9.40 7.60 9.63 9.38 9.62 9.46 SD K2O 0.02 0.13 0.08 0.08 0.06 0.12 0.25 n/a 0.11 0.09 0.07 Cl 0.07 0.06 0.08 0.02 0.11 0.10 0.07 0.05 0.05 0.04 0.03 SD Cl 0.00 0.01 0.02 0.01 0.01 0.01 0.00 n/a 0.01 0.00 0.01 F 0.65 0.54 0.69 0.57 0.84 0.38 0.00 0.52 0.26 0.46 0.66 SD F 0.03 0.05 0.09 0.05 0.02 0.04 0.00 n/a 0.03 0.04 0.08 Total 98.52 99.22 99.25 98.90 99.11 98.51 97.45 98.84 97.85 98.53 98.29 Ke 0.953 0.970 0.969 0.976 0.916 0.931 0.781 0.962 0.948 0.961 0.953 Na 0.008 0.012 0.011 0.011 0.011 0.020 0.003 0.016 0.009 0.013 0.012 Fe 1.445 1.456 1.440 1.590 1.479 1.321 1.790 1.434 1.516 1.428 1.490 Mg 0.877 0.768 0.891 0.285 0.937 0.843 0.423 0.838 0.546 0.675 0.571 Mn 0.028 0.027 0.026 0.093 0.020 0.022 0.091 0.033 0.049 0.034 0.039 Ti 0.225 0.194 0.200 0.156 0.214 0.201 0.176 0.194 0.178 0.199 0.172 Ca 0.002 0.001 0.000 0.001 0.004 0.001 0.001 0.003 0.001 0.001 0.000 VIAl 0.283 0.382 0.309 0.608 0.237 0.415 0.438 0.338 0.488 0.451 0.529 IVAl 1.242 1.268 1.244 1.233 1.154 1.271 1.408 1.256 1.287 1.280 1.272 Si 2.758 2.732 2.756 2.767 2.846 2.729 2.592 2.744 2.713 2.720 2.728 XF 0.161 0.134 0.169 0.141 0.208 0.094 0.000 0.128 0.065 0.114 0.166 XCl 0.009 0.007 0.010 0.003 0.015 0.013 0.010 0.006 0.006 0.005 0.004 XOHf 0.830 0.859 0.820 0.856 0.777 0.893 0.990 0.865 0.929 0.881 0.830 Fe# = Fe/(Fe+Mg) 0.62 0.65 0.62 0.85 0.61 0.61 0.81 0.63 0.74 0.68 0.72 ∑ Al = 2*(VIAl+IVAl) 3.05 3.30 3.11 3.68 2.78 3.37 3.69 3.19 3.55 3.46 3.60 log aHCl/aHFg,h,i –0.62 –0.61 –0.52 –1.31 –0.41 –0.16 2.46 –0.62 –0.47 –0.76 –1.11 SD log aHCl/aHF 0.01 0.11 0.15 0.09 0.03 0.06 0.02 n/a 0.04 0.05 0.14 log aHOH/aHFi 1.73 1.80 1.73 1.60 1.65 1.99 4.82 1.84 2.05 1.84 1.63 log aHOH/aHF(corr.)j 0.89 1.09 0.90 0.53 0.66 1.61 9.32 1.21 1.78 1.20 0.61 SD log aHOH/aHF 0.04 0.10 0.15 0.12 0.04 0.13 0.04 n/a 0.15 0.11 0.12 400 Pluton: Cantung 03 M -2 03 03 M -2 07 03 M -2 12 K R -0 5- 21 5 K R -0 3_ 09 -5 K R -0 3_ 31 -7 K R -0 3_ 07 -1 8 K R -0 3_ 07 -1 8k K R -0 3_ 07 -2 2 na 5 4 3 5 5 3 5 3 3 Stageb P P M P P/M? M P M M SiO2c 34.14 34.72 37.10 34.59 37.13 36.08 35.97 35.89 35.37 SD SiO2d 0.91 0.29 0.08 0.36 0.62 0.14 0.35 0.18 0.27 TiO2 3.02 3.16 2.55 2.91 1.77 2.43 3.33 2.85 2.80 SD TiO2 0.15 0.32 0.14 0.16 0.78 0.21 0.16 0.14 0.12 Al2O3 18.71 18.43 16.97 18.70 15.65 16.77 15.72 15.91 17.08 SD Al2O3 0.35 0.28 0.13 0.13 1.00 0.34 0.14 0.15 0.12 Cr2O3 0.00 0.02 0.00 0.01 0.01 0.00 0.02 0.01 0.00 SD Cr2O3 0.00 0.01 0.00 0.01 0.01 0.00 0.01 0.02 0.00 FeOT 25.67 22.79 14.33 22.73 17.84 16.79 21.86 21.81 22.48 SD FeOT 2.04 0.33 0.54 0.20 0.45 0.46 0.29 0.19 0.17 MnO 0.74 0.45 0.31 0.44 0.64 0.91 0.48 0.45 0.65 SD MnO 0.27 0.03 0.01 0.02 0.04 0.08 0.05 0.02 0.03 MgO 4.00 6.33 12.76 5.99 12.22 11.75 8.51 8.65 7.20 SD MgO 1.14 0.14 0.09 0.08 0.43 0.05 0.24 0.18 0.06 CaO 0.03 0.02 0.04 0.01 0.02 0.05 0.01 0.00 0.05 SD CaO 0.03 0.02 0.02 0.01 0.02 0.03 0.01 0.01 0.01 Na2O 0.10 0.11 0.06 0.09 0.11 0.20 0.06 0.06 0.07 SD Na2O 0.02 0.02 0.01 0.01 0.03 0.00 0.03 0.01 0.01 K2O 9.65 9.53 9.83 9.53 9.59 9.45 9.60 9.67 9.32 SD K2O 0.19 0.07 0.10 0.10 0.15 0.13 0.10 0.11 0.06 Cl 0.12 0.09 0.07 0.11 0.08 0.09 0.08 0.08 0.10 SD Cl 0.03 0.01 0.01 0.01 0.02 0.01 0.02 0.00 0.03 F 0.24 0.49 1.11 0.60 1.57 1.17 0.98 0.89 0.66 SD F 0.11 0.04 0.09 0.04 0.16 0.04 0.05 0.02 0.05 Total 99.95 99.50 98.03 98.94 99.07 98.48 99.54 99.27 98.97 Ke 0.968 0.948 0.959 0.955 0.947 0.933 0.957 0.966 0.933 Na 0.015 0.017 0.009 0.014 0.017 0.030 0.009 0.008 0.011 Fe 1.691 1.486 0.916 1.493 1.156 1.087 1.429 1.428 1.476 Mg 0.469 0.736 1.455 0.701 1.411 1.356 0.992 1.010 0.843 Mn 0.049 0.030 0.020 0.029 0.042 0.059 0.032 0.030 0.043 Ti 0.179 0.185 0.147 0.172 0.103 0.142 0.196 0.168 0.165 Ca 0.003 0.002 0.003 0.001 0.002 0.004 0.001 0.000 0.004 VIAl 0.422 0.403 0.366 0.447 0.306 0.322 0.260 0.279 0.355 IVAl 1.313 1.292 1.163 1.284 1.124 1.208 1.188 1.189 1.224 Si 2.687 2.708 2.837 2.716 2.876 2.792 2.812 2.811 2.776 XF 0.061 0.121 0.267 0.148 0.384 0.286 0.242 0.221 0.164 XCl 0.016 0.013 0.010 0.014 0.010 0.012 0.011 0.010 0.013 XOHf 0.924 0.867 0.723 0.837 0.605 0.702 0.747 0.768 0.822 Fe# = Fe/(Fe+Mg) 0.78 0.67 0.39 0.68 0.45 0.44 0.59 0.59 0.64 ∑ Al = 2*(VIAl+IVAl) 3.47 3.39 3.06 3.46 2.86 3.06 2.90 2.94 3.16 log aHCl/aHFg,h,i –0.08 –0.36 –0.45 –0.41 –0.64 –0.46 –0.60 –0.58 –0.42 SD log aHCl/aHF 0.06 0.08 0.06 0.06 0.12 0.06 0.10 0.02 0.09 log aHOH/aHFi 2.19 1.83 1.70 1.72 1.47 1.62 1.60 1.72 1.52 log aHOH/aHF(corr.)j 1.84 1.17 0.81 0.87 0.19 0.60 0.51 0.63 0.87 SD log aHOH/aHF 0.39 0.12 0.10 0.09 0.11 0.05 0.07 0.01 0.10 (cont.) k altered margin of intrusive phase. 401 Table 6.4. Apatite saturation temperatures (ASTs) for samples with available whole rock chemistry. ASTs for metaluminous samples are after Harrison and Watson (1984), and for peraluminous samples are after Bea et al. (1992). Pluton Sample AST (°C) C. Nahanni KR-05-130 870 Rifle Range KR-05-148 819 Lened KR-05-175 860 Cantung 03M-203 786 Cantunga 03M-204 874 Cantung 03M-207 803 Cantunga 03M-212 801 Cantung KR-05-215 836 Notes: a intensely altered. 402 7 CHAPTER: PETROGENETIC, THERMAL, AND FLUID EVOLUTION OF MAGMATISM IN THE EASTERN SELWYN BASIN: INTEGRATING GEOCHRONOLOGICAL AND MINERAL COMPOSITIONAL DATASETS 7.1 INTRODUCTION A regional, reconnaissance-scale study was carried out on intrusions associated with a variety of styles of mineralization in the eastern Selwyn Basin, southwest Northwest Territories. This work was undertaken as one component to a Mineral and Energy Resource Assessment in the Selwyn and Mackenzie mountains. The study was designed to evaluate the magmatic thermal and fluid evolution of intrusions throughout the study area, within the context of their overall potential for associated mineralization. A preliminary characterization of the magmatic thermal and fluid evolution of ~25 intrusions in the eastern Selwyn Basin was used to speculate which intrusions in the region may be more prospective for related mineralization (Rasmussen et al., 2007). This chapter reports the final 40Ar/39Ar age determinations and magmatic fluid compositions for the intrusions that were examined during the Mineral and Energy Resource Assessment, along with more detailed interpretations of these data. 40Ar/39Ar ages determined from igneous biotites are integrated with U–Pb ages (presented in Chapter 2) to provide information on magmatic thermal timescales, from emplacement to sub-solidus conditions. Magmatic apatite and biotite compositions add information to the petrogenetic framework proposed for the plutonic suites in Chapter 2, and fluorine, chlorine ± sulfur contents in both minerals are examined from magmatic evolution and metallogenic perspectives. Apatite and biotite compositions form the basis for a regional characterization of magmatic fluid compositions using the thermodynamic calculations of Zhu and Sverjensky (1992). This chapter concludes with an assessment of the utility of the geochronological and fluid compositional data in identifying intrusions that are prospective for related mineralization, while acknowledging the limits imposed by reconnaissance-scale sampling, integration of different geochronological datasets, and uncertainties in estimating magmatic fluid compositions. 7.1.1 Thermal Evolution of Magmatic–(Hydrothermal) Systems It is generally accepted that deeper emplacement of magmas allows for prolonged cooling, extensive magma fractionation, and the development of late-stage magmatic fluids―particularly in regions of higher heat flow and (or) metamorphic grades (e.g., the lower to middle crust). Relatively shallowly emplaced (i.e., <10 km depth), slower cooling magmatic systems, however, require multiple episodes of magma recharge to remain thermally active for extended periods (e.g., Hart, 2004). The influx of hotter, incompatible element enriched, and potentially volatile-rich magmas may work to increase the metal budget of the system and to substantially change the composition and evolution of volatile material within the magma chamber (e.g., Christiansen and Keith, 1996; Vigneresse, 2006, 2007). Furthermore, the longer cooling periods experienced by multi-pulse intrusions may provide sufficient time for the segregation and evolution of a magmatic fluid that is capable of extracting and concentrating a large proportion of incompatible metals from the magma (e.g., Newberry and Swanson, 1986). Magmatic–hydrothermal systems with multiple pulses of magma and (or) fluid influx are generally thought to have 403 a greater capacity to form sizable ore deposits compared to short-lived and (or) single-pulse magmatic–hydrothermal systems (e.g., Meinert et al., 2005). Extended magmatic thermal histories, therefore, may have important implications with respect to the size of resulting mineral deposits. The duration of a magmatic system emplaced into the upper crust is considered to be the amount of time in which the intrusion cools from emplacement temperatures to sub-solidus conditions. The former is generally taken as the saturation or crystallization temperature of zircon or monazite in the magma, which typically occurs at temperatures of 950–750°C (e.g., Harrison and Watson, 1984). The latter is usually taken as either the closure temperature for argon exchange for biotite (~300–360°C; Blanckenburg et al., 1989; Hodges, 2003)57, or the closure temperature for helium exchange in zircon (~200°C; Reiners et al., 2004). The most common measure for the duration of a magmatic thermal system is the difference between U–Pb crystallization age for magmatic zircon or monazite and the 40Ar/39Ar cooling age of magmatic mica (e.g., Coulson et al., 2002; Hart, 2004; 2008; Kouzmanov et al., 2008; Baumgartner et al., 2009; Kuşcu et al., 2011; Li et al., 2011; Chiaradia and Schütte, 2011). When evaluating magmatic systems, 40Ar/39Ar ages of unaltered igneous micas provide the low temperature constraint on the duration of magmatic cooling alone. When evaluating magmatic–hydrothermal systems, however, the micas are typically the hydrothermal products of alteration in intrusions, veins, associated mineralization (e.g., biotite-facies skarn), and (or) in country rocks hosting mineralization. The hydrothermal micas provide constraints on the cooling of the magmatic–hydrothermal system, from emplacement to the end of mineralization. In this chapter the focus will be on the duration of magmatic systems, although much of the following discussion on magmatic–hydrothermal systems (on which more work has been done) is still applicable. The vast majority of research on the duration of magmatic–hydrothermal mineralizing systems has focused on copper porphyry and gold–silver epithermal systems (e.g., Kouzmanov et al., 2005, 2008; Baumgartner et al., 2009; Li et al., 2011; Chiaradia and Schütte, 2011), which are emplaced at relatively shallow depths (e.g., <3 km). Convective cooling processes are directly linked to the permeability of the intrusion and its host rocks and are, therefore, particularly relevant in shallow magmatic–hydrothermal settings. Modeling of convective cooling suggests that shallowly emplaced plutons cool quickly, over several 10’s of k.y., and that this short duration is controlled by the shallow depth of emplacement (Cathles, 1997). Although there may be a correlation between longer cooling periods and mineralization, short cooling periods do not preclude mineralization in these shallower magmatic–hydrothermal settings. There are many examples of large deposits that formed in the porphyry– epithermal environment that have experienced multiple pulses of magma and fluid over timeframes of less than 0.5 m.y. (Ronacher et al., 2002; examples in Hart, 2004; Kouzmanov et al., 2005; Simmons and Brown, 2008; Kuşcu et al., 2011; Li et al., 2011; Chiaradia and Schütte, 2011; Bineli Betsi et al., 2012). In shallow epithermal or porphyry settings, a magmatic–hydrothermal system does not have to be particularly long-lived to produce economic mineralization. That said, there are also many large porphyry or epithermal deposits that have experienced longer thermal histories of ~1.5–4 m.y. and more (Hart, 2004; Kouzmanov et al., 2008; Baumgartner et al., 2009; Wainwright, 2009; Sillitoe and Mortensen, 2010; Kuşcu et al., 2011; Li et al., 2011). These longer-lived systems are characterized by multiple batches of magma separated by several 100’s of k.y. to several m.y. (e.g., Sillitoe and 57 For biotite containing 36–70% annite from intermediate to felsic intrusions. 404 Mortensen, 2010). Conductive cooling is the predominant mode of cooling in impermeable, or deeper settings. As with convective cooling, however, modeling of conductive processes suggests that at moderate depths of emplacement (~6 km) a single magma pulse will still cool within several tens of k.y. (e.g., <65 k.y.; Hart, 2004). Observations from natural systems (i.e., intrusive textures, contact aureole minerals, fluid inclusion studies, etc.) suggest that many moderately deep (i.e., 6–10 km) intrusions, however, experienced more protracted thermal histories than the shallower systems. To produce longer cooling periods in the upper crust, where country rock temperatures are typically limited to ~200–300°C, heat influx by the episodic replenishment of magma chambers with hotter magmas must occur (e.g., Hart, 2004). The building of a composite intrusion with sequential batches of magma is consistent with newer models of magma generation, ascent, and emplacement (e.g., Vigneresse, 2007). This scenario allows for prolonged magma fractionation and fluid evolution in one magma chamber (or several closely spaced magma chambers). It also allows for the addition of metals and volatile species to the magma chamber (e.g., Christiansen and Keith, 1996; Vigneresse, 2006, 2007). Additional heat can also be derived from spatially and temporally distinct magmatic ± hydrothermal systems with partially overlapping thermal aureoles. Differences in the timing of such events may be resolvable to within several 100’s of k.y., but the overlapping (hydro)thermal regimes commonly reset the argon isotopic system in micas associated with the earlier events (Chiaradia and Schütte, 2011)58. Both sources of external heat―the influx of hotter magmas, and nearby magmatic ± hydrothermal systems―could enhance the likelihood for associated mineralization by extending the cooling period of a magma. Regardless of the initial temperature of the magma, modeling of both convective and conductive magmatic systems demonstrates that a single magma pulse emplaced at less than 3 kbar pressure will cool to about 25% of the pre- emplacement temperature (i.e., sub-solidus conditions) within 800 k.y. of emplacement (Cathles, 1997). This limits all single-pulse magmatic events in the upper crust to timeframes of less than 1 m.y. Multi-pulse magmatic systems are thermally longer-lived, and are more likely to be associated with significant mineralization (e.g., Meinert et al., 2005), even if shallower levels of emplacement and shorter-lived systems do not preclude mineralization. The data presented in this chapter, as part of the larger PhD investigation, focuses on the magmatic–thermal evolution of intrusions—specifically, the identification of plutons in the Selwyn Basin that were thermally active for longer than 1 m.y. Herein, the term “absolute” cooling period refers to the difference between the reported U–Pb age (from magmatic zircon) and the 40Ar/39Ar age (from magmatic biotite) without including the 2-sigma errors for both ages. “Minimum” and “maximum” cooling periods are defined as the quantitative difference between the U–Pb crystallization age and the 40Ar/39Ar cooling age when the 2-sigma error ranges are included (i.e., the maximum and minimum possible differences between the two ages, within the 95% confidence interval). The overall duration of the magmatic systems examined here will be qualitatively referred to as the “thermal” or “cooling” history of the intrusion in question. 58 In these cases, the timing of events may only be resolved with high precision U–Pb geochronology on robust minerals (e.g., Sillitoe and Mortensen, 2010). 405 7.1.1.1 Limits Associated with Dating Magmatic Events The 238U/206Pb and 235U/207Pb isotopic systems provide the most accurate estimates for the timing of events, mainly because of their precisely known decay constants for the parent isotopes (e.g., Chiaradia and Schutte, 2011) and the calibration of these ages against tightly constrained and widely accepted standard materials. Robust minerals containing trace amounts of uranium, thorium, and lead—such as zircon, monazite, xenotime, and titanite—are particularly effective in preserving the U–Pb isotopic systems because of their resistance to re-equilibration during alteration and (or) re-heating. These accessory minerals, however, are typically only found in intermediate to felsic igneous rocks and high-grade (amphibolite facies and greater) metamorphic rocks, with few exceptions. External errors for the U–Pb age determinations range from 0.1% to over 1%, depending on the analytical method employed (Chiaradia and Schutte, 2011); however, many of the techniques that produce higher precisions sometimes have a poorer accuracy. The highly precise ages determined by chemically abraded ID–TIMS (Isotope Dilution – Thermal Ionization Mass Spectrometry; Mattinson, 2005) techniques for whole grains of zircon and monazite, for example, may incorporate older inherited cores or antecrystic (older magmatic) zircon in the final age determination (Hodges, 2003; Hart, 2004; Miller et al., 2007; Chiaradia and Schutte, 2011). These inherited components may only be a few m.y. older than the magmatic rims, but are thought to be largely responsible in cases where a significant range of ages (i.e., outside of 2-sigma errors constraints) is produced from several concordant analyses (Miller et al., 2007). Careful grain selection and single-grain (vs. multi-grain) analysis improves on the precision of ID–TIMS age determinations, but there are still instances where the effects of inheritance are unavoidable (e.g., Hodges, 2003; Miller et al., 2007). Selective analysis of magmatic rims for individual grains can be done using in situ techniques, such as the SHRIMP (Sensitive High Resolution Ion Microprobe) or LA–ICP–MS (Laser Ablation – Inductively Coupled Plasma – Mass Spectrometry) methods. In situ analysis makes it possible to avoid dating most or all of any inherited zircon sectors, but its precision may be reduced and 2-sigma errors increase up to 1% with these methods. Timescales of less than ~2 m.y. for 100 Ma intrusions, therefore, may not be differentiated at this upper estimate of precision, even if the age is more accurately determined. Similarly, timescales of several 10’s to 100’s of k.y. for magmatic systems are well beyond the reach of the 2.5% external error previously recommended for comparisons between 40Ar/39Ar ages and U–Pb ages (e.g., Hart, 2004; Mattinson, 2005; Kuiper et al., 2008; Chiaradia and Schutte, 2011). The large external error for the argon isotopic system results from uncertainties in decay rates and the ages of standards (Renne et al., 1998a), as well as the observation that argon ages are consistently ~1% younger than U–Pb ages that date the same event (Schoene et al., 2006)59. For magmatic systems alone, there are no appropriate measures for determining the length of their duration from emplacement to sub-magmatic temperatures without employing more than one isotopic system. Recently, however, a new and more precise age for the widely used Fish Canyon sanidine argon standard has demonstrated that much of the difference between U–Pb and 40Ar/39Ar ages was a product of an inaccurate previously reported age for this standard, and the external error for comparisons between these two isotopic systems may now be as low as 59 The slightly younger 40Ar/39Ar ages were thought to indicate prolonged cooling of events; this inference, however, has since been shown to be incorrect in some notable instances (Kuiper et al., 2008). 406 0.25% (Kuiper et al., 2008). With this smaller estimate of external error, distinguishing magmatic–hydrothermal events that are only separated by about 250 k.y. (for 100 Ma intrusions) may now be possible for many magmatic– hydrothermal systems, particularly if the more precise U–Pb methods is also used (e.g., chemically abraded ID– TIMS). In magmatic systems alone, 40Ar/39Ar ages record the time of igneous mica closure to argon exchange, which will occur later and at lower temperatures than the closure time for uranium and lead exchange in igneous zircon or monazite. Thus, the U–Pb age should not be replicated by the 40Ar/39Ar age if the intrusion experienced a thermal history of more than 250 k.y. In an attempt to acknowledge the limits that may be imposed by integrating the LA–ICP–MS U–Pb and 40Ar/39Ar geochronological datasets presented here, and because cooling periods of less than 1 m.y. indicate rapid cooling (i.e., the emplacement of a single magma pulse), a greater emphasis is placed on identifying intrusions that have longer (>1 m.y.) minimum cooling periods. As stated earlier, minimum cooling periods incorporate the 2-sigma errors and are calculated as the difference between the youngest possible U–Pb age and the oldest possible 40Ar/39Ar age, as constrained by the 2-sigma error estimates. 7.1.2 Magmatic Fluid Compositions from Mineral Chemistry As with the identification of intrusions that are thermally and hydrothermally active for longer periods of time, the study of volatile species (e.g., F–, Cl–, HS–, OH–) in cooling magmas is also of interest in metallogenic studies involving intrusive rocks. These ionic species act as ligands in magmatic ± hydrothermal systems, which form soluble compounds with metals in the magma, and facilitate the transport of these metals in solution (e.g., Seward and Barnes, 1979; Crerar et al. 1985; Piccoli and Candela, 1994). The ability of a solution to transport metals is strongly dependent on the activity or thermodynamic concentration of appropriate ligand(s); indeed, the presence of suitable ligand species is thought to be equally, if not more important than the overall metal content of the magma (e.g., Crerar et al. 1985; Seward and Barnes, 1997). The accumulation of a ligand- and metal-rich fluid, therefore, and its timely extraction from the magma chamber, has important implications regarding the fluid’s capacity to sequester, concentrate, transport, and ultimately precipitate metals. Herein, a magma body is defined as comprising crystals, melt, and dissolved volatile species (or “magmatic volatiles”). An exsolved magmatic volatile phase (MVP), is defined as an immiscible, buoyant, aqueous and (or) vapourous solution that saturated in the magma and was in equilibrium with the melt and crystals for at least a short period of time after saturation and before it was separated from the magma. Here, the terms “magmatic fluid” and “magmatic fluid phase” are used interchangeably with “MVP”. Relative F:Cl:OH contents in a magma should not vary if these constituents are concentrated in an isothermal, closed-system magma (Piccoli and Candela, 1994). Once MVP exsolution occurs, however, fluorine is partitioned preferentially into minerals and melt over fluid (e.g., Zhu and Sverjensky, 1991), and chlorine and H2O partition preferentially into the fluid (e.g., Candela, 1986; Piccoli and Candela, 1994, Webster, 1997a,b). Two approaches to using F:Cl:OH variations in magmatic minerals or fluid to predict associated mineralization are: (1) to detect the early loss of a mineralizing MVP (e.g., first boiling), or 407 (2) to detect the buildup of a late-stage mineralizing MVP (e.g., ore-second boiling, or a non-boiling segregation of immiscible saline fluid). In the first case, the earlier exsolution of an MVP lowers Cl/F (and Cl/H2O at > 0.5 kbar) and H2O/F in the overall magmatic system; this in turn is thought to lower chlorine and hydroxyl contents in hydrous magmatic minerals (and subsequent MVP; Candela, 1986; Christiansen and Keith, 1996; Sheets et al., 1996; Webster, 1997a,b; Lang, 1998; Piccoli et al., 1999; Villemant and Boudon, 1999; Piccoli and Candela, 2002). The early loss of an MVP is typical of intrusions associated with copper-rich, porphyry-style mineralization (Newberrry and Swanson, 1986), but would presumably also be important in other intrusion-related deposit types where metals behave compatibly in the magma. In the second case, the late exsolution of an MVP might not deplete chlorine and H2O (relative to fluorine) in the magmatic system, as well as in hydrous magmatic minerals. Relative fluorine depletions will occur if hydrous minerals are removed by crystal fractionation, which will lead to an enrichment in dissolved chlorine and H2O, as long as off-gassing (MVP exsolution) is not also occurring and these volatiles are allowed to accumulate in the magma. These processes will lead to higher Cl/F, Cl/H2O, and H2O/F in the magma (Zhu and Sverjensky, 1991)―and presumably in magmatic minerals crystallizing in equilibrium with the magma―at least until MVP exsolution occurs. The late loss of an MVP is particularly applicable to tungsten mineralizing systems where mineralization is generally attributed to late-stage, non-boiling MVP segregation from a highly crystallized intrusion (e.g., Newberry and Swanson, 1986; Chapter 6), but it could also be relevant to other intrusion-related deposit types where metals behave incompatibly in the magma. 7.1.2.1 Limits Associated with Studies of Magmatic Volatile and Fluid Composition Many studies have attempted to document major and minor element, halogen, and (or) sulphur concentrations in hydrous magmatic or hydrothermal minerals, and to link variations in mineral compositions to an association (or lack thereof) with mineralization. Several studies have used biotite to try and distinguish intrusions associated with mineralization (e.g., Kesler et al., 1975; Mason, 1978; Parry et al., 1978; Munoz, 1984; Lang, 1998). These biotite- based studies have generally met with limited success, commonly reporting contradictory results or compositional parameters that could not be reproduced by other studies elsewhere (Lang, 1998). Studies incorporating apatite compositions (or of apatite composition alone) have also typically met with contradictory results. For example, chlorine enrichment in apatite ± fluorine or chlorine enrichment in biotite has been observed in intrusions associated with copper-porphyry mineralization (Loferski and Ayuso, 1995; Sheets et al., 1996). Depletions in concentrations of fluorine and chlorine have been observed in apatites from intrusions with related mineralization in proximal country rocks, whereas enrichments in apatite have been observed in intrusions hosting the mineralization (e.g., Loferski and Ayuso, 1995; Sheets et al., 1996; Lang, 1998; Piccoli and Candela, 2002). For example, Lang (1998) found high F/Cl in apatite and biotite from intrusions associated with economic copper-porphyry mineralization (with some exceptions). Other studies have shown that F/Cl in apatite increases with the overall degree of magmatic differentiation (e.g., Candela, 1986; Lang, 1998; Sha and Chappell, 1999; Belousova et al., 2001), although not all 408 studies of apatite demonstrate this (e.g., Hoskin et al., 2000). Similarly, variations in fluorine and chlorine abundances in apatite from plutons in the northern Cordillera have also been linked to the composition of the intrusion (i.e., degree of differentiation and (or) aluminicity vs. alkalinity; Coulson et al., 2001; Lang, 2001; Rasmussen et al., 2011a). One explanation for contradictory inter- and intra-study results may be that the timing of MVP exsolution is rarely considered or is poorly constrained. Although the timing of fluid saturation is strongly controlled by the initial H2O content, it is also subject to many other factors. Some of these include: the amount of volatile material lost during depressurization, the concentration of other volatile species in the magma (e.g., CO2), whether recharging of volatile concentrations by magma influx and mixing occurred, and the interplay between volatile degassing versus build-up in the magma. As described above, intrusions that undergo early depressurization-induced MVP exsolution typically lose chlorine and H2O―and any metals in solution―to the exsolved volatile phase (e.g., Tingle and Fenn, 1984; Candela, 1986; Webster, 1997a,b; Loferski and Ayuso, 1995; Lang, 1998; Piccoli and Candela, 2002). Intrusions that undergo later crystallization-induced MVP exsolution will be rich in chlorine and H2O (relative to fluorine) if exsolution occurs after crystallization is nearly complete (Piccoli and Candela, 2002; Chapter 6). Halogen concentrations and MVP compositions (calculated from apatite and biotite compositions) may also be decoupled, or show opposing trends (Chapter 6). This decoupling could result if degassing occurs after early higher temperature saturation of apatite―typically between 750–950oC in felsic systems (Bea et al., 1992) and before the saturation of biotite―which may be as low as 400–600oC in magmas at the magnetite–hematite buffer (Wones and Eugster, 1965). The closure temperature for halogen exchange in biotite is unknown, but it could be even lower than the saturation temperature (perhaps 300–350°C; van Middelaar and Keith, 1990). Despite its obvious importance, the timing of MVP exsolution is just one of many facets in the complicated petrogenesis of magmatic systems. Comparisons between, and even within, biotite and (or) apatite composition datasets may result in contradictory results if differences in the petrogenesis of the magma and associated MVP are not similarly considered (Ague and Brimhall, 1988a,b; Loferski and Ayuso, 1995; Lang, 1998; Selby and Nesbitt, 2000). Much of the compositional work done on biotite in intrusions associated with mineralization also pre-dates the publication of well-constrained cationic controls on halogen partitioning from fluid to biotite (Lang, 1998). Mutual Fe–F and Mg–Cl avoidances related to the reciprocal effects of octahedral and hydroxyl site occupancy in biotite will also affect the partitioning of halogens into micas. This typically causes annitic biotites to be fluorine depleted and phlogopitic micas to be chlorine depleted (e.g., Zhu and Sverjensky, 1992). The partitioning of halogens between hydrous minerals and melt or fluid is also dependant on pressure, temperature, pH, and fluid composition conditions (e.g., Zhu and Sverjensky, 1991, 1992; Brenan, 1993a; Piccoli and Candela, 1994). Furthermore, generally fluorine preferentially partitions into apatite, and generally chlorine preferentially partitions into biotite— regardless of magma, mineral, or fluid compositions (Zhu and Sverjensky, 1991). It is because of these internal and external controls on the partitioning of fluorine and chlorine between minerals, melt, and fluid that the abundance of halogens in biotite and apatite do not directly reflect their abundance in associated melts and fluids60. Once the 60 This may also be why earlier studies that attempted to link biotite and (or) apatite compositions to mineralization did not identify any widely applicable compositional parameters that are indicative of associated mineralization. 409 controls on halogen partitioning into apatite and biotite are considered and quantitatively accounted for, these minerals may provide independent estimates of F:Cl:OH in the MVP with which they equilibrated (Coulson et al., 2001). In cases where these independent estimates of MVP composition do not agree (i.e., apatite–biotite disequilibrium), apatite may approximate MVP compositions at higher temperature magmatic conditions, and biotite may approximate MVP compositions at lower temperature magmatic―or even hydrothermal―conditions. 7.1.3 The Importance of Magmatic Petrogenesis and Differentiation Continental granitoids are commonly subdivided into three petrogenetic classes (e.g., Chappell and White, 1974; Loiselle and Wones, 1979): • “S-type” calc-alkaline granites, derived from partial melting of supracrustal (sedimentary) rocks; • “I-type” calc-alkaline granites, derived from partial melting of infracrustal (igneous) rocks; and • “A-type” calc-alkaline to alkaline granites, thought to have been derived from partial melting of an anhydrous source―although alkaline examples are almost certainly sourced from the mantle and (or) mantle-derived rocks (e.g., Frost and Frost, 2011). This ‘alphabetic’ classification scheme for continental granitoids has fallen into disfavour due to the extensive overlap between these three end-member magma types that can be observed, as well as the petrogenetic implications of using these terms. It does, however, provide a well known scheme for characterizing granitoids on the basis of simple compositional observations. Furthermore, much of the discussion of apatite composition in the literature has been interpreted in terms of the S-, I-, and A-type terminology. These classes are used here as a simple, first-order approximation of plutonic suite composition―more detailed discussions of granitoid type and composition in the eastern Selwyn Basin have already been described in previous chapters (2, 4, and 5). Traditionally S-type plutons have been associated with tungsten–tin mineral occurrences; however, it has been observed that economically significant tungsten deposits do not have contain appreciable amounts of tin, nor are they associated with S-type plutons (e.g., Newberry and Swanson, 1986; Newberry, 1998). Instead, intrusions associated with large tungsten occurrences, at least in the form of skarn mineralization, are mineralogically, compositionally, and isotopically more similar to highly fractionated, reduced, I-type granitoids (e.g., Newberry, 1998). Similarly, the intrusions that are associated with known gold mineralization in the northern Cordillera are also typically I-type granitoids, although volumetrically smaller, associated late-stage alkalic A-type intrusive phases may also be important with respect to intrusion-related gold (Lang and Baker, 2001). Gold-associated granitoids may not be as fractionated or reduced as tungsten-associated plutons as a whole. However, gold mineralization is in some cases directly associated with highly fractionated intrusive phases (e.g., Fort Knox, Alaska), and in most cases with plutons that evolved highly fractionated phases (Lang and Baker, 2001). It would, therefore, be important to determine whether or not a pluton is of supracrustal (S-type) or infracrustal (I-type) origin in assessing its potential for associated tungsten or gold mineralization. Although this has been done on the basis of whole rock geochemistry and mineralogy in Chapter 2, and zircon studies in Chapter 5, it is of general interest to ascertain whether or not apatite (and biotite) compositions in these plutonic suites are consistent with the interpreted petrogenesis. 410 Mineralogical studies have revealed that elevated MnO (>0.2 wt.%) and FeO (>0.15 wt.%) are typical of apatites from S-type, and fractionated I-type, granitoids (Sha and Chappell, 1999; Broska et al., 2004). Elevated manganese and iron may be inherited from supracrustal rocks containing a higher component of lithic material, such as greywacke or pelitic protoliths. However, granitoids that fractionate along a calc-alkaline evolutionary trend will become enriched in iron with increasing degrees of differentiation (Frost et al., 2007), and FeO may be inversely correlated to CaO in apatite during magma differentiation (Piccoli and Candela, 2002). MnO is also known to increase in apatite as the aluminicity of a magma increases or the oxygen fugacity of a magma decreases (Sha and Chappell, 1999; Hoskin et al., 2000) ―both of which are associated with mineral fractionation and overall magmatic differentiation. Although partitioning preferences ensure that fluorine is always enriched in apatite relative to chlorine and hydroxyl (e.g., Zhu and Sverjensky, 1991; Pan and Fleet, 2002), elevated F/Cl in both apatite and magma is also a feature of many crustally derived magmas (e.g., Sallet, 2000; Broska et al., 2004). High F/Cl may be inherited from supracrustal melt sources, which are composed of chlorine depleted detritus (Blevin, 1993; Loferski and Ayuso, 1995; Sha and Chappell, 1999; Belousova et al., 2001). High F/Cl may also be inherited from highly metamorphosed supracrustal to infracrustal protoliths―such as granulite-grade metamorphic rocks―due to the loss of chlorine to metamorphic fluid and(or) melt (Collins et al., 1982; Creaser et al., 1991; Skjerlie and Johnston, 1992, 1993; Loferski and Ayuso, 1995). Other causes of relative fluorine enrichment in apatite are early or continual off- gassing of chlorine and H2O (e.g., Coulson et al., 2001), the fractionation of anhydrous minerals as a magma cools (e.g., Sha and Chappell, 1999), or even just a reduction in magma temperature (Piccoli and Candela, 1994). In some instances, elevated fluorine concentrations are the product of the 1:1 coupled substitution of F– + CO32– for PO43– in the apatite crystal (McArthur, 1990; Lang, 2001). In other cases fluorine concentrations are greater than what may be accommodated in the halogen site of apatite because of improper apatite grain orientations during EMP analysis, which will produce falsely elevated fluorine measurements (Stormer, et al., 1993; Raudsepp, 1995). Clearly high F/Cl measurements in apatite are influenced by several factors and could be associated with more than one of the above processes. Sulfur-poor apatite (i.e., <0.04 oxide wt.%) is also typical of both S-type and differentiated I-type magmas (Sha and Chappell, 1999). Low sulfur concentrations in apatite may simply reflect low sulfur content in an oxidized magma (e.g., Parat and Holtz, 2005). However, in a reduced magma, minimal sulphate is available to substitute for PO43– in apatite, which leads to low sulfur contents in the apatite, regardless of the sulfur content of the magma (Peng et al., 1997; Pan and Fleet, 2002; Sato et al., 2004). 7.2 GEOLOGICAL BACKGROUND The study area is located in the eastern Selwyn Basin, and focuses on intrusions emplaced in the Northwest Territories (Fig. 7.1). In this region, intermediate to felsic intrusions of ‘mid’-Cretaceous age were emplaced in an ‘inboard’ post-arc setting. Plutonic suites defined in the study area include the 99–96 Ma Tay River suite, which typically forms moderate to large intrusions, and a long narrow belt known as the Tombstone–Tungsten Belt (TTB; 411 Figs. 7.1 and 7.2). The TTB comprises mainly moderate-sized to very small plutons and stocks, and includes the 98– 96 Ma Tungsten, 98–95 Ma Mayo, and 93–89 Ma Tombstone suites (Figs. 7.1 and 7.2). Although the TTB is located north of the Tay River suite across most of the northern Cordillera, in the eastern Selwyn Basin the TTB and the Tay River suite overlap geographically (Figs. 7.1 and 7.2). The Tungsten and Tombstone suite plutons have been further sub-divided on the basis of geography (i.e., southeast vs. northeast Selwyn Basin, at ~62°N), which corresponds to minor intra-suite chemical differences across the eastern Selwyn Basin. Most of the ‘mid’-Cretaceous granitoids in the eastern Selwyn Basin have high-K calc-alkaline, I-type compositions (Chapter 2). These plutons are thought to have been derived from predominantly infracrustal melt sources of intermediate composition. The Tombstone suite, which comprises the youngest plutons in the study area, is also associated with volumetrically minor, (per)alkalic (A-type) intrusive phases, however, that indicate the involvement of enriched mantle-derived magmas in the petrogenesis of this suite. One example of a high-K calc-alkaline S-type plutonic suite is represented by the Hess–MacMillan sub-suite, which is of Mayo to Tombstone suite age (Fig. 7.2; Chapter 2). This sub-suite is interpreted to have been derived from partial melting of a quartzo–feldspathic protolith of possible Neoproterozoic age (Chapters 4 and 5). Early to ‘mid’-Cretaceous, syn- to post-accretionary magmatism was the most metallogenically important igneous event in the northern Cordillera (e.g., Hart et al., 2004a), and this region is known to be one of the richest tungsten districts in the world (e.g., Dawson and Dick, 1978; Dick, 1979; Dick and Hodgson, 1982; Dawson, 1995). The eastern Selwyn Basin is particularly notable as it hosts two world-class tungsten skarn deposits (Cantung and Mactung; cf. Menzie and Jones, 1986). In addition to world-class examples, numerous other intrusion-related tungsten skarns and precious and base metal occurrences are also present in the study area (NORMIN, 2013). LENED61 is the third largest tungsten resource in the eastern Selwyn Basin, with an indicated resource of 0.75 Mt grading 1.17% WO3 and 0.15% Cu (Figs. 7.1 and 7.2; NORMIN, 2013). Some of these occurrences have undergone advanced exploration and drilling, although the difficulty in accessing these occurrences may be partly responsible for their sub-economic status (e.g., RUDI, CAC, LENED, IVO, etc.; Fig. 7.2; NORMIN, 2013). Several of these occurrences would also have been good targets for future exploration if most of this region had not been permanently closed to exploration activity for socio–environmental reasons. All of the ‘mid’-Cretaceous plutonic suites in the eastern Selwyn Basin are associated with abundant intrusion-related mineral occurrences of variable metal tenor and mineralization style (Fig. 7.2). In addition to tungsten skarn mineralization, many of the intrusions in the eastern Selwyn Basin fit the compositional profile of plutonism that is typically associated with intrusion-related gold mineralization. Indeed, the intrusion-related gold deposit model was largely developed on observations of the auriferous members of the TTB in Yukon and Alaska (Fig. 7.1; e.g., Lang and Baker, 2001). Several plutons in the eastern Selwyn Basin, for example, also have compositions that are typical of reduced intrusion-related gold (Hart et al., 2004a, 2005; Chapter 2). These characteristics include slightly oxidized, metaluminous, sub-alkalic, intermediate to felsic I-type compositions, and are associated with volumetrically small, strongly reduced and peraluminous, or strongly oxidized and alkalic phases 61 Mineral occurrences and showings are referred to by fully capitalized names, whereas only the first letter for mineral deposits (Mactung and Cantung) and pluton names is capitalized. 412 (Lang and Baker, 2001; Hart, 2005). The Mayo and Tombstone plutonic suites of the TTB fit these criteria, although several members of the Tay River plutonic suite might also fulfill most these criteria (except for the association with alkalic magmas). Although there are no known examples of intrusion-related gold occurrences in the study area, there is evidence that this type of mineralization could be present, including: • gold-rich mineral occurrences and placer gold showings associated with Tombstone suite plutons (Fig. 7.2); • sheeted gossanous fractures, or en echelon quartz veins, containing visible arsenopyrite or scordierite, galena, chalcopyrite or copper oxide, and (or) sphalerite (± anomalous Au, Bi, Mo, Ag, Pb, Zn, Sb, or Hg; Rasmussen et al., 2005, 2007) hosted in the Mt. Christie pluton (Mayo suite), in a surface exposure of the Mine Stock at Cantung (Tungsten suite), and as stockwork veins in a silicified dolostone at the CHUCK showing (Fig. 7.2); and • a plethora of small proximal tungsten and (or) copper skarns with bismuth–gold enrichments, as well as distal silver ± gold-rich base metal veins and mantos―typically found distal to intrusion-related gold (e.g., Hart, 2005)―throughout the study area (Fig. 7.2). Therefore, despite the lack of documented occurrences of this type in the eastern Selwyn Basin, the continuation of the TTB into the eastern Selwyn Basin, where local gold-rich showings also occur (Fig. 7.2), suggests that intrusions prospective for related gold mineralization are present in the southwestern Northwest Territories. 7.2.1 Thermal and Magmatic Fluid Conditions for Intrusion-related Tungsten Skarns In most mineral deposit models, tungsten skarns are interpreted to form at moderate depths that correspond to lithostatic pressures of 1.5–2.5 kbar (e.g., Newberry and Swanson, 1986). These moderate levels of emplacement are thought to be responsible for longer cooling periods in the associated intrusions. This allows tungsten―which behaves incompatibly in reduced, ilmenite-poor magmas (Candela and Bouton, 1990; Newberry, 1998)―to concentrate in the residual melt as crystallization of the magma progresses to near-solidus conditions (e.g., Newberry and Swanson, 1986). At this point, the exsolution of a saline (NaCl-rich) aqueous solution is thought to extract and transport the tungsten from the residual melt, through the plutonic carapace to prospective sites of mineralization (Newberry and Swanson, 1986; Keith et al., 1989; Candela and Piccoli, 1995). Fluid inclusion studies of late-stage veins in TTB intrusions indicate that magma emplacement and crystallization for tungsten-associated intrusions (including those associated with the Mactung and Cantung deposits) occurred at moderate upper crustal depths corresponding to pressures of 2–3.8 kbar (e.g., G.R. Thompson, unpublished data, 2000; Baker and Lang, 2001; Marshall et al., 2003; Yuvan et al., 2007; Rasmussen et al., 2011b). The upper limit of 3.8 kbar is probably too high given that andalusite is observed in hornfelsed contact aureoles for most of the plutons in the study area (e.g., Anderson, 1982, 1983; Mathieson and Clark, 1984; Gordey and Anderson, 1993; this study). Furthermore, the pyroxene–hornfels facies that is proximal to the Mine Stock (at Cantung; Mathieson and Clark, 1984) is compatible with emplacement pressures nearer the low end of this range. Most of the felsic, peraluminous intrusions in the eastern Selwyn Basin that are prospective for associated reduced tungsten skarn mineralization were probably emplaced at about 2–2.5 kbar. 413 Previously published U–Pb (ID–TIMS and SHRIMP) and 40Ar/39Ar ages demonstrate that both the Cantung and Mactung magmatic–hydrothermal systems experienced relatively long cooling periods (Selby et al., 2003; Hart, 2004; Hart et al., 2004b; Gebru and Lentz, 2008). Absolute cooling periods for the main ~97 Ma granite phase at Mactung range from about 2.5–3.0 m.y., with a minimum magmatic cooling period of 1.5 m.y. Skarn mineralization at Mactung occurred within ~1 m.y. of granite emplacement and appears to have terminated with the ~97 Ma magmatic event (Gebru and Lentz, 2008). A ~92 Ma equigranular leucogranite has also thermally disrupted some of the 40Ar/39Ar cooling ages in the skarn and earlier intrusions (Gebru and Lentz, 2008), extending the overall absolute duration of the Mactung magmatic system to about 5–6 m.y. Similarly, the main underground granite phase exposed in underground mine workings at Cantung was emplaced at least 0.8 m.y. (absolute) prior to the formation of the magmatic–hydrothermal mineralized biotite skarn facies (Hart, 2004; Hart et al., 2004b), although the difference in the U–Pb and 40Ar/39Ar age determinations is within the range of 2-sigma errors corresponding to these two ages. Data presented in Section 7.4.1.1 will demonstrate that Cantung may have experienced a more protracted magmatic thermal history, similar to that at Mactung. A study of the evolution of magmatic fluid compositions for reduced intrusion-related tungsten skarn systems was done on a series of composite TTB plutons in the study area (Chapter 6). A gradual decrease in the activity of HCl/HF has been observed from the least differentiated (metaluminous, intermediate-composition) intrusive phases to the most differentiated (peraluminous, felsic) intrusive phases―and MVP were dominated by HF relative to HCl in felsic intrusions associated with tungsten skarns. This differentiation trend was disrupted in greisenized or metasomatized samples, in which lower HCl/HF was universally observed. The differentiation trend was also disrupted in intrusive phases with anomalously high HCl/HF and H2O/HF that were associated with tungsten skarn mineralization, including one intrusive phase (each) from the Rudi, Lened, and Mine Stock (Cantung) plutons. These specific intrusive phases were also the most (or one of the most) proximal to the tungsten mineralization. The evidence that an anomalously HCl and H2O enriched MVP may have characterized intrusive phases that are spatially associated with tungsten mineralization―particularly given the high degree of differentiation for the samples from Lened and Cantung, which would normally have produced an HF/HCl enrichment―led to the inference that these specific intrusive phases exsolved a supercritical saline aqueous MVP at near-solidus magmatic conditions. It was suggested in Chapter 6 that this method of identifying intrusive phases that may have produced and exsolved a metal- and chlorine-rich aqueous MVP (particularly at near-solidus conditions) could be of use in the exploration for tungsten skarns in the northern Cordillera. 7.2.2 Thermal and Magmatic Fluid Conditions for Intrusion-related Gold Mineralization The intrusion-related gold setting in the northern Canadian Cordillera has been well-characterized by Hart (2005). The causative pluton is typically exposed as a small, elongate body of metaluminous, sub-alkalic, intermediate to felsic composition (Lang and Baker, 2001; Hart, 2005, 2007), and is commonly cut by (or otherwise closely associated with) aplite–pegmatite and lamprophyre dykes. Associated mineralization is zoned, from a central or proximal gold- (Bi–Te ± W) rich zone, through an intermediate arsenic- (Au ± Sb) rich zone, to an outer or distal 414 silver–lead–zinc-rich zone. Intrusion-related gold typically occurs as sulphide-poor, sheeted veins hosted by the intrusion or nearby country rocks, and small gold-rich tungsten- or copper-bearing skarns may occur where the pluton is in contact with calcareous country rocks. Replacement-style or disseminated gold mineralization is found in the thermal aureole developed in country rocks overlying buried portions of the pluton. Shallowly emplaced plutons may evolve a more saline and H2O-rich mineralizing fluid (Thompson et al., 1999; Baker and Lang, 2001; Hart, 2005, 2007); however, CO2-rich mineralizing fluids are more typical and reflect the importance of this species in the parental magmas to the intrusions that are associated with gold mineralization (Lang and Baker, 2001). In the intrusion-related gold environment, the associated or host granitoids are thought to have been emplaced at pressures ranging from 0.5–3 kbar (Lang and Baker, 2001), which correspond to a wide range of upper crustal depths (~1.5–9 km). Near the lower end of the pressure range (~1.3 kbar; Baker and Lang, 2001), the composite Emerald Lake pluton (Fig. 7.1) is enriched in gold, although this mineralization occurs as sub-economic disseminations throughout the intrusion and in miarolitic cavities associated with sheeted veins (Duncan, 1999). There may be other examples of shallowly emplaced metaluminous intrusions in the study area. Preliminary aluminum–in–hornblende thermometry for the Central Nahanni pluton, for example, suggest pressures of emplacement between ~0.8–1.6 kbar (although large hornblende phenocrysts may have crystallized at deeper levels corresponding to ~2.3–4.2 kbar; this study, unpub. data). Despite evidence for the emplacement of multiple phases of compositionally variable magmas in the Emerald Lake pluton (e.g., Duncan, 1999; Coulson et al., 2002), geochronological studies of this intrusion show that it experienced short absolute cooling periods of ~1.1 m.y. (Duncan, 1999). This translates to minimum and maximum cooling periods of approximately 0.2 and 2.0 m.y., respectively, but thermal modeling for this body produced significantly shorter estimates of cooling time (70–100 k.y; Coulson et al., 2002). This is in spite of composite character of this pluton, which required several episodes of compositionally variable magma emplacement (e.g., Duncan, 1999; Coulson et al., 2002). It appears that the Emerald Lake magmas cooled too rapidly to have evolved a long-lived magmatic–hydrothermal system capable of concentrating gold as economic bodies of mineralization (Coulson et al., 2001). The Dublin Gulch deposit (Fig. 7.1), which is the largest known body of intrusion-related gold mineralization in Yukon (Hart et al., 2000; 2002; Hart, 2005), was emplaced at similarly shallow crustal levels corresponding to ~1.1 kbar (Baker and Lang, 2001) 62. Other gold-associated intrusions of the TTB that were emplaced at depths greater than 1.5 kbar record somewhat longer absolute magmatic cooling periods of ~2 m.y. (e.g., Clear Creek and Scheelite Dome; Baker and Lang, 2001; Marsh et al., 2003; Selby et al., 2003; Fig. 7.1). When 2-sigma errors are accounted for, these timescales correspond to minimum cooling periods of essentially 0 m.y. and maximum cooling periods of about 4 m.y. Although the longer thermal histories suggested by the absolute and maximum cooling periods for Clear Creek and Scheelite Dome plutons may be significant with respect to their associated economic concentrations of gold mineralization, clearly the depth of magma emplacement is not the only determining factor in the formation of intrusion-related gold mineralization. 62 Although Selby et al. (2003) determined a magmatic U–Pb age of 94.0 ± 0.3 Ma and a hydrothermal Re–Os age of 93.2 ± 0.3 Ma for the well-mineralized Dublin Gulch deposit, there are no 40Ar/39Ar age constraints from magmatic or hydrothermal micas that would allow the total duration of the Dublin Gulch magmatic-hydrothermal system to be compared with other intrusion-related gold systems in the TTB. 415 A study of the evolution of magmatic fluid compositions on the composite Emerald Lake pluton (Coulson et al., 2001) served as inspiration for similar work on intrusion-related tungsten mineralization in the study area (Chapter 6). Coulson et al. (2001) first demonstrated that HCl/HF gradually decreases from the least differentiated (earliest) intrusive phases through to the most differentiated (latest) intrusive phases. This change in MVP composition over time was attributed to the loss of chlorine from the magma as it evolved, by a process of continual or periodic exsolution and off-gassing of gold-bearing fluids (Coulson et al., 2001). The low-grade gold mineralization throughout the Emerald Lake pluton, and the gradual decrease in HCl/HF with time, were both attributed to the inability of this magmatic system to retain its volatile species until the later stages of crystallization, when fluid exsolution might have concentrated more gold in brittle fracture openings and miarolitic cavities (Coulson et al., 2001). 7.3 ANALYTICAL METHODS The sample set selected for this study was chosen based on the following criteria: • the sampled phase is representative of the overall plutonic composition, mineralogy, and textures; • there are complete U–Pb geochronological and geochemical datasets available for the sample (Chapters 2); and • all plutonic suites in the region (representing a range of compositions, ages, sizes, and metallogenic characteristics) are examined. The representative sample in most cases was a hypidiomorphic–granular, moderately to strongly K-feldspar porphyritic, biotite ± hornblende- or muscovite-bearing monzogranite or granodiorite. Sampling of leucocratic and (or) late-stage phases for this study was avoided whenever possible so that comparisons between samples would be more relevant. Two limits to this study are: (1) the collection only a single sample from each intrusion for analysis; and, (2) the assumption that nearby intrusion-related mineralization was related to the intrusive phase sampled. Predetermined U–Pb ages from the same intrusive phase were necessary to approximate the duration of cooling periods in each of the intrusions, and available geochemical data were used to estimate the relative degree of differentiation for each sample using the differentiation index (D.I.) of Thornton and Tuttle (1960; Table 7.1). The D.I. is the sum of the percentage of the three normative minerals quartz, K-feldspar, and albite, and is used as a proxy for the degree of fractional crystallization undergone by a magma as it approaches a pure SiO2–NaAlSiO4– KAlSiO4 system (Thornton and Tuttle, 1960). The D.I. has been rounded up or down to within one unit of its actual value in several samples (with the exception of Pelly River, which was rounded up 1.5; see Table 7.1) to make the data distribution in compositional plots better for viewing purposes. All sample preparation (with the exception of irradiation for biotite separates; see Section 7.3.1) and data collection were completed at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia (UBC), with the assistance of T. Ullrich (PCIGR) and M. Raudsepp (Electron Microbeam/X–ray Diffraction Facility). Apatite and biotite were analyzed following the methods of Coulson et al. (2001) and Raudsepp (1995). Related data reduction and magmatic fluid compositions (calculated using the equations of Zhu and Sverjensky, 1992) were done with the assistance of G. Dipple and L. Groat, after methods described by I. 416 Coulson (2006, pers. comm.). A more detailed description of analytical methods employed in analyzing apatite and biotite, and calculations used for determining mineral and magmatic fluid compositions are provided in Chapter 6. A preliminary interpretation of data presented here was summarized in GSC Open File 5344 (Rasmussen et al. 2007), and an earlier interpretation of the magmatic fluid compositions was also published in the Gangue Newsletter of the Mineral Deposits Division of the Geological Association of Canada (Rasmussen et al., 2011a). 7.3.1 40Ar/39Ar Geochronology Biotite mineral separates picked from a clean fraction of lightest material off a wet shaking table were hand-picked, washed in acetone, dried, wrapped in aluminum foil and stacked in an irradiation capsule with similar-aged samples and neutron flux monitors (Fish Canyon Tuff sanidine standard, 28.02 ± 0.56 Ma; Renne et al., 1998b). The samples were irradiated at the McMaster Nuclear Reactor in Hamilton, Ontario, for 90 MWH, with a neutron flux of approximately 3x1016 neutrons/cm2. Analyses (n = 57) of 19 neutron flux monitor positions produced errors of less than 0.5% in the J value. At PCIGR, the mineral separates were then step-heated at incrementally higher powers in the defocused beam of a 10W CO2 laser (New Wave Research MIR10) until fused. The gas exsolved at each step was analyzed by a VG5400 mass spectrometer equipped with an ion-counting electron multiplier. All measurements were corrected for total system blank, mass spectrometer sensitivity, mass discrimination, radioactive decay during and subsequent to irradiation, as well as interfering argon from atmospheric contamination and the irradiation of Ca, Cl, and K. Isotope production ratios were: (40Ar/39Ar)K = 0.0302 ± 0.00006, (37Ar/39Ar)Ca = 1416.4 ± 0.5, (36Ar/39Ar)Ca = 0.3952 ± 0.0004, Ca/K = 1.83 ± 0.01(37ArCa/39ArK). The plateau and inverse correlation ages were calculated using Isoplot v.3.09 (Ludwig, 2008). Errors are quoted at the 2-sigma (95% confidence) level and are propagated from all sources except mass spectrometer sensitivity and the age of the flux monitor. The best statistically-justified plateau and plateau age were picked based on the following criteria: • three or more contiguous steps comprising more than 50% of the 39Ar; • probability of fit of the weighted mean age greater than 5%; • slope of the error-weighted line through the plateau ages equals zero at 5% confidence; • ages of the two outermost steps on a plateau are not significantly different from the weighted-mean plateau age (at 1.8 sigma, six or more steps only); and • outermost two steps on either side of a plateau do not have non-zero slopes with the same sign (at 1.8 sigma, nine or more steps only). Plateau ages (determined using an age of 28.02 Ma for the Fish Canyon Tuff sanidine standard; Renne et al., 1998b) and 2-sigma internal errors for 24 plutonic samples are reported in Table 7.2 and the isotopic ratios, ages, associated uncertainties for individual steps, and plateau plots are presented in Appendices C.1 and C.2. Re-calculated 40Ar/39Ar ages and 2-sigma internal errors, using the new age of 28.201 Ma for the Fish Canyon Tuff sanidine standard (Kerr et al., 2008; Kuiper et al., 2008), which was accepted by the geochronological community in 2009 (cf. McClean, 2009), are presented in Table 7.2. 417 7.3.2 Mineral Analysis 7.3.2.1 Apatite Apatite grains from 24 intrusive samples were separated from crushed material using conventional heavy liquid separation methods. Grains were typically prismatic with well-defined crystal faces and pyramidal terminations, although commonly broken perpendicular to the c-axis. The separation of individual grains from crushed rock mixed together apatite grains that were hosted in hydrous and anhydrous minerals. This is considered by some to be a problem as halogens in apatite are thought to diffuse and partially equilibrate with those in hydrous mineral hosts (e.g., Brenan, 1993b). The diffusion of halogens between associated biotite and apatite, however, should not markedly affect magmatic halogen concentrations since both minerals are assumed, here, to equilibrate with similar magmatic fluids at similar thermodynamic conditions (see Section 7.4.3). Furthermore, a comparison of data collected from in situ apatite grains in polished thin sections also showed there are no obvious variations in composition between grains hosted in hydrous or anhydrous minerals (Chapter 6). Between 5–20, whole and unfractured, inclusion-free, apatite grains from each sample were selected using a binocular microscope. The grains were mounted in epoxy with the long (c-) axis of each grain parallel to the surface of the mount so that it would be perpendicular to the electron beam during analysis, in order to minimize fluorine diffusion. The epoxy mounts were polished with a 1 μm diamond paste until the approximate centres of the apatite grains were exposed, and each grain was imaged in back-scattered electron (BSE) mode on the scanning electron microscope (SEM). The apatite grains were then analyzed on a fully automated CAMECA SX–50 electron microprobe (EMP), as described in Chapter 6 (Section 6.3.1). Detection limits for fluorine, chlorine, and sulphur were 0.09, 0.06, and 0.05 wt.% oxide, respectively detection limits for all other elements are listed in Table 7.2). Due to the orientation of the apatite grains mounted in epoxy, the fluorine concentration increased minimally, if at all, with time during EPM analysis. For samples displaying a fluorine increase with time, ‘initial’ fluorine concentrations were calculated using a linear regression (Section 6.3.1, Chapter 6). Due to the minimal increase in fluorine concentrations with time in most cases, however, it seems that an average of all measurements (as opposed to a linear regression) would have been equally appropriate. Isolated analyses with inexplicably high or low fluorine concentrations, or with inconsistent linear regressions (tentatively interpreted to represent either restitic/xenocrystic apatite grains, or apatites that locally re-equilibrated with fluid(s) of unknown origin), were discarded. Atomic proportions and mole fractions of fluorine, chlorine, and hydroxyl (by stoichiometry) were normalized to 12 anions (excluding H2O, F, and Cl). Analyses with totals between 97.0–100.9 wt.% were plotted and interpreted, although analyses with totals of 97.0–98.0 wt.% were monitored closely and used if they were indistinguishable from analyses with higher totals. Apatite from intrusive phases with compositions consistent with true end-member fluorapatite compositions were not used for interpretation due to the impossibility of calculating HF/HCl and HF/HOH activity ratios from these grains. Averages and 1-sigma standard deviations of all apatite analyses used for interpretation are presented for each sample in Table 7.3, whereas the full datasets used for interpretations and figures are provided in Appendix F.3. 418 7.3.2.2 Biotite Magmatic biotite in polished thin sections for 22 samples was examined petrographically, first in transmitted light and then in BSE mode on the SEM. Only euhedral to subhedral, primary grains of magmatic origin (in accordance with the criteria set by Miller et al., 1981) were analyzed. Analytical methods and conditions are as described in Chapter 6 (Section 6.3.2). Detection limits for fluorine and chlorine, respectively, were 0.08 and 0.014 wt%. Atomic proportions were normalized to 11 anions (excluding H2O, F, and Cl) and only analyses with the total weight percent of the components between 98.0–100.99% were plotted and used for interpretation. Averages and calculated 1-sigma standard deviations of all of the analyses used for interpretation, with 1-sigma standard deviations, are presented for each sample in Table 7.4, whereas the datasets used for interpretations and figures are provided in Appendix F.4. The initial textural classification of biotite grains as primary and magmatic was consistent with their high TiO2, which is typical of igneous biotite (i.e., >2 wt.%; although some altered biotites do have high TiO2 contents), and their low molar magnesium versus total iron (i.e., XMg < 0.55; Table 7.4; e.g., Moore and Czamanske, 1973; Giudotti, 1976; Robert, 1976; Mason, 1978; Parry et al., 1978; Anderson and Rowley, 1981; Miller et al., 1981; Zen, 1988; Selby and Nesbitt, 2000; Nachit et al., 2005). 7.4 DISCUSSION 7.4.1 Magmatic Thermal Timescales: emplacement to sub-solidus conditions Calculated absolute, minimum, and maximum cooling periods in Table 7.1 are based on the difference between the U–Pb zircon age and the corrected 40Ar/39Ar biotite age. Over half of the intrusions for which U–Pb and 40Ar/39Ar ages are available record short absolute cooling periods. Either 40Ar/39Ar ages for the plutonic rocks are within the reported 2-sigma error for the U–Pb crystallization age, and (or) the U–Pb and 40Ar/39Ar ages have an absolute difference of less than 1 m.y. (Fig. 7.6). Several samples have had more prolonged thermal histories and record absolute cooling periods longer than 2 m.y., two of also have minimum cooling periods over 1 m.y. (Fig. 7.6). All of these intrusions are associated with tungsten skarns, including the two most economically important examples in the study area: • 2.2 (0.7–3.6) m.y. in the Cac pluton; • 2.2 (0.8–3.6) m.y. in the Lened pluton; • 2.9 (1.4–4.4) m.y. in the Coal River batholith near the CRAT tungsten skarn occurrence(s); and • 2.5 (1.7–3.3) m.y. in the Mine Stock at the Cantung tungsten skarn deposit. An additional four intrusions also appear to have been thermally active for unrealistically long periods of time: • 6.9 (5.8–8.1) m.y. in the Little Hyland pluton; • 10.1 (9.2–11.1) m.y. in the Nahanni Range Road pluton; • 25.3 (24.0–26.6) m.y. in the Coal River batholith near the ROYALL occurrence; and • 30.6 (29.3–32.0) m.y. in the Ivo pluton. 419 7.4.1.1 Magmatic Thermal Evolution of the Intrusions Most of the plutons in the study area experienced moderately rapidly cooling, from approximately emplacement temperatures of 830 ± 30°C (ranging from 760–913°C; ZSTs in Table 7.1) to sub-solidus temperatures marked by the closure temperature of argon exchange in biotite (~300°C). This suggests that most of the intrusions in the eastern Selwyn Basin were emplaced as either a single magma pulse, or as several closely spaced magma pulses. Regardless, relatively rapid cooling to sub-solidus temperatures typifies most of the intrusions in the eastern Selwyn Basin. Samples with ‘negative’ calculated absolute cooling periods (up to –1.8 m.y.) are still within 2-sigma error of the associated U–Pb age, with the exceptions of KR-05-126 and KR-05-148. The samples with the largest ‘negative’ cooling periods (>–1.5) have poor argon plateaus (e.g., KR-05-126) that are suggestive of argon loss or diffusion from the sample (Table 7.2; Appendix C.1). The biotites and zircons used for analysis came from the same hand samples, so negative values associated with good argon plateaus require either very fast cooling of the intrusion, below the resolution when comparing 40Ar/39Ar and LA–ICP–MS U–Pb ages (the latter has an average associated error of ~0.75%; Table 7.1). In other cases (e.g., KR-05-143) the negative values might indicate that the U–Pb ages are not as accurate as the 40Ar/39Ar ages, which have a higher precision and good accuracy when calibrated to the new Fish Canyon standard age (~0.25%; Kuiper et al., 2008). Two of the intrusions with long (>2 m.y.) minimum cooling periods are the Coal River batholith sample near―but not adjacent to―the CRAT tungsten skarn occurrences, and the apical phase of the Mine Stock (Fig. 7.6). The Coal River sample is unlikely to be associated with significant tungsten mineralization based on its composition. This strongly oxidized, barely peraluminous (ASI = 1.04), biotite–hornblende, monzogranite (Table 7.1; Appendix A.1), however, does fit some of the criteria for plutons associated with gold mineralization (e.g., Thompson et al., 1999; Lang and Baker, 2001).63 The extended thermal history in the apical Mine Stock at Cantung is not surprising given its proximity, and probable genetic relationship (Rasmussen et al., 2011b), to the one of the highest-grade tungsten skarns in the world (cf. Menzie and Jones, 1986). Mafic magmatism as young as ~91 Ma (Table 2.1; Chapter 2) also occurs at Cantung in the form of post-mineralization kersantitic lamprophyre dykes, and may extend the absolute duration of this magmatic system to ~7 m.y.. A 95.8 ± 0.6 Ma age for the mineralized biotite-skarn facies (Hart, 2000), thought to be the latest mineralizing event (Mathieson and Clark, 1984), indicates that the magmatic– hydrothermal system was closer to 2.5 ± 0.8 m.y. in length. Recent fluid inclusion studies combined with field observations suggest pressures of emplacement of circa. 2–2.5 kbar, which correspond to mineralization at depths of ~6.5–8 km (Marshall et al., 2003; Yuvan et al., 2007; Rasmussen et al., 2011b). Although the slightly higher pressures of emplacement may be consistent with the longer thermal history detected by the 40Ar/39Ar age data, this system was clearly not emplaced at the mid- to lower crustal levels where higher country rock temperatures would prolong magmatic cooling. These moderate emplacement depths require that the Mine Stock (or deeper equivalent) magma chamber was re-charged by hotter (mafic to intermediate) magma batches (Hart, 2004; Rasmussen et al., 63 This Coal River–CRAT (KR-05-196) sample is also unique because it contains titanite with niobium–fluorine– ytterbium enriched rims, although it is not known whether the titanite is of primary igneous, late-magmatic, or hydrothermal origin. As an aside, fluorine-rich (but not Nb- or Yb-rich) titanite occurs in felsic intrusions associated with tungsten skarns in the study area (e.g., the Mine Stock at Cantung, the Lened pluton, and the Coal River sample near the ROYALL occurrence; Appendix A.1). 420 2011b). The confirmation of ~98–91 Ma magmatic processes at Cantung, and a similar duration at Mactung―another world-class tungsten skarn deposit (Fig. 7.1; Gebru and Lentz, 2008)―suggests that prolonged thermal histories related to magma re-charging may be important with respect to formation of economically significant tungsten mineralization (Vigneresse, 2006; Rasmussen et al., 2011b). Only two other intrusions within the study area have similarly prolonged absolute thermal histories, and both are also associated with known tungsten mineralization. These are the Lened pluton, associated with the second largest tungsten resource within the study area (Figs. 7.1 and 7.2), and the Cac pluton, associated with several small tungsten skarn occurrences (Fig. 7.2; Allan, 1978). Overall, however, all four of the examined intrusions known to be associated with tungsten mineralization were identified as having prolonged cooling periods, or absolute cooling periods of more than 2 m.y.. 7.4.1.2 Reset 40Ar/39Ar Ages Anomalously long thermal activities (i.e., 7–31 m.y.) for four intrusions in the study area are not related to a long- lived magmatic–hydrothermal system, particularly as only one of these intrusions is associated with known mineralization (Ivo). All of these samples display good plateau ages lacking saddle-shaped spectra (Appendix C.1), which is interpreted to be consistent with the complete resetting of the 40Ar/39Ar isotopic system in these micas (Alexandrov et al., 2002). Thus, these ages record the timing of two distinct, but possibly local, Late Cretaceous thermal perturbations at 89–86 Ma and 67–65 Ma. The earlier event (89–86 Ma) has affected biotites from the Little Hyland and Nahanni Range stocks, southeast of the Cantung mine (Fig. 7.2). These two intrusions were emplaced near the inferred trace of the Hyland Fault, a dextral strike-slip structure that is known to have been active during the Early Cretaceous (e.g., Murphy et al., 1995; Gabrielse et al., 2006; Colpron et al., 2011). These younger biotite ages may have been reset by heating during Late Cretaceous movement along splays related to the Hyland fault, extending the known range of activity on this structure to ~85 Ma. This is evidenced by a strong foliation in the Little Hyland sample (KR-05-208; Appendices A.1 and A.3). The Nahanni Range Road sample is not deformed; instead, its young biotite age may be more consistent with a thermal perturbation related to nearby, undiscovered Late Cretaceous plutonism. The later event (ca. 67–65 Ma) affects biotites from the undeformed Ivo pluton and Marion plug (the latter is exposed in a streambed near the headwaters of the Coal River; Fig. 7.2). The 40Ar/39Ar data confirm similar K–Ar ages that were previously reported (Main and Cathro, 1982) for two other nearby intrusions associated with tungsten skarn occurrences in this area of Yukon (Fig. 7.2): • the Salivo pluton (60.3 ± 1 Ma), associated with the Salivo tungsten skarn showing and immediately adjacent to the Ivo pluton; and • the Cream stock (69.4 ± 1.2 Ma), associated with the ROYALL tungsten skarn showing and several small plugs near the Coal River batholith. Despite latest Cretaceous biotite ages, more recent U–Pb dating of samples from the Salivo and Cream intrusions 421 has confirmed that they are of earliest ‘mid’-Cretaceous age (ca. 98 Ma; Mortensen and Bezzola, 2010, unpub. data). It is notable that the sample of the Coal River batholith near the ROYALL tungsten showing did not return an anomalously young 40Ar/39Ar age, despite its proximity to the thermally disturbed Cream Stock; instead it is the Marion plug located to the west near the Coal River headwater that has an anomalously young 40Ar/39Ar age. Near the Marion plug there is a tin-bearing base metal vein (containing frankeite and geocronite; Evans, 1957), which is one of the few mineral occurrences in the northern Cordillera associated with sulphosalt minerals (T. Liverton, 2012, pers. comm.; Yukon MINFILE, 2003). This plug is an oxidized biotite monzogranite that does not appear to contain any unusual accessory minerals (Appendix A.1); therefore, the tin–antimony–sulphosalt mineral assemblage might imply a source in magmas that are compositionally distinct from most plutonic rocks in the eastern Selwyn Basin. As with the 86–89 Ma thermal resetting of intrusions near the inferred Hyland fault, there is also a structure that parallels the eastern edge of the Coal River valley (the Skonseng fault; Fig. 7.2; Pigage et al., 2011). Immediately west of this structure is a narrow corridor of higher regional metamorphic grade rocks (biotite–garnet–staurolite mineral assemblage; indicated by diagonal stippling; Fig. 7.2). This area of higher metamorphic grade has been mapped near the south end of this fault in Figure 7.2 (Pigage et al., 2011), but it is not shown on earlier geologic maps at the northern end of the Skonseng Fault (Gabrielse et al., 1973b); thus, the extent of this belt of higher metamorphic grade is uncertain. There is also a north-trending, 40-km long and ~1-km deep, fault-bounded basin of Late Cretaceous to Eocene age located south of the Ivo pluton in the Rock River drainage (Fig. 7.2; Yorath, 1991). The timing of movement along these structures is entirely unconstrained, except perhaps by the plutonic 40Ar/39Ar ages presented here. However, in the absence of textures indicative of deformation in the plutons with latest Cretaceous biotite ages, it seems more likely that thermal perturbations related to latest Late Cretaceous magmatism must be responsible for the resetting of plutonic mica ages in this region. The biotites from the four samples with reset 40Ar/39Ar ages have TiO2 concentrations (2–4 wt%) and Mg/(Fe+Mg) values (0.20–0.47) that are also consistent with a thermal disruption, rather than alteration by hydrothermal or metamorphic fluids. Other resetting of K–Ar and Rb–Sr isotopic ages (by Eocene magmatic activity) have occurred in many Early to Late Cretaceous intrusions throughout the northern Cordillera (Anderson, 1988). Therefore, a model of conductive re-heating in the absence of externally derived fluid—which can only be accomplished if ~67–65 Ma magma bodies were emplaced in close proximity to the ~98 Ma plutons—is the most realistic explanation for the re-setting for the 40Ar/39Ar ages in undeformed ‘mid’-Cretaceous plutons. This inferred, latest Cretaceous magmatic event in the eastern Selwyn Basin must be buried as the only known coeval granitoids in the Selwyn Basin are located several 100’s of km to the northwest (the 67–64 Ma McQuesten plutonic suite, and the 63 Ma Rackla pluton; Fig. 7.1; Murphy, 1997; Thiessen et al., 2012). It is uncertain whether or not buried examples of the latest Cretaceous magmatic event are also responsible for base metal mineralization near the Ivo, Marion, and Cream plutons. Field evidence suggests that at least the IVO tungsten skarn is directly related to the ~96 Ma Ivo pluton, and the ROYALL tungsten skarn is related to the ~98 Ma Cream, Jox, and (or) Bean stocks. The stanniferous vein at the MARION showing is also thought to be related to the ~97 Ma nearby plug exposed near the Coal River headwaters. 422 7.4.2 Mineral Compositions as Monitors of Magmatic Petrogenesis and Metallogenesis When examining the composition of apatite and biotite, one would expect variations in some of the components to be related to not only the degree of differentiation of the magma sampled, but also the composition of the melt source. Variations in halogen contents in apatite and biotite may be particularly significant with respect to associated mineralization, as some studies have demonstrated (e.g., Lang, 1998; Piccoli and Candela, 2002). Therefore, several compositional variables in both minerals are discussed in terms of their petrogenetic and metallogenic significance. Molar fractions for the main components of the hydroxyl sites of apatite and biotite display several trends relative to D.I. (Fig. 7.3a). With the exception of the least differentiated samples (D.I. < 65) and the two most differentiated samples (‘evolved’ members of the Tay River; Chapters 2 and 4), XF increases and XCl decreases with differentiation in both apatite and biotite (Fig. 7.3a). Also with the exception of the least and most differentiated samples, stoichiometrically determined XOH in apatite also decreases with D.I., but remains static in biotite (Fig. 7.3a). The molar ratios XCl/XF and XOH/XF display similar trends in both apatite and biotite, with XCl/XF broadly decreasing and XOH/XF clearly decreasing relative to D.I. (Fig. 7.3b). These elemental and molar fraction trends in apatite and biotite might in some, or all cases actually reflect changing fluorine, chlorine, or H2O contents in the magma and (or) magmatic fluid as the magma becomes increasingly differentiated. Or, these trends could instead be controlled by factors that affect mineral–fluid partitioning, and that may or may not relate to the magma and (or) magmatic fluid composition (e.g., changes to temperature or pressure conditions, etc.). 7.4.2.1 Apatite Compositions Low MnO (<0.2 wt.%) and FeO (<0.15 wt.%) in apatite characterize most of the Tombstone, Mayo, and Tay River plutonic suites (with the exception of several ‘evolved’ members of the Tay River suite; Fig. 7.4a,b). This is consistent with previous interpretations (Chapters 2, 4, and 5) that these three plutonic suites were derived from an infracrustal melt source. Higher MnO and FeO in apatites from the Hess–MacMillan sub-suite (circled; Fig. 7.4a,b) also consistent with the supracrustal origin interpreted for this sub-suite (Chapter 4). Apatites with the highest MnO and FeO, however, are from samples of ‘evolved’ members of the Tay River suite and the plutonic Tungsten suite (especially in the southeast; Fig. 7.4a,b). Both are interpreted to be either small partial melt fractions or highly evolved plutons derived from the infracrustal rocks inferred to be the source of high-K calc-alkaline magmas belonging to the Tay River, Mayo, and Tombstone plutonic suites (Chapters 2, 4, and 5). MnO and FeO contents in apatite from other highly fractionated examples of I-type suites with similar D.I. (>80), however, are typically at, or lower, than the EMP detection limits for these components (Fig. 7.4b). If even highly fractionated I-type plutons are not enriched in manganese and iron in the study area, then apatite compositions from the Tungsten (“SE”) and ‘evolved’ Tay River suite plutons might require a different explanation. The Coal River sub-suite of the Tay River suite has relatively elevated FeO ± MnO for its low D.I. (Fig. 7.4b,d), but these comparatively sodic plutons have been interpreted to have incorporated a significant proportion of magma from mafic rocks (Chapters 2 and 4). Apatite from the majority of samples in Figure 7.4c have high fluorine contents (typically >3 wt.%) that approach, 423 and in some cases exceed, the stoichiometric limit of 3.77 wt.% for fluorine in the hydroxyl site of apatite (indicated by the dotted line; Fig. 7.4c). Chlorine is negatively correlated with fluorine in apatite (Fig. 7.4c). Apatite from the least differentiated members of the Tay River plutonic suite (including the Coal River sub-suite plutons) and the Mayo plutonic suite have the highest chlorine and the lowest fluorine concentrations (Fig. 7.4c). The Tungsten plutonic suite has some of the lowest chlorine and highest fluorine concentrations (Fig. 7.4c), although chlorine is present in slightly higher abundances in apatites from several samples in the southeast Selwyn Basin (Fig. 7.4c). The Tombstone suite plutons also have low chlorine and high fluorine concentrations, particularly the samples from the southeast Selwyn Basin (Fig. 7.4c). Overall, the high fluorine concentrations for the Tombstone suite (regardless of their degree of differentiation) may be more consistent with an anhydrous melt source or one that was otherwise deficient in chlorine (e.g., Skjerlie and Johnston, 1992, 1993; Sallet, 2000). However, apatites in the Tombstone suite could also have high fluorine concentrations if they contain abundant CO2. Fluorine content is known to increase beyond what can be accommodated in the hydroxyl site of apatite if the apatites contain more than ~3 wt.% CO2 (McArthur, 1990)64. True end-member fluorapatite has 100*F/P2O5 ≈ 8.8 (McArthur, 1990); therefore, values greater than nine could be the result of coupled F– and CO32– substitution into apatite. Most of the highly differentiated samples are characterized by high fluorine relative to phosphorous (e.g., D.I. > 85 and 100*F/P2O5 > 9; Table 7.3; Fig. 7.d), however, three less-evolved Tombstone suite samples (and two Tungsten suite samples) have anomalously high fluorine relative to phosphorous, despite their lower degree of differentiation (e.g., D.I. < ~85 and 100*F/P2O5 > 9; Table 7.3; outlined with red box in Fig. 7.d). If the elevated fluorine in the Tombstone suite is not an artifact of analytical issues, then apatites from these plutonic suites could contain up to ~3 wt.% CO2. It is not possible to ascertain with these data whether or not apatites from any of the plutonic suites are CO2- bearing―particularly for samples with lower fluorine overall, but still high fluorine relative to phosphorus (100*F/P2O5 ≈ 5.5 corresponds to zero wt.% CO2; McArthur, 1990). However, if CO2 is present in apatite from plutons in the eastern Selwyn Basin, particularly members of the Tombstone suite, then these magmas may have an increased potential for related gold mineralization (Lang and Baker, 2001). Sulfur content in apatite decreases with increasing fluorine (although SO3 and F are not as closely correlated as Cl and F; Fig. 7.4c). As fluorine in apatite generally increases with D.I. (Fig. 7.3a), decreasing SO3 in the intrusions in the Selwyn Basin is probably related to mineral fractionation. SO3 is below the detection limit of 0.05 wt.% in all samples, except for the Coal River sub-suite and one member each of the Mayo and Tay River plutonic suites (Fig. 7.4c). Even when present in detectable amounts in samples containing a mineral assemblage of quartz–magnetite– titanite that indicates a high magmatic oxygen fugacity (e.g., KR-05-43, -191; Appendix A.1)—and is, therefore, consistent with a strong enrichment in SO42– relative to S2– in the magma (Carroll and Rutherford, 1988)―SO3 is still low when compared to similarly oxidized intrusions elsewhere (e.g., Yerington, Nevada; Streck and Dilles, 1998). Thus, the low SO3 in apatite strongly suggests that the magmas in the eastern Selwyn Basin were derived from a particularly sulfur-poor source, and were not significantly contaminated by sulfur-rich sedimentary rocks during ascension and emplacement. Low sulfur contents in apatite and the overall magma are also consistent with predominantly low whole rock sulfur contents determined for the same sample set (<250 ppm in most felsic to 64 It is not known where this additional fluorine resides in the crystallographic structure (McArthur, 1990), but apatites with over 3.77 wt.% fluorine must have accommodated the excess fluorine outside of the hydroxyl site. 424 intermediate intrusions; Chapter 4). In addition to providing details on magma sources (or magma types), compositional data for apatite also reveal information on the degree of differentiation experienced by the magma. The same apatite compositions that are typical of S-type magmas are also indicative of highly differentiated or fractionated magmas. In highly fractionated rocks, therefore, it is almost impossible to determine which features of S-type magmas (elevated Mn, Fe, F, and depleted S) are related to the composition of the melt source, and which trends are related to differentiation of the magma. Just as MnO and FeO in apatite broadly increase with D.I. in Figure 7.4b, CaO generally decreases with differentiation (Fig. 7.4a,d). Similarly, differentiation also increases F/Cl in magmas by the fractionation of anhydrous minerals and periodic saturation and off-gassing of chlorine-rich fluid (Webster, 1997a,b; Sha and Chappell, 1999; Sallet, 2000). Although this increase in magmatic F/Cl is not always reflected in the apatite composition (e.g., Hoskin et al., 2000), most studies of apatite composition have shown that F/Cl increases in apatite with magmatic differentiation (e.g., Tsusue et al., 1981; Coulson et al., 2001; Rasmussen et al., 2011a). Overall, apatite compositions in Figure 7.4 for intrusions in the eastern Selwyn Basin indicate that the Tay River (including Coal River samples) and Mayo plutonic suites are the least differentiated I-type plutons, and the Tungsten and ‘evolved’ Tay River suite samples are the most differentiated plutonic suites. The Tombstone suite, and especially the O’Grady batholith (KR-05-97b), which has a lower DI, contain apatites with consistently high F/Cl regardless of their relative degree of differentiation. This high F/Cl is probably related to the composition of the melt source for the Tombstone suite. The Hess–MacMillan sub-suite, which is the only supracrustally derived plutonic suite in the area, has unusually high MnO and FeO, low CaO, and high F/Cl, despite a relatively low D.I. (~71 for sample KR- 05-68; Fig. 7.4; Table 7.1). 7.4.2.2 Biotite Compositions Biotite compositions in the eastern Selwyn Basin plutons may also be evaluated as potential petrogenetic indicators. Biotites in the study area range from iron-poor annitic biotite (Fe# ≈ 0.4) to iron-rich siderophyllitic biotite (Fe# ≈ 0.8; using the International Mineralogical Association classification of Rieder et al., 1999), forming a discontinuous positively sloping trend that is typical of calc-alkaline rocks (Fig. 7.5a; Mason, 1978; Frost et al., 2001). Crystal– liquid fractionation in the magma is responsible for the increase in Fe# with D.I. in Figure 7.5 (Clarke, 1981); however, there is a notable break in the range of biotite compositions exhibited by the samples examined at Fe# ≈ 0.5 and D.I. ≈ 75–80 (Fig. 7.5b). All of the Tungsten suite intrusions, Tombstone suite plutons in the southeast Selwyn Basin, and ‘evolved’ members of the Tay River plutonic suite lie on the higher Fe# side of this break. The Tombstone suite plutons could be more ferroan because they are interpreted to be closely associated with mantle- derived A-type (or “ferroan”; Frost and Frost, 2011) magmas, especially as some of the high-K calc-alkaline (I-type) intrusive phases appear to have been modified by mixing with alkaline magmas (Chapter 2; Frost and Frost, 2011). In the calc-alkaline Tungsten and ‘evolved’ Tay River suite plutons, however, the increase in iron content is probably associated with differentiation (cf. Frost et al., 2001). The marked change from Tay River and Mayo suite samples with Fe# less than 0.4, to Tungsten suite and ‘evolved’ Tay River suite samples with Fe# greater than 0.6 (for the same D.I.) might suggest that the latter are not more fractionated examples of the Tay River or Mayo 425 plutonic suites (proposed by Mortensen et al., 2000, and postulated in Chapter 2). This also supports a similar interpretation based on Eu/Eu* in zircon (Chapter 5). Total aluminum in biotite is also correlated to D.I. for the eastern Selwyn Basin plutons, although there are two apparent trends in Figure 7.5b. Lower total aluminum contents overall form a shallowly sloping negative correlation with D.I., whereas higher aluminum contents (most samples with D.I. > 77) have a steeply sloping positive correlation with D.I. The low aluminum trend is defined mainly by the I-type Tay River (excluding ‘evolved’ examples) and Mayo plutonic suites (Fig. 7.5b). The low aluminum in these samples overall is consistent with the inferred infracrustal melt source for these suites (cf. Chappell and White, 2001). By definition, a pluton with a D.I. of 100 comprises the mineral assemblage quartz–albite–K-feldspar (similar to a pure aplite), and because the Aluminum Saturation Index (ASI) of feldspar is 1.0 (and the ASI of quartz is zero), a pluton with a D.I. of 100 should have an ASI of 1.0. Therefore, a negative correlation between total aluminum and ASI (outlined with dashed line, Fig. 7.5b) reflects a decrease in the magmatic concentration of aluminum with increasing degree of differentiation. Conversely, the strong positive correlation between total aluminum in biotite and D.I. in the Tungsten and ‘evolved’ Tay River samples (Fig. 7.5b) demonstrates that these magmas become increasingly aluminous with differentiation. This may occur during contamination by host rocks, or by mineral fractionation involving mainly feldspars rather than micas and other aluminous minerals. Elevated aluminum in biotite is linked to the presence of other aluminous minerals: muscovite, garnet, cordierite, and particularly the Al2SiO5 polymorphs (Clarke, 1981). Most of the Tungsten and ‘evolved’ Tay River suite samples do contain minor amounts of muscovite, garnet, and (or) rare andalusite or sillimanite (Anderson, 1988; Gordey and Anderson, 1993; Chapter 2). The higher total aluminum in the Tungsten and ‘evolved’ Tay River biotites―compared to most of the Tay River and Mayo suite plutons―correlates well with elevated Fe# as well (Fig. 7.5a). Together, higher aluminum and iron contents in the Tungsten and ‘evolved’ Tay River magmas would imply that they were either extensively contaminated by pelitic host rocks, or that they were derived from a pelitic melt source. A high degree of contamination by, or a melt source in pelitic rocks for these samples would also be consistent with the apatite compositions (e.g., high Mn, Fe, F). This is interesting because geochemical and mineralogical evidence presented in Chapter 2 suggested that the Tungsten suite plutons (at least in the southeast Selwyn Basin, and probably in the northeast Selwyn Basin) and ‘evolved’ members of the Tay River suite, are highly fractionated (although possibly country rock contaminated) I-type magmas. Alternatively, these biotite and apatite compositions could also have resulted from small-fraction partial melting involving only biotite, resulting in a highly peraluminous, iron-rich magma (although an enrichment in Cl would also be expected in this scenario). As with apatites from intrusions with high D.I.’s, it appears impossible to determine how strongly biotite compositions are controlled by the melt source (as well as anatectic, contamination, assimilation, or magma mixing processes) as opposed to a high degree of magmatic differentiation through contamination, assimilation, fractional crystallization processes. 426 7.4.2.3 Variations in Fluorine and Chlorine ± SO3 and Associated Metallogeny The partitioning of fluorine and chlorine into biotite is controlled by numerous factors (i.e., a preference for chlorine vs. fluorine, as well as both internal cationic controls and external thermodynamic controls). As a result, biotite is likely a poor predictor of halogen contents in a magma or MVP. Although apatite compositions may be better at predicting the halogen contents of a magma or an MVP, its use is also limited. This is partly because apatite is always enriched in fluorine, and partly because fluorine content in apatite also varies with thermodynamic conditions (Zhu and Sverjensky, 1991). Unlike biotite, however, strong internal compositional controls on the partitioning of halogens into apatite are nonexistent or have little effect on what appears to be ideal F–Cl–OH mixing in the hydroxyl site of apatite (Tacker and Stormer, 1989; Zhu and Sverjensky, 1992; Lang, 1998). Concentrations of fluorine, chlorine, and hydroxyl (and sulphur) in apatite may, therefore, indirectly reflect the relative abundances of these species coexisting in melt or fluid in a thermodynamically stable environment when apatite crystallizes (e.g., Piccoli and Candela, 1994; Parat and Holtz, 2005). Despite the limited utility of directly using apatite and biotite compositions to describe magma or magmatic fluid compositions, a brief examination of halogen and sulfur contents in both minerals is warranted because of the many earlier studies that attempted to link both mineral compositions directly to the presence (or lack) of intrusion-related mineralization (Kesler et al., 1975; Mason, 1978; Parry et al., 1978; Munoz, 1984; Lang, 1998; Piccoli and Candela, 2002). When measured concentrations of chlorine and SO3, and calculated initial fluorine concentrations in apatite are plotted with the D.I., it is apparent that many of the intrusions with a D.I. greater than 70 have near end-member fluorapatite compositions (e.g., F = 3.77 wt.%; Fig. 7.7a). Chlorine and SO3 are typically at, or below their respective detection limits of 0.16 wt.% and 0.05 wt.% in these moderately to highly differentiated intrusions (Table 7.3; Fig. 7.7a). Elevated chlorine (up to ~1 wt.%) ± SO3 contents are more typical of the least evolved intrusions (i.e., D.I. < 65), but also occur in a small number of the more evolved Tay River, Mayo, and Tungsten suite intrusions that are also associated with mineralization (Fig. 7.7a). These plutons are associated with silver–lead–zinc or copper mineralization (Mt. Christie, North Nahanni, Roy, and Marion), and four of the seven plutons associated with tungsten mineralization are also distinguished by elevated chlorine (the Coal River–CRAT, South Nahanni, and Lened intrusions as well as the apical Mine Stock intrusion at Cantung). Although the elevated chlorine in apatite from these samples could be linked to an infracrustal melt source, it is interesting that these apatites are from plutons that are associated with mineralization. Many of the samples with associated mineralization and elevated chlorine have correspondingly lower fluorine, even samples with higher D.I. where higher F/Cl would be expected instead (Fig. 7.7a). Thus, fluorine and chlorine contents in apatites from intrusive phases spatially associated with mineralization might be controlled by magma or MVP compositions, that in turn might be linked to nearby mineralization. Several plutons that lack anomalous chlorine concentrations in apatite, but have elevated fluorine concentrations (>3.77 wt.%) are also associated with mineralization (Fig. 7.7a). These include tungsten skarns adjacent to the Cac, Ivo, and Coal River–ROYALL plutons, and an antimony skarn associated with the highly metallogenic O’Grady batholith (Table 6.1; Figs. 7.2 and 7.7a). Elevated F/Cl in apatites from the intrusive phases associated with tungsten 427 skarns are probably related to the relatively high degree of magmatic differentiation experienced by these magmas, particularly the Ivo sample. It is uncertain, however, that the intrusive phase sampled is genetically related to the nearby mineralization, and high F/Cl in these samples may be no more anomalous than high F/Cl in highly fractionated samples (i.e., D.I. > 80) that are not associated with nearby mineralization. The elevated F/Cl in the Tombstone suite intrusions (O’Grady and the Coal River–ROYALL samples) was attributed to an enrichment in CO2 for this plutonic suite (Section 7.4.2.1). Elevated fluorine and CO2 in apatite appears to be a feature of plutons that have potential for associated gold mineralization, particularly when combined with favourable whole rock compositions (e.g., Lang and Baker, 2001). As with apatite, measured concentrations of chlorine and fluorine in biotite were plotted against D.I. to highlight trends exhibited by the halogens (Fig. 7.7b). With few exceptions, fluorine contents clearly increase in biotite with D.I., although it may decrease in the most differentiated samples (two ‘evolved’ members of the Tay River suite). Chlorine concentrations are much more variable and do not display any overall trends related to D.I., possibly due to the ease with which chlorine may be partitioned into a fluid phase and lost to the magmatic system at any stage of magmatic evolution. Biotite from samples of the Tungsten and ‘evolved’ Tay River plutonic suites (as well as the Hess–MacMillan sub-suite) are typically less enriched in chlorine (e.g., <0.15 wt.%) than samples from Tombstone suite and Tay River suite plutons (Fig. 7.7b). The Mayo plutonic suite has both chlorine-poor and chlorine-rich biotite and varies on a sample–by–sample basis (Fig. 7.7b). Unlike the apatite compositions, however, fluorine and chlorine concentrations in biotite are not distinct in plutons associated with mineralization. This is probably because the partitioning of halogens into biotite is more complex than into apatite. 7.4.3 Regional Characterization of Magmatic Fluid Compositions Estimates of MVP compositions using thermodynamic schemes based on apatite and biotite compositions can be used to determine whether an intrusive phase has produced and exsolved a saline aqueous fluid (possibly metal- bearing), depleting the magma and possibly subsequent fluids in HCl and H2O (e.g., Candela, 1986; Sheets et al., 1996; Lang, 1998; Selby and Nesbitt, 2000; Coulson et al., 2001; Piccoli and Candela, 2002). However, it seems that in mineralizing systems where the metals behave incompatibly, an enrichment in HCl/HF and H2O/HF may be important if the fluid was not exsolved or lost to the magmatic system until the latest stages of crystallization (e.g., Piccoli and Candela, 2002). This was demonstrated in Chapter 6 for intrusions associated with reduced tungsten skarn mineralization. In order to estimate the composition of a magmatic fluid with respect to fluorine, chlorine, and hydroxyl from apatite and biotite compositions, the internal and external controls on halogen partitioning must be accounted for. Several thermodynamic schemes based on the experimental data and examples of natural systems have been developed to approximate MVP compositions, which are typically expressed as fugacity ratios (e.g., Munoz and Swenson, 1981; Korzhinsky, 1981; Munoz, 1984, 1992; Speer, 1984; Candela, 1986; Piccoli and Candela, 1994; Piccoli et al., 1999; Aksyuk, 2000; Sallet, 2000). These fugacity ratios may then be used to evaluate the evolution of fluorine, chlorine, and hydroxyl species in fluids from mineralizing magmatic–hydrothermal systems (e.g., Loferski 428 and Ayuso, 1995; Selby and Nesbit, 2000; Yang and Lentz, 2005; Idrus et al., 2007). The use of fugacity ratios to represent the relative abundances of HF, HCl, and H2O species in an MVP, however, is not fully applicable to systems involving solutions (Zhu and Sverjensky, 1992). Instead, the use of fluid activities may be better for estimating MVP compositions for many mineralizing systems, especially when the exsolved MVP is predicted or known to comprise a single-phase aqueous fluid. Furthermore, most thermodynamic schemes consider the partitioning of F–Cl–OH between minerals and fluid, which is strongly affected by external factors (e.g., pressure, temperature, pH, and of course fluid composition; Piccoli and Candela, 1984; Zhu and Sverjensky, 1991, 1992; Brenan, 1993a). To eliminate or minimize the external thermodynamic controls on halogen partitioning into minerals, a thermodynamic treatment that takes advantage of the solid solution between end-member mineral compositions in natural systems was developed by Zhu and Sverjensky (1992), known as the reciprocal solid solution model (RSSM). The RSSM uses apatite–biotite mineral pairs―chosen for the quantity of pre-existing experimental and natural data available for these minerals―to approximate the composition of MVP that these minerals equilibrated with. This eliminates the effects of fluid composition and pressure, and significantly reduces the effect of temperature on the partitioning of halogens into these minerals, particularly biotite. The RSSM also employs one of the most comprehensive treatments concerning crystal chemical effects on the partitioning of F–Cl– OH into biotite. Although the partitioning of fluorine and chlorine in the RSSM is not significantly influenced by pressure, it is still somewhat dependent on temperature; therefore, the approximate temperature for the crystallization of apatite and biotite must be determined. Apatite saturation temperatures (ASTs) were calculated from whole rock geochemical compositions (Harrison and Watson, 1984; Bea et al., 1992) and are considered to approximate, within 30–100oC, the crystallization temperature for over 80% of the apatite in a magma (Piccoli and Candela, 1994). ASTs for the intrusions in the study area range from 732–943°C (Tavg = 835°C; Table 7.1). Temperature estimates for biotite are less accurate, but a range of 750–900°C (averaging ~800–850°C) for magmatic biotite in the study area is suggested65. An approximate temperature of formation of 850°C―and a simple pressure estimate of 2 kbar based on the regional presence of andalusite in contact aureole rocks―were assumed for calculations involving both apatite and biotite. This use of the same thermodynamic conditions for all of samples is justified by the negligible effect of pressure on partitioning between two solid phases, and by minimal variation within calculated activity ratios associated with temperature deviations (e.g., Zhu and Sverjensky, 1992; Piccoli and Candela, 1994; Kifer and Wolf, 2000; Coulson et al., 2001; Chapter 6). Furthermore, apatite and biotite typically crystallize close in time and space in felsic magmas, as is evident in their common petrographic association. If apatite saturated at higher temperatures than biotite―which textural relationships generally suggest (e.g., euhedral apatites are included in all major rock- forming minerals, although apatites are much larger and more abundant in association with hydrous minerals)―it will provide an estimate of an earlier MVP composition (Piccoli and Candela, 2002; Yang et al., 2004; Chapter 6). This earlier MVP may or may not be substantially different from a later estimate of MVP composition determined from biotite; however, significant variations between apatite and biotite estimates of MVP composition will indicate 65 This estimate is based on the experimentally and theoretically determined QFM buffer equilibria curves of Wones and Eugster (1965), using the quartz–magnetite–fayalite to graphite–methane buffer conditions determined for the Mine Stock (Zaw and Clark, 1978; van Middelaar, 1988) and the Fe / [Fe + Mg] compositional of all of the biotites in the study area. 429 a change in magmatic fluid conditions occurring after apatite has fully crystallized, but before biotite has crystallized and closed to halogen exchange. The use of the same temperature for calculations involving both minerals reduces, but does not entirely remove, differences in calculated magmatic fluid compositions related to equilibration with different fluids at different times. Calculated fluid activities for HCl, HF, and H2O species determined from the RSSM are only relevant as ratios; therefore, log[aHCl/aHF] and log[aHOH/aHF] are used to estimate relative HF, HCl, and H2O abundances for the fluids with which apatite and biotite equilibrated. Although log[aHCl/aHF]AP,BT and log[aHOH/aHF]AP do not rely on any assumptions involving the activity of H2O, log[aHOH/aHF]BT was calculated assuming a pure, or dilute aqueous fluid at unity (aHOH = 1). Averages of the calculated log[aHCl/aHF] and log[aHOH/aHF] values and respective 1-sigma standard deviations for both minerals in each sample are tabulated in Table 7.2. All of the calculated activity ratios, however, are plotted in Figure 7.8, and the complete datasets are tabulated in Appendices F.3 and F.4. Activity ratios for end-member fluorapatite, typically in samples with a D.I. greater than 85, could not be calculated in cases because of the absence of chlorine in these apatites, and the lack of stoichiometric space for hydroxyl in the halogen site. Primary magmatic biotite provides the only estimates of magmatic fluid composition for samples containing end-member fluorapatite, however, these data must be interpreted with caution because biotite compositions are easily modified by sub-solidus fluid (e.g., Dilles, 1987; Piccoli and Candela, 1994; Sallet, 2000). The independent apatite- and biotite-based estimates of MVP composition are used below to evaluate the intrusions from a metallogenic perspective. Several trends defined by the activity ratios in Figure 7.8 (upper) are apparent. Initially log[aHCl/aHF] increases with D.I.; however, for samples with D.I. greater than 65, log[aHCl/aHF] steadily decreases with increasing D.I. from positive values indicative of a chlorine-rich MVP to negative values indicative of a fluorine-rich MVP (upper Fig. 7.8). These trends were also illustrated by molar ratios (and fractions) for samples with D.I. < 65 in Figure 7.3, and a similar gradual decrease in HCl/HF with increasing differentiation was also observed by Coulson et al. (2001) in the multi-phase Emerald Lake pluton, where it was attributed to gradual off-gassing of chlorine-rich MVP over time. A decrease in HCl/HF may be also compounded by fractional crystallization of anhydrous minerals and the tendency for fluorine to partition into the melt rather than into a MVP (e.g., Sallet, 2000). Apatite- and biotite-based estimates of log[aHCl/aHF] for most of the samples in Figure 7.8 (upper) overlap closely, and few samples have significantly different estimates of fluid composition (see Section 7.4.3.2). Calculated log[aHOH/aHF] in Figure 7.8 (middle) also initially increase for samples with a D.I. of less than 65, then progressively decrease as differentiation progresses. Again, this decrease in the H2O/HF activity ratio is very similar to the trends exhibited by molar fractions and ratios in Figure 7.3. Despite the decrease in H2O relative to fluorine, log[aHOH/aHF] remains positive in nearly all cases, which is consistent with H2O > HF in the MVP. Clearly magmatic fluid compositions vary systematically with the degree of differentiation exhibited by the plutons in the eastern Selwyn Basin, although they appear to be fairly independent of the plutonic suite (with the exception of the Tombstone suite, which arguably has lower H2O/HF due to a petrogenetically related fluorine enrichment). Unlike the generally good overlap between apatite- and biotite-based estimates of log[aHCl/aHF], however, log[aHF/aHOH] 430 determined from biotite are consistently higher than those determined from apatite for all but one of the intrusions that have both apatite- and biotite-based estimates of MVP composition (middle Fig. 7.8). Although the late loss of an HF-rich fluid, or overprinting by an H2O-rich fluid may be inferred from the biotite compositions, it is extremely unlikely that all of these samples were affected to a similar degree by late-stage H2O-rich. Nor can the regular H2O/HF variation in Figure 7.8 (middle) be the result of improper pressure or temperature estimates used for the biotite-based calculations; otherwise, HCl/HF from apatite and biotite would not produce such similar estimates of MVP composition for most of the samples in Figure 7.8 (upper). The only remaining possibility is that the assumption of a pure fluid (i.e., aHOH in biotite is at unity) required for the biotite-based log[aHF/aHOH] calculations was incorrect. If it is assumed that log[aHOH/aHF] for apatite and biotite should be roughly equivalent (i.e., both minerals are closely associated in time and space in most samples, and neither mineral has re-equilibrated with later fluids), then the linear relationship exhibited between log[aHOH/aHF]AP and log[aHOH/aHF]BT in Figure 7.9a (inset) allows the biotite-based estimates of fluid composition to be corrected for the initial assumption of a pure fluid. Log[aHOH/aHF]BT was, thus, re-calculated using the equation: log[aHOH/aHF]BT–corrected = (log[aHOH/aHF]BT + 1.2484) / 0.4533 (Table 7.4). The re-calculated log[aHOH/aHF]BT are plotted in Figure 7.8 (lower). For the majority of samples the new estimates of log[aHOH/aHF] in Figure 7.8 (lower) appear to be an improvement on the original, uncorrected log[aHOH/aHF] data. Typically the corrected log[aHOH/aHF] values for biotite overlap more closely with those of apatite; however, there are also examples where the reverse is evident and where the biotite data now display more scatter (lower Fig. 7.8). Assuming that this correction is valid for most samples, however, it also allows for the estimation of the aHOH for the fluid that equilibrated with biotite. While not a pure fluid, aHOH is still generally less than 0.2, ranging from average values for each sample of 0.01 to 0.46 for all samples (except for the hydrothermally altered ‘plug’ southwest of O’Grady; Table 7.4; Fig. 7.9b). This suggests that, in most cases, the assumption of a pure fluid (i.e., aHOH = 1) is incorrect, although the amount of error that may be attributed to this assumption varies, but appears to increase with the degree of differentiation. Just as log[aHOH/aHF] decreases with differentiation, in Figure 7.9b it is apparent that aHOH increases exponentially with D.I. However, as with the calculated activities of HF and HCl, the value of aHOH has no real significance for the concentration of H2O in the magmatic fluid unless it is compared to HF or HCl in the form of an activity ratio. 7.4.3.1 Small Variations in Apatite- and Biotite-based Estimates of Fluid Composition Small variations in apatite and biotite estimates of HCl/HF and HF/H2O content in the magmatic fluid might be tied to slight differences in temperature (and possibly pressure) conditions specific to each pluton―this is particularly the case for biotite, which may have a lower closure temperature for halogen exchange with fluid or melt than apatite (van Middelaar and Keith, 1990; Chapter 6). Apatite- and biotite-based estimates of log[aHCl/aHF] in many samples, however, agree so closely that the effects of small variations in temperature and pressure conditions between samples across the study area―even between apatite and biotite themselves―must be almost negligible in studies of this kind. Larger variations in temperature and pressure conditions or estimates could be more problematic and will produce more variation in apatite- and biotite-based MVP estimates. For example, an earlier presentation of 431 these data calculated magmatic fluid compositions based on apatite and biotite compositions at temperatures of 850°C and 650°C, respectively. Larger variations in log[aHCl/aHF] were observed and discussed from a metallogenic perspective (Rasmussen et al., 2007, 2011a). The work presented here―in which the same temperature of equilibration is used for apatite and biotite―produced a better overlap in estimated activity ratios, which suggests that an estimate of 850°C is more valid for the biotite-based calculations, and that apatite and biotite do crystallize at similar thermodynamic conditions and, thus, in equilibrium with the same or a similar magmatic fluid. 7.4.3.2 Significant Variations in Apatite- and Biotite-based Estimates of Fluid Composition For intrusive samples in which activity ratios determined from apatite and biotite differ significantly, a change in the magmatic fluid composition—after apatite was fully crystallized and (or) closed to halogen exchange and before biotite was fully crystallized and (or) closed to halogen exchange—may have occurred. A notable decrease in HCl/HF from apatite- to biotite-based estimates of MVP composition might indicate that the magma exsolved a large quantity of fluid before biotite crystallization or equilibration was completed. This fluid would have been a saline aqueous (Cl–H2O-rich) fluid capable of transporting metals due to the common ability of chlorine species to form soluble complexes with many metals (e.g., Candela and Piccoli, 1995; Seward and Barnes, 1997). Therefore, the decrease in HCl/HF from apatite to biotite data for the Cac pluton, for example, might suggest a role for an exsolved saline aqueous fluid in the genesis of the associated tungsten skarns (upper Fig. 7.8). Although one would expect evidence for fluid exsolution to be indicative of a greater potential for associated mineralization, variations in apatite- and biotite-based estimates of fluid composition in Figure 7.8 do not characterize all of the intrusions known to be associated with mineralization. Nor are these variations limited only to intrusions known to be associated with mineralization. It was mentioned earlier that a major weakness of any reconnaissance-scale study is the limited numbers of samples actually analyzed from each composite intrusion. Any interpretations of these fluid compositional data from a metallogenic perspective must, therefore, be made cautiously. Another possible explanation for significant variations between the apatite- and biotite-based estimates of MVP composition is that biotite (or apatite; cf. Parry et al., 1978) might have re-equilibrated with a sub-solidus magmatic or hydrothermal fluid. A good example of probable re-equilibration is the pervasively altered ‘plug’ located southwest of the O’Grady batholith. Intense and pervasive sericitization of the feldspars in this sample (Appendix A.1) provide evidence for alteration by a late fluid. Similarly, HCl/HF estimates from biotite are much lower than those from apatite in this pluton (upper Fig. 7.8), which is consistent with either (a) the crystallization of magmatic biotite in the presence of an HF enriched magmatic fluid (MVP), or (b) the re-equilibration of biotite with a sub- solidus HF-rich hydrothermal fluid. The biotite compositions for this pluton are similar to other magmatic biotites in the study area with respect to TiO2 content and XMg (Table 7.4) and petrographic examination shows that the biotites are not altered (Appendix A.1). Therefore, the scenario of biotite equilibration with a late-stage HF enriched MVP is more applicable to this pluton. This is the only example of an intensely altered intrusion in this study that was investigated, although biotites from other, less altered, intrusions could also have exchanged with late-stage magmatic or hydrothermal fluids. Differences in apatite- and biotite-based estimates of magmatic fluid composition may be the only evidence for this process in the absence of petrographic evidence for biotite alteration. 432 There are several samples that plot above the general range of log[aHCl/aHF] and log[aHOH/aHF] compositions shown in Figure 7.8. These four samples are all enriched in chlorine and H2O relative to fluorine, and are all associated with mineralization: two tungsten skarns (CRAT and BLUE) and two argentiferous lead–zinc mantos (ROY and MARION). The biotites associated with the IVO tungsten occurrence are also enriched with respect to log[aHOH/aHF]; however, this may be due to re-setting of the 40Ar/39Ar ages for these magmatic biotites in the latest Cretaceous (Section 7.4.1.2). Apatites and biotites from the nine other intrusive phases associated with mineralization in the study area do not appear to have equilibrated with a magmatic fluid containing anomalously high HCl and H2O (Fig. 7.8). It seems that the magmatic fluid compositional component to this study, therefore, was significantly less successful at identifying intrusions that are prospective for mineralization in the eastern Selwyn Basin. One explanation for this discrepancy may be because anomalous magmatic fluid compositions will only be found in the specific intrusive phases that were directly associated with mineralization (despite the earlier statement that intrusive phases sampled “near” known mineralization were assumed to be related to the mineralization event), whereas anomalously long cooling periods may be detectable in any of the intrusive phases in a composite intrusion that is associated with mineralization. 7.4.4 Integrated Magmatic Thermal and Fluid Compositional Datasets Interestingly, intrusions with longer cooling periods are not necessarily anomalous with respect to calculated magmatic fluid compositions (Fig. 7.8). The poor correlation between the thermal and magmatic fluid composition datasets may be because the magmatic fluid composition interpretations assumed that ‘nearby’ intrusion-related mineralization was genetically related to the representative intrusive phase sampled. If the composite plutons that have been mapped (Gordey and Anderson, 1993) and observed during traverses (this study) are typical of the eastern Selwyn Basin, it is likely that most (and perhaps all) plutons and batholiths in this region are products of multiple intrusive phases. However, there are still some potentially significant observations that may be drawn from the integration of the thermal and magmatic fluid composition datasets in Figure 7.8. To give an example, magmatic fluids for the samples of the Coal River batholith from near the MARION and CRAT showings are anomalously enriched in HCl/HF ± H2O/HF overall compared to other samples with similar D.I. (Fig. 7.8). The extended thermal history indicated by 40Ar/39Ar ages for at least the CRAT sample (this cannot be determined for the MARION sample because of its reset 40Ar/39Ar age) is compatible with prolonged cooling. Taken together, these thermal and magmatic fluid compositional features indicate a higher potential for a long-lived magmatic–hydrothermal system that may have produced a saline aqueous fluid capable of extracting, transporting, and concentrating metals (e.g., Candela and Piccoli, 1995). Despite extended cooling periods, magmatic fluid compositions for representative phases from the two intrusions associated with the largest tungsten resources in the study area (Lened and Cantung) are not anomalous with respect to HCl/HF and H2O/HF. Although this may appear initially discouraging, at least at Cantung the intrusive phase(s) thought to have equilibrated with the main body of mineralizing fluids was not examined by this reconnaissance study (Rasmussen et al., 2011b). Chapter 6, however, demonstrates that there are intrusive phases at both Lened and Cantung that do display extreme enrichments in HCl ± H2O relative to HF. These are the intrusive phases that have 433 also been linked to mineralization on the basis of proximity (at Lened) and their degree of differentiation (at Cantung; Rasmussen et al., 2011b; Chapter 6). This study failed to distinguish intrusions associated with a world- class tungsten skarn deposit from other intrusions in the study area on the basis of a single sample from the pluton that underlies the orebody, which highlights a central weakness to all reconnaissance-scale studies. If the correct intrusive phase is not sampled, in other words, estimates of magmatic fluid composition may not provide much useful information. Although impractical, one method of increasing the likelihood of identifying intrusions with greater potential for associated mineralization would be to sample and analyze as many intrusive phases as possible. Another is to identify plutons with extended thermal histories, then follow-up with more detailed magmatic fluid compositional work (similar to that presented in Chapter 6) on the prospective bodies. 7.5 CONCLUSIONS The compositions of apatite and biotite from the plutons examined here are mostly consistent with the previously interpreted melt sources for the plutonic suites (Chapters 2, 4, and 5). Low aluminum, iron, and (or) manganese, and elevated chlorine in the least differentiated samples characterize apatite ± biotite for the majority of the I-type Tay River and Mayo suite samples. A distinction cannot be made on the basis of apatite and biotite composition alone, however, between plutons previously interpreted to be S-type granitoids (i.e., the Hess–MacMillan sub-suite) and plutons previously interpreted to be highly fractionated I-type granitoids (i.e., the Tungsten suite, and ‘evolved’ members of the Tay River suite). The Tombstone suite apatites are characterized by relatively high F/Cl for their given degree of differentiation, which is interpreted to reflect a petrogenesis involving CO2-rich magmas and (or) an anhydrous melt source, possibly in the form of alkaline mantle-derived magmas (cf. Hart et al., 2005). It is not clear whether source composition, magmatic fractionation, the presence of CO2, or even the effects of analytical difficulties (or some combination of these and other factors listed in Section 7.4.2.1) are the driving forces behind the high F/Cl that characterizes more differentiated members of the other plutonic suites. In some instances, however, relatively elevated F/Cl, and elevated chlorine ± SO3 contents occurs in apatites from samples associated with mineralization. Unlike apatite, attempts to use biotite compositions alone as a proxy for the composition of a magma and (or) the fluid that exsolved from the magma may have significant uncertainties and errors. This study did not identify any compositional variations or trends in magmatic biotite that may be used to predict an association with intrusion-related mineralization. Tungsten skarn mineralization is by far the most significant deposit type known in the eastern Selwyn Basin, but there is significant potential for intrusion-hosted gold mineralization deposits as well as more distal base metal mineralization associated with both tungsten and gold mineralization (e.g., Hart, 2005; Rasmussen et al., 2007). Although shallower emplacement levels (e.g., <1.3 kbar; Coulson et al., 2002) may not be conducive to the prolonged cooling histories (e.g., Hart, 2004; Hart et al., 2004b), any intrusion that experienced multiple episodes of magma recharge, regardless of its depth of emplacement, has the potential to generate significant mineralization. These intrusions may also have longer cooling periods, although shorter cooling periods do not necessarily preclude an association with intrusion-related mineralization (Hart, 2004). The data presented here show that longer cooling periods are predictive of associated mineralization in the case of reduced tungsten skarns (e.g., the >2 m.y. absolute 434 thermal histories for the Mine Stock, Lened, and Cac plutons, and possibly the Coal River batholith near the CRAT showing). The detection of extended cooling all four intrusions associated with tungsten skarns in the study area demonstrates this to be an efficient exploration method for tungsten. Magmatic fluids associated with intrusions in the study area appear to have typically evolved to lower HCl/HF and H2O/HF compositions as the plutons become more differentiated. This study failed, however, to identify anomalous HCl/HF and H2O/HF ratios―indicative of the evolution of a saline aqueous fluid capable of transporting most metals (e.g., Candela and Piccoli, 1995)―in the majority of the intrusions associated with known mineralization. This may be largely due to the limited number of samples analyzed from each pluton. For example, an HCl-rich fluid was not detected in the Lened pluton and Mine Stock (Cantung) in this study, whereas intrusive phases that did evolve potentially mineralizing fluids in these bodies were identified in the more detailed work presented in Chapter 6. Therefore, as a regional exploration strategy the evaluation of single samples from intrusions may not be effective in highlighting intrusions that have a greater potential for associated mineralization, even though more detailed studies could be helpful once an intrusion has been identified that is prospective for related mineralization (cf. Chapter 6). The poor correlation between the thermal and magmatic fluid composition datasets may be because the magmatic fluid composition interpretations assumed that ‘nearby’ intrusion-related mineralization was genetically related to the representative intrusive phase sampled, whereas longer thermal histories likely characterize the entire pluton and not just the intrusive phase that was sampled. Despite the limitations of the fluid compositional study, the integration of the thermal and fluid compositional datasets did produce a couple interesting prospects. At least two intrusive phases that experienced both prolonged cooling to sub-solidus temperatures, and evolved a HCl- and H2O-rich fluid that could have been capable of transporting metals (e.g., Candela and Piccoli, 1995), were identified. 435 Figure 7.1. Distribution of Early to ‘mid’-Cretaceous, syn- to post-accretionary plutonic suites in the northern Cordillera (AK = Alaska; YK = Yukon; NWT = Northwest Territories; BC = British Columbia). The study area (Fig. 7.2) is outlined in black. Plutonic suites discussed in the text are labeled: To = Tombstone, M = Mayo, Tu = Tungsten, TR = Tay River. The Tombstone, Mayo, and Tungsten plutonic suites for the Tombstone–Tungsten Belt (TTB) referred to in the text. Plutons associated with the three largest tungsten skarn deposits in or near the study area are labelled (Mactung, Cantung, and Lened); other plutons discussed in the text and are located outside the study area are labelled. The location of the latest Late Cretaceous–Paleocene McQuesten plutonic suite and Rackla pluton are indicated by bright pink polygons. Tectonostratigraphic units and map lines are as in Figure 1.1 (Chapter 1). 436 Figure 7.2. (next page) Distribution of syn- to post-accretionary, Early to ‘mid’-Cretaceous magmatism in the eastern Selwyn Basin study area. Sample locations are indicated by bright green circles, and plutons mentioned in the text are labeled. All of the mineral occurrences and deposits (labeled with capitalized and italicized text) shown are inferred to be intrusion-related on the basis of metal tenor, mineralization style, spatial associations, and (in some cases; bolded labels) based on the similarity between lead isotopic compositions of sulphides from occurrences and of feldspars from intrusions (Heffernan et al., 2004; Rasmussen et al., 2007). Mineral occurrence locations are from NORMIN (2013) and Yukon MINFILE (2013), and Gordey and Anderson (1993). The approximate southeast–northeast boundary between the Tungsten and Tombstone plutonic suites in the southeast Selwyn Basin is at a latitude of 62°N. The diagonally stippled area represents a region of higher metamorphic grade indicated by a mapped biotite- garnet-staurolite mineral assemblage (Pigage et al., 2011), and the northern half of the Late Cretaceous– Eocene Rock River fault-bounded basin is outlined by the bolded dashed line (labeled RRB). Tectonostratigraphic elements and map lines are as in Figure 1.1 (Chapter 1), and abbreviations are: Ft = fault, FS = Fault System, and H–M = Hess–MacMillan. 437 438 Figure 7.3. Logarithmic molar concentrations and ratios of fluorine, chlorine, and hydroxyl in the hydroxyl site of apatite and biotite, versus differentiation index (D.I.; after Thornton and Tuttle; 1960). Apatite data are indicated by grey and black symbols, and biotite data are indicated by red symbols with equivalent D.I. Anomalous samples discussed in the text are circled in red. Note that where XCl is <0.001, analyses are plotted at a value of 0.001 along the x-axis. a) Molar fluorine, chlorine, and H2O versus D.I.. b) Molar Cl/F and OH/F versus D.I.. 439 Figure 7.4. (next page) Apatite chemistry. Symbols are as in Figure 7.3. Differentiation trends are indicated with grey arrows, and granitoid type trends are indicated with black arrows. Detection limits are indicated by the dashed lines where applicable, and the maximum amount of fluorine that may be stoichiometrically accommodated in the hydroxyl site of apatite (3.77 wt.%) is represented by the dotted line. The S-type Hess–MacMillan sub-suite is circled in blue. a) CaO versus MnO and versus FeO (wt.%). b) The differentiation index (D.I.) of Thornton and Tuttle (1960) versus MnO and versus FeO. c) Fluorine versus chlorine and versus SO3 (wt.%). d) D.I. versus CaO (wt.%) and versus 100 * F / P2O5; anomalous samples are outlined with the red box. 440 441 Figure 7.5. (next page) Biotite chemistry. Symbols are as in Figure 7.3. Anomalous samples discussed in the text are circled in red, and differentiation trends are indicated with grey arrows. a) Fe/(Fe+Mg), or Fe#, versus total aluminum (apfu) biotite classification after Rieder et al. (1998) where biotite end-members are: Ann = annite, Sid = siderophyllite, Phl = phlogopite, and Eas = eastonite. b) The differentiation index (D.I.) of Thornton and Tuttle (1960) versus Fe/(Fe+Mg) (upper), total aluminum (middle), and the Aluminum Saturation Index (ASI; lower). 442 443 Figure 7.6. Summary of cooling periods. Symbols are as in Figure 7.3. Absolute cooling periods are indicated by the symbol for the plutonic suite, and minimum and maximum cooling periods are indicated by the lower and upper limits to the black bars, respectively. The dashed lines at 1 m.y. and 2 m.y. distinguish intrusions with anomalously long cooling histories (i.e., minimum > 1 m.y.; absolute >2 m.y.). 444 Figure 7.7. (next 2 pages) Fluorine, chlorine, (and SO3) concentrations versus differentiation index (D.I.; after Thornton and Tuttle; 1960). Symbols are as in Figure 7.3. Samples linked by bolded grey lines are associated with mineralization. a) Fluorine (upper half), chlorine (lower half), and SO3 (lower half, red symbols) concentrations in apatite. Detection limits are indicated by the dashed bolded black line (F = 0.09 wt.%), the dashed black line (Cl = 0.16 wt.%), and the dashed red line (SO3 = 0.05 wt.%); the maximum amount of fluorine that can be stoichiometrically accommodated into the hydroxyl site in apatite (3.77 wt.%) is indicated by the dotted line. Note the change in scale on the y-axis at 1.5 wt.%. b) Fluorine (upper half) and chlorine (lower half) concentrations in biotite. Detection limits for fluorine (0.08 wt.%) and chlorine (0.014 wt.%) are indicated by the dashed black lines. Note the two different scales for the y-axis. 445 446 447 Figure 7.8. (next page) Log[aHCl/aHF] and log[aHOH/aHF] estimates for the magmatic fluid composition determined from apatite and biotite (red symbols) compositions, versus differentiation index (D.I.; Thornton and Tuttle, 1960). Symbols are as in Figure 7.3. Samples linked by bolded grey lines are associated with mineralization, samples linked by bolded black lines have anomalously long cooling periods (as described in the text), and samples circled in red have anomalous HCl/HF and (or) H2O/HF contents. 448 449 Figure 7.9. Corrected aHOH. Symbols are as in Figure 7.3. a) The linear relationship between log[aHF/aHOH] for apatite and biotite that was used to correct log[aHF/aHOH]. b) Calculated aHOH versus the differentiation index (D.I.) of Thornton and Tuttle (1960). Sample KR-05- 126 has anomalously high aHOH values of 14–25 that plot off the chart (as indicated by the red arrow). 450 Table 7.1. Sample information and final U–Pb crystallization ages (summarized from Table 2.2, Chapter 2) for plutons in the eastern Selwyn Basin. Sample Plutonic Suite (sub- suite/region)a Plutonb General Composition ASIc D.I. (plotted)c Associated Mineral Showing ZSTd (°C) ASTd,e (°C) U–Pb Age (Ma)f 2-δ error (Ma)g KR-05-97b Tombstone (NE) O'Grady oxidized, qtz monzonite 0.86 70.9 (70) KAREN Sb skarn 790 943 92.8 0.6 KR-05-198 Tombstone (SE) Hole-in-the-Wall reduced, monzogranite 1.08 78.4 (78.5) 860 827 91.3 2.1 KR-05-194 Tombstone (SE) Coal River reduced, monzogranite 1.10 83.4 (83) ROYALL W skarn 870 852 92.8 0.9 KR-05-76 Mayo (Hess–MacMillan) Pelly River reduced qtz diorite 1.08 53.5 (55) 786 861 98.0 1.1 KR-05-62 Mayo Mount Christie reduced qtz monzonite 0.84 61.1 (61) 'U" Pb–Zn–Au vein 786 929 95.1 1.2 KR-05-130 Mayo Central Nahanni reduced granodiorite 1.13 64.2 (65) 813 842 97.1 0.8 KR-05-136 Mayo North Nahanni reduced monzogranite 1.14 69.2 (69) RISBY Cu skarn 827 794 96.9 0.7 KR-05-68 Mayo (Hess–MacMillan) Mount Wilson reduced monzogranite 1.14 71.3 (71) 851 794 97.7 0.8 KR-05-113 Mayo South Nahanni reduced monzogranite 1.05 78.2 (78) BLUE W skarn 831 908 nd nd KR-05-08 Tay River (Coal River) Jorgensen reduced qtz monzodiorite 1.18 61.9 (62) 781 781 95.9 0.8 KR-05-43 Tay River (Coal River) Park oxidized granodiorite 1.14 63.2 (63) 800 854 96.1 1.2 KR-05-10 Tay River (Coal River) Powers oxidized, qtz monzodiorite 1.17 63.6 (64) 781 836 98.9 0.7 KR-05-196 Tay River Coal River oxidized hbl-bt granodiorite 1.06 75.4 (75) CRAT W skarn 827 882 98.4 1.0 KR-05-22 Tay River Roy oxidized monzogranite 1.07 81.1 (81) ROY Ag–Pb–Zn manto 843 896 96.1 0.9 KR-05-191 Tay River Marion oxidized monzogranite 1.07 81.7 (82) MARION Sn–Pb–Zn and Sn– Sb–sulfosalt vein 836 846 96.7 1.4 KR-05-26 Tay River Fish oxidized monzogranite 1.04 87.7 (87) 842 822 96.1 0.6 0.3M-216 Tay River Faille reduced monzogranite 1.11 75.9 (n/a) - 845 97 (ID–TIMS) n/a 0.3M-217 Tay River Mount Sir James McBrien (Cirque) oxidized monzogranite 1.09 74.5 (n/a) - 893 96.8 (ID–TIMS) 0.5 KR-05-164 Tay River ('evolved') Hole-in-the-Wall reduced monzogranite 1.00 88.8 (89) 826 816 95.3 1.1 KR-05-32 Tay River ('evolved') Ivo reduced monzogranite 1.05 89.6 (90) IVO W skarn 871 735 96.0 1.0 KR-05-110 Tungsten (NE) Cac reduced monzogranite 1.11 77.0 (77) CAC W skarn 833 851 97.4 1.0 KR-05-175 Tungsten (NE) Lened reduced monzogranite 1.11 78.5 (79) LENED W skarn 843 860 97.6 0.8 KR-05-126 Tungsten (NE) (plug SW of O'Grady) reduced monzogranite (altered) 1.11 88.3 (88) 798 734 95.9 1.4 03M-207 Tungsten (SE) Mine Stock-apical, Cantung pluton reduced monzogranite 1.20 78.6 (79.5) Cantung W skarn - 803 98.2 (ID–TIMS) 0.4 KR-05-143 Tungsten (SE) Circular Stock reduced monzogranite 1.11 80.1 (80) 832 827 96.9 0.6 KR-05-208 Tungsten (SE) Little Hyland reduced monzogranite 1.15 84.2 (84) 832 816 96.8 0.6 KR-05-148 Tungsten (SE) Rifle Range reduced monzogranite 1.01 84.9 (85) 835 819 96.3 1.0 KR-05-210 Tungsten (SE) Nahanni Range reduced monzogranite 1.20 86.9 (86) 845 732 97.2 0.5 Notes: a bracketed geographic region abbreviations are: NE = northeast Selwyn Basin, SE = southeast Selwyn Basin; b italicized = names given in this study; c ASI (Aluminum Saturation Index, after Frost et al., 2001) and D.I. (Differentiation Index, after Thornton and Tuttle, 1960) are calculated from the geochemical dataset presented in Appendix A.1; d zircon and apatite (for metaluminous samples) saturation temperatures (°C) assuming magma temperature of 850°C; Harrison and Watson, 1984); ZSTs were calculated using the ALS Chemex geochemical dataset (Appendix A.1); note that in all KR-05 samples listed above where geochemical data was available from PCIGR, the ZSTs calculated from ALS Chemex and PCIGR labs were within 3% of each other; anomalously high Zr in the three 03M samples is related to analytical (digestion) issues (Appendix A.1) and ZSTs are not reported for these samples. e AST (Apatite Saturation Temperature, after Bea et al., 1992; for peraluminous samples and assuming magma temperature of 850°C). f U–Pb ages were determined by LA–ICP–MS methods unless otherwise noted as ID– TIMS (the italicized age lacks precision); g 95% confidence interval (reported as internal 2-sigma error). 451 Table 7.2. 40Ar/39Ar ages and associated 2-sigma internal errors for biotite, and calculated absolute, minimum, and maximum cooling periods. Uncorrected ages were previously reported in Rasmussen et al. (2007). Plutonic Suite (sub- suite/region)a Sample Intrusion (MD = mafic dyke)b 40Ar/39Ar age (Ma)a 2-δ error (Ma)b 40Ar/39Ar age (Ma)c 2-δ error (Ma)b,c Notesd Absolute Cooling Periode Absolute Cooling Periodcalcf Minimum Cooling Periodcalcg Maximum Cooling Periodcalch Tombstone (NE) KR-05-97b O'Grady nd nd nd nd nd Tombstone (SE) KR-05-198 Hole-in-the-Wall 92.5 0.5 93.1 0.5 35.3% of 39Ar, steps 5–7 (excess Ar: forced plateau is within error of the I.I. age for plateau steps 8–14) –1.2 –1.8 –4.4 0.8 Tombstone (SE) KR-05-194 Coal River 67.0 0.4 67.5 0.4 93.4% of 39Ar, steps 3–8 (step 9 is slightly younger) 25.7 25.3 24.0 26.6 Mayo (Hess– MacMillan KR-05-76 Pelly River 97.1 0.5 97.7 0.5 78.9% of 39Ar, steps 5–14 0.9 0.3 –1.3 1.9 Mayo (Hess– MacMillan KR-05-77 Pelly River (MD) 95.7 0.7 96.3 0.7 91.9% of the 39Ar, steps 3–9 Mayo KR-05-62 Mount Christie 94.4 0.5 95.0 0.5 72.2% of 39Ar, steps 5–10 0.7 0.1 –1.7 1.8 Mayo KR-05-130 Central Nahanni 96.0 0.5 96.6 0.5 78.5% of 39Ar, steps 6–15 1.2 0.5 –0.7 1.8 Mayo KR-05-136 North Nahanni 96.8 0.5 97.4 0.5 83.4% of 39Ar, steps 5–14 0.1 –0.5 –1.8 0.7 Mayo (Hess– MacMillan KR-05-68 Mount Wilson 96.4 0.5 97.0 0.5 71.1% of 39Ar, steps 4–9, poor I.I. from plateau steps 1.4 0.7 –0.5 2.0 Mayo KR-05-113 South Nahanni 95.8 0.5 96.5 0.5 56.9% of 39Ar, steps 7–13 Tay River (Coal River) KR-05-08 Jorgensen nd nd nd nd Tay River (Coal River) KR-05-43 Park 97.0 0.5 97.6 0.5 66.5% of 39Ar, steps 6–14 –0.9 –1.5 –3.2 0.2 Tay River (Coal River) KR-05-10 Powers nd nd nd nd Tay River KR-05-196 Coal River 94.9 0.5 95.5 0.5 100% of 39Ar, steps 1–11 3.5 2.9 1.4 4.4 Tay River KR-05-22 Roy 96.2 0.5 96.8 0.5 71.9% of 39Ar, steps 7–15 –0.1 –0.7 –2.2 0.7 Tay River KR-05-157 Roy (MD) 96.5 0.7 97.2 0.7 99.12% of the 39Ar, steps 2–8 Tay River KR-05-158 Roy (MD) 97.6 0.7 98.3 0.7 94.1% of the 39Ar, steps 3–8 Tay River KR-05-191 Marion 95.5 0.5 96.1 0.5 95.9% of 39Ar, steps 4–12 1.3 0.6 –1.3 2.5 Tay River KR-05-26 Fish 97.2 0.6 97.8 0.6 59.3% of 39Ar, steps 5–9, slightly humped profile –1.1 –1.7 –2.9 –0.6 Tay River 03M-216 Faille 96.4 0.5 97.0 0.5 89.2% of 39Ar, steps 3–10 (step 11 is slightly older) 0.6 –0.0 –0.5 0.5 Tay River 03M-217 Mount Sir James McBrien (Cirque) 95.3 0.5 95.9 0.5 70.8% of 39Ar, steps 5–17 1.5 0.9 1.7 1.9 Tay River ('evolved') KR-05-164 Hole-in-the-Wall 95.2 0.5 95.8 0.5 44.8% of 39Ar, steps 5–11 (step 12 is slightly younger) 0.1 –0.5 –2.2 1.1 Tay River ('evolved') KR-05-32 Ivo 65.0 0.4 65.4 0.4 83.3% of 39Ar, steps 3–9 31.1 30.6 29.3 32.0 Tungsten (NE) KR-05-110 Cac 94.6 0.5 95.2 0.5 69.8% of 39Ar, steps 5–11 (step 12 is slightly older) 2.8 2.2 0.7 3.6 452 Plutonic Suite (sub- suite/region)a Sample Intrusion (MD = mafic dyke)b 40Ar/39Ar age (Ma)a 2-δ error (Ma)b 40Ar/39Ar age (Ma)c 2-δ error (Ma)b,c Notesd Absolute Cooling Periode Absolute Cooling Periodcalcf Minimum Cooling Periodcalcg Maximum Cooling Periodcalch Tungsten (NE) KR-05-175 Lened 94.8 0.6 95.4 0.6 39% of 39Ar, steps 4–10 2.8 2.2 0.8 3.6 Tungsten (NE) KR-05-126 (plug SW of O'Grady) nd nd nd nd Tungsten (SE) 03M-207 Mine Stock-apical, Cantung pluton 95.1 0.4 95.7 0.4 94.4% of 39Ar, steps 3–16 (steps 17 and 18 were not included) 3.1 2.5 1.7 3.3 Tungsten (SE) KR-05-204 Mine Stock (MD) 91.0 0.8 91.6 0.8 63.4% of the 39Ar, steps 4–8 Tungsten (SE) KR-05-143 Circular Stock 97.6 0.5 98.2 0.5 70.7% of 39Ar, steps 5–12 –0.7 –1.3 –2.5 –0.2 Tungsten (SE) KR-05-208 Little Hyland 89.3 0.5 89.9 0.5 79.4% of 39Ar, steps 3–9 7.5 6.9 5.8 8.0 Tungsten (SE) KR-05-148 Rifle Range 96.6 0.5 97.2 0.5 77% of 39Ar, steps 4–11 –0.3 –0.9 –2.4 0.6 Tungsten (SE) KR-05-213 Rifle Range (MD) 96.7 0.8 97.4 0.8 99.99% of the 39Ar, steps 1–7 Tungsten (SE) KR-05-210 Nahanni Range 86.5 0.5 87.1 0.5 69.5% of 39Ar, steps 4–11 10.7 10.1 9.1 11.1 Notes: a bracketed geographic region abbreviations are: NE = northeast Selwyn Basin, SE = southeast Selwyn Basin; b italicized = names given in this study; a 40Ar/39Ar plateau ages for biotite separates with problematic analyses italicized (see notes for description of problem); b 95% confidence interval (reported as internal 2-sigma error); c ages and errors recalculated using the revised age of 28.201± 0.046 Ma for the Fish Canyon standard (Kuiper et al., 2008); d I.I. = inverse isochron; e absolute difference between U–Pb and 40Ar/39Ar ages without accounting for 2-sigma errors (negative values are within 2-sigma errors); f absolute difference between U–Pb and 40Ar/39Arcalc ages without accounting for 2-sigma errors (negative values are within error); h cooling period as defined in text, and using 40Ar/39Arcalc ages. 453 Table 7.3. Average apatite compositions for each sample, with 1-sigma deviations and relevant calculated parametres. Atomic proportions are ormalized to 13 anhydrous oxygen. Sub-suites are denoted: H–M = Hess–MacMillan, and CR = Coal River. Plutonic Suite: Tombstone Mayo Tungsten (sub-suite/region) (NE) (SE) (SE) (H–M.) (NE) (NE) (NE) de te ct io n lim it (o xi de w t.% ) K R -0 5- 97 b ± 1 -σ K R -0 5- 19 4 ± 1 -σ K R -0 5- 19 8 ± 1 -σ K R -0 5- 13 0 ± 1 -σ K R -0 5- 62 ± 1 -σ K R -0 5- 13 0 ± 1 -σ K R -0 5- 13 6 ± 1 -σ K R -0 5- 68 ± 1 -σ K R -0 5- 11 0 ± 1 -σ K R -0 5- 17 5 ± 1 -σ K R -0 5- 12 6 ± 1 -σ na 6 5 9 2 4 3 6 4 4 4 3 P2O5 0.11 41.13 0.42 41.74 0.45 41.68 0.42 41.47 0.21 41.35 0.31 42.30 0.25 41.95 0.55 41.94 0.46 42.00 0.36 41.72 0.34 41.69 0.15 CaO 0.07 54.99 0.26 54.66 0.16 54.44 0.47 55.00 0.46 55.63 0.08 54.60 0.81 54.73 0.59 53.40 0.32 53.91 0.25 53.48 0.30 55.46 0.03 Al2O3 0.03 0.03 0.02 0.02 0.02 0.02 0.02 0.05 0.01 0.06 0.01 0.02 0.02 0.03 0.02 0.01 0.01 0.01 0.01 0.02 0.01 0.05 0.01 FeOT 0.12 0.05 0.03 0.05 0.01 0.06 0.02 0.05 0.03 0.07 0.02 0.07 0.05 0.07 0.02 0.39 0.08 0.08 0.04 0.24 0.03 0.15 0.05 MnO 0.12 0.07 0.03 0.13 0.05 0.11 0.03 0.10 0.02 0.06 0.03 0.09 0.03 0.13 0.05 0.54 0.05 0.20 0.05 0.66 0.05 0.13 0.03 MgO 0.04 0.05 0.04 0.03 0.02 0.05 0.04 0.10 0.02 0.09 0.02 0.04 0.06 0.06 0.03 0.09 0.02 0.01 0.01 0.06 0.01 0.11 0.02 Cl 0.06 0.13 0.05 0.03 0.01 0.03 0.01 0.71 0.37 0.13 0.03 0.17 0.04 0.24 0.02 0.09 0.01 0.02 0.01 0.04 0.02 0.05 0.01 Fb 0.09 3.73 0.06 3.86 0.03 3.84 0.03 2.60 0.04 3.32 0.03 3.42 0.07 3.18 0.07 3.47 0.07 3.76 0.03 3.63 0.02 3.92 0.07 S 0.05 0.02 0.01 0.03 0.00 0.01 0.02 0.02 0.00 0.06 0.05 0.03 0.02 0.03 0.02 0.02 0.01 0.04 0.03 0.01 0.01 0.03 0.01 Total 98.60 0.56 98.92 0.37 98.62 0.65 99.19 0.35 99.50 0.27 99.37 0.60 99.22 0.45 98.54 0.75 98.44 0.54 98.33 0.43 99.92 0.24 100*F/P2O5 9.06 0.15 9.25 0.05 9.20 0.13 6.27 0.07 8.02 0.05 8.08 0.20 7.57 0.12 8.28 0.11 8.96 0.10 8.70 0.05 9.40 0.18 cFe2+ 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.03 0.01 0.01 0.00 0.02 0.00 0.01 0.00 Mg2+ 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.01 0.01 0.00 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.00 Mn2+ 0.00 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.04 0.00 0.01 0.00 0.05 0.00 0.01 0.00 Ca2+ 5.01 0.02 4.97 0.03 4.96 0.02 5.00 0.03 5.02 0.01 4.95 0.04 4.97 0.04 4.89 0.01 4.93 0.00 4.90 0.02 4.98 0.01 P5+ 2.96 0.02 3.00 0.02 3.00 0.03 2.98 0.02 2.95 0.01 3.03 0.04 3.01 0.04 3.03 0.01 3.04 0.01 3.02 0.02 2.96 0.01 F- (XF) 1.00 0.02 1.04 0.01 1.03 0.01 0.70 0.01 0.88 0.01 0.91 0.01 0.85 0.02 0.94 0.01 1.02 0.01 0.98 0.00 1.04 0.02 Cl- (XCl) 0.02 0.01 0.00 0.00 0.00 0.00 0.10 0.05 0.02 0.00 0.03 0.01 0.03 0.00 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.00 XOHd –0.02 0.02 –0.04 0.01 –0.04 0.01 0.20 0.04 0.10 0.01 0.06 0.01 0.12 0.02 0.05 0.02 –0.018 0.01 0.01 0.00 –0.047 0.02 XCl/XF 0.02 0.01 0.00 0.00 0.00 0.00 0.15 0.08 0.02 0.00 0.03 0.01 0.04 0.00 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.00 XF/XOH 196 n/a n/a n/a n/a n/a 3.53 0.63 9.03 0.70 15.5 3.22 7.60 1.54 20.4 6.28 n/a n/a 86.3 9.73 n/a n/a nHCl/Hfe 6 5 8 2 4 3 6 4 1 3 3 log aHCl/aHFf –0.18 0.14 –0.75 0.12 –0.80 0.11 0.69 0.24 –0.15 0.09 –0.02 0.10 0.16 0.04 –0.32 0.06 –0.84 n/a –0.60 0.10 –0.62 0.11 nHF/HOHe 1 0 0 2 4 3 6 4 0 4 0 log aHOH/aHFg 0.41 n/a n/a n/a n/a n/a 2.16 0.08 1.75 0.03 1.52 0.09 1.83 0.09 1.41 0.14 n/a n/a 0.77 0.05 n/a n/a Notes: a n = number of analyses averaged (SrO was analyzed but always <0.01 wt.%); b calculated initial F concentration; c atomic Al was negligible and was removed from the table (typically = 0.00; ) d determined from stoichiometry; e number of analyses accepted for calculations; f log a(HCl) = –ΔG0R,2 / (2.302585 * RT) – log a(Hap) + log a(ClAp), where ΔG0R,2 = –5860 (850°C, 2 kbar) for the reaction: HCl + HAp = H2O + ClAp; g log a(HF) = –ΔG0R,1 / (2.302585 * RT) – log a(Hap) + log a(FAp), where ΔG0R,1 = –13874 (850°C, 2 kbar) for the reaction: HF + HAp = H2O + Fap. 454 Plutonic Suite: Tungsten Tay River (sub-suite/region) (SE) (SE) (SE) (SE) (SE) (CR) (CR) (CR) K R -0 5- 14 3 ± 1 -σ 03 M -2 07 ± 1 -σ K R -0 5- 20 8 ± 1 -σ K R -0 5- 21 0 ± 1 -σ K R -0 5- 14 8 ± 1 -σ K R -0 5- 10 ± 1 -σ K R -0 5- 43 ± 1 -σ K R -0 5- 08 ± 1 -σ K R -0 5- 19 6 ± 1 -σ K R -0 5- 19 1 ± 1 -σ K R -0 5- 26 ± 1 -σ K R -0 5- 22 ± 1 -σ na 4 7 1 3 4 5 2 9 2 4 9 6 P2O5 41.43 0.24 41.54 0.56 41.73 n/a 41.53 0.54 41.79 0.51 41.77 0.27 41.94 0.04 41.75 0.33 41.42 0.51 41.31 0.58 41.42 0.49 41.74 0.15 CaO 53.67 0.20 54.38 0.50 53.23 n/a 53.21 0.60 53.41 0.55 54.29 0.37 54.28 0.66 55.05 0.36 53.76 0.20 54.19 0.15 54.14 0.42 54.07 0.31 Al2O3 0.02 0.01 0.06 0.06 0.04 n/a 0.02 0.01 0.03 0.02 0.01 0.01 0.01 0.00 0.04 0.02 0.02 0.01 0.01 0.00 0.02 0.01 0.01 0.01 FeOT 0.28 0.04 0.29 0.07 0.19 n/a 0.22 0.16 0.35 0.06 0.20 0.03 0.05 0.01 0.09 0.10 0.10 0.02 0.10 0.05 0.11 0.02 0.04 0.02 MnO 0.43 0.03 0.67 0.07 1.47 n/a 1.33 0.62 0.87 0.06 0.15 0.04 0.17 0.01 0.09 0.02 0.23 0.04 0.22 0.02 0.20 0.04 0.16 0.03 MgO 0.08 0.01 0.13 0.09 0.09 n/a 0.07 0.02 0.08 0.03 0.11 0.02 0.03 0.01 0.07 0.05 0.05 0.00 0.04 0.01 0.06 0.02 0.04 0.03 Cl 0.07 0.02 0.08 0.02 0.03 n/a 0.02 0.01 0.07 0.03 0.72 0.03 0.19 0.04 0.34 0.16 0.51 0.05 0.30 0.08 0.07 0.02 0.22 0.03 Fb 3.81 0.04 3.60 0.06 3.75 n/a 3.80 0.05 3.78 0.05 1.90 0.05 2.82 0.01 2.84 0.30 3.15 0.01 3.38 0.09 3.87 0.04 3.42 0.03 S 0.03 0.02 0.02 0.01 0.01 n/a 0.02 0.01 0.02 0.01 0.14 0.03 0.09 0.00 0.08 0.06 0.02 0.03 0.14 0.12 0.04 0.02 0.04 0.03 Total 98.18 0.37 99.28 0.92 98.95 n/a 98.61 0.94 98.80 0.35 98.99 0.58 98.71 0.67 99.42 0.56 97.93 0.65 98.26 0.49 98.30 0.56 98.34 0.32 100*F/P2O5 9.18 0.12 8.66 0.11 8.99 n/a 9.14 0.13 9.03 0.22 4.56 0.13 6.71 0.01 6.80 0.70 7.61 0.06 8.18 0.10 9.35 0.15 8.19 0.09 cFe2+ 0.02 0.00 0.02 0.00 0.01 n/a 0.02 0.01 0.03 0.00 0.01 0.00 0.00 0.00 0.01 0.01 0.01 0.00 0.01 0.00 0.01 0.00 0.00 0.00 Mg2+ 0.01 0.00 0.02 0.01 0.01 n/a 0.01 0.00 0.01 0.00 0.01 0.00 0.00 0.00 0.01 0.01 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00 Mn2+ 0.03 0.00 0.05 0.01 0.11 n/a 0.10 0.04 0.06 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.02 0.00 0.02 0.00 0.01 0.00 0.01 0.00 Ca2+ 4.92 0.01 4.94 0.03 4.85 n/a 4.87 0.06 4.87 0.04 4.95 0.01 4.95 0.02 4.98 0.02 4.95 0.01 4.97 0.03 4.95 0.03 4.95 0.01 P5+ 3.00 0.01 2.98 0.02 3.01 n/a 3.00 0.01 3.01 0.04 3.01 0.01 3.02 0.03 2.99 0.01 3.01 0.02 2.99 0.03 2.99 0.03 3.02 0.01 F- (XF) 1.03 0.01 0.97 0.01 1.01 n/a 1.02 0.01 1.02 0.01 0.51 0.01 0.76 0.01 0.76 0.08 0.86 0.00 0.91 0.02 1.05 0.01 0.92 0.01 Cl- (XCl) 0.01 0.00 0.01 0.00 0.00 n/a 0.00 0.00 0.01 0.00 0.10 0.00 0.03 0.01 0.05 0.02 0.07 0.01 0.04 0.01 0.01 0.00 0.03 0.00 XOHd –0.04 0.01 0.01 0.00 –0.01 n/a –0.03 0.01 –0.03 0.01 0.38 0.01 0.21 0.00 0.19 0.06 0.07 0.01 0.04 0.02 –0.06 0.01 0.04 0.01 XCl/XF 0.01 0.00 0.02 0.01 0.00 n/a 0.00 0.00 0.01 0.00 0.20 0.01 0.04 0.01 0.07 0.04 0.09 0.01 0.05 0.01 0.01 0.00 0.03 0.00 XF/XOH n/a n/a 51.1 29.7 n/a n/a n/a n/a n/a n/a 1.33 0.09 3.53 0.04 4.31 1.41 12.4 1.78 23.7 7.55 n/a n/a 22 3.86 nHCl/Hfe 3 7 1 3 3 5 2 9 2 4 9 6 log aHCl/aHFf –0.49 0.11 –0.38 0.11 –0.84 n/a –1.13 0.27 –0.45 0.17 0.86 0.02 0.11 0.11 0.34 0.22 0.49 0.04 0.22 0.14 –0.47 0.13 0.10 0.06 nHF/HOHe 0 7 1 0 0 5 2 9 2 4 0 6 log aHOH/aHFg n/a n/a 1.05 0.23 n/a n/a n/a n/a n/a n/a 2.58 0.03 2.15 0.01 2.09 0.17 1.61 0.06 1.34 0.16 n/a n/a 1.36 0.08 455 Plutonic Suite: Tay River (sub- suite/region) ('evolved') ('evolved') K R -0 5- 16 4 ± 1 -σ K R -0 5- 32 ± 1 -σ na 3 1 P2O5 41.58 0.27 41.61 n/a CaO 51.66 0.50 53.45 n/a Al2O3 0.02 0.01 0.00 n/a FeOT 0.58 0.19 0.07 n/a MnO 1.36 0.11 0.44 n/a MgO 0.08 0.05 0.03 n/a Cl 0.13 0.06 0.03 n/a Fb 3.63 0.11 3.81 n/a S 0.02 0.02 0.01 n/a Total 97.50 0.46 97.84 n/a 100*F/P2O5 8.72 0.22 9.16 n/a cFe2+ 0.04 0.01 0.01 n/a Mg2+ 0.01 0.01 0.00 n/a Mn2+ 0.10 0.01 0.03 n/a Ca2+ 4.79 0.02 4.92 n/a P5+ 3.05 0.01 3.03 n/a F- (XF) 0.99 0.02 1.04 n/a Cl- (XCl) 0.02 0.01 0.00 n/a XOHd –0.01 0.02 –0.04 n/a XCl/XF 0.02 0.01 0.00 n/a XF/XOH 121 n/a n/a n/a nHCl/Hfe 3 1 log aHCl/aHFf –0.21 0.25 –0.85 n/a nHF/HOHe 1 0 log aHOH/aHFg 0.62 n/a n/a n/a 456 Table 7.4. Average biotite compositions for each sample, with 1-sigma deviations and relevant calculated parametres. Atomic proportions are normalized to 11 anhydrous oxygen. Sub-suites are denoted: H–M = Hes–MacMillan, and CR = Coal River. Plutonic Suite: Tombstone Mayo (sub-suite/region) (NE) (SE) (SE) d et ec tio n lim it (w t.% o xi de ) K R -0 5- 97 b ± 2 -σ K R -0 5- 19 4 ± 2 -σ K R -0 5- 19 8 ± 2 -σ K R -0 5- 13 0 ± 2 -σ K R -0 5- 62 ± 2 -σ K R -0 5- 11 3 ± 2 -σ K R -0 5- 13 6 ± 2 -σ n 4 3 1 3 2 4 4 SiO2 0.017 37.69 0.33 36.32 0.13 35.70 n/a 37.06 0.43 36.96 0.48 36.44 0.28 36.74 0.05 TiO2 0.021 4.53 0.09 3.41 0.15 4.13 n/a 3.58 0.15 3.54 0.41 3.58 0.08 3.39 0.11 Al2O3 0.013 12.86 0.03 16.34 0.20 13.42 n/a 14.10 0.13 14.36 0.01 13.71 0.13 14.11 0.09 Cr2O3 0.034 0.02 0.02 0.02 0.02 0.06 n/a 0.00 0.00 0.07 0.07 0.03 0.03 0.04 0.03 FeOT 0.041 20.37 0.15 23.10 0.49 27.37 n/a 20.18 0.50 20.14 0.67 22.22 0.27 21.13 0.51 MnO 0.036 0.26 0.04 0.36 0.06 0.52 n/a 0.21 0.08 0.21 0.01 0.32 0.04 0.37 0.08 MgO 0.012 10.78 0.16 6.69 0.09 6.04 n/a 11.58 0.21 10.84 0.48 10.02 0.14 10.93 0.06 CaO 0.018 0.03 0.02 0.02 0.01 0.05 n/a 0.02 0.01 0.03 0.04 0.01 0.01 0.04 0.03 Na2O 0.017 0.05 0.03 0.09 0.03 0.04 n/a 0.04 0.01 0.07 0.04 0.05 0.04 0.08 0.02 K2O 0.021 9.69 0.09 9.68 0.16 9.37 n/a 9.56 0.08 9.68 0.09 9.70 0.10 9.46 0.03 Cl 0.014 0.28 0.03 0.23 0.02 0.22 n/a 0.21 0.01 0.09 0.01 0.23 0.04 0.11 0.01 F 0.078 1.31 0.03 1.25 0.01 1.02 n/a 0.28 0.03 0.83 0.04 0.74 0.03 0.35 0.02 Total 100.43 0.35 100.11 0.65 100.68 n/a 100.41 0.33 99.95 0.61 100.13 0.19 100.29 0.50 bK+ 0.96 0.01 0.97 0.02 0.95 n/a 0.93 0.00 0.95 0.02 0.96 0.01 0.93 0.00 Na+ 0.01 0.00 0.01 0.00 0.01 n/a 0.01 0.00 0.01 0.01 0.01 0.01 0.01 0.00 Fe2+ 1.32 0.01 1.51 0.03 1.82 n/a 1.29 0.04 1.30 0.06 1.45 0.02 1.36 0.03 Mg2+ 1.24 0.02 0.78 0.01 0.72 n/a 1.32 0.02 1.25 0.04 1.16 0.01 1.25 0.01 Fe/(Fe+Mg) 0.51 0.00 0.66 0.01 0.72 n/a 0.49 0.01 0.51 0.02 0.55 0.01 0.52 0.01 Mn2+ 0.02 0.00 0.02 0.00 0.03 n/a 0.01 0.01 0.01 0.00 0.02 0.00 0.02 0.00 Ti4+ 0.26 0.01 0.20 0.01 0.25 n/a 0.21 0.01 0.21 0.02 0.21 0.00 0.20 0.01 Al (VI) 0.09 0.01 0.35 0.02 0.11 n/a 0.09 0.01 0.15 0.01 0.10 0.00 0.10 0.01 Al (IV) 1.09 0.01 1.15 0.01 1.15 n/a 1.17 0.02 1.15 0.01 1.16 0.02 1.18 0.01 Al (apfu) 2.34 0.01 3.02 0.03 2.52 n/a 2.53 0.03 2.61 0.03 2.52 0.03 2.55 0.01 Si4+ 2.91 0.01 2.85 0.01 2.85 n/a 2.83 0.02 2.85 0.01 2.84 0.02 2.82 0.01 F- 0.32 0.01 0.31 0.00 0.26 n/a 0.07 0.01 0.20 0.01 0.18 0.01 0.08 0.00 Cl- 0.04 0.00 0.03 0.00 0.03 n/a 0.03 0.00 0.01 0.00 0.03 0.01 0.01 0.00 XFe 0.53 0.00 0.71 0.00 0.73 n/a 0.51 0.01 0.54 0.02 0.57 0.01 0.54 0.01 XMg 0.47 0.00 0.29 0.00 0.27 n/a 0.49 0.01 0.46 0.02 0.43 0.01 0.46 0.01 XF 0.16 0.00 0.16 0.00 0.13 n/a 0.03 0.00 0.10 0.00 0.09 0.00 0.04 0.00 XCl 0.02 0.00 0.02 0.00 0.01 n/a 0.01 0.00 0.01 0.00 0.02 0.00 0.01 0.00 XOHc 0.82 0.01 0.83 0.00 0.86 n/a 0.95 0.00 0.89 0.00 0.89 0.00 0.95 0.00 XCl/ XF 0.11 0.01 0.10 0.01 0.11 n/a 0.41 0.06 0.06 0.01 0.17 0.04 0.17 0.01 XF/ XOH 0.19 0.01 0.19 0.00 0.15 n/a 0.04 0.00 0.11 0.00 0.10 0.00 0.04 0.00 log aHCl/ aHFd,e –0.04 0.04 –0.36 0.03 –0.33 n/a 0.54 0.06 –0.33 0.07 0.06 0.09 0.14 0.03 log aHOH/ aHFe,f 1.54 0.02 1.38 0.01 1.46 n/a 2.30 0.03 1.76 0.00 1.78 0.01 2.18 0.03 log aHOH/ aHF(corr.)e,g 0.64 0.03 0.31 0.01 0.47 n/a 2.32 0.08 1.14 0.00 1.18 0.03 2.05 0.06 aHOH 0.11 0.01 0.17 0.02 0.15 n/a 0.01 0.00 0.05 0.00 0.05 0.00 0.01 0.00 Notes: a SD = standard deviation; n = number of analyses averaged or analyses accepted for calculations; b Ca atomic proportions are negligible and were removed from the table (typically = 0.00); c determined from stoichiometry; d log a(HCl) = –( 0.5(ΔG0R,c / 2.303RT) * XMg + 0.5logKa – log (XCl/XOH)bt), where ΔG0R,c = –4862 (850°C, 2 kbar) for the reaction: Chphl + Ann = Phl + Clann; e log a(HF) = –( 0.5(ΔG0R,b / 2.303RT) * XMg + 0.5logKa – log (XF/XOH)bt), where ΔG0R,b = –10099 (850°C, 2 kbar) for the reaction: Phl + Fann = Fphl + Ann; assumed pure H2O fluid, or log a(H2O) = 0; f assumed a pure H2O fluid where aHOH = 1 (Zhu and Sverjensky, 1992); g *log aHOH/aHF(corr.) is corrected assuming log aHOH/aHF-AP = log aHOH/aHF-BT and using the linear relationship (R2 = 0.73): *log aHOH/aHF = (log aHOH/aHF-BT+1.246) / 0.4533. 457 Plutonic Suite: Mayo Tungsten (sub-suite/region) (HM) (HM) (NE) (NE) (NE) (SE) (SE) K R -0 5- 68 ± 2 -σ K R -0 5- 76 ± 2 -σ K R -0 5- 11 0 ± 2 -σ K R -0 5- 17 5 ± 2 -σ K R -0 5- 12 6 ± 2 -σ K R -0 5- 14 3 ± 2 -σ 03 M -2 07 ± 2 -σ n 2 5 4 4 4 2 4 SiO2 36.58 0.39 37.12 0.12 35.40 0.24 35.66 0.16 36.39 0.16 35.77 0.00 34.72 0.29 TiO2 3.41 0.21 3.52 0.08 3.68 0.08 3.37 0.15 3.25 0.25 3.30 0.18 3.16 0.32 Al2O3 15.25 0.33 14.64 0.08 15.94 0.30 17.47 0.22 18.11 0.89 17.84 0.05 18.43 0.28 Cr2O3 0.04 0.06 0.02 0.02 0.03 0.02 0.01 0.02 0.03 0.05 0.03 0.03 0.02 0.01 FeOT 20.42 0.39 19.33 0.22 23.69 0.58 22.70 0.29 21.96 0.21 22.96 0.12 22.79 0.33 MnO 0.30 0.04 0.30 0.06 0.51 0.08 0.47 0.06 0.56 0.12 0.47 0.01 0.45 0.03 MgO 10.70 0.02 11.44 0.14 7.12 0.10 7.25 0.04 5.69 0.64 6.59 0.13 6.33 0.14 CaO 0.01 0.02 0.00 0.01 0.01 0.01 0.02 0.02 0.03 0.02 0.00 0.00 0.02 0.02 Na2O 0.10 0.02 0.05 0.05 0.06 0.01 0.08 0.02 0.31 0.03 0.10 0.00 0.11 0.02 K2O 9.60 0.05 9.75 0.08 9.53 0.05 9.74 0.14 9.66 0.09 9.67 0.16 9.53 0.07 Cl 0.07 0.01 0.05 0.01 0.11 0.02 0.05 0.01 0.07 0.02 0.11 0.02 0.09 0.01 F 0.53 0.08 0.68 0.04 0.60 0.04 0.60 0.04 6.19 0.25 0.85 0.00 0.49 0.04 Total 100.43 0.67 100.20 0.34 99.93 0.67 100.74 0.30 100.32 0.27 100.74 0.06 99.50 0.77 bK+ 0.94 0.00 0.95 0.01 0.95 0.01 0.96 0.01 1.01 0.01 0.95 0.02 0.95 0.01 Na+ 0.01 0.00 0.01 0.01 0.01 0.00 0.01 0.00 0.05 0.00 0.01 0.00 0.02 0.00 Fe2+ 1.30 0.02 1.23 0.01 1.55 0.03 1.46 0.02 1.51 0.02 1.48 0.01 1.49 0.02 Mg2+ 1.22 0.01 1.30 0.02 0.83 0.01 0.83 0.00 0.70 0.08 0.76 0.02 0.74 0.02 Fe/(Fe+Mg) 0.52 0.01 0.49 0.01 0.65 0.00 0.64 0.00 0.68 0.02 0.66 0.00 0.67 0.01 Mn2+ 0.02 0.00 0.02 0.00 0.03 0.01 0.03 0.00 0.04 0.01 0.03 0.00 0.03 0.00 Ti4+ 0.20 0.01 0.20 0.00 0.22 0.00 0.20 0.01 0.20 0.02 0.19 0.01 0.19 0.02 Al (VI) 0.17 0.02 0.15 0.01 0.24 0.02 0.33 0.01 0.75 0.07 0.39 0.01 0.40 0.03 Al (IV) 1.20 0.04 1.17 0.00 1.23 0.01 1.25 0.01 1.01 0.01 1.24 0.00 1.29 0.01 Al (apfu) 2.75 0.05 2.63 0.02 2.94 0.06 3.17 0.04 3.51 0.16 3.25 0.01 3.39 0.04 Si4+ 2.80 0.04 2.83 0.00 2.77 0.01 2.75 0.01 2.99 0.01 2.76 0.00 2.71 0.01 F- 0.13 0.02 0.16 0.01 0.15 0.01 0.15 0.01 1.61 0.07 0.21 0.00 0.12 0.01 Cl- 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00 XFe 0.55 0.00 0.52 0.00 0.68 0.00 0.68 0.00 0.76 0.02 0.71 0.00 0.72 0.00 XMg 0.45 0.00 0.48 0.00 0.32 0.00 0.32 0.00 0.24 0.02 0.29 0.00 0.28 0.00 XF 0.06 0.01 0.08 0.00 0.07 0.00 0.07 0.00 0.81 0.04 0.10 0.00 0.06 0.01 XCl 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.01 0.00 XOHc 0.93 0.01 0.91 0.00 0.92 0.00 0.92 0.01 0.19 0.04 0.89 0.00 0.93 0.01 XCl/ XF 0.07 0.00 0.04 0.01 0.10 0.02 0.05 0.01 0.01 0.00 0.07 0.02 0.10 0.02 XF/ XOH 0.07 0.01 0.09 0.01 0.08 0.01 0.08 0.01 4.40 1.10 0.12 0.00 0.06 0.01 log aHCl/ aHFd,e –0.31 0.00 –0.48 0.06 –0.34 0.07 –0.66 0.08 –1.70 0.16 –0.54 0.10 –0.36 0.08 log aHOH/ aHFe,f 1.97 0.07 1.89 0.03 1.77 0.03 1.78 0.03 -0.04 0.08 1.58 0.00 1.83 0.04 log aHOH/ aHF(corr.)e,g 1.60 0.15 1.42 0.06 1.15 0.06 1.19 0.07 -2.83 0.18 0.75 0.00 1.29 0.10 aHOH 0.03 0.01 0.04 0.00 0.05 0.01 0.05 0.01 18.0 4.94 0.10 0.00 0.04 0.01 458 Plutonic Suite: Tungsten Tay River (sub-suite/region) (SE) (SE) (SE) (CR) K R -0 5- 20 8 ± 2 -σ K R -0 5- 21 0 ± 2 -σ K R -0 5- 14 8 ± 2 -σ K R -0 5- 43 ± 2 -σ K R -0 5- 19 1 ± 2 -σ K R -0 5- 26 ± 2 -σ K R -0 5- 22 ± 2 -σ n 2 1 2 4 5 3 3 SiO2 35.45 0.00 35.14 n/a 35.62 0.38 37.25 0.28 36.87 0.08 35.86 0.12 37.23 0.52 TiO2 2.65 0.78 2.07 n/a 3.31 0.28 3.86 0.11 4.29 0.23 3.70 0.26 4.25 0.09 Al2O3 17.78 0.68 18.13 n/a 16.37 0.13 14.12 0.23 13.31 0.06 12.89 0.15 13.73 0.22 Cr2O3 0.00 0.00 0.00 n/a 0.00 0.00 0.02 0.03 0.02 0.04 0.02 0.02 0.01 0.02 FeOT 24.39 0.15 24.74 n/a 24.34 0.00 18.84 0.13 20.98 0.71 26.53 0.36 18.17 0.83 MnO 1.13 0.14 0.89 n/a 0.52 0.01 0.42 0.05 0.40 0.04 1.28 0.07 0.47 0.03 MgO 5.30 0.07 5.37 n/a 6.75 0.24 11.98 0.12 10.85 0.75 6.49 0.17 12.66 0.62 CaO 0.01 0.00 0.02 n/a 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.03 0.02 Na2O 0.05 0.03 0.06 n/a 0.14 0.06 0.06 0.02 0.13 0.03 0.10 0.02 0.08 0.04 K2O 9.78 0.17 9.79 n/a 9.45 0.06 9.69 0.13 9.54 0.11 9.21 0.04 9.71 0.03 Cl 0.07 0.00 0.03 n/a 0.20 0.00 0.05 0.03 0.27 0.06 0.33 0.07 0.19 0.03 F 1.08 0.06 1.12 n/a 1.09 0.04 0.21 0.01 0.68 0.22 1.96 0.38 1.20 0.08 Total 100.50 0.13 100.15 n/a 100.54 0.49 100.27 0.26 100.48 0.11 100.18 0.46 100.48 0.26 bK+ 0.98 0.02 0.99 n/a 0.94 0.00 0.94 0.01 0.94 0.01 0.95 0.01 0.95 0.00 Na+ 0.01 0.00 0.01 n/a 0.02 0.01 0.01 0.00 0.02 0.00 0.02 0.00 0.01 0.01 Fe2+ 1.60 0.01 1.63 n/a 1.60 0.01 1.19 0.01 1.35 0.05 1.80 0.03 1.16 0.06 Mg2+ 0.62 0.01 0.63 n/a 0.79 0.02 1.35 0.01 1.25 0.09 0.78 0.02 1.44 0.06 Fe/(Fe+Mg) 0.72 0.00 0.72 n/a 0.67 0.01 0.47 0.00 0.52 0.03 0.70 0.00 0.45 0.02 Mn2+ 0.07 0.01 0.06 n/a 0.03 0.00 0.03 0.00 0.03 0.00 0.09 0.00 0.03 0.00 Ti4+ 0.16 0.05 0.12 n/a 0.20 0.02 0.22 0.01 0.25 0.01 0.23 0.02 0.24 0.00 Al (VI) 0.43 0.06 0.46 n/a 0.31 0.02 0.08 0.02 0.05 0.01 0.14 0.01 0.08 0.00 Al (IV) 1.22 0.00 1.23 n/a 1.21 0.01 1.18 0.02 1.16 0.01 1.09 0.01 1.15 0.02 Al (apfu) 3.29 0.12 3.37 n/a 3.02 0.01 2.52 0.05 2.42 0.01 2.46 0.03 2.47 0.06 Si4+ 2.78 0.00 2.77 n/a 2.79 0.01 2.82 0.02 2.84 0.01 2.91 0.01 2.85 0.02 F- 0.27 0.01 0.28 n/a 0.27 0.01 0.05 0.00 0.17 0.05 0.50 0.10 0.29 0.02 Cl- 0.01 0.00 0.00 n/a 0.03 0.00 0.01 0.00 0.04 0.01 0.04 0.01 0.03 0.00 XFe 0.77 0.00 0.77 n/a 0.71 0.01 0.49 0.00 0.53 0.02 0.71 0.00 0.46 0.02 XMg 0.23 0.00 0.23 n/a 0.29 0.01 0.51 0.00 0.47 0.02 0.29 0.00 0.54 0.02 XF 0.13 0.01 0.14 n/a 0.14 0.00 0.03 0.00 0.08 0.03 0.25 0.05 0.14 0.01 XCl 0.00 0.00 0.00 n/a 0.01 0.00 0.00 0.00 0.02 0.00 0.02 0.00 0.01 0.00 XOHc 0.86 0.01 0.86 n/a 0.85 0.00 0.97 0.00 0.90 0.02 0.73 0.04 0.84 0.01 XCl/ XF 0.03 0.00 0.02 n/a 0.10 0.00 0.12 0.07 0.24 0.12 0.09 0.04 0.09 0.02 XF/ XOH 0.16 0.01 0.16 n/a 0.16 0.01 0.03 0.00 0.09 0.03 0.35 0.08 0.17 0.01 log aHCl/ aHFd,e –0.93 0.05 –1.24 n/a –0.37 0.01 –0.01 0.23 0.24 0.20 –0.41 0.18 –0.07 0.09 log aHOH/ aHFe,f 1.40 0.03 1.38 n/a 1.45 0.01 2.46 0.02 1.88 0.13 1.11 0.11 1.66 0.01 log aHOH/ aHF(corr.)e,g 0.35 0.07 0.30 n/a 0.45 0.02 2.68 0.04 1.40 0.29 -0.29 0.25 0.91 0.03 aHOH 0.17 0.02 0.19 n/a 0.01 0.00 0.06 0.00 0.03 0.00 0.46 0.15 0.08 0.00 459 Plutonic Suite: Tay River (sub-suite/region) ('evolved') ('evolved') K R -0 5- 16 4 ± 2 -σ K R -0 5- 32 ± 2 -σ n 2 3 SiO2 35.32 0.24 34.10 0.21 TiO2 3.02 0.05 2.77 0.05 Al2O3 18.37 0.07 17.71 0.11 Cr2O3 0.00 0.00 0.01 0.01 FeOT 23.69 0.48 26.13 0.37 MnO 0.31 0.04 0.73 0.04 MgO 5.92 0.05 4.38 0.05 CaO 0.02 0.02 0.01 0.01 Na2O 0.07 0.01 0.08 0.02 K2O 9.90 0.06 9.64 0.12 Cl 0.08 0.01 0.07 0.01 F 1.00 0.03 0.50 0.03 Total 100.62 0.44 99.42 0.71 bK+ 0.98 0.01 0.98 0.01 Na+ 0.01 0.00 0.01 0.00 Fe2+ 1.54 0.03 1.74 0.01 Mg2+ 0.69 0.01 0.52 0.00 Fe/(Fe+Mg) 0.69 0.00 0.77 0.00 Mn2+ 0.02 0.00 0.05 0.00 Ti4+ 0.18 0.00 0.17 0.00 Al (VI) 0.43 0.02 0.38 0.01 Al (IV) 1.25 0.02 1.28 0.01 Al (apfu) 3.37 0.01 3.33 0.02 Si4+ 2.75 0.02 2.72 0.01 F- 0.25 0.01 0.13 0.01 Cl- 0.01 0.00 0.01 0.00 XFe 0.74 0.00 0.80 0.00 XMg 0.26 0.00 0.20 0.00 XF 0.12 0.00 0.06 0.00 XCl 0.01 0.00 0.01 0.00 XOHc 0.87 0.00 0.93 0.00 XCl/ XF 0.04 0.00 0.08 0.01 XF/ XOH 0.14 0.00 0.07 0.00 log aHCl/ aHFd,e –0.76 0.03 –0.59 0.04 log aHOH/ aHFe,f 1.47 0.01 1.73 0.03 log aHOH/ aHF(corr.)e,g 0.49 0.03 1.07 0.07 aHOH 0.14 0.01 0.06 0.01 460 8 CHAPTER: SUMMARY OF CONCLUSIONS AND A GEODYNAMIC MODEL FOR SYN- TO POST-ACCRETIONARY MAGMATISM IN THE NORTHERN CORDILLERA The large amount of geochronological, geochemical, isotopic, and magmatic fluid composition data generated by this study, and its integration with compiled datasets, has substantially advanced our knowledge of the timing, composition, and petrogenesis for syn- to post-accretionary igneous rocks in the northern Cordilleran miogeocline. The conclusions of the main aspects of this dissertation are summarized as follows. Section 8.1 focuses on the age, composition, magma sources, mechanisms of partial melting, and metallogenic associations for these igneous rocks, drawing on data presented in Chapters 2 to 5. Data from Chapters 6 and 7 are summarized in Section 8.2, in which the evolution of magmatic fluids and the duration of cooling for intrusions at the eastern end of the most metallogenically prospective magmatic belt in the Selwyn Basin (e.g., Hart et al., 2004a, 2005) are reviewed. This section specifically discusses a strategy for identifying intrusions with associated tungsten skarn mineralization. In Section 8.3, a tectono–magmatic summary puts the findings of this study into context with respect to plate movements and the changing Mesozoic to Cenozoic deformational regime in northwest North America. Lastly, suggestions for additional work that would help to further our understanding of syn- to post-accretionary magmatism in the northern Cordillera, and of the age and composition of the basement to the northern Cordilleran miogeocline—which remains a largely unknown aspect of Cordilleran geology—are given in Section 8.4. It is hoped that additional research will not only broaden our understanding of this exceptionally large and well-exposed, intermediate–felsic continental granitoid belt (and its associated mineralization), but also our understanding of similar magmatic–metallogenic belts worldwide. 8.1 AGE, COMPOSITION, MAGMA SOURCES, MECHANISMS OF PARTIAL MELTING, AND METALLOGENY 8.1.1 Cassiar Terrane Prior to this study, all undeformed igneous rocks in the Cassiar Terrane had been classified as the 110–99 Ma Cassiar plutonic suite. This study has not only demonstrated that this plutonic suite is slightly older than previously thought (115–104 Ma), but it has also distinguished three other compositionally, temporally, and geographically distinct magmatic episodes in the Cassiar Terrane. These are the 103–94 Ma Seagull plutonic suite, the 82–75 Rancheria plutonic suite, and the 49–47 Ma Black River plutonic suite, all of which are of A-type affinity. The high-K calc-alkaline S- > I-type Cassiar plutonic suite has historically been attributed to partial melting of over- thickened metasedimentary rocks including minor layers of amphibolite (e.g., Driver et al., 2000; Morris and Creaser, 2008). This is still considered to be the most appropriate lithological source of these magmas. The probable melt source, as indicated by all of the isotopic data for this plutonic suite, is inferred here to be Paleozoic in age, 461 which is younger than previously suggested. Although Paleozoic rocks are present at the surface in this part of Yukon, they have not previously been interpreted to be present at the depths required for fluid-absent partial melting involving mica and amphibole dehydration reactions. These melt-producing reactions require depths greater than 25 km—corresponding to a present-day depth of over 15 km—which requires modifications to be made to current interpretations concerning the age of the lower–middle crustal basement to the Cassiar Terrane. Even at these depths, however, additional heat is required. Both internal (i.e., the decay of radioactive elements in U–Th–K-rich supracrustal rocks) and external (i.e., the mantle) sources of heat are invoked. The latter is necessary to generate partial melt in the subordinate, interlayered mafic to intermediate metaigneous rocks. Vigorous asthenospheric upwelling in a transpressional back-arc setting may have been responsible for conductively, and possibly advectively, transferring mantle heat to the crust (Fig. 8.1a). The alkali-calcic (to peralkaline) A-type Seagull, Rancheria, and Black River plutonic suites are interpreted here to have been predominantly sourced in the mantle, although varying degrees of infracrustal input is also apparent in all three suites. The Seagull and Black River plutonic suites have geochemical signatures that are suggestive of an undepleted asthenospheric melt source (although a derivation soely from crustal rocks cannot be ruled out for the Black River plutonic suite), whereas the Rancheria plutonic suite has a geochemical signature that is suggestive of an enriched or upper lithospheric mantle melt source. Isotopic compositions, however, indicate that a significant component of crustal neodymium, strontium, and oxygen isotopes have been added to these magmas through crustal contamination, assimilation, and (or) mixing processes. The spatial association of the A-type plutonic suites with major structural discontinuities, and their emplacement in the form of narrow, linear belts (Seagull and Rancheria suites; Fig. 8.1) require extremely focused magma generation. Two possible mechanisms for the localization of these magmas are: (1) the long, linear Seagull plutonic suite may be attributed to rifting or tearing of normal-dipping subducted oceanic crust below the western Cassiar Terrane at a depth of ~150 km (Fig. 8.1b); and, (2) the Rancheria and Black River suites may be attributed to local tapping of mantle and (or) crustal magmas along lithosphere-scale structures (i.e., a terrane boundary, a regional strike-slip fault, or a steep-dipping normal fault related to older basement rifting and extension; Fig. 8.1d,e). 8.1.2 Selwyn Basin The results of this study have not substantially changed previous plutonic suite divisions across the Selwyn Basin. They have, however, better defined the temporal and geographic boundaries of the plutonic suites in the eastern Selwyn Basin, where this work was primarily focused, and identified some inter- and intra-suite compositional differences across the Selwyn Basin. The Early Cretaceous Anvil (112–106 Ma) and Hyland (106–100 Ma) plutonic suites are interpreted to be derived from partial melting of supracrustal rocks in the Selwyn Basin. The mineralogical, chemical, and isotopic compositions of the Anvil plutonic suite are compatible with a derivation by low-temperature partial melting of a compositionally homogeneous, mainly pelitic lithology. Conversely, the Hyland plutonic suite appears to have been generated by higher-temperature partial melting of a compositionally heterogeneous, mainly greywacke lithology. 462 They protolith to the Hyland plutonic suite is isotopically less evolved than its migmatitized Neoproterozoic supracrustal host rocks; therefore, basinal equivalents to the late Mesoproterozoic Mackenzie Mountains Supergroup are suggested as one possibility. The differences between these two lithologically distinct supracrustal protoliths to these S-type plutonic suites may be related to their respective distance from the inferred ancestral margin of North America, with finer-grained and more homogeneous lithologies generally deposited in westerly (distal) position, and coarser-grained and more lithologically variable lithologies generally deposited in a more easterly (proximal) position. The supracrustal source rocks were structurally thickened during Mesozoic thin-skinned deformation, and probably contained a relatively high proportion of radioactive heat-producing elements (e.g., U, Th, K), both of which would have been important factors in Early Cretaceous anatexis. If local occurrences of hornblende-bearing porphyry phases associated with the Hyland plutonic suite are cogenetic (which has not yet been temporally confirmed), amphibolitic sills or flows within the source rocks and a role for added heat (±magma) from the mantle would have been required for this particular plutonic suite. Partial melting of supracrustal rocks in the western Selwyn Basin to form the Early Cretaceous Anvil and Hyland plutonic suites may have been aided by nearby structures, possibly during decompression in the hanging wall to thrust (Anvil suite) and strike-slip (Hyland suite) faults. Finally, as with the Cassiar Terrane back-arc, the western Selwyn Basin was also a transpressional, Early Cretaceous back-arc that experienced high heat flow related to vigorous asthenospheric convection above a down- going oceanic plate (Fig. 8.1a; cf. Currie and Hyndman, 2006). The largest contribution made by this study is the generation of an extensive geochemical and isotopic dataset for igneous rocks in the eastern Selwyn Basin. These data have allowed previous interpretations regarding the origin of the high-K calc-alkaline (to alkaline) ‘mid’-Cretaceous (i.e., 100–90 Ma) plutonic suites in this region to be revised. The Tungsten plutonic suite was previously interpreted to represent magmas derived from the anatexis of supracrustal rocks (e.g., Mortensen et al., 2000; Hart et al., 2004a), but may instead be of infracrustal anatectic origin. There is some mineral compositional (apatite, biotite, zircon) evidence for a supracrustal origin for some of the northeast examples of the Tungsten plutonic suite (which is similar to the supracrustally derived Hess– MacMillan sub-suite of Mayo and Tombstone suite age5). Other isotopic evidence, however, is more compatible with a shared petrogenesis with the I-type Mayo plutonic suite. As an origin by fractionation of the Mayo suite magmas was deemed unlikely for various reasons (including cf. Clemens and Stevens, 2012), the Tungsten suite is interpreted to comprise smaller melt fractions from the same (or a slightly more felsic) metaigneous protolith. The Tay River plutonic suite is also interpreted to be of infracrustal origin, partly due to the common presence of hornblende in these rocks and an association with coeval volcanic rocks and intermediate–mafic phases. Although this study cannot conclusively establish that the Tay River plutonic suite is derived from metaigneous rocks in the crust (mainly because the observed high positive δ18O are atypical of infracrustal magmas; e.g., Taylor and Sheppard, 1986), nearly all of the geochemical and isotopic evidence presented here are compatible with an origin by high temperature partial melting of intermediate-composition infracrustal rocks. The post-arc, predominantly I- type plutonic suites (Tay River, Tungsten, Mayo, and high-K calc-alkaline members of Tombstone) were most likely sourced from metaigneous basement of mainly intermediate composition (SiO2 = ~60–65 wt.%). The change to a predominantly I-type magmatic affinity northeast of the Anvil and Hyland plutonic suites implies that there is either a change in the basement domain (i.e., mainly supracrustal to mainly infracrustal), or that post-arc anatexis 463 occurred at greater depths (i.e., in the metaigneous lower crust thought to underlie this region; Cook et al., 2004). There also appear to be several changes in the composition of the infracrustal basement across the Selwyn Basin. The eastern Selwyn Basin, for example, is less isotopically evolved with respect to neodymium and strontium than the western Selwyn Basin, as evidenced by the Tay River plutonic suite. Another change in the infracrustal basement composition, from the relatively isotopically primitive south (Nd, Sr, O) to the more isotopically evolved north (Nd, Sr, O) in the eastern Selwyn Basin, is also reflected in whole rock chemistry and zircon compositions in the Tay River, Tungsten, Mayo, and Tombstone plutonic suites. Finally, the radiogenic isotopes in the Mayo and Tombstone plutonic suites indicate that the infracrustal basement to the northwest is the most isotopically primitive (Nd, Sr, and O) and possibly of oceanic affinity. Lead isotopes, however, are not necessarily correlated to neodymium, strontium, or oxygen isotopes. In the Selwyn Basin, lead isotopic compositions of intrusive rock units are the least radiogenic in the northwest (mainly with respect to 207Pb and 208Pb), generally less radiogenic in the northeast (mainly with respect to 206Pb), and are the most radiogenic in the southeast. One sample from the western Selwyn Basin indicates that this region may be underlain by the least radiogenic basement, with respect to lead, although this needs to be confirmed with additional isotopic work in this region. Since more radiogenic lead compositions are not necessarily correlated with more evolved neodymium and strontium (as well as ‘heavier’ oxygen), the lead isotopes probably reflect differing metamorphic histories experienced by these basement regions of the Selwyn Basin (although differences in the age, and (or) LILE concentrations in the basement could also be important controls on this isotopic system). The association between ‘mid’-Cretaceous igneous rocks and regional structures throughout much of the eastern and northern Selwyn Basin (e.g., the Hess–MacMillan fault system; Fig. 8.1b) also appears to have been important in the localization (and possibly genesis) of at least some of the I-type magmas. Partial melting of hornblende- ± biotite- bearing, infracrustal rocks also necessitates heat from the mantle, but the progression of ‘mid’-Cretaceous magmatism farther inboard of the continental margin with time requires a similar progression of heat flux with time. This additional heat up to 700 km farther southwest of the former (Early Cretaceous) continental magmatic arc cannot be attributed to vigorous asthenospheric convection above a northeast-dipping, subducting plate if arc magmatism—and, therefore, subduction—have ceased. The increased distance of ‘mid’-Cretaceous igneous rocks from the continental margin, the high temperatures required to general partial melting of intermediate crustal rocks, and the preceding termination of continental arc magmatism to the west are compatible with a cessation to normal subduction west of the Selwyn Basin at ~100 Ma. Furthermore, uplift in the Cordilleran orogen between ~100–93 Ma has also been indicated by foreland basin analysis (Thomson et al., 2011). Therefore, a flattening of the subducting plate is suggested here as one possible mechanism for moving the region of mantle hydration above the subducting slab farther inboard (Fig. 8.1b). Convection in the asthenospheric tongue or mantle wedge above a northeast-migrating flat slab would have aided heat transfer from the hot asthenosphere to the upper mantle, and ultimately into the (lower?) crust, both conductively and advectively (as late-stage high-K calc-alkaline to shoshonitic mafic dykes associated with the Tay River, Tungsten, and Mayo suite intrusions). From ~93–90 Ma, a subordinate but significant volume of mantle-derived magmas associated with the Tombstone plutonic suite mark the end of ‘mid’-Cretaceous magmatism in the Selwyn Basin. This final pulse of both mantle and infracrustally 464 sourced magma may have been related to tearing or delamination at the leading edge of the flat slab (Fig. 8.1c), although the Tombstone plutonic suite could also have been focused by lithosphere-scale structures near the basin– platform transition. Although they were undoubtedly locally important, these structures do not appear to have been capable of producing the widespread ‘mid’-Cretaceous intrusions across the entire Selwyn Basin. Decompression and local pull-apart segments along these transpressional faults (allowing magma generation and ascension; e.g., Glazner, 1991) can only have been a local feature. 8.1.3 Summary of Metallogenic Associations Early Cretaceous, S-type plutonic suites across the northern Cordilleran miogeocline are typically associated with proximal tungsten ± molybdenum, and relatively distal argentiferous lead–zinc occurrences. Few, if any, of these occurrences appear to be economic, with the possible exception of the Northern Dancer tungsten–molybdenum porphyry (associated with the Cassiar plutonic suite). Conversely, infracrustally derived ‘mid’-Cretaceous igneous rocks (Tay River, Tungsten, and Mayo plutonic suites) have a strong base metal signature, particularly in the eastern Selwyn Basin where two world-class tungsten–(copper) skarns are associated with the Tungsten plutonic suite. Other I-type suites, specifically the Mayo and Tombstone plutonic suites, have a well-defined gold association. Bismuth–tellurium–arsenic–antimony enrichments have also been observed in association with I-type intrusions from both Tungsten (e.g., Cantung; Rasmussen et al., 2007) and Mayo suite intrusions (Hart et al., 2004a, 2005). The uranium–thorium–REE–fluorine signature of the Tombstone plutonic suite, however, is unique to this plutonic suite in the Selwyn Basin, and probably reflects the mantle derivation of alkaline phases belonging to this plutonic suite. Although the A-type plutonic suites in the Cassiar Terrane are associated with mineral occurrences of similar metal tenor to the Cassiar plutonic suite, there are some distinctions. The Seagull plutonic suite, for example, has an important association with tin mineralization that is typical of other A-type granites elsewhere (Chappell et al., 2000). Molybdenum–tungsten mineralization is found in association with Late Cretaceous plutons of the Rancheria plutonic suite in the Cassiar Terrane, although one U–REE mineral occurrence associated with these intrusions is similar to the metal tenor of the Tombstone plutonic suite in the Selwyn Basin. Eocene magmatism in east–central Yukon is typically associated with precious and base metal-bearing vein or replacement styles of mineralization. 8.2 MAGMATIC FLUID COMPOSITIONS AND THERMAL HISTORIES IN THE SELWYN BASIN Magmatic apatite and biotite compositions also provided some information on the petrogenesis and degree of differentiation of plutonic rocks in the eastern Selwyn Basin. Increasing MnO (± FeO, XF:XCl) with decreasing CaO in apatite appears to broadly correspond to an increasing degree of differentiation, although Fe# versus total aluminum in biotite is a better constrained indicator of inter- and intra-plutonic differentiation. Relatively elevated MnO, FeO, and fluorine, and depleted SrO, chlorine, and sulfur, in magmatic apatites also indicated that most of the plutons examined in the Selwyn Basin are of crustal origin—either derived from supracrustal rocks as mainly felsic magmas, or from less siliceous infracrustal rocks and fractionated to felsic magma compositions. It was apparent that less siliceous members of the Mayo and Tay River plutonic suites were derived from chlorine-rich infracrustal 465 sources, which implies that more siliceous members of these plutonic suites could represent highly fractionated magmas, or smaller melt fractions (in accordance with newer interpretations of felsic granite petrogenesis; Clemens and Stevens, 2012), from the same infracrustal rocks. Overall, however, it was not possible to distinguish between S-type and I-type affinities in highly siliceous granitoids using apatite and biotite compositions alone. High F/Cl also typifies apatite compositions for more differentiated intrusions in the eastern Selwyn Basin. Whether source composition, magmatic fractionation, the presence of CO2, or even the effects of analytical difficulties (or some combination of these and other factors listed in Section 7.4.2.1) are the driving forces behind the high F/Cl is uncertain in most cases. In some instances, however, relatively elevated F/Cl ± SO3 contents were observed in apatites from samples associated with mineralization. Perhaps the retention of volatiles until later stages of crystallization in some of these intrusive phases allowed apatite to equilibrate with a chlorine- ± SO3-rich(er) MVP capable of transporting metals. The Tombstone suite intrusions are also typified by high F/Cl, which is typical of A- type magmas and could be related to either an F/Cl enriched (i.e., restitic) melt source, or the presence of abundant CO2 in these magmas. Unlike apatite, attempts to use biotite compositions alone as a proxy for the composition of a magma and (or) the fluid that exsolved from the magma may have significant uncertainties, nor did this study identify any specific compositional features in magmatic biotite that may be used to predict an association with intrusion-related mineralization. The use of magmatic fluid compositions and thermal histories of plutons (based on the difference in U–Pb and 40Ar/39Ar ages in zircons and biotites, respectively) was first done as one component of a regional resource assessment study (Rasmussen et al., 2007). The objective was to identify intrusions in the area that had a heightened potential for associated mineralization. Although the reconnaissance-scale nature of the study precluded detailed analysis of each intrusion, a representative sample from each pluton was dated by U–Pb and 40Ar/39Ar methods, and apatites and biotites from the sample were analyzed mainly to determine their fluorine, chlorine, and hydroxyl (by stoichiometry) contents. Given the limits associated with comparing LA–ICP–MS U–Pb and 40Ar/39Ar ages, intrusions with minimum and absolute cooling periods >1 and >2 m.y., respectively, were considered to be ‘extended’ or prolonged. The thermal histories of these intrusions were remarkably effective at identifying intrusions associated with known tungsten skarn mineralization, as all of those examined exhibited prolonged cooling based on the minimum and absolute differences in the U–Pb and 40Ar/39Ar ages66. The mineral compositional data also allowed estimates of magmatic fluid composition―approximated by the activity ratios log[aHCl/aHF] and log[aHOH/aHF] calculated from apatite and biotite compositions―for each of these samples to be determined following the thermodynamic scheme of Zhu and Sverjensky (1991, 1992). Magmatic fluids associated with intrusions in the study area typically evolved to lower HCl/HF and H2O/HF compositions with differentiation. Although the regional estimates of magmatic fluid composition were easily correlated with the composition of the intrusion, these data by themselves were not particularly helpful in terms of identifying metallogenically prospective intrusions. This is probably due to the reconnaissance-scale of this study, and the incorrect assumption that one representative intrusive sample is associated with nearby mineralization. 66 The calculated minimum cooling periods take into account the smallest possible difference between the U–Pb and 40Ar/39Ar ages when 2-sigma errors are also taken into account, and absolute cooling periods are the difference between the U–Pb and 40Ar/39Ar ages when 2-sigma errors are ignored. 466 More detailed fluid compositional work from five composite plutons―including both barren and mineralized bodies―in the eastern Selwyn Basin, however, has demonstrated that in tungsten-mineralizing systems there may be distinct magmatic fluid compositional parameters that are compatible with the buildup of a hydrous saline brine capable of segregating tungsten from the magma and transporting it to sites of mineralization. Two barren plutons demonstrate that fluids in silicic magmas become HF enriched relative to HCl with increasing differentiation (as was also observed regionally in Chapter 7), although re-equilibration with late-stage sub-solidus or hydrothermal fluids may obscure this trend. The three peraluminous intrusions with associated tungsten skarn mineralization―including the world-class Cantung deposit―also show a relationship between increasing HF/HCl and magmatic evolution. However, magmatic apatite and biotite in individual intrusive phases that are proximal to mineralization equilibrated with an earlier HCl- and H2O-rich magmatic fluid and a later HCl- and (particularly) H2O-rich magmatic fluid (which is consistent with the progression of magmatic crystallization at supercritical conditions). These anomalous fluid compositions were only observed in magmas directly associated with mineralization; none of the other intrusive phases in the same composite plutons had similarly anomalous apatite- and biotite-based activity ratios. Therefore, the identification of intrusive phases that did exsolve anomalous magmatic fluids enriched in HCl and H2O could, therefore, be a predictor of nearby tungsten skarn mineralization. Overall, given that magmatic fluid compositions were only successfully used to identify intrusive phases associated with tungsten mineralization in the detailed study, it appears that the regional identification of prospective intrusions would be best done by identifying plutons with extended cooling periods first. Follow-up apatite- and biotite-based fluid compositional work on intrusions that experienced prolonged cooling could then help to identify specific intrusive phases near which tungsten mineralization is more likely to be found. 8.3 PLATE INTERACTIONS, DEFORMATION, AND MAGMATISM: A PROPOSED TECTONO–MAGMATIC MODEL FOR SYN- TO POST-ACCRETIONARY IGNEOUS ROCKS IN THE NORTHERN CORDILLERAN MIOGEOCLINE None of the models previously proposed for the tectonic, magmatic, and metallogenic evolution of the northern Cordillera (summarized in Chapter 1) fully account for all of the observed features exhibited by syn- to post- accretionary magmatic suites in the inboard regions. With the new geochronological, geochemical, and isotopic datasets generated and compiled by this study, an updated or modified interpretation of the tectonic regime into which these igneous rock suites were emplaced is now possible. The timing of changes in relative plate motion vectors between the western margin of North America and oceanic plates to the west appears to have controlled changes in the deformational regime of the northern Cordillera, which in turn closely coincide with shifts in the pattern and the composition of syn- to post-accretionary magmatism in the northern Cordillera. All references to direction in the following discussion are relative to modern day coordinates. 8.3.1 Early Cretaceous (~115–100 Ma) The dynamics of plate convergence along the western margin of North America in the Early Cretaceous created a 467 mainly transpressional regime along the length of the northern Cordillera (at least as far south as the northwest United States; Wyld and Wright, 2001). Oblique convergence was probably responsible for initiating dextral strike- slip deformation along margin-parallel structures, beginning as early as 140 Ma and well underway by ~120 Ma (Gabrielse et al., 2006; Colpron et al., 2011). An increased rate of convergence also occurred at ~120 Ma (Kelley, 1993), and may be related to a nearly coincident flare-up in Early Cretaceous continental arc magmatism resulting from east-dipping subduction below the Cordilleran margin (Fig. 8.1a). At ~120 Ma, therefore, oceanic crust is interpreted to have been moving relatively rapidly to the east–northeast, and colliding approximately orthogonally with the margin of the north–northwest moving North American plate (Fig. 8.1a; e.g., Engebretson et al., 1985; Kelley, 1993; Evenchick, 2001)67. The earliest arc magmas may have been generated locally at 123–115 Ma (the Teslin plutonic suite), but (north)east-dipping and related continental subduction arc magmatism were well underway by 115 Ma when the voluminous Whitehorse, Dawson Range, and Casino plutonic suites were emplaced (Fig. 8.1a). More ‘inboard’ plutonism related to high heat flow in the transpressional back-arc was also initiated at ~115 Ma and is represented by the supracrustally sourced Cassiar, Anvil, and Hyland plutonic suites (Fig. 8.1a). Dextral faults appear to have had a significant role in the localization of the back-arc plutonic rocks, where deformation along these structures appears to have been synchronous with much of the syn- to post-accretionary plutonism (Gabrielse, 1991; Mihalynuk, 1999; Gabrielse et al., 2006; Mair et al., 2006b; Colpron et al., 2011). Arc and back-arc magmatism appear to have terminated synchronously, at ~104 Ma in the Cassiar Terrane region, and at ~100 Ma in the Selwyn Basin region. 8.3.2 ‘Mid’-Cretaceous (~103–94 Ma and ~99–90 Ma) Following the cessation of subduction-related arc–back-arc magmatism, a significant change to the composition and style of igneous rocks in the northern Cordilleran miogeocline occurred, but manifested differently in the Cassiar Terrane and Selwyn Basin regions. In the Cassiar Terrane, a long, narrow belt of A-type plutons (the 103–94 Ma Seagull plutonic suite) was emplaced east of the Teslin Tectonic Zone and the Big Salmon Fault near the plate boundary between the autochthonous Yukon–Tanana terrane and the allochthonous Cassiar Platform (Fig. 8.1b). In the Selwyn Basin, the locus of magmatism shifted inboard, and changed to a dominantly I-type affinity comprising the ~99–93 Ma Tay River, Tungsten, and Mayo plutonic suites (Fig. 8.1b). The youngest and most inboard plutonic rocks emplaced in the Selwyn Basin, however, is the ~93–90 Ma Tombstone plutonic suite, which is of mixed I- and A-type affinity and requires the involvement of enriched mantle derived magmas (Fig. 8.1c). One unresolved question that is important in terms of Cordilleran plate dynamics and syn- to post-accretionary magmatism in the northern Cordillera is the timing of the accretion of the Insular Superterrane. Although it may have occurred relatively early in the southern Cordillera (i.e., Middle Jurassic; McClelland et al., 1992), this accretionary event apparently migrated north with time (e.g., Trop and Ridgway, 2007). In the northern Cordillera, estimates for the timing of its accretion range from the Early Jurassic (e.g., Plafker and Berg, 1994) to the Late Jurassic (e.g., Trop and Ridgway, 2007) to the latest Cretaceous (e.g., Cole et al., 1999; Mezger et al., 2001). Here, 67 Pre-84 Ma oceanic plate vectors are, however, poorly constrained due to limitations imposed by a locally preserved and long-lived (120–84 Ma) seafloor magnetic anomaly (e.g., Wyld and Wright, 2001). 468 based on the timing and changes in composition documented in the Early to ‘mid’-Cretaceous intrusive rocks, the closure of the Gravina ocean basin (separating the Insular Superterrane from northwest North America) may have coincided with the end of subduction-related, arc–back-arc magmatism at circa 104 Ma west of the Cassiar Terrane. West of the Selwyn Basin, this accretionary event may have marked the end of post-arc, I-type magmatism at circa 93 Ma (Fig. 8.1). In both regions, this event would have pre-dated or been synchronous with the emplacement of post-collisional mantle-derived A-type plutonic suites. These estimates of the exact timing for the Insular Superterrane accretion in the northern Cordillera are consistent with a detailed review of sedimentation patterns and timing in southern Alaska, which suggested that if it had not yet accreted to North America by the Early Cretaceous (i.e., ~140 Ma), this superterrane was in “close proximity” to the North American margin at this time (Trop and Ridgway, 2007). Thus, the 103 Ma and 93 Ma onset of A-type plutonism in the Cassiar Terrane and Selwyn Basin, respectively, may have been linked to the accretion of the Insular Superterrane. If so, these post-arc—and, therefore, post-collisional— plutonic suites occurred in response to the end of subduction (due to the buoyancy of the collided Insular Superterrane). Regardless of the cause, however, the ‘mid’-Cretaceous magmatic events reflect an important change to plate dynamics west of North America, as well as a major deviation in the distribution and composition of magmatism in the southern (Cassiar Terrane) and northern (Selwyn Basin) regions of the northern Cordilleran miogeocline. 8.3.3 Late Cretaceous (~85–67 Ma) Plate vector models suggest the Kula plate detached from the northern Farallon plate at circa 85, and exhibited a higher rate and more oblique angle of dextral convergence along the Laurentian margin than the Farallon plate (Engebretsen et al., 1985; Gabrielse and Yorath, 1991; Kelley, 1993; Miller et al., 2002). The timing of Kula plate formation coincides with the emplacement of the A-type Rancheria plutonic suite as a narrow plutonic belt associated with a regional structure (Fig. 8.1d). The timing of this plutonic suite also immediately post-dates inboard’ deformation along the Hyland fault system in the Selwyn Basin (Chapter 7), and generally coincides with the gradual change from a primarily transpressional to transtensional deformational regime affecting the northern Cordillera (e.g., Woods and Davies, 1982; Gabrielse and Yorath, 1991; Kelley, 1993). Despite the emplacement of post-orogenic or anorogenic A-type plutonic suites, subduction and associated continental arc processes and (or) lithospheric delamination may have occurred at the western margin of North America immediately after this time (e.g., the Carmacks Group; Fig. 8.1d). 8.3.4 Latest Cretaceous to Eocene (~67–47 Ma) Changes in the northern Cordilleran stress field through the Paleocene–Eocene were the cause of large transverse offsets along major dextral structures, beginning with the most easterly structures (e.g., the Tintina Fault) and migrating to the most westerly structures (e.g., Denali Fault; Fig. 8.1; Gabrielse, 1991; Hart et al., 2004a). Localized heating in the southeast Selwyn Basin at this time was likely to have been related to the emplacement of unexposed, 469 latest Cretaceous intrusions (perhaps similar in composition to the coeval McQuesten plutonic suite and Rackla pluton in the northwest Selwyn Basin; Chapter 7). Also at this time, convergent margin processes and continental arc magmatism (e.g., the Coast Plutonic Complex, or Sloko–Skukum volcanic suite; Fig. 8.1e) probably lasted until ~45 Ma, although the angle and rate of convergence likely varied (e.g., Engebretson et al., 1985; Kelley, 1993; Madsen et al., 2006). By 45–40 Ma, oblique subduction of the Kula―and related hypothesized plates, including Resurrection ± Eshamy (formed at ca. 65 and 45 Ma, respectively)―ended with the subduction of the Kula–Farallon spreading centre (Fig. 8.1e; e.g., Haeussler et al., 2003; Madsen et al., 2006). North–northwest motion of the Pacific plate (and unsubducted portions of the Kula plate) relative to the North American plate—at a much reduced rate— resulted in a highly oblique to transform plate boundary ca. 40 Ma (Engebretson et al., 1985; Thorkelson and Taylor, 1989; Gabrielse and Yorath, 1991; Bradley et al., 1993; Kelley, 1993; Miller et al., 2002; Madsen et al., 2006). Increased plate coupling at the transform plate boundary was likely responsible for movement along major dextral shear structures and localized extension in Yukon and northern British Columbia along Cenozoic strike-slip structures (e.g., Dewey, 1980; Beck, 1983; Gabrielse and Yorath, 1991; Miller et al., 2002; Haeussler et al., 2003; Madsen et al., 2006). During Paleocene–Eocene uplift, bimodal volcanic and sub-volcanic rocks occurred mainly in fault-controlled basins, particularly near with the Tintina Fault (Fig. 8.1e; Gabrielse, 1985; Jackson et al., 1986; Yorath, 1991; Gabrielse et al., 2006). Small, isolated, felsic plutons of A-type affinity—including the Black River plutonic suite—were emplaced near structural discontinuities (strike-slip faults or basement structures; Fig. 8.1e; Gabrielse, 1991). 8.4 SUGGESTIONS FOR FUTURE WORK Despite the difficulties with accessibility, the northern Cordillera provides a world-class example of well-exposed, voluminous, syn- to post-accretionary, intermediate–felsic, magmatism (and mineralization). Further studies of the age, composition, and petrogenesis of this granitoid province will not only lead to an imporved understanding of the tectonic and metallogenic evolution of this region, but will also contribute to the general knowledge of continental granitoid provinces world-wide. Future investigations will help to address the remaining questions regarding granitoid petrogenesis and basement compositions, many of which were raised in this dissertation. 8.4.1 The Petrogenesis of Syn- to Post-Accretionary Magmatism Additional geochronological, geochemical, and isotopic data will provide more information and constraints on syn- to post-accretionary igneous rocks, and their origin, throughout the northern Cordillera. Some of this work has recently been undertaken, with a focus on the western portion of the northern Cordillera where primarily continental magmatic arc rocks were emplaced (e.g., Mortensen and Hart, 2010; Ciolkiewicz et al., 2012). Additional regional studies, however, are required for plutonic rocks in the northern Cordilleran miogeocline, particularly in east–central Alaska, the north–central and western Selwyn Basin, and north–central British Columbia. More specifically, there is a dearth of information on the Anvil and Hyland plutonic suites in the western Selwyn 470 Basin, despite the large regional extent and magma volumes associated with these two plutonic suites. Although the Anvil batholith (of the Anvil plutonic suite) has been well described by previous workers (e.g., Pigage and Anderson, 1985; Pigage, 2004), petrogenetic interpretations for the Anvil and Hyland plutonic suites are based on extremely limited data. Additional sampling, petrographic, geochemical, and isotopic work would help to better constrain the compositional range exhibited by these two S-type plutonic suites. Similarly, additional isotopic data for the Tay River plutonic suite (and associated South Fork volcanic rocks) in the western Selwyn Basin would help to confirm whether the differences detected between this suite in the western and eastern Selwyn Basin are real and related to distinct basement domains below these regions. Zircon compositional investigations discussed in Chapter 5 throughout the eastern Selwyn Basin were only reconnaissance in-scale, but the results of this work did help to confirm (and challenge) earlier petrogenetic interpretations based on mineralogical, geochemical, and isotopic datasets. A regional study of zircon compositions (including Hf and O isotopes) for all of the crustally derived syn- to post-accretionary plutonic suites in the northern Cordilleran miogeocline, therefore, could also prove to be useful. In the Cassiar Terrane, syn- to post-accretionary plutonic rocks have been sub-divided into four temporally and compositionally well constrained suites. However, the interpreted petrogenesis of the three A-type plutonic suites is largely incompatible with the isotopic compositions exhibited by these plutons, even though it is consistent with current thoughts on the genesis of chemically similar A-type magmas. Unfortunately, the available isotope data for these A-type suites are too limited to draw major conclusions regarding the petrogenesis of these rocks. There are no lead or neodymium isotopic data available for the Seagull plutonic suite, and isotopic data for the Rancheria Black River plutonic suites are restricted to only 1–3 samples. Again, a study of zircon texture, composition, inheritance (if present), and oxygen (as the O isotopes in quartz could have been compromised by late-stage fluids in these suites) isotopic compositions would provide additional petrogenetic constraints. Zircon compositional work may help to confirm or refute the suggestion that the parental magmas to the A-type plutonic suites in the Cassiar Terrane originated in the mantle (as opposed to the crust). This mantle versus crustal origin is an ongoing issue in studies of post-orogenic or post-collisional A-type magmatism, largely due to the ease with which the isotopic data are compromised by crustal contamination or assimilation processes (e.g., Turner et al., 1992; Bonin et al., 1998; Frost and Frost, 2011). 8.4.2 Age and Composition of Miogeoclinal Basement The above suggested collection of more extensive, ideally widespread and appropriately spaced, isotopic data for the syn- to post-accretionary plutonic suites would also help to constrain the age, composition, and geographic extent of basement domains beneath the northern Cordilleran miogeocline, particularly the Selwyn Basin. However, zircon compositional work, if done systematically, could also be used to help characterize the composition—as well as the age!—of potential crustal magma sources. The basement is not exposed anywhere in the northern Cordillera; therefore, information on its age and composition, particularly the lower to middle crust from which the syn- to post- accretionary igneous rocks are thought to have been largely derived, is extremely limited. Better temporal and compositional constraints on this buried crust would prove useful for both geophysical interpretations and for reconstructions of pre-Cretaceous tectonostratigraphic settings. Zircon compositions may best define areas of the 471 northern Cordilleran miogeocline that are underlain by mainly infracrustal basement of Laurentian affinity, and could be used to help delineate distinct basement domains, including the western extent of crystalline Laurentian basement. It is possible that metaigneous basement below most of Selwyn Basin is thicker than previously thought (e.g., Cook et al., 2004). This, combined with the apparent isotopically distinct domains of metaigneous rock sampled by magmatism across the Selwyn Basin, add previously unrecognized complexity to the possible geometry and composition of this “lower-plate” (Cecile et al., 1997) rifted margin. If zircon inheritance also indicates that a metaigneous crust typifies some or all of this basement, the age and composition information of such a crust that could be derived from inherited zircons would allow basement domains to be mapped and characterized (similar to the results of geophysical and drill core mapping of basement below the western Canada Sedimentary Basin through most of Alberta and northeast British Columbia; e.g., Pana, 2002; see Fig. 3.2b). A detailed study of the age and distribution of zircon inheritance in the crustally derived S-type plutonic suites, and a comparison of these results to provenance studies for sedimentary rocks in the northern Cordillera, would also provide new information on the age and composition of basement rocks. This type of work might identify specific supracrustal package(s) that were involved in partial melting, and that were likely present in the lower to middle crust (i.e., at depths > 25 km) during the Cretaceous. This type of work would be particularly interesting for the Cassiar plutonic suite, which is interpreted to have a mainly supracrustal protolith of Paleozoic age. Unlike the Selwyn Basin, however, the Cassiar Platform was formed on an “upper plate” rifted margin (Cecile et al., 1997) that should not be characterized by a thick package of Paleozoic sedimentary rocks. Regardless, it is inferred here that supracrustal rocks of probable Paleozoic age are currently present at depths of at least 15 km in this region. 8.4.3 A Final Comment on ‘Mid’-Cretaceous Magmatism in the Selwyn Basin The U–Pb and 40Ar/39Ar geochronological data generated by this study, as well as a compilation of published and unpublished U–Pb and 40Ar/39Ar age determinations for syn- to post-accretionary igneous rocks in the northern Cordillera, have substantially added to our understanding of Early Cretaceous to Eocene magmatic events in this region. Similar transitions from magmatism with medium- to high-K calc-alkaline I- and (or) S-type compositions, to more alkaline A-type compositions have been documented in many orogens (e.g., King et al., 1997; Bonin et al., 1998; Liégeois et al., 1998; Jung et al., 1998; Soesoo and Nicholls, 1999; Paquette et al., 2003; Bonin, 2004). In all of these instances, this transition occurs in response to a changing geodynamic regime controlled by plate interactions (e.g., Rogers and Greenberg, 1990; Liégeois et al., 1998; Bonin, 2004), and in some examples it follows the accretion of buoyant continental crust (e.g., Davies and von Blanckenburg, 1995). Igneous rocks in the Selwyn Basin comprise a broad continuum of plutonic suites with overlapping ages and geographic distribution. Here, older, medium- to high-K calc-alkaline, back-arc, S-type plutonic rocks transition to mainly high-K calc-alkaline, post-arc, I-type plutonic (and minor volcanic) rocks with time, and towards the northeast. A mixture of high-K calc-alkaline I- type, and alkali-calcic to alkaline A-type plutons emplaced farthest to the northeast marks the termination of syn- to post-accretionary magmatism in this region (at least until the latest Cretaceous). With few exceptions (e.g., the Hess–MacMillan and Coal River sub-suites, and A-type magmas associated with the Tombstone plutonic suite), the lack of major geochemical differences between ‘mid’-Cretaceous plutonic suites across the Selwyn Basin has made 472 conclusive distinctions between plutonic suites difficult, particularly in the eastern Selwyn Basin where the plutonic suites overlap temporally and geographically. Despite apparent isotopic variations related to geography, the gradual compositional changes and the short (~10 m.y.) period of magmatic activity in this region means these igneous cannot be satisfactorily compartmentalized into individual plutonic suites that are distinct in terms of their timing, geographic distribution, composition, and melt source. Although our understanding of igneous rocks in the Selwyn Basin has been much advanced by this study (and previous work; e.g., Lang, 2001; Hart et al., 2005; Heffernan, 2004), it may be that the use of an all-inclusive designation—perhaps a modified definition of the “Selwyn Plutonic Suite” first introduced by Anderson (1983) to encompass all ‘mid’-Cretaceous intrusions in the eastern Selwyn Basin—would be more appropriate for compositionally similar, mainly infracrustally derived, post-arc plutonic rocks across much of the Selwyn Basin. 473 Figure 8.1. (next 5 pages) Tectonic setting of syn- to post-accretionary magmatism in the northern Cordillera. Tectonostratigraphic elements, map lines, and abbreviations are as in Figure 1.1. Black shading represents oceanic crust (labeled with the inferred plate name), and the orientation and shape of the Insular Superterrane (IST) is unknown and represented schematically by an ellipsoid. The red dashed line(s) in the map view shows the extent of the cross-sectional view(s). In the cross-sectional view(s): CC = continental crust; MOHO = Mohorovic Discontinuity; LM = lithospheric mantle; EM = enriched mantle; AS = asthenospheric mantle. Forces acting on the subduction slab are indicated by black arrows, and the direction of mantle-flow related to viscous coupling with the subducting slab and thermal buoyancy are indicated by solid and dashed red arrows, respectively (after Currie and Hyndman, 2006). Depths are in kilometres, the extent of upper crustal Neoproterozoic to Paleozoic rocks are based on the interpretation of Cook et al. (2004), and schematic normal faults in the Laurentian basement are illustrative deformation styles along the rifted margin (after Hansen et al., 1993; Cecile et al., 1997). In the map view: vectors for the direction of plate travel and the relative magnitude of plate travel are indicated by the large white (oceanic plate) and black (continental plate) arrows (after Kelley, 1993), the proto-Tintina and Denali faults are indicated by the heavy dashed lines, and dextral and thrust faults that are known to have been active in each of the stages below are illustrated with thick black lines (after Gabrielse et al., 2006). Note that a large amount of dextral strike-slip displacement occurred on dextral faults known to have been active by the late Early Cretaceous, and the displacement of continental margin and terrane rocks in the ‘mid’-Cretaceous has not been quantified, nor shown here. a) The Early Cretaceous (~115–104 Ma). Pre-collisional arc magmatism is represented by the red plutonic suites (W = Whitehorse, and DR = Dawson Range, including Casino) and back-arc magmatism is represented by the purple plutonic suites (C = Cassiar, A = Anvil, H = Hyland, and YTU = Yukon– Tanana Uplands). b) The ‘mid’-Cretaceous, from ~100–94 Ma. Post-arc (and possibly post-collisional magmatism) is represented by the blue plutonic suites (S = Seagull, TR = Tay River, Tu = Tungsten, M = Mayo, F–S = Fairbanks–Salcha). c) The ‘mid’-Cretaceous, from ~93–90 Ma. Post-arc (and possibly post-collisional magmatism) is represented by the yellow Tombstone (To) and Livengood (L) plutonic suites. d) The Late Cretaceous (~84–67 Ma). Shades of green denote plutono–volcanic suites of differing age and composition (eC = ‘early’ Carmacks, lC = ‘late’ Carmacks, SL = Surprise Lake, and R = Rancheria). e) The latest Cretaceous to Eocene (~67–47 Ma). 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Annual Review of Earth and Planetary Sciences 14, 493–571. 516 APPENDICES 517 Appendix A – Sample Descriptions 518 A.1 Geochronological samples Sample locations, rock names, and brief textural and mineralogical descriptions for all samples that were dated using U–Pb and (or) 40Ar/39Ar geochronological analysis (as well as geochemical and isotopic analysis). New names assigned to unnamed plutons and batholiths by this study are italicized. 519 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host Cassiar Terrane and western Selwyn Basin (Geochronology Samples): KR-07-01 - Dycer Creek pluton (61.4461, 133.9822) f-m-gr mc-phyric K=P=Q bt>ms monzogranite 0.13 I R Ce-La-Th-(Ca-Ta- Nb) mnz, Ti-Ca- Pb-(Fe) uran, rt, zrn, ap bt is moderate alt'd to chl (blue); well-developed TT or minor bleb perth in un-TT mc; mc ph-xst up to 1x2cm; moderate ser-ms in pl; minor-moderate clay speckling in mc proximal: 105F097-KIDLARK and 105E031-HITCHENS (W skarns) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-02 - Quiet Lake batholith (61.1526, 133.7503) m-gr equigr (mc- phyric) Q>K=P bt>ms monzogranite 0.13 M R Ce-La-Th-(Ca) mnz, Fe>Mg>Mn grt, Mg>Fe px, rt, zrn, ap some myrm; qtz-rich zones made up of composite grains; ms can be relatively c-gr; perth and local TT in mc; mc sub-ph-xst <1cm in size; moderate ser in pl; minor to moderate clay speckling in mc; ap is c- gr, rt is in mc proximal: 105F003/004- WOPUS/GOPHER (Ag-Pb ± Mo vein), 105F104-ALVIN (W porphyry), 105E030-SALMON (W skarn) Devono-Mississippian Y-T (Snowcap) - crustal block KR-07-03 - Quiet Lake batholith (61.0436, 133.3706) m-gr equigr K>/=Q>P bt>ms leucogranite 0.05 nd R Ta xeno, Ce-La- Th-(Ca) mnz, Fe>Mn grt, zrn perth and TT in mc; minor to moderate ser in pl; minor clay speckling in mc; ms is relatively f-gr but primary; moderately abun zrn; xeno as small inclusions in bt cut by leucocratic dykelets proximal: 105F105-CRATER (W vein), 105F104-ALVIN (W porphyry), 105F079-MURPHY (U porphyry), 105F002-MOBS (Ag-Pb vein), 105C022-LINDSAY (Cu unknown) Devono-Mississippian Y-T (Snowcap) - crustal block KR-07-04 - Nisutlin batholith (61.2118, 133.6426) f-gr equigr Q>/=P>K ms>bt leucogranite 0.03 I R toz, xeno (inclusions in bt), zrn TT and perth in mc (may be zoned); minor to moderate ser in pl; strong clay speckling in mc proximal: 105F001-MOLLY (Mo skarn) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-05 - Nisutlin batholith (61.4994, 133.1271) f-m-gr foliated equigr Q>/=K>P bt leucogranite 0.08 nd R Ce-La-Th-(Ca) mnz, Fe>Mg>Mn grt, zrn some myrm; diffuse perth and local TT in mc; lots of evidence for late or sub-s fluid flow (ie, patches of v f-gr qtz-fsp); v f-gr to f-gr foliated bt w/minor chl; v f- gr 2ndary ms veins/clots (1%); moderate sauss in pl; moderate clay specking in mc and local pervasive clay alt'n of fsp lots metamorphic rafts/fingers; pegmatite dykelets proximal: 105F098-FIRST (Cu-Pb-Zn skarn), 105F032-PONY (Ag-Pb vein), 105F094-OBVIOUS and 105F129-HID and 105F092-AYDUCK (W skarns) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-06 - Nisutlin batholith (61.4545, 132.8480) c-gr mc-phyric Q>K>P bt>hbl monzogranite 16.80 M-I > T O Mg>Fe-Ce-La aln, Fe>Nb ttn, U-(Ca- F) thor, betafite, zrn, ap c-gr aggregates of qtz (rounded; 1cm); perth and local TT in mc; abun mc ph-xst averaging 1x2cm; minor to moderate chl (blue) in bt; hbl may be zoned?; moderate ser-ms in pl; v abun and v c-gr ttn (may be zoned); v abun and c-gr mag proximal: 105F011-STORMY (Mo-W skarn) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-07 - Young pluton (61.4332, 132.1898) f-gr equigr (mc-phyric) K=Q=P bt monzogranite 0.16 nd R Fe>Mg-Ce-La-(Th) aln, Ce-La-Th-Ta- Nb-(Ca) mnz, zrn, ap minor myrm; perth in mc; mc ph-xst are inclusion- rich and 0.5x1cm; bt moderate alt'd to chl (blue) ± 2ndary ms and cal; moderate to strong ser in pl; minor to moderate clay speckling in mc; moderate abun and frequently v c-gr zrn and ap cut by aplites and rusty fractures proximal: 105F072-YOUNG (Cu skarn), 105F059-HOGG (Cu vein) Paleozoic NA (Cam. to Triassic) - platform and clastic influx 520 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-07-08 - St. Cyr pluton (61.9970, 133.4387) f-(m)-g equigr P=Q>K bt granodiorite 12.80 M O Fe>Mg-Ce-La aln, Ca grt, zrn, ap ~50% of bt is chl (blue); moderate to intermediate ser in pl; minor to moderate clay specking in mc; abun c-gr mag small plug proximal: 105F037-EVA (W skarn) Paleozoic NA (Cam. to Triassic) - platform and clastic influx KR-07-09 - Fox Mountain pluton (61.9513, 133.4974) f-gr qtz-fsp-bt leucogranitic porphyry 0.21 nd R Ce-La-Th-Ta-Nb- Nd mnz, Mg>Fe grt, Ca>Fe grt, zrn, ap some bt laths up to 12cm but most (~5%) bt in gr- m, up to 50% alt'd to chl (blue) ± ep; some to abun myrm (esp in gr-m); rare mc m-xst up to 0.5x1.2cm; minor to moderate ser in pl in gr-m (ph-xst are pristine) homogenous; shallowly- dipping aplite swarms proximal: 105F037-EVA and 105F034- RISBY (W skarns) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-10 - Big Salmon batholith (61.8400, 133.6943) f-m-gr (mc-phyric) K=Q>/=P bt monzogranite 13.90 M-T > I O Ce-(La) aln, Ca grt, Ce-La-Th-Ta-Nb- (Ca) mnz, ttn, zrn, ap up to 50% of bt is chl (brown > blue) w/minor ep; perth in mc; slightly mc-phyric but <05cm (rare 1x2cm ph-xst); minor ser in pl; minor to moderate clay speckling in mc; abun c-gr mag heterogenous: pegmatitic patches, foliated, xenoliths proximal: 105F114-BIG SAM (W skarn) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-11 - Glenlyon batholith (61.7967, 134.0323) f-gr equigr P=K=Q bt leucogranite 0.08 nd R Fe>Mg grt, (Ta) xeno, Th-(Ca-Ta- Nb) mnz, zrn up to 50% chl (blue) ± cal-ep; minor to moderate ser-ms in pl; minor clay speckling in mc n/a Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-12 - Tintina pluton (61.1701, 131.1821) f-m-gr mc-crystic qtz- phyric Q>K>P bt (sub- phyric) leucogranite porphyry 0.12 I R Ta-Nb-(Fe) rt, Ce- La-Th-(Ca-Ta-Nb) mnz, tur, zrn, ap gr-m is v f-gr and contains abun scattered f-gr clots of 2ndary/sub-s ms (possibly after peraluminous minerals such as grt or crd); rounded qtz grains are <1cm; perth in mc, esp in clay-alt'd grains; mc m- xst range from 1x2 to 1.5x3cm; up to 50% of bt is alt'd to chl; moderate ser in pl cores, occassionally moderate ser-ms in pl; moderate clay speckling in mc (esp larger grains) small plug in hornfelsed shales w/local dykes proximal: 105G006-TINTINA (Ag-Pb- Zn replacement) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-13 - Square Lake pluton (61.1908, 130.7315) m-gr qtz-phyric Q>/=P>/=K bt-ms-tur leucogranite 0.08 nd R Ta-Nb-(Fe) rt, Ce- La-Th-(Ca-Ta-Nb) mnz, tur, zrn, ap minor myrm; perth in mc; scattered c-gr primary ms (<5%); almost all bt (~5%) is chl (blue?) w/some f- gr op; moderate ser-ms in pl; moderate clay speckling in mc n/a Devono-Mississippian YTT inlier (Snowcap) - crustal block KR-07-14 - Black River (F1) batholith (61.0636, 130.6214) m-gr mc-phyric Q=K>>P bt-(ms) monzogranite 0.13 I > M R (Ta) xeno, (Th) uran, Ce-La-Th- (Ca-Ta-Nb) mnz, Fe>Mg grt, zrn, ap trace ms in gr-m, occassionally assoc'd w/2ndary ser; perth in mc, esp ph-xst; mc ph-xst are typically <1x2cm (up to 2x3cm); minor chloritization of bt; moderate ser-ms in pl cores; minor (to moderate) clay speckling in mc; relatively abun and c-gr zrn prominent E-W jointing/fabric proximal: 105G001-MONT (Cu vein), 105G003-BLUEBERRY (Ag-Pb vein), 105G005-SLAM (Zn vein) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related 521 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-07-15 - Meister Lake pluton (60.3380, 130.2167) m-gr equigr K=Q>P ms>bt leucogranite 0.09 M R (Ce-La) aln, Ce- La-Th-(Ca-Ta-Nb) mnz, Nb-Y-Ta-U- Dy-Yb-Th oxide (possibly betafite or formanite- fergusonite or kobeite), zrn, ap perth and abun TT in mc; some ms ragged and f-gr (2ndary/sub-s?) but typically c-gr and magmatic; ~50% bt is alt'd to chl (blue?); moderate ser-ms in pl; minor clay speckling in mc local oxidized fratures proximal: 105B135-GOLDEX (Cu-Mo- Zn vein); distal: 105B114-MEISTER (Pb-Zn-Ag replacement), 105B138/9- HAIRSINE/FAIRFIELD (Pb-Zn unknwn); 105B109-BRODHAGEN (Cu vein) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-16 - Cassiar batholith (60.1564, 130.6306) m-(c)-gr mc-phyric Q>K>P bt>>ms leucogranite 1.06 M > I O Fe>(Mg-Ce-La) aln, Ce-La-Th-(Ca- Ta-Nb) mnz, Ca grt, Ca>>Fe px, zrn, ap perth and TT in mc; mc-phyric typically 1x2cm but up to 15x3cm; trace ms, some apparently f-gr primary and assoc'd w/bt but also some sub-s f-gr patches; bt (greenish) is variably (minor to strongly) alt'd to chl ± ep w/op; moderate ser-ms in pl, esp cores; moderate clay speckling in mc cm-scale f-(m)-gr bt-rich xenoliths proximal: 105B097-URP (W-Mo-Cu skarn), 105B012-SHILSKY and 105B006-LENA and 105B059-FAG and 105B133-SCHELLENBURG (Ag- Pb veins), 105B122-ABBOTT and 105B005-ANNE (Ag veins), 105B060- DEATH (Pb-Zn vein), 105B004- FIDDLER (W vein); distal: 105B003- LUCK (Ag-Pb-Zn replacement), 105B017/016/018- HARDTACK/KODIAK/KERNS (Pb-Ag- Zn veins) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-17 - Cassiar batholith (60.3306, 130.9213) f-(m)-gr equigr (mc- phyric) Q=K>P bt>ms monzogranite 1.90 I > M R Ce-La-Th-(Ca-Ta- Nb) mnz, zrn, ap some myrm; perth and TT in mc; rare mc ph-xst (<05cm); minor ms (~5%), some apparently f-gr primary and assoc'd w/bt but also some sub-s; bt (5-7%) moderate chloritizited (blue); minor ser in pl; anh op 1-cm clots of (v) f-gr bt-rich xenoliths proximal: 105B022-AURORA (Pb-Zn- Ag intrusion-relativelyated), 105B021- HART (Ag-Pb-Zn skarn) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-18 - Cassiar batholith (60.4531, 131.4942) f-(m)-gr equigr (mc-bt- phyric) Q=K>P bt>ms monzogranite 0.13 I > M R Ce-La-Th-(Ca-Ta- Nb) mnz, zrn, ap v rare bt ph-xst<0.5cm; some myrm; perth and TT in mc; rare mc ph-xst <0.5cm; minor ms (~5%), some apparently f-gr primary and assoc'd w/bt, but also some sub-s ms; bt (~5%) is moderately alt'd to chl (blue) w/minor ep; moderate ser-ms in pl; minor clay-speckling in mc; anh op proximal: 105B042-TROUT (F vein), 105B103-THRALL (Mo-Ag porphyry), 105B056-ZAK (Cu-Pb-Zn vein), 105B024-ALMOST (W-vein), 105B077-POOLE (W skarn) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-19 - Marker Lake batholith (60.5113, 130.9552) c-gr mc-crystic K=Q<P bt granodiorite 23.30 M-T > I O Fe>Mg-Ce-La-(Th) aln, ttn, Ca>>Fe px, zrn, ap perth in mc; mc m-xst are 1x2cm; local chl (blue) and ep alt'n of bt; minor to moderate ser-(ms) in pl; minor clay speckling in mc; ap can be v c-gr; v abun and v c-gr ttn; abun euh mag proximal: 105B058-BINGY (Ag-Pb unknown), 105B099-LOGAN (Zn-Ag vein), 105B044-IRVINE (Pb-Zn-Ag skarn), 105B124-SILVER CREEK (Ag vein) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related 522 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-07-20 - Cabin Creek pluton (60.6247, 130.4842) m-c-gr equigr Q=K<P bt granodiorite 4.25 M > I O Fe>Mg-(Ce-La-Th) aln, Ce-La-Th-(Ta- Nb-Ca) mnz, zrn, ap some myrm; perth and TT in mc; rare mc phenocrusts are rounded and ~1.2cm; trace ms is mostly sub-s; minor chl (blue) alt'n of bt; minor ser in pl; minor to no clay-speckling in mc; abun large euh mag proximal: 105B126-ELECTRICITY (U porphyry); 105B046-TUNG? and 105B047-CABIN (W skarns), 105B126-ELECTRICITY (U-porphyry), distal: 105B101-CORDILLERAN (Pb- Zn replacement) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-21 - Allen Stock (60.6232, 130.4115) c-gr mc-crystic K>Q>P bt monzogranite 18.50 M > I O (Ce-La) aln, Ce- La-Th-(Ta-Nb) mnz, F>Fe ttn,Ca- Al-Ta-Nb rt, Nb-Y- Ta-U-Dy/As-Yb- Th-Ca oxide (possibly betafite or formanite- fergusonite or kobeite), Ca grt, zrn, ap perth and TT in mc; trace sub-s ms; minor chl (blue) alt'n of bt; minor to moderate ser ± ms in pl and clay speckling in mc; mc m-xst generally 1x2cm but up to 5x7cm; abun mag up to 0.5cm sampled near U-showing; v green mineral (fuchsite?) in mc m-xst fractures proximal: 105B126-ELECTRICITY (U porphyry); 105B046-TUNG and 105B047-CABIN? (W skarns); distal: 105B101-CORDILLERAN? (Pb-Zn replacement) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-22 - Cassiar batholith (60.0772, 130.8386) v f-gr mylonitic mc- augen K=Q>P bt-ms monzogranite 0.09 I R (Ta) xeno, Ce-La- Th-(Ta-Nb-Ca) mnz rt, zrn v f-gr gr-m w/foliated intermediate ser-ms-(op) alt'n; ms may be locally c-gr but is metamorphic, and minor f-gr ms is more common; perth and TT in mc; mc augen are rounded and 0.5-1cm; trace chl (after bt w/ms-op?); moderate ser ± ms in blocky pl; minor clay-speckling in mc proximal: 105B041-POULIN (Ag-Pb vein), 105B002-STERLING and 105B008-DALE and 105B007- HOLLIDAY (Ag-Pb-Zn veins), 105B001-STAR (Au-Ag-Pb-Zn vein), 105B102-FREER (Ag-Pb vein) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-23 - Gravel Creek pluton (60.7406, 130.8541) c-gr equigr (sub-mc- phyric) Q=K=P bt>ms monzogranite 0.09 I > M > T R (Ta) xeno, Ce-La- Th-(Ta-Nb-Ca) mnz, Ca>Fe>Mg grt, Fe-Mg ttn, Mg px, zrn, ap perth and TT in mc; mc sub-ph-xst are up to 1x2cm (only slightly larger than surrounding qtz-fsp-bt gr- m); ms has 1st order birefringence but is c-gr and magmatic; local chl alt'n of bt; minor to moderate ser-ms in pl; minor (to mod) clay speckling in mc proximal: 105B104-TEAM and 105B115-STONEAXE (W skarns) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related KR-07-24 - Cassiar batholith (60.0928, 130.6972) m-gr equigr Q=K>P bt monzogranite 4.58 M > I O Fe>Mg-Ce-La-Th aln, Ce-La-Th-Ta- Nb-Ca mnz, Ca grt, zrn, ap minor myrm; perth and TT in mc; trace sub-s ms; minor chl (blue) alt'n of bt; minor to moderate ser in pl; minor clay-speckling in mc; abun mag proximal: 105B096-LICK (Ag-Pb-Zn vein); distal: 105B129-LOWEY (Zn vein) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related 07M-150 - Carolyn batholith (62.1527, 131.8455) pl-hbl-qtz-phyric (chl- ep ± ms-cal) granodiorite porphyry 0.34 I > M R ttn, Ca>>Fe px, zrn, ap v f-gr chl-ep-qtz-cal gr-m; pl ph-xst are (2-4mm); gr- m and ph-xst are v strongly to moderately alt'd to chl (blue); hbl ph-xst are 3x5 to 5x10mm and appear to be aggregates of f-gr material or poikilitic n/a Paleozoic NA (Cam- Ord Road River Gp.) - basin 523 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host 07M-151 - (volcanic) (62.0736, 131.9884) welded dark grey pl- qtz-bt-(hbl) lapilli tuff w/aphanitic to microcrystalline groundmass 0.23 I R Fe>Mg-Ce-La-(Th) aln, crn, zrn, ap some ep in the matrix (high relief brown mineral in matrix similar to ttn but w/v low birefringence); bt are deformed and full of small black inclusions; matrix is welded; hbl (cc+chl) and pl (ser) are moderately alt'd n/a Paleozoic NA (Cam- Ord Road River Gp.) - basin 07M-152 - Buttle Creek pluton (62.1822, 133.4155) foliated m-gr equigr Q>P>K bt granodiorite 0.19 nd R Fe>Mg-(Ce-La-Th) aln, Fe>Mg>Mn grt, rt, zrn, ap bt is strongly foliated, interstitial to Q-K-P, f-gr, and contains abun thick blades of dark-brown rt; some large composite qtz; no TT/perth (orthoclase?); minor ser but otherwise pristine; ap is abun and c- gr cut by undeformed aplites up to 3cm wide; local grt-ms pegmatites (Gordey, 2011, in prep) n/a Paleozoic NA (Cam. to Triassic) - platform and clastic influx 07M-153 - Glenlyon batholith (62.1730, 134.0078) f-m-gr equigr P=Q>K bt granodiorite 0.19 I R Fe>Mg-Ce-La aln, Ce-La-Th-(Ca-Ta- Nb) mnz, Ca>>Fe px, zrn, ap local TT and perth in mc; bt locally alt'd to chl; minor to moderate ser abundant dark grey f-gr xenoliths and aplites proximal: 10500003-LITTLE SALMON (Ag-Pb-Zn skarn), 105L001-LOKEN and 05K013-THOMAS (Zn skarns), 105L004-MOULE (Cu unknown) Proterozoic (Windemere SG) to lower Cambrian NA - rift-related 07M-154 - Little Salmon pluton (62.1730, 134.0078) (m)-c-gr equigr P=Q>K bt monzogranite 0.13 I R Fe>Mg-Ce-La aln, Fe-Ca thor, Fe>Nb ttn, Ca>(Fe) grt, Ca>(Fe) px, hbl, zrn, ap local TT and perth in mc; bt is green and locally alt'd to chl; moderate ser ± ms in pl and moderate clay speckling in mc; relatively abun and c-gr zrn nd n/a Devono-Mississippian Y-T (Snowcap) - crustal block FL-46 - Square Lake pluton (61.1597, 130.5137) nd nd nd nd *mnz, zrn, ap nd nd n/a Devono-Mississippian YTT inlier (Snowcap) - crustal block MP-18 - Money pluton (61.3614, 130.2537) nd nd nd nd *mnz, zrn, ap nd nd n/a Devono-Mississippian YTT inlier (Snowcap) - crustal block ARG-33 - Tsa Da Glisza pluton (61.2699, 130.5685) nd nd nd nd *mnz, zrn, ap nd nd proximal: 105G147-REGAL RIDGE (emerald), 105G102-HOWDEE (W skarn) Devono-Mississippian YTT inlier (Snowcap) - crustal block KG-31 - Tsa Da Glisza pluton (61.3414, 130.5187) nd nd nd nd *mnz, zrn, ap nd nd proximal: 105G071-MYDA (W skarn) Devono-Mississippian YTT inlier (Snowcap) - crustal block KM-07 - Mink pluton (61.5344, 131.1168) nd nd nd nd *mnz, zrn, ap nd nd proximal: 105G097-FOG (W skarn) Devono-Mississippian YTT inlier (Snowcap) - crustal block 524 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host MC-CL-01 - dyke (Clear Lake) (62.7843, 135.1568) nd nd nd nd *zrn, ap nd nd n/a Devono-Mississippian YTT inlier (Snowcap) - crustal block Easterm Selwyn Basin (Geochronology Samples): KR-05-08 - Jorgensen pluton (60.8717, 126.2921) f-gr pl-qtz-hbl-phyric qtz monzodiorite porphyry 0.19 ? R Fe-Mg-Ce-La aln, zrn, ap hbl ph-xst replaced by chl (blue/brown) + mag (small round to elongate) + ep; pl ph-xst intensely ser'd w/abun cal and rare relict ph-xst display osc zoning; one large dark honey-brown zoned elongate-prismatic ph-xst/x-xst (aln/phl); large cal pseudomorphs (of pl?); trace ttn assoc'd w/ep (2ndary); abun blocky py sampled felsenmeer n/a Paleozoic NA (Cambrian-Devonian) - basin KR-05-10 - Powers pluton (60.7962, 126.0797) f-gr pl-(hbl)-phyric qtz monzodiorite porphyry 8.32 ? O Fe-Mg-Ce-La aln, zrn, ap hbl ph-xst strongly alt'd to ep in cores or to chl (blue) + mag (small equant-euh) + ep; pl ph-xst moderate to intensely ser'd, and sub-ph-xst v intensely ser'd-sauss'd; trace ttn assoc'd w/ep (2ndary); mc ph-xst up to 1cm n/a Paleozoic NA (Cambrian-Devonian) - basin KR-05-22 - Roy pluton (61.5818, 127.6647) c-gr equigr to pl-phyric P>K>Q bt monzogranite 17.30 T O Fe-Mg-Ce-La-(Th) aln, ttn, zrn, ap perth in mc; large aggregates of pl w/complicated zoning, including osc; albite and penetration twins common in pl; moderate clay-speckling in mc; minor ser in pl except in intensely ser'd cores in gr- m; well-zoned aln; some euh blocky mag aplites and pegmatites assoc't w/gossan; sulphide-bearing veins in finer-grained phase distal: 095ENW0012-ROY (Pb-Zn-Ag- Au manto) Paleozoic NA (Cambrian-Devonian) - basin KR-05-26 - Fish pluton (61.6275, 127.4917) f-m-gr pl-phyric P>Q=K bt monzogranite 4.82 nd O Fe-Mg-Ce-La-(Th) aln, Ca-(Fe) grt, Fe>Mg-(Mn) grt, zrn, ap perth in mc; pl ph-xsts have osc zoning and are <1x2cm blocks or 1.5x4cm laths; bt partially to fully alt'd to chl (sometimes bluish) w/ep and op; moderate to intense clay speckling in mc; minor to mod/intense ser in pl, esp cores of gr-m grains; well-zoned aln; some euh blocky mag n/a Paleozoic NA (Cambrian-Devonian) - basin KR-05-32 - Ivo pluton (61.0519, 127.1072) f-gr equigr K>/=Q>P bt>ms grt leucogranite 0.06 I R Ce-La-Th-(Nd-Ta- Nb) ± Ca mnz, Fe>Mn-(Mg-Ca) grt, trm, xeno, zrn myrm; pl does not have good albite twins; TT present in mc; ~3-5% bt and <2% ms; some bt alt'd to chl and op, esp proximal to primary ms; minor to locally moderate ser in pl; mc ph-xst <0.5cm but may be up to 2cm cut by aplite and qtz-tur veins proximal: 095ESE0002-TUANIPEL and 095ESE0001-78 (W skarns) Proterozoic NA (Windemere SG) - rift- related KR-05-43 - Park pluton (61.4714, 126.4309) f-m-gr equigr P>>Q>K hbl>bt granodiorite 16.4 T O Fe-Mg-Ce-La-(Th), (Nb-F-Y) ttn, zrn, ap almost all pl has osc zoning; minor ser in pl, w/intensely alt'd zones; minor clay speckling in mc; ttn and/or ep assoc'd w/hbl; moderate euh mag n/a Proterozoic NA (Windemere SG) - rift- related 525 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-05-62 - Mount Christie pluton (62.9935, 129.4825) m-gr mc-phyric K>P>Q hbl>>bt qtz monzonite 0.34 T R Fe-Mg-Ce-La-(Th) aln, Pb-P-Ca-F thor, zrn, ap perth and TT present in mc and some large zoned mc (1x2cm); some bt alt'd to chl-ep; minor sercite in pl, w/some intensely alt'd cores; minor to no clay speckling in mc; v c-gr and v abun ttn cut by min'd qtz-apy-gn-py- azurite and graphic qtz-tur veins; rusty tur-filled miarolitic vugs proximal: "U" (Pb-Zn vein) Paleozoic NA (Mississippian- Triassic) - basin KR-05-68 - Mount Wilson pluton (62.8915, 129.6860) (foliated) f-gr mc- crystic Q=K=P bt monzogranite 0.12 I R Fe-Mg-Ce-La aln, Fe>Ca>Mg-(Mn) grt, zrn, ap perth rare or not present; mc m-xst 0.5cm to 1x2cm; pl has complicated zoning; some late intergranular chl (brown) and ms; ~50% bt alt'd to chl (brown>blue) w/op; moderate ser-ms in pl; minor to moderate clay-speckling in mc bt locally foliated around mc- ph-xst; large rounded qtz ph- xst (xenocrysts?) in some phases distal: 105INW0021-CAM (Pb-Ag-Cu vein) Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-76 - Pelly River pluton (62.7649, 129.6424) foliated f-gr equigr Q=K=P bt qtz diorite 0.1 I R Fe-Mg-Ce-La-(Th) aln, scheelite, zrn, ap perth rare or not present; mc up to 0.5cm; pl has complicated zoning; some late intergranular chl (brown) and ms; minor chl (brown) after bt; minor (to mod) ser-(ms) in pl n/a Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-77 - mafic dyke, Pelly River pluton (62.7649, 129.6424) altered bt-phyric kersantite (lamprophyric) dyke 0.23 nd R nd f-gr; euh bt laths w/up to 50% felted chl; pl is interstitial to bt ph-xst, sub-ph-xst, and in gr-m; minor qtz in gr-m; pl in gr-m is moderate to highly sauss'd, and pl sub-ph-xst (~7%) are sauss'd or replaced by euh qtz and anh cal; trace to minor interstitial op cutting mc-porphyritic granite n/a Cretaceous granitoid - Mayo Suite KR-05-97b - O'Grady batholith (62.9459, 128.7133) f-m-gr mc-phyric (crowded) K>>P>>Q tur-hbl>cpx>bt qtz monzonite 4.06 T>M+I O Fe-Mg-Ce-(La) aln, Mg>Fe-(Ca) grt (prp core, alm rim?), tur, zrn, ap perth in mc ph-xst (1x2cm); bt frequently as cores in hbl; minor to moderate clay speckling in mc; cpx is Mg-Ca-(Fe-Al) rich; minor ser in pl; v abun and c- gr ttn graphic tur-infilled miarolitic cavities in many phases; cut by tur-aplite and qtz veins; qtz- fsp porphyry at KAREN showing proximal: 105INE0022-STIB1 (Sb- skarn), 105INE0011-KAREN (Pb-Zn- Sn-Sb skarn), 105INE0025-HAN (Zn unkn), 105INE0028-STIB3 (Cu-skarn), 105INE0029-STIB4 and 105INE0030- STIB5 (Pb-Zn skarn), 105INE0026- GRUMP (diss. Au), 105INE0027- STIB2 (Zn skarn), 105INE0023- O'GRADY (rubellite tur), "V" (Pb-vein) Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-110 - Cac pluton (62.3692, 128.5493) m-gr mc-phyric K=P=Q bt monzogranite 0.15 I R Fe-(Mg-Ce-La-Th) aln, Ce-Nd-La-Th- (Ta-Nb) mnz, Fe>Ca>Mg-(Mn) grt, zrn, ap perth and TT present in mc; some bt alt'd to chl (blue) w/ep; moderate ser-ms in pl cores; minor clay speckling in mc; mc ph-xst up to1.5x6cm laths and 2x4cm blocks; Ti-poor ilm may be mantled by Fe-S sulphide-bearing aplite (distal), cross-cutting zoned ms-tur-grt aplites, and f-gr leucogranite to bt-rich phases present proximal: 105ISE0019-CAC (W skarn) Paleozoic NA (Devono- Mississippian) - clastic influx 526 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-05-113 - South Nahanni pluton (62.5526, 128.6084) m-c-gr mc-crystic (pl- qtz sub-phyric) K=Q=P bt>hbl>>cpx monzogranite 0.13 T R Fe-(Mg-Ce-La) aln, zrn, ap one normally-zoned pl has a rounded x-xst core of older pl full of f-gr bt inclusions; perth and TT present in mc; minor alt'd to chl bt; moderate clay speckling in mc; minor (to mod) ser ± ms in pl; cpx is Ca-Fe-(Al-Mn)-rich (augite?); ttn may be Nb-rich; mc m-xst 1x2cm but up to 2x6cm cut by altered mafic dyke proximal: "BLUE" (W skarn) Paleozoic NA (Cambrian-Devonian) - basin KR-05-126 - (plug SW of O'Grady) (62.7624, 129.0990) altered qtz-mc-pl-bt- phyric leucogranitic porphyry 0.04 ? R tur, zrn 2ndary ragged ms throughout gr-m; pl ph-xst ser'd w/lots of cal; mc ph-xst intensely clay speckled; bt ph-xst mostly alt'd to chl (blue) w/op; locally zoned and poikilitic tur pervasively altered; aplitic tur- mottled phase present n/a Paleozoic NA (Mississippian- Triassic) - basin KR-05-130 - Central Nahanni pluton (62.6869, 128.6281) m-gr equigr bt granodiorite 0.22 nd R Fe-Mg-Ce-La-(Th) aln, zrn, ap nd xenoliths of mafic to intermediate phases; cut by mafic qtz diorite dyke and sheeted gossanous fractures/qtz veins n/a Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-136 - North Nahanni pluton (62.7739, 128.6443) f-gr equigr P=K>Q bt>hbl monzogranite 0.19 nd R Fe-Mg-Ce-La-(Th) aln, zrn, ap perth in mc; complicated zoning in pl; hbl is 'rotten' (partially alt'd to chl) and has a greeny-blue pleochroism; some v f-gr chl-ser-ms mantling bt; minor to moderate clay speckling in mc; minor to moderate ser in pl, esp in cores; ap can be relatively c-gr; aln is well-zoned f-gr leucogranite phases present; cut by aplites and tennannite?-filled fractures; ccp-endoskarn developed near RISBY showing proximal: 105INE0002-RISBY and "PR" (Cu-skarns) Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-143 - Circular Stock (61.9642, 128.2477) m-gr mc-phyric K>P>Q bt monzogranite 0.11 I R Ca-Fe-Mg-Ce-La aln, Ce-La-Th-Nd- Ta-Ca mnz, Fe>Mg-(Ca-Mn) grt, zrn, ap perth on ph-xst (0.5x2cm) and TT in gr-m mc; rarely bt alt'd to chl (blue), and sub-s v f-gr ms-chl-ser mantling bt; minor to intense ser-(ms) in pl; moderate (to intense) clay speckling in mc same phase 2003 sample: 28- 1; cut by sulphide-bearing grey-purple qtz veins and bt or ms±grt aplites; fsp-bt porphyry phases and a marginal late breccia pipe are present n/a Paleozoic NA (Cambrian-Devonian) - basin KR-05-148 - Rifle Range pluton (62.0035, 128.1661) c-gr mc-phyric (sparse) Q>/=K>P bt monzogranite 0.12 nd R Fe-Mg-Ce-La-(Th) aln, Ce-La-Th-Nd- Ta mnz, zrn, ap large qtz grains are composite; perth and TT present in mc; mc up to 1x2.5cm; rarely bt alt'd to chl (blue); minor to moderate ser-(ms) in pl; minor clay speckling in mc cut by qtz-tur veins w/scheelite-mineralized greisen and scheelite-bearing fractures; f-gr porphyry and aplites present n/a Proterozoic NA (Windemere SG) - rift- related 527 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-05-157 - mafic dyke, Roy pluton (61.6107, 127.6580) v f-gr pl-qtz-phyric quartz diorite dyke 0.68 nd O nd pl sub-ph-xst (10%) up to 0.7cm and sieve-textured; rare qtz ph-xst (1%); minor opx sub-ph-xst w/wormy-sieve texture in core and poorly developed lamellar twins; brown bt is present (~25%) but most is pale brown and shreddy bt-cal- ms (irregular distribution); pl in gr-m is v fresh and more abun than bt; trace op; cal clots are mantled by shreddy pale brown bt; rare bt ph-xst up to 1cm sulfide-bearing (py+galena) proximal: 095ENW0012-ROY (Pb-Zn- Ag-Au manto) Paleozoic NA (Cambrian-Devonian) - basin KR-05-158 - mafic dyke, Roy pluton (61.6113, 127.6588) v f-gr pl-qtz-phyric quartz diorite dyke 1.88 nd O nd irregularly resorbed pl sub-ph-xst (5%) up to 0.4cm and sieve-textured; minor qtz assoc'd w/subhedral- euh blocky pseudomorphs (10%; after px?) of interstitial op-cal and mantled by felted bt and chl (radiating brown); clots and networks of pale brown/green shreddy bt-chl ± cal throughout ~65% of gr-m); pl in gr-m (~35%) is fresh; trace opx sub- ph-xst proximal: 095ENW0012-ROY (Pb-Zn- Ag-Au manto) Paleozoic NA (Cambrian-Devonian) - basin KR-05-164 - Hole-in-the-Wall batholith (61.7361, 127.3551) f-m-gr mc-crystic K>P>Q bt monzogranite 0.11 I R Ce-La-Th-Nd-(Ta) mnz, Fe>Mg-(Ca- Mn) grt, wolframite, Ta xeno, zrn, ap perth in m-xsts and TT in gr-m mc; clots of ms-chl and v f-gr intergranular ms-chl mantling bt; minor ser in pl; minor clay speckling in mc; f-gr granitic bt- rich xenolith in half of TS; mc m-xst of 2x3cm blocks and 1x3cm laths contains abundant large tur- filled and aplite-rimmed miarolitic cavities and rounded bt-ap rich mc-phyric xenoliths; ribboned tur±bt aplite- pegmatite phase present; cut by open fractures partially infilled w/tur±qtz (may be gossanous) n/a Paleozoic NA (Cambrian-Devonian) - basin KR-05-175 - Lened pluton (62.3803, 128.6466) f-m-gr mc-crystic (pl sub-phyric) K>/=Q=P bt monzogranite 0.14 I R Mg-Ce-La-(Fe) aln, Ce-La-Th-Nd-Ta- Ca mnz, Ca-(Ti- Fe) grt, F ttn, zrn, ap perth and rare TT present in mc; minor to moderate ser-ms in pl; minor clay speckling in mc; minor chl in bt; mc m-xst generally 1x3cm but up to 2x6cm multiple phases of bt-ms ± grt leucogranite; cut by aplite, and qtz/tur veins and dykelets; tur- infilled pods in some phases proximal: 105ISE0003-LENED (W skarn), 105ISE0070-LENED-BERYL (emerald vein) Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-191 - Marion pluton (61.3927, 127.3413) (m)-c-gr mc-phyric (sparse) Q>/=K=P bt monzogranite 5.5 T O Ca-Fe-Mg-Ce-La- (Th) aln, Ca-(Fe) grt, zrn, ap bt generally in f-m-gr clots; intergranular f-gr chl-bt- ms in gr-m; ~50% bt fully to partially alt'd to chl (blue); intense clay speckling in mc; minor to intense ser in pl; abun well-zoned aln; abun mag assoc'd w/bt; mc-ph-xst are 0.5x1cm sampled from a boulder in valley floor streambed proximal: 095E006-MARION (Pb-Zn vein) Paleozoic NA (Cambrian-Devonian) - basin KR-05-194 - Coal River batholith (61.3705, 127.2392) m-gr mc-phyric K=P>/=Q bt monzogranite 0.14 I R Ca-Fe-Mg-Ce-La aln, Ce-La-Th-Nd- (Ta) mnz, Pb-P thor, F ttn, zrn, ap perth and TT present in mc; mc ph-xst are 0.5x1cm; minor chl in bt; minor clay speckling in mc; moderate to intense ser-ms in pl; abun op assoc'd w/bt bt leucogranite also present; cut by tur-ms and qtz veins/dykelets proximal: 095E036-ROYALL and 095E005-ISO and 095E039-KOMISH (W skarns); distal: 095E040-VNER (W skarn) Proterozoic NA (Windemere SG) - rift- related 528 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-05-196 - Coal River batholith (61.5434, 127.2377) m-gr equigr K=P>Q bt>hbl monzogranite 9.01 T O Ca-Fe-Mg-Ce-La- (Th) aln, (Nb-F-Y) ttn, ep, zrn, ap diffuse perth and abun TT present in mc; most hbl is chl; hbl up to 2cm; minor chl in bt; moderate to intense clay speckling in mc; minor ser in pl; possible c-gr magmatic ep in pl cores; abun mag assoc'd w/bt; ttn has Nb-Y rich rims similar phase to 99-SH-006 and 99-SH-008 proximal: 095ENW0027/28/29- CRAT1/2/33 (W ±Zn ±Mo skarns) Paleozoic NA (Cambrian-Devonian) - basin KR-05-198 - Hole-in-the-Wall batholith (61.6150, 127.1816) c-gr mc-phyric (sparse) Q>>P>K bt monzogranite 0.18 I R Ca-Fe-Mg-Ce-La- (Te-Th) aln, zrn, ap perth in mc; minor chl ± ms alt'n of bt; minor ser in pl, esp cores; minor to moderate clay speckling in mc; mc ph-xst are 1.5x3cm mapped as a large round and distinct phase of the Hole-in- the-Wall batholith by Gabrielse et al. (1973); cut by bt leucogranite dykes n/a Paleozoic NA (Cambrian-Devonian) - basin KR-05-204 - mafic dyke, Cantung (approximately 61.9642, 128.2477) altered foliated f-gr bt- phyric kersantite (lamprophyric) dyke nd nd ? nd sparse (5-7%) bt ph-xst (partially to completely alt'd to ep-ms); gr-m appears to be clinozoisite (possibly pl), ep, ms, and trace qtz; large (2mm) interstitial to euh aggregates of op, ms, felted bt, ep (piedmontite) ± rt ± mc 118 Lift; appears to intrude fault zone proximal: 05HNE0007-E-ZONE (W skarn) Cretaceous granitoid - Tungsten Suite KR-05-208 - Little Hyland pluton (61.6150, 127.1816) strongly foliated (mc- augen) f-m-gr equigr Q=K>P bt>>ms monzogranite 0.11 nd R Ca-Fe-Mg-Ce-La- (Th) aln, Th-Ce-La- (Ca-Nd) mnz, Fe>Mg-Mn grt, uran, zrn, ap qtz forms f-gr aggregates (annealed?); intergranular veinlets of v f-gr qtz-fsp-ms ± chl throughout gr-m between nearly all grains; perth and TT present in mc; albite and pericline twins in pl; ms is primarily v f-gr intergranular and 2ndary mantling bt; moderate ser in pl ms appears to be fracture/vein related greisenization; ribboned tur-qtz vein cutting greisenized granite with scheelite mineralization n/a Proterozoic NA (Windemere SG) - rift- related KR-05-210 - Nahanni Range pluton (61.9167, 128.4072) m-c-gr equigr Q=K>P bt>>ms monzogranite 0.09 nd R Ca-Fe-Ce-(La) aln, Th-Ce-La-(Ca) ± Nd mnz, Fe-Ca- Mn-Mg aluminosilicate (grt or piedmontite), zrn, ap qtz forms large rounded f-gr aggregates (annealed?); perth and TT present in mc; ms is primarily v f-gr intergranular and 2ndary mantling bt, but rare possible primary grains are assoc'd w/bt; moderate to intense ser-ms in pl not v fresh; local greisenization assoc'd w/ small qtz veins n/a Proterozoic NA (Windemere SG) - rift- related KR-05-213 - mafic dyke (61.9772, 128.2103) altered v f-gr bt-phyric kersantite (lamprophyric) dyke 0.62 nd O nd ~75% of bt alt'd to felted chl>bt (1/2 of bt forms large aggregates after an unknown ph-xst mineral); pl and bt are present in equal amounts in gr-m; pl is interstitial to bt and unalt'd to slightly sauss'd; minor rounded/irregular qtz (3%) as ph-xst or in gr-m is mantled by felted bt-chl; minor interstitial op sampled from a cluster of mafic dykes found on the south side of Dutchman's Peak n/a Paleozoic NA (Cambrian-Devonian) - basin 529 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-05-215 - Mine Stock/Cantung pluton (61.9662, 128.2288) f-m-gr mc-phyric (small and sparse) K=Q>P bt>>>ms monzogranite 0.11 nd R Fe>Ca-(Mg-Mn) grt, F ttn, zrn, ap qtz forms large anh grains and aggregates; perth and well developed TT in mc but ph-xst (up to 1cm) are moderate to intensely clay-speckled; pl may be well-zoned and cores are moderate to intensely sauss'd; trace sub-s ms; minor myrm same phase as KR sample in 2003: 27-1; cut by gossanous scordierite-qtz vein and sheeted qtz-tur veins proximal: 05HNE0007-E-ZONE (W- Cu-Zn-Bi skarn), 05HNE0006- CANTUNG (W-Cu-Bi-Au-Zn skarn), 105HNE0002-BAKER (Cu-W-Zn-F skarn, "REDPATH" (W skarn) Paleozoic NA (Cambrian-Devonian) - basin CL-06-33A - Dechen'La pluton (63.3739, 129.9027) m-gr mc-crystic hbl- phyric K>P>/=Q hbl>bt>>cpx monzogranite 0.29 T R Ca-Fe-Mg-Ce-La- (Th) aln, zrn, ap perth in mc; mc ph-xst up to 1x2cm; hbl up to 1cm; minor to moderate ser in pl; bt is pale green-brown; cpx is Mg-Ca-(Fe>Al)-bearing (augite?); relatively c- gr zrn <5% rounded cm-scale f-gr bt- hbl rich xenoliths (~50% mafics) n/a Proterozoic NA (Windemere SG) - rift- related CL-06-34A - Natla pluton (63.0493, 129.3941) f-m-gr mc-crystic K>P>Q bt monzogranite 0.21 T R Ca-Fe-Mg-Ce-La- (Th) aln, zrn, ap most qtz v f-gr in matrix; mc m-xst 1x2cm; bt locally alt'd to chl and forms both c-gr euh laths and f-gr clots; moderate ser-ms in pl; mc is perthitic spotted w/round 1cm clots of bt ± ms-chl; pegmatitic vein and aplitic phases also present n/a Paleozoic NA (Devono- Mississippian) - clastic influx CL-06-35A - Logan pluton (63.0822, 129.3648) f-gr equigr P>K>Q bt>hbl>>cpx qtz monzodiorite 0.31 T R Ca-Fe-Mg-Ce-La- (Th) aln, zrn, ap minor ser in pl and minor clay speckling in mc; large euh pl and v abun f-gr euh pl inclusions (in pl!) have v complicated zoning patterns; cpx is Mg-(Ca-Al- Fe)-bearing (augite?) sub-cm rounded clots of finer- grained bt-rich xenoliths (?) n/a Paleozoic NA (Devono- Mississippian) - clastic influx CL-06-36A - Mount Christie pluton (63.0060, 129.4231) (f)-m-gr mc-phyric (crowded) K>P>Q hbl>bt>>cpx qtz monzonite 1.6 T O Ca-Fe-Mg-Ce-La- (Th) aln, zrn, ap mc ph-xst are <0.5cm; mc is perthitic and has TT; bt moderate alt'd to chl (blue); minor to moderate ser in pl; minor to moderate clay speckling in mc; v c-gr and v abun ttn; cpx is Ca-Mg-(Al-Fe)-bearing (augite?) v similar to O'Grady but mc p- xst are smaller and less crowded; aplite-pegmatite dykes are present proximal: "U" (Pb-Zn vein) Paleozoic NA (Devono- Mississippian) - clastic influx CL-06-37A - Mount Christie pluton (63.0267, 129.5817) f-m-gr mc-phyric (near- crowded) K>>P=Q hbl>>bt>>cpx qtz monzonite 0.96 T O Ca-Fe-Mg-Ce-La- (Th) aln, zrn, ap mc ph-xst up to 1cm and may be zoned; bt mostly alt'd to chl (blue) unless is an inclusion in unalt'd minerals; minor ser and moderate cal in pl; perth in mc; v c-gr and v abun ttn; cpx is Mg-Ca-Al-(Fe)- bearing (augite?) tur hosted breccia (rich in sulphides, granite clasts) proximal: "U" (Pb-Zn vein) Paleozoic NA (Devono- Mississippian) - clastic influx CL-06-38A - Ross River pluton (63.0474, 129.9700) f-m-gr mc-phyric (scattered) Q=P>K bt (leuco)monzogranite 0.15 I R zrn, ap mc ph-xst are <1x2cm; qtz is typically c-gr and inclusion-free (samples is almost qtz-phyric); perth in mc; moderate chl and minor ms after bt; moderate ser-ms in pl; moderate clay speckling in mc; in addition to magmatic Ti-bearing ilm, Ti-poor ilm is after Fe-S; apy ± Fe-S in veins/fractures n/a Paleozoic NA (Devono- Mississippian) - clastic influx 530 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host CL-06-39A - Keele River pluton (63.1262, 129.8556) f-gr mc-crystic K>P>/=Q hbl>bt >>cpx qtz monzonite 0.27 T R Ca-Fe-Mg-Ce-La- (Th) aln, zrn, ap mc m-xst average 1cm and are zoned; complicated zoning in pl; perth in mc; minor-moderate chl (blue) in bt; zoning-controlled moderate ser in pl; minor to moderate clay speckling in mc; cpx is Mg-Ca-(Fe- Al)-bearing (augite?) contains v f-gr bt-hbl-rich (~50% mafic) cm-scale xenoliths; cut by 20cm thick mafic dyke n/a Paleozoic NA (Devono- Mississippian) - clastic influx CL-06-40A - Mile 222 pluton (63.2143, 129.8647) f-m-gr mc-phyric K=P>Q hbl>bt>>cpx qtz monzonite 0.49 T>M R Ca-Fe-Mg-Ce-La- (Th) aln, tur, zrn, ap mc ph-xst are ~0.5cm; bt is moderate alt'd to chl (blue); perth and v local TT present in mc; minor to moderate ser in pl along zoning (zoning can be complicated); ap can be c-gr; ttn is abun; cpx is Mg- Ca-(Fe-Al)-bearing (augite?) some oxidized mafics along fractures n/a Paleozoic NA (Devono- Mississippian) - clastic influx 03M-203 - felsic dyke, Cantung (approximately 61.9642, 128.2477) f-m-gr equigr K=P>Q bt>ms leucogranite dyke nd nd R *mnz, zrn, ap qtz forms anh grains in the gr-m; perth and TT present in mc; pl has moderate sauss-ms alt'd cores; ms is present as ragged magmatic and sub-s grains, and is assoc'd w/bt and trace chl; minor myrm; minor spherulitic qtz-fsp intergrowths emplaced in No. 5 fault and cuts unmineralized massive pyrrhotite skarn proximal: 05HNE0007-E-ZONE (W skarn) Paleozoic NA (Cambrian-Devonian) - basin 03M-204 - aplite, Cantung (approximately 61.9642, 128.2477) altered f-m-gr pl-phyric aplitic dyke; calcic-1 and calcic-sericitic alt'd (cf. Rasmussen et al., 2011a) nd nd R *zrn, ap equal amounts of f-gr qtz and pl; most pl in gr-m is boxy-cellular; some magmatic pl ph-xsts and grains w/relict albite twins that are strongly overprinted by cal-ser; pl is moderately to strongly alt'd to cal-ser- ms; cut by chl (brown) + ep-qtz vein 3900’ level: cuts Mine Stock- apical; w/disseminated and patchy po, locally abundant scheelite proximal: 05HNE0007-E-ZONE (W skarn) Paleozoic NA (Cambrian-Devonian) - basin 03M-207 - Mine Stock-apical, Cantung pluton (approximately 61.9642, 128.2477) f-m-gr (mc-phyric) K=P=Q bt>>ms monzogranite 0.04 I R *grt, mnz, zrn, ap qtz as large aggregate clots and large anh grains in the gr-m; perth and TT present in mc; pl may be v well zoned and have sieve-textured cellular cores in larger (sub-ph-xst) grains; ms is present as ragged sub-s grains assoc'd w/bt; minor to moderate myrm from mechanic shop off main haulway proximal: 05HNE0007-E-ZONE (W skarn) Paleozoic NA (Cambrian-Devonian) - basin 03M-214 - Shelf Lake pluton (62.1168, 128.1591) f-m-gr mc-crystic K>P>Q bt>ms monzogranite 0.03 nd R *mnz, zrn, ap qtz forms large anh grains; perth and well developed TT present in mc; mc m-xst up to 2x5cm; pl may be well-zoned and cores are moderate ser'd where bt is also alt'd to chl (blue) and ep; some spherulitic fsp-qtz intergrowths widespread tur veins and a couple aplite dykes; greisenized sulphide-bearing fractures; bt-pl-phyric xenoliths proximal: 105ISE0014-JAY (Pb-Zn- Ag-Au vein), 105ISE0013-monA (Pb- Zn-Ag-Au skarn), "D" (W skarn); distal: "B" (Pb-vein) Proterozoic NA (Windemere SG) - rift- related 03M-215 - Mount Appler pluton (62.2212, 128.0925) f-m-gr mc-crystic K>P>Q bt>ms monzogranite 0.09 nd R *mnz, zrn, ap moderate abun myrm; some spherulitic fsp-qtz intergrowths; qtz forms large anh grains; perth and well developed TT present in mc; mc m-xst up to up to 5x5 cm; pl cores are moderate sauss'd in areas where bt is alt'd to chl (blue) and ep rare aplites; rusty weathering holes (miarolitic cavities?) in float w/traces of py-ccp- scordierite; "spotted" with abundant small tur-filled and aplite-rimmed miarolitic cavities; mc-cumulate phase observed; pl-phyric xenoliths proximal: 105ISE0060-SANDY (W-Zn- PB skarn) Paleozoic NA (Devono- Mississippian) - clastic influx 531 Summarya MSb I/M/Tc R/Od Accessory Mineralse Petrographic Notes a Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host 03M-216 - Faille pluton (62.1913, 127.6591) f-m-gr mc-crystic K>P=Q bt>ms monzogranite 0.29 nd R *zrn, ap minor myrm; qtz forms large anh grains; mc ph-xsts are 2x3cm and moderately clay-speckled; perth and well developed TT in ph-xsts present in mc; many pl grains are well-zoned and cores are strongly sauss'd ± ms; c-gr (zoned?) ep assoc'd w/alt'd to chl (blue) bt no aplites/veins; rounded dark brown bt rich xenoliths up to 40cm diameter n/a Paleozoic NA (Cambrian-Devonian) - basin 03M-217 - Mount Sir James McBrien (Cirque) pluton (62.0981, 127.6913) m-c-gr mc-phyric K>P>Q bt>hbl monzogranite 2.73 T-M O *aln, zrn, ap qtz forms large anh grains; mc ph-xsts are up to 2.5x5cm and moderately clay-speckled; perth and TT present in mc; many pl grains are well-zoned and cores are moderate ser'd; minor bt and hbl alt'n to chl; well-zoned aln no aplites/veins; abundant bt- rich bt-qtz xenoliths to 50cm diameter n/a Paleozoic NA (Cambrian-Devonian) - basin 03M-218 - Hole-in-the-Wall batholith (north) (61.8905, 127.6800) m-c-gr mc-phyric K>Q=P bt>hbl monzogranite 7.28 M O *zrn, ap qtz forms large anh grains; mc in gr-m is moderately to intensely clay-speckled, and only moderate in ph-xsts (1x3cm); many pl grains are well-zoned and skeletal scattered mafic xenoliths up to 20cm in diameter; one aplite dyke w/a miarolitic cavity n/a Paleozoic NA (Cambrian-Devonian) - basin 03M-219 - Pyramid Mountain pluton (61.8426, 127.8698) m-gr mc-phyric bt>>ms K=Q>P monzogranite 0.07 nd R *grt, mnz, zrn, ap qtz forms large rounded f-gr aggregates; perth w/rare and poorly developed TT present in mc; mc ph-xst are 1x2cm; rare ms is c-gr and primary; moderate to intense ser in pl; rt needles in bt where alt'd to minor chl; grt assoc'd w/v f-gr bt and minor myrm narrow (<2 cm) qtz-tur veins proximal: 095ENW0015-PYRAMID MTN (W-Cu-Ag skarn), 095ENW0011- HOUSE OF LORDS, 095ENW0025- SImonS, 095ENW0005-WHISTLER NORTH, 095ENW0026-FREDERICK CREEK, 095ENW0014-MB (W skarns) Paleozoic NA (Cambrian-Devonian) - basin 03M-220 - Raggged pluton (61.9838, 128.0806) f-gr mc-phyric (sparse) bt monzogranite 0.12 nd R *mnz, zrn, ap missing; mc ph-xst up to 1cm rare aplites; some local brecciation w/f-gr dark matrix; narrow apy/scordierite veins n/a Proterozoic NA (Windemere SG) - rift- related Notes: Rock-forming mineral abbreviations are after Kretz, 1983, or are: feldspar = fsp, px = pyroxene, cpx = clinopyroxene, ser = sericite, sauss = saussurite; xenotime = xeno; uran = uraninite, thor = thorite, op = opaque. Descriptive and textural abbreviations are: cm = centimetre, w/ = with, esp = especially, v = very, abun = abundant, assoc'd = assoc'd, alt'n = alteration, alt-d = altered, f-gr = fine-grained, m-gr = medium-grained, c-gr = coarse-grained, equigr = equigranular, gr-m = groundmass, crystic = megacrystic, ph-xst = phenocryst, m-xst = megacryst, x-xst = xenocryst, euh = euhedral, anh = anhedral, myrm = myrmekite, TT = tartan twins, perth = perthite, osc = oscillatory, sub-s = sub- solidus, 2ndary = secondary. a brackets indicate a minor component; leucogranite is <3% mafic and (leuco)granite is <7% mafic; P = plagioclase component, Q = quartz component, K = K-feldspar component; b magnetic susceptibility (x10-3SI units); c M = magnetite; I = ilmenite; T = titanite; nd = not detected; d R = reduced; O = oxidized; e relative elemental abundances listed in order of decreasing SEM peaks (brackets indicated a very small peak); * = no compositional information. 532 A.2 Geochemical samples Sample locations, rock names, and brief textural and mineralogical descriptions for all samples that underwent geochemical and isotopic analysis (but not U–Pb and 40Ar–39Ar geochronological analysis). 533 Summarya MSb R/Od Petrographic Notes Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host Eastern Selwyn Basin (Isotopic + Geochemistry Samples): KR-05-207 - West Tuna pluton (61.8392, 128.3003) mylonitic c-gr mc- phyric bt monzogranite 0.11 R bt and v f-gr sub-s ms are wrapped around larger grains, gr-m of seriate v f-gr to extremely f-gr qtz, local moderate ser in pl cores Proterozoic NA (Windemere SG) - rift-related KR-05-212 - mafic dyke (61.9769, 128.2076) v f-gr leucocratic monzonitic porphyry dyke 1.22 O interlocking/radiating zoned and twinned pl laths comprise most of gr-m, w/abun ms and lesser ep-op (py/po), rounded/embayed qtz x/ph-xsts (or feldspathoid mineral) are mantled by ms, cores of larger pl are moderately sauss'd/ser'd south side of Dutchman's Peak Paleozoic NA (Cambrian-Devonian) - basin Eastern Selwyn Basin (Geochemistry Samples): KR-05-33 - felsic dyke, Ivo (61.0519, 127.1072) v f-gr equigr ms aplitic dyke 0.03 R pl is moderately ser'd; P+Q>K cuts KR-05-32 granite proximal: 095ESE0002-TUANIPEL and 095ESE0001-78 (W skarns) Proterozoic NA (Windemere SG) - rift-related KR-05-47b - Cirque Lake/Mactung pluton (63.2871, 130.1399) moderately foliated m-gr pl-bt-qtz-mc- phyric monzogranitic porphyry 0.13 R v f-gr primarily qtz matrix w/larger bt and fsp, ph-xsts are broken and fractures filled w/f-gr qtz, pl cores are moderately ser'd, minor chl after bt; bt foliated around ph-xsts proximal: 105OSE0001-MACTUNG (W-Cu-Zn skarn) Proterozoic NA (Windemere SG) - rift-related KR-05-53 - aplite, Rockslide Mountain pluton (63.2703, 130.1843) v f-gr equigr (sparse mc ph-xsts) ms-bt aplitic dyke 0.05 R gr-m is seriate, pl is moderately ser'd, micas are v f-gr and ragged (possibly sub-s ms) aplite dyke in valley, some chips of aplite material w/sulphides (dark grey, slightly magnetic, partially oxidized) Paleozoic NA (Cambrian-Devonian) - basin KR-05-59 - Rockslide Mountain pluton (63.2621, 130.1863) m-gr equigr (sparse mc ph-xsts) bt monzogranite 0.12 R nd; mc ph-xsts < 0.5cm Paleozoic NA (Cambrian-Devonian) - basin KR-05-60 - aplite, Rockslide Mountain pluton (63.2621, 130.1863) v f-gr equigr ms-bt aplitic dyke 0.05 R gr-m is seriate, pl is moderately ser'd, micas are v f-gr and ragged (possibly sub-s ms), single ph-xst of unknown mineral (coarsely zoned, colourless, 1st order grey) aplite from large dyke intruding country rock near (~100m below) granite contact Paleozoic NA (Cambrian-Devonian) - basin KR-05-71 - Gun pluton (62.8247, 129.9211) m-gr bt monzogranite 0.16 R some myrm, boxy-cellular pl cores and sieve textures, minor chl after bt, large high-relief colourless 1st order grey to yellow grain w/coarse-lamellar twins and inclusions of andalusite/chiastolite and zoned aln/mnz? Paleozoic NA (Cambrian-Devonian) - basin KR-05-87 - O'Grady batholith (62.9666, 128.6589) f-gr (qtz)-bt-fsp granodioritic porphyry 0.15 R pervasive ser (and shreddy bt?) alt'n overprinting everything, embayed and rounded qtz ph-xsts (possibly rimmed by myrm/granophyre) proximal: 105INE0022-STIB1 (Sb- skarn), 105INE0011-KAREN (Pb-Zn- Sn-Sb skarn) Paleozoic NA (Devono- Mississippian) - clastic influx 534 Summarya MSb R/Od Petrographic Notes Field Notes Mineral Prospects, Showings, and Deposits in Vicinity Main Stratigraphic Host KR-05-106 - felsic dyke, Cac pluton (62.3692, 128.5493) f-m-gr equigr ms-bt leucogranite dyke 0.03 R some chl after bt, moderate ser in pl and clay speckling in mc, spherulitic intergrowths of pl and qtz aplite dyke in skarnified country rock (incl. bt-poor margin, leucocratic tur-grt pockets, and dyke w/glassy green mineral: ap?) proximal: 105ISE0019-CAC (W skarn) Paleozoic NA (Devono- Mississippian) - clastic influx KR-05-112 - mafic dyke, South Nahanni pluton (62.5526, 128.6084) alt'd v f-gr equigr mafic dyke (lamprophyre?) 0.41 R mostly extremely f-gr chl/chd? (colourless, radiating to shreddy morphology, 1st order grey-brown) w/~5% ep mafic/alt'd aplite dyke(s) w/small fractures throughout cutting hornfelsed (min'd) and bleached dolostone proximal: "BLUE" (W skarn) Paleozoic NA (Cambrian-Devonian) - basin KR-05-115 - felsic dyke, South Nahanni pluton (62.5521, 128.5712) f--m-gr qtz-pl-bt monzogranitic porphyry dyke 0.59 O intensely developed granophyric and micrographic qtz-fsp intergrowths on margins of ph-xsts and throughout gr-m, ph-xsts may be rounded/embayed, pl is well-zoned and may have sieve-textures, trace ttn and aln, no mc observed (unless as granophyres), up to 10% clots of interstitial po dyke/plug proximal: "BLUE" (W skarn) Paleozoic NA (Cambrian-Devonian) - basin KR-05-154 - felsic dyke, Rifle Range pluton (62.0001, 128.1624) slightly foliated f-gr equigr bt leucogranite dyke 0.40 R nd about 10m across, and cut by qtz and Fe- CO3 veins; mineralized w/py or po Proterozoic NA (Windemere SG) - rift-related KR-05-156 - aplite, Rifle Range pluton (62.0035, 128.1661) v f-gr equigr aplitic dyke 0.03 R minor bt is all alt'd to pale chl ± rt needles, moderate ser in pl and minor- moderate clay speckling in mc, seriate texture cutting mc-porphyritic granite Proterozoic NA (Windemere SG) - rift-related KR-05-211 - felsic dyke (61.9815, 128.2034) v f-gr foliated (mc)- (ms)-pl-qtz-bt monzogranitic porphyry dyke 0.08 R qtz ph-xsts are rounded, pl is well-zoned but almost entirely replaced w/intense ser, bt may be alt'd to chl + rt needes, rare rounded mc ph-xsts, fibrolitic gr-m (qtz-fsp-ms?) south side of Dutchman's Peak Paleozoic NA (Cambrian-Devonian) - basin 03M-212 - felsic dyke (61.9699, 128.2420) f-gr pl-bt-mc-qtz- phyric leucocratic porphyry dyke 0.52 O qtz ph-xst are euh w/rare embayments; intensely developed granophyric- micrographic-spherulitic qtz-fsp intergrowths on margins of phenocrysts and throughout g-ms; perth + TT present in mc; pl ph-xst cores moderately to intensely sauss'd; bt mostly alt'd to ms-chl N side of the pit wall proximal: Paleozoic NA (Cambrian-Devonian) - basin Notes: Rock-forming mineral abbreviations are after Kretz, 1983, or are: feldspar = fsp, px = pyroxene, ser = sericite, sauss = saussurite; op = opaque. Descriptive and textural abbreviations are: cm = centimetre, w/ = with, v = very, abun = abundant, alt'n = alteration, f-gr = fine-grained, m-gr = medium-grained, c-gr = coarse-grained, equigr = equigranular, gr-m = groundmass, ph-xst = phenocryst, euh = euhedral, myrmekite = myrm, perth = perthite, TT = tartan twins, sub-s = sub-solidus. a brackets indicate a minor component; leucogranite is <3% mafic and (leuco)granite is <7% mafic; b magnetic susceptibility (x10-3SI units); c R = reduced; O = oxidized. 535 A.3 Images of geochronological and geochemical samples Images of polished slabs (left), and stained slabs (right). K-feldspar is bright yellow in the stained slabs. Samples are labeled by sample ID, name of the magmatic body (italicized if assigned by this study), and for plutonic samples, suite (bracketed). Plutonic suite abbreviations are: C = Cassiar, A = Anvil, TR = Tay River, Tu = Tungsten, M = Mayo, To = Tombstone, R = Rancheria, BR = Black River, HM = Hess–MacMillan sub-suite, CR = Coal River sub-suite, and ‘e’ = ‘evolved’. The white scale bar represents 1 cm, and “N/A” = not available. a) Cassiar Terrane and western Selwyn Basin samples (“KR-07” and “07M”). b) Eastern Selwyn Basin samples (“KR-05” and “CL-06”). 536 537 538 539 540 541 542 543 544 Appendix B – Supporting Materials for Reported U–Pb Ages 545 B.1 U–Pb geochronological methods. a) Laser Ablation – Inductively Coupled Plasma – Mass Spectrometry (LA–ICP–MS). b) Isotope Dilution – Thermal Ionization Mass Spectrometry (ID–TIMS). 546 a) Laser Ablation – Inductively Coupled Plasma – Mass Spectrometry (LA–ICP–MS). Sixty–two samples of plutonic rocks and one sample of a volcanic rock were analyzed by LA– ICP–MS methods at the Pacific Centre for Geochemical and Isotopic Research (PCIGR), University of British Columbia (UBC). Accessory minerals for each sample were picked from the heavy mineral separate and are tabulated in Table 2. A representative population of approximately 40–70 zircon grains from the heavy mineral separates for each sample was mounted in heat-curing epoxy with several grains from internationally accepted standard material (listed below), then ground and polished to a 1μm diamond paste until polished zircons faces were visible. The grain mounts were then carbon-coated and imaged in back- scattered electron (BSE) mode on the SEM in order to obtain an image of internal grain structure prior to LA–ICP–MS analysis. Typical zircon textures (zoning, core and rim morphologies, inclusions, etc.) for each sample are tabulated and summarized in Appendix A3. The carbon coat was then removed by re-polishing with the 1μm diamond paste, and the polished surfaces of the grain mounts were submerged in warm ultrapure 3N HNO3 for ten minutes then rinsed with ultrapure water and sub-boiled acetone. Only one grain mount was placed in the holding cell at a time, and after sample insertion at least 20 minutes was allowed for the removal of air from the sample chamber. Helium was used as the carrier gas for all experiments and gas flow rates, together with other parameters such as torch position, are optimized prior to beginning a series of analyses. Line scans rather than spot analyses were used in order to minimize within-run elemental fractionation, and each analysis consists of a 7 second background measurement of the carrier gas (at which time the laser is not ablating) followed by a ~28 second data acquisition period with the laser ablating the sample material. Generally, a minimum of 16 ablation lines were done for each intrusive sample for which appropriate grains existed (i.e., grains with minimal inheritance and well-developed magmatic rims), using a New Wave 213 nm Nd–YAG laser with a “Supercell” ablation chamber designed to achieve very high efficiency entrainment of aerosols into the carrier gas, and a Thermo Finnigan Element 2 High Resolution Inductively Coupled Plasma – Mass Spectrometer (ICP–MS). Each analysis consisted of a line scan at a speed of 2μm/s, 35–60% power, and a repetition rate of 20 Hz; a 5–10 μm-deep and 25–30 μm-wide spot size were used to ablate zircon grains for approximately 50 seconds resulting in a ~100 μm-long ablation path. Although an effort was made to analyze only the magmatic 547 rims of the zircon grains, frequently older magmatic and xenocrystic cores were also analyzed due to very thin magmatic rims in samples with a significant amount of inheritance. Typically four (although in some cases 3) beginning and end analyses each, as well as regularly interspersed analyses of accepted standard material (e.g., 1–2 standard analyses for every 3–4 unknown analyses) were used to monitor and correct for isotopic drift, mass discrimination inherent in the ICP–MS, and any elemental fractionation during ablation and transport of ablated material. These included the 115+6/–3 Ma “Sackinaw” in-house zircon standard, the 1099.1 +/– 0.2 Ma “FC-1” (AS3: Duluth Complex, anorthite series; Schmitz et al., 2003), and the 337.13 +/– 0.37 Ma “PL” (Bohemian Massif, Plešovice zircon; Sláma et al., 2008) standards, as well as two bracketing analyses of the 196.6 +/– 0.2 Ma (Pb206/ U238 ID–TIMS age; R. Friedman, pers. comm., 2009) internal monitor “KL” (Kloochlimmis). A minimum of 6–8 standard analyses were thus included in any dataset generated for each plutonic sample of unknown age, allowing the identification and subsequent rejections of anomalous analyses (e.g., Jackson et al., 2004). Many of these samples were first analyzed during the developmental stages of the LA–ICP–MS technique at PCIGR and some early technical issues resulted in problematic age determinations for several of the samples. These included ongoing decreasing laser strength, the initial use of the ion-counting (IC) mode whereby the 232Th and 238U isotopes frequently exceeded the capability of the detector, poor 207 count statistics due to extremely low 207Pb concentrations in the in-house standard Sackinaw due to its relatively young age, and difficulties with the FC1 standard due to low 235U. Ultimately, stronger ion beams and much improved counting statistics, leading to better precision and accuracy on individual analyses, was obtained after a lengthy period of troubleshooting and repair work that ultimately resulted in the replacement of the ‘laser’, changing the data collection to a mixed IC–analogue mode whereby the strongest isotopic peaks (238U and 232Th) were counted in analog mode and the remaining weaker peaks were counted in ion counting mode, and using PL as the standard reference material. However, it was not always possible to re-analyze some of the earlier samples under such these conditions, therefore some of the ages reported have larger errors (lower analytical precision) or less than ideal counting statistics. Follow-up imaging by cathodoluminescence (CL) on the SEM for analyzed grains was done for each sample. Subtraction of background isotope measurements, propagation of analytical errors, and calculated isotopic ratios and ages (with 1-sigma internal precisions) were produced for individual line scans by initial data reduction using GLITTER™ 4.4 software 548 package developed by the GEMOC group at Macquarrie University (e.g., Van Achterbergh et al., 2001; Jackson et al., 2004). This application generates a time-resolved record of each laser ablation line; therefore, if evidence existed in the time-resolved isotopic ratios or isotopic count/second signals of individual line scans for lead loss, inheritance, or the interference of common lead, the line scans were either modified in GLITTER™ 4.4 to remove portions of the analysis with problematic data, or were not used for age determinations. Further details on the GLITTER™ software package and error propagation are described by Jackson et al. (2004, and references therein). Examination of BSE and CL images were also used to discriminate between magmatic analyses and analyses of significantly older magmatic and xenocrystic inheritance and zones of resorption and probable lead loss. Additionally, discordant analyses were typically rejected unless they were well within the range of error ellipses of concordant analyses, and means of measured 206Pb/238 U ages weighted by assigned errors were calculated using Isoplot 3.7 (Ludwig, 2008). Accuracy of weighted mean age determinations averaged 0.84%, which is close to the typical 0.8–1% error for reported LA– ICP–MS U–Pb ages from zircon (e.g., Tiepolo, 2003; Jackson et al., 2004; Chang et al., 2006), and ranges from 0.42–2.30%. Due to the Cretaceous (and younger) age of the samples and the consequent lack of 207Pb in the zircons, only 206Pb/238 U ages were used to determine the crystallization ages and 2-sigma uncertainties (within the 95% confidence interval, unless otherwise noted). The results of standard analyses are not reported with the data for the unknown samples; however, these are available on the accompanying DVD containing raw analytical data for this thesis. Weighted mean plots used for calculating final 206Pb/238 U crystallization ages, and concordia plots used to monitor for lead loss or inheritance are presented in Appendix B.5. All isotopic ages and counts per second for the individual analyses used for final age calculations, as well as conditions of analysis, are tabulated in Appendix B.6. 549 b) Isotope Dilution – Thermal Ionization Mass Spectrometry (ID–TIMS). Samples from five Cretaceous plutons emplaced into the Yukon–Tanana terrane inlier northwest of the TF, one felsic dyke cross-cutting the Clear Lake SEDEX deposit in drill core immediately southwest of the TF, and several plutons and felsic dykes in the Cantung mine area of the eastern SB (Table 1) were analyzed by ID–TIMS methods at the Pacific Centre for Geochemical and Isotopic Research (PCIGR), University of British Columbia (UBC). Mineral fractions for analysis were selected based on grain quality, size, magnetic susceptibility and morphology. All zircon fractions were air abraded prior to dissolution to minimize the effects of surface-correlated lead loss, using the technique of Krogh (1982). Zircons were then washed with warm ultrapure 3N HNO3, monazites were washed with 1 N HNO3, and both were rinsed with ultrapure water and sub-boiled acetone and weighed (to ± 2mg). Zircons were dissolved in 300 mL PTFE or PFA microcapsules with ~100 mL of sub- boiled 29N HF and ~15 mL of sub-boiled 14N HNO3 in the presence of a mixed 233–235U– 205Pb tracer for 40 hours at 240°C. Dissolution took place in stainless steel Parr bombs with 250 mL Teflon PTFE liners. Sample solutions were then dried to salts at ~125°C and re- bombed in ultrapure ~200 mL 3.1N HCl for 12 hours at 210°C. Monazites were dissolved in ~1 mL of sub-boiled 6.2 N HCl for a minimum of 72 hours in 3.5 ml PFA screw top beakers on a hot plate at ~125°C, also in the presence of a mixed 233–235U–205Pb tracer. These solutions were then dried to salts on a hot plate at ~125°C and re-dissolved in ~0.5 mL 3.1 HCl on a hot plate at ~125°C. Lead and uranium separation for both zircon and monazite employed ion exchange column techniques similar to those described by Parrish et al. (1987). Lead and uranium were eluted sequentially into the same beaker followed by the addition of ~10μL of 0.6N ultra pure phosphoric acid. Each sample was loaded onto a single zone refined Re- filament using a phosphoric acid-silica gel emitter (SiCl4). Isotopic ratios were measured using a modified single collector VG–54R thermal ionization mass spectrometer equipped with an analogue Daly photomultiplier. Both uranium and lead, respectively, were run at 1450°C, in peak-switching mode on the Daly detector. Uranium fractionation was determined directly on individual runs using the 233–235U tracer, and lead isotopic ratios were corrected for fractionation of 0.37%/amu, based on replicate analyses of the NBS–981 Pb standard and the values recommended by Thirlwall (2000). Uranium and lead analytical blanks were less than 1 pg and 3 pg, respectively. Common lead isotopic compositions are derived from the model 550 of Stacey and Kramers (1975), and all analytical errors were numerically propagated through the entire age calculation using the technique of Roddick (1987). Concordia intercept ages and associated errors were calculated using a modified version the York–II regression model (wherein the York–II errors are multiplied by the MSWD) and the algorithm of Ludwig (1980). Accuracy of age determinations averaged 0.31% and ranges from 0.19–0.62%. Isotopic ratios, ages, and associated uncertainties for all the mineral fractions analyzed and concordia plots used to determine the final ages are presented in Appendices B.7 and B.8. 551 B.2 Zircon separation details for samples analyzed by LA–ICP–MS methods. a) Heavy liquid and Franz separation details. b) Zircon fraction descriptions. 552 a) Heavy liquid and Franz separation details. Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) 07M-150 14- Mar- 09 moderate amounts of dark grey material moderate amount of pale yellow-green and black material 15- Mar- 09 moderate amount (~1/3) black > green moderate amount (~1/3) green >> black very small amount green and pale grey small amount pale brown- cream n/a n/a ttn, Ca>>Fe px, zrn, ap 07M-151 14- Mar- 09 very little pale grey material small amount of black and pale yellow material 15- Mar- 09 moderate amount (~1/3) black moderate amount (~1/3) black very small amount black >> brown small amount cream n/a n/a Fe>Mg-Ce-La-(Th) aln, crn, zrn, ap 07M-152 14- Mar- 09 very little pinkish-grey material (iron filings) moderate amount of red- pink, pale yellow, and black material 15- Mar- 09 moderate amount (~1/2) red garnet concentrate moderate amount (~1/4) black, some red and brown small amount black > brown small amount brown >> black small amount cream-grey n/a Fe>Mg-(Ce-La-Th) aln, Fe>Mg>Mn grt, rt, zrn, ap 07M-153 14- Mar- 09 very little pale grey (pinkish?) material very small amount of pale yellow and black material 15- Mar- 09 small amount (~1/4) black, green, and brown small-moderate amount (<1/2) yellow > green > black small amount (~1/4) pale pinkish cream n/a n/a n/a Fe>Mg-Ce-La aln, Ce- La-Th-(Ca-Ta-Nb) mnz, Ca>>Fe px, zrn, ap 07M-154 14- Mar- 09 very little dark grey material moderate amount of black, pale yellow, and pale pink- brown material 15- Mar- 09 moderate to large amount (~1/2) black >> brown moderate amount brown small amount brown small amount pale brown- grey moderate amount pale brown-grey n/a Fe>Mg-Ce-La aln, Fe- Ca thor, Fe>Nb ttn, Ca>(Fe) grt, Ca>(Fe) px, hbl, zrn, ap KR-05-08 26- Jan- 06 moderate amount brown material small- moderate amount black = brown material 26- Jan- 06 small amount black > red-brown (oxidized black) small amount black > red brown glassy and grey glassy small amount black > red brown glassy and grey glassy very small amount glassy red-brown >> black very small amount glassy red-brown and golden tiniest amount golden blonde Fe-Mg-Ce-La aln, zrn, ap KR-05-10 25- Jan- 06 moderate amount material moderate amount green > black-brassy 26- Jan- 06 small-moderate amount green (ep?) >> black small-moderate amount green (ep?) >> black small amount brassy (pyrite) > green VERY small amount brassy (pyrite) very small amount brassy (pyrite) tiniest amount pyrite > grey Fe-Mg-Ce-La aln, zrn, ap KR-05-110 14- Feb- 06 little handmag small amount of black material 21- Feb- 06 very small amount black > grey- brown small amount black VERY small amount black > grey VERY small amount black = blonde-grey n/a n/a Fe-(Mg-Ce-La-Th) aln, Ce-Nd-La-Th-(Ta-Nb) mnz, Fe>Ca>Mg-(Mn) grt, zrn, ap KR-05-113 14- Feb- 06 very little handmag small amount of black material 21- Feb- 06 (very) small amount black small amount black very small amount brown VERY small amount brown VERY small amount black VERY small amount grey- blonde Fe-(Mg-Ce-La) aln, zrn, ap 553 Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) KR-05-126 14- Feb- 06 very little handmag very small amount red- brown black (oxidized) 21- Feb- 06 VERY small amount black w/brown oxide tiniest amount colourless > mix of other material n/a n/a n/a n/a tur, zrn KR-05-130 14- Dec- 05 very small amount small- moderate amount black 22- Jan- 06 moderate amount black (hornblende) very small amount black tiniest amount black tiniest amount colourless and black tiniest amount colourless > black tiniest amount colourless Fe-Mg-Ce-La-(Th) aln, zrn, ap KR-05-136 9- Nov- 05 very small amount (metal filings) small amount black > pale yellow-brown 22- Jan- 06 small amount black (hornblende) very small amount black (hornblende) tiniest amount colourless > black (rare pyrite) tiniest amount pale brown very small amount pale brown >/= black very small amount pale brown >>> black Fe-Mg-Ce-La-(Th) aln, zrn, ap KR-05-143 26- Jan- 06 very little handmag (steel filings?) VERY small amount black = golden 26- Jan- 06 tiniest amount black > red (garnet) VERY small amount pale yellow > black tiniest amount of colourless and yellow zircon n/a n/a n/a Ca-Fe-Mg-Ce-La aln, Ce-La-Th-Nd-Ta-Ca mnz, Fe>Mg-(Ca-Mn) grt, zrn, ap KR-05-148 14- Dec- 05 very small amount (more after second handmag) small- moderate amount black = blonde 22- Jan- 06 small amount black w/minor oxide small amount yellow (monazite?) >>> black VERY small amount pale grey very small amount pale grey-brown very small amount pale grey-brown > black n/a Fe-Mg-Ce-La-(Th) aln, Ce-La-Th-Nd-Ta mnz, zrn, ap KR-05-164 16- Dec- 05 small amount small amount blonde = black 22- Jan- 06 small amount bright pink garnet > black small amount yellow (monazite?) >>> colourless tiniest amount colourless VERY small amount colourless VERY small amount colourless n/a Ce-La-Th-Nd-(Ta) mnz, Fe>Mg-(Ca-Mn) grt, wolframite, Ta xeno, zrn, ap KR-05-175 14- Feb- 06 very little handmag very small amount blonde-brown > black 21- Feb- 06 very small amount blonde-brown >> black VERY small amount black VERY small amount blonde-(brown) VERY small amount blonde > black VERY small amount blonde > black VERY small amount yellow brown Mg-Ce-La-(Fe) aln, Ce-La-Th-Nd-Ta-Ca mnz, Ca-(Ti-Fe) grt, F ttn, zrn, ap KR-05-191 16- Dec- 05 lots handmag moderate amount black > green > medium brown 22- Jan- 06 moderate amount black > glassy green> glassy brown (some zircon may have been trapped here) small-moderate amount glassy brown (titanite) >> green and black VERY small amount glassy green > brassy and colourless VERY small amount brassy (pyrite) > colourless VERY small amount colourless to pale brown > brassy VERY small amount colourless to pale brown > brassy Ca-Fe-Mg-Ce-La-(Th) aln, Ca-(Fe) grt, zrn, ap KR-05-194 14- Dec- 05 VERY little material small- moderate amount of black >>> grey-blonde material 22- Jan- 06 moderate amount black VERY small amount grey- blonde > black VERY small amount grey- blonde > black VERY small amount grey- blonde small amount grey-blonde n/a Ca-Fe-Mg-Ce-La aln, Ce-La-Th-Nd-(Ta) mnz, Pb-P thor, F ttn, zrn, ap KR-05-196 16- Dec- 05 lots handmag moderate amount black > golden material 24- Jan- 06 moderate amount black >>> lime green small amount pale brown >>> black and lime green VERY small amount pale brown >>> black and lime green tiniest amount colourless > medium brown VERY small amount colourless > medium brown VERY small amount colourless Ca-Fe-Mg-Ce-La-(Th) aln, (Nb-F-Y) ttn, ep, zrn, ap 554 Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) KR-05-198 14- Dec- 05 very little handmag moderate amount black >> blonde 24- Jan- 06 moderate amount black tiniest amount black > colourless tiniest amount black = blonde VERY small amount blonde and colourless small amount blonde and colourless small amount blonde and colourless Ca-Fe-Mg-Ce-La-(Te- Th) aln, zrn, ap KR-05-208 14- Feb- 06 very little handmag small amount pale yellow- grey > black 21- Feb- 06 very small amount black > pale pink garnet pale yellow grey >> black tiniest amount colourless = medium brown tiniest amount colourless = pale brown n/a n/a Ca-Fe-Mg-Ce-La-(Th) aln, Th-Ce-La-(Ca-Nd) mnz, Fe>Mg-Mn grt, uran, zrn, ap KR-05-210 16- Dec- 05 little handmag small amount black = pale brown 25- Jan- 06 small amount black > grey >> pink garnet small amount opaque yellow-pink-brown (monazite?) >>> black tiniest amount pale brown > pink garnet and colourless tiniest amount pink-brown tiniest amount pale brown- yellow and brown-pink tiniest amount pale brown-yellow and colourless Ca-Fe-Ce-(La) aln, Th-Ce-La-(Ca) ± Nd mnz, Fe-Ca-Mn-Mg aluminosilicate (grt or piedmontite), zrn, ap KR-05-215 nd VERY minimal black material VERY small sample nd some black material some yellow(monazite?)>black material some yellow(monazite?)>black material small amount ash-blonde material w/some black and grey material n/a n/a Fe>Ca-(Mg-Mn) grt, F ttn, zrn, ap KR-05-22 14- Dec- 05 lots magnetite moderate amount black = brown material 12- Jan- 06 moderate amount black >> medium red-brown moderate amount medium red-brown (titanite) >>> black tiniest amount yellow = colourless tiniest amount yellow > colourless tiniest amount yellow, pale grey-brown, and colourless very small amount pale grey-brown to colourless Fe-Mg-Ce-La-(Th) aln, ttn, zrn, ap KR-05-26 25- Jan- 06 large amount small amount black >> blonde 25- Jan- 06 small amount black tiniest amount blonde tiniest amount blonde tiniest amount blonde tiniest amount blonde n/a Fe-Mg-Ce-La-(Th) aln, Ca-(Fe) grt, Fe>Mg- (Mn) grt, zrn, ap KR-05-32 25- Jan- 06 very small amount garnet- rich material very small amount garnet-rich 25- Jan- 06 very small amount red >> black very small amount golden brown tiniest amount blonde? tiniest amount blonde? VERY small amount golden brown tiniest amount blonde Ce-La-Th-(Nd-Ta-Nb) ± Ca mnz, Fe>Mn- (Mg-Ca) grt, trm, xeno, zrn KR-05-43 14- Dec- 05 moderate amount material moderate amount black > pale brown material 19- Jan- 06 small amount black >> glassy green and glassy med-brown (grains got stuck; may have lost some v-f-gr pale yellow grains) moderate amount pale honey-brown (titanite) >> black tiniest amount colourless > brown and black tiniest amount colourless > brown and black tiniest amount colourless > brown and black tiniest amount colourless to pale yellow Fe-Mg-Ce-La-(Th), (Nb-F-Y) ttn, zrn, ap KR-05-62 14- Dec- 05 moderate magnetite moderate- large amount red-brown > black material 19- Jan- 06 small-moderate amount black >> red-brown moderate amount red- brown (titanite) > black moderate amount red- brown (titanite) >> black VERY small amount colourless > dark brassy (pyrrhotite?) tiniest amount colourless > black VERY small amount blonde Fe-Mg-Ce-La-(Th) aln, Pb-P-Ca-F thor, zrn, ap 555 Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) KR-05-68 14- Dec- 05 VERY small amount small amount black material 19- Jan- 06 very small amount black (stubby) >>> grey brown VERY small amount grey- brown > black tiniest amount colourless > green-brown (elongate) tiniest amount colourless tiniest amount blonde n/a Fe-Mg-Ce-La aln, Fe>Ca>Mg-(Mn) grt, zrn, ap KR-05-76 14- Dec- 05 very small amount tiniest amount black material 19- Jan- 06 tiniest amount black tiniest amount colourless = dark brown (zircon?) tiniest amount colourless = dark brown (zircon?) n/a n/a n/a Fe-Mg-Ce-La-(Th) aln, scheelite, zrn, ap KR-05-97b 25- Jan- 06 large amount large amount red-brown > black material 26- Jan- 06 small amount black = brown large amount (~4/5) brown > black (titanite concentrate) small amount red-brown small amount brassy (pyrite) > black small amount brassy (pyrite) > black very small amount pale grey-blonde > black Fe-Mg-Ce-(La) aln, Mg>Fe-(Ca) grt (prp core, alm rim?), tur, zrn, ap KR-07-01 14- Mar- 09 essentially no material (grey) small to moderate amount black and pale brown material 15- Mar- 09 large amount (~2/3) black moderate amount (~1/4) rusty brown >> black very small amount black and grey small- moderate amount cream n/a n/a Ce-La-Th-(Ca-Ta-Nb) mnz, Ti-Ca-Pb-(Fe) uran, rt, zrn, ap KR-07-02 14- Mar- 09 little pale grey material (very) small amount black and pale yellow-brown material 15- Mar- 09 small amount black small amount brown small amount white-cream n/a n/a n/a Ce-La-Th-(Ca) mnz, Fe>Mg>Mn grt, Mg>Fe px, rt, zrn, ap KR-07-03 14- Mar- 09 little pale grey material small amount red and pale brown material 15- Mar- 09 moderate amount (~1/4) red garnet concentrate; rare black small amount (>1/2) yellow small amount cream n/a n/a n/a Ta xeno, Ce-La-Th- (Ca) mnz, Fe>Mn grt, zrn KR-07-04 14- Mar- 09 little to some reddish grey material (iron filings?) small amount cream and black material 15- Mar- 09 moderate amount (~1/2) black small amount black and brown small amount brown > black very small amount brown >> black moderate amount (~1/2) cream n/a toz, xeno (inclusions in bt), zrn KR-07-05 15- Mar- 09 very little reddish grey (iron filings?) moderate amount red material 22- Mar- 09 moderate amount (almost all of sample) red-pink garnet concentrate very small amount brown- yellow small amount pale blonde n/a n/a n/a Ce-La-Th-(Ca) mnz, Fe>Mg>Mn grt, zrn KR-07-06 15- Mar- 09 ~1/2 sample black magnetite after second handmag, large amount of black < brown material 22- Mar- 09 large amount (~2/5) black >>> brown large amount (~2/5) med- brown > black small amount light-brown >> black small amount pale brown >> red small- moderate amount (~1/2) blonde n/a Mg>Fe-Ce-La aln, Fe>Nb ttn, U-(Ca-F) thor, betafite, zrn, ap KR-07-07 15- Mar- 09 very little red- grey (iron filings?) small amount black < brown material 22- Mar- 09 moderate amount black moderate amount light- brown >/= black very small amount light- grey-brown > black very small amount light- grey-brown >> black small amount light-brown n/a Fe>Mg-Ce-La-(Th) aln, Ce-La-Th-Ta-Nb- (Ca) mnz, zrn, ap 556 Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) KR-07-08 20- Mar- 09 ~1/2 sample black magnetite, with some brown material after second handmag (~1/4 removed), moderate amount black and brown material 22- Mar- 09 moderate amount black > brown moderate amount black > brown small amount brown = black small amount brown > black small amount grey-blonde > black n/a Fe>Mg-Ce-La aln, Ca grt, zrn, ap KR-07-09 20- Mar- 09 small amount silvery-black and brown material small- moderate amount black > brown material 22- Mar- 09 small-moderate amount black small amount black = brown small amount blonde n/a n/a n/a Ce-La-Th-Ta-Nb-Nd mnz, Mg>Fe grt, Ca>Fe grt, zrn, ap KR-07-10 20- Mar- 09 ~1/3-1/2 black magnetite, with some brown material after second handmag (~2/3-3/4 removed), small- moderate amount light- brown >> black 22- Mar- 09 small amount black >> brown small amount med-brown > black small amount yellow- blonde n/a n/a n/a Ce-(La) aln, Ca grt, Ce-La-Th-Ta-Nb-(Ca) mnz, ttn, zrn, ap KR-07-11 20- Mar- 09 very little reddish grey (iron filings?) small amount mostly yellow- blonde material 22- Mar- 09 very small amount black = brown very small amount brown- yellow very small amount blonde n/a n/a n/a Fe>Mg grt, (Ta) xeno, Th-(Ca-Ta-Nb) mnz, zrn KR-07-12 20- Mar- 09 very little reddish grey (iron filings?) moderate amount black = brown material 22- Mar- 09 moderate amount black smll amount yellow > black small amount blonde > black n/a n/a n/a Ta-Nb-(Fe) rt, Ce-La- Th-(Ca-Ta-Nb) mnz, tur, zrn, ap KR-07-13 20- Mar- 09 very little reddish grey (iron filings?) small amount yellow-brown > black 5- Apr- 09 next to nothing; combined with 0.6-1.8 small amount golden- brown > black very small amount black very small amount grey- blonde n/a n/a Ta-Nb-(Fe) rt, Ce-La- Th-(Ca-Ta-Nb) mnz, tur, zrn, ap KR-07-14 3- Apr- 09 very little reddish grey (iron filings?) and black material small- moderate amount black > yellow-brown and blonde material 5- Apr- 09 small amount black small amount yellow-gold small amount blonde n/a n/a n/a (Ta) xeno, (Th) uran, Ce-La-Th-(Ca-Ta-Nb) mnz, Fe>Mg grt, zrn, ap KR-07-15 3- Apr- 09 very little reddish grey (iron filings?) small amount black > pale yellow-brown 5- Apr- 09 very small amount black = brown v(Very) small amount yellow-blonde > brown/black VERY small amount pale grey > brown very small amount grey- blonde > brown n/a n/a (Ce-La) aln, Ce-La- Th-(Ca-Ta-Nb) mnz, Nb-Y-Ta-U-Dy-Yb-Th oxide (possibly betafite or formanite- fergusonite or kobeite), zrn, ap 557 Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) KR-07-16 3- Apr- 09 ~1/3 sample black magnetite after second handmag (~5% removed), large amount black material 5- Apr- 09 large amount black (first run at 0.3 Amps) >>> brown (second run at 0.6 Amps) small-moderate amount green-yellow > brown small amount blonde n/a n/a n/a Fe>(Mg-Ce-La) aln, Ce-La-Th-(Ca-Ta-Nb) mnz, Ca grt, Ca>>Fe px, zrn, ap KR-07-17 3- Apr- 09 moderate amount sooty black material moderate amount black >>> green- brown-yellow material 5- Apr- 09 moderate amount black (first run at 0.3 Amps) >>> brown (second run at 0.6 Amps) moderate amount black = brown very small amount pale- grey and brassy (pyrite) small amount brassy (pyrite) n/a n/a Ce-La-Th-(Ca-Ta-Nb) mnz, zrn, ap KR-07-18 3- Apr- 09 small amount brown-black material large amount black >>> red- brown material 5- Apr- 09 large amount black (first run at 0.3 Amps) = red- brown (second run at 0.6 Amps) small amount coarse- grained red-brown, and yellow small amount blonde >> brown n/a n/a n/a Ce-La-Th-(Ca-Ta-Nb) mnz, zrn, ap KR-07-19 4- Apr- 09 ~1/2 sample black magnetite, with moderate amount of brown material after second handmag (~5% brown + black removed), large amount blonde- brown/red >> black material 5- Apr- 09 moderate amount black = red-brown moderate-large amount light-red-brown/gold >>> black very small amount blonde >>> brown and black small amount blonde >>>brown n/a n/a Fe>Mg-Ce-La-(Th) aln, ttn, Ca>>Fe px, zrn, ap KR-07-20 4- Apr- 09 ~1/2 sample black magnetite after second handmag (removed ~10% black + brown), moderate amount yellow- brown = black material 5- Apr- 09 small amount black >>> brown small amount yellow-brown >> black (very) small amount blonde n/a n/a n/a Fe>Mg-(Ce-La-Th) aln, Ce-La-Th-(Ta-Nb- Ca) mnz, zrn, ap KR-07-21 4- Apr- 09 ~1/3 black magnetite(?), with some light- brown material after second handmag (~5% black > brown removed), large amount black > brown material 5- Apr- 09 moderate-large amount black (first run at 0.3 Amps) >> dull-brown (second run at 0.6 Amps) small-(moderate) amount black = red-brown very small amount blonde = dull-brown small amount blonde n/a n/a (Ce-La) aln, Ce-La- Th-(Ta-Nb) mnz, F>Fe ttn,Ca-Al-Ta-Nb rt, Nb- Y-Ta-U-Dy/As-Yb-Th- Ca oxide (possibly betafite or formanite- fergusonite or kobeite), Ca grt, zrn, ap 558 Sample Date Handmag Recovery Date 0.0-0.6 (garnet, pyroxene, hornblende, epidote, rock fragments) 0.6-1.8 (most titanite, most monazite, most rutile, some non-magnetic epidote) M10 (non-magnetic titanite and possibly monazite; some non-magnetic rutile) M5 (magnetic zircon; some v. weakly magnetic titanite) M2 (weakly magnetic zircon; chalcopyrite) NM2 (non- magnetic zircon, pyrite, barite, apatite) Heavy Minerals detected by SEM (mineral abbreviations from Kretz, 1983) KR-07-22 4- Apr- 09 very little reddish grey (iron filings?) moderate amount green- yellow-grey material >> black material; very pyrite-rich 5- Apr- 09 very small amount reddish-grey small amount pale-green- blonde-grey >> black small amount pyrite, some blonde n/a n/a n/a (Ta) xeno, Ce-La-Th- (Ta-Nb-Ca) mnz rt, zrn KR-07-23 4- Apr- 09 very little reddish grey (iron filings?) small amount yellow- brown/red > black material 5- Apr- 09 very small amount dark-brown (very small amount yellow by accident, probably 0.6-1.8) small amount yellow- orange VERY small amount blonde-grey (some yellow- black material by accident) n/a n/a n/a (Ta) xeno, Ce-La-Th- (Ta-Nb-Ca) mnz, Ca>Fe>Mg grt, Fe-Mg ttn, Mg px, zrn, ap KR-07-24 4- Apr- 09 ~1/4 black magnetite, with little brown material after second handmag (~5% black removed), moderate- large amount black > brown material 5- Apr- 09 moderate amount black (first run at 0.3 Amps) >> dull- browngreen-grey- brown (second run at 0.6 Amps) very small amount yellow >> black very small amount blonde- grey n/a n/a n/a Fe>Mg-Ce-La-Th aln, Ce-La-Th-Ta-Nb-Ca mnz, Ca grt, zrn, ap CL-06-33 nd very little black material small amount material nd black black > brown very small amount yellow > black black (most stuck on ramp) moderate ash blonde zircon n/a Ca-Fe-Mg-Ce-La-(Th) aln, zrn, ap CL-06-34 nd nd nd nd nd nd nd nd nd nd Ca-Fe-Mg-Ce-La-(Th) aln, zrn, ap CL-06-35 nd nd nd nd nd nd nd nd nd nd Ca-Fe-Mg-Ce-La-(Th) aln, zrn, ap CL-06-36 nd nd nd nd nd nd nd nd nd nd Ca-Fe-Mg-Ce-La-(Th) aln, zrn, ap CL-06-37 nd nd nd nd nd nd nd nd nd nd Ca-Fe-Mg-Ce-La-(Th) aln, zrn, ap CL-06-38 nd VERY minimal black small- moderate amount material nd lots black material small amount white < black very small amount white < black small amount white > black material lots ash blonde zircon n/a zrn, ap CL-06-39 nd minimal black material large amount material nd large amount black non- prismatic mineral yellow titanite > abundant black yellow titanite lots ash blonde zircon n/a (lots black material clogging up ramp dumped into 0.0-0.6 fraction Ca-Fe-Mg-Ce-La-(Th) aln, zrn, ap CL-06-40 nd small-moderate amount black material large amount material nd abundant black abundant black and pale glassy green yellow titanite w/minimal black some yellow titanite abundant ash blonde zircon n/a Ca-Fe-Mg-Ce-La-(Th) aln, tur, zrn, ap 559 b) Zircon fraction descriptions. Sample Mineral Fraction Code Colour Clarity Morphology Zoning Inclusions Cores % Broken Degree of Fracturing Comments 07M-150 zircon M5 colourless good elongate-prismatic (lesser small-equant and long-skinny) not obvious round black- opaques abundant; apatite noted not ovious (rare turbid cores) <50 most unfractured to moderately fractured very similar to 07M-151 07M-151 zircon M5 colourless good elongate-prismatic (lesser small-equant and long-skinny) not obvious round black- opaques abundant; apatite noted not ovious (rare turbid cores) <50 most unfractured very similar to 07M-150 07M-152 zircon M2 colourless to yellow good; large yellow grains are slightly cloudy equant to prismatic (lesser long-skinny) not obvious rare black-opaques; apatite noted some turbid cores visible >50 most unfractured lots of pyrite; generally very small grains (yellow grains are larger) 07M-153 zircon M10 colourless to yellow good; turbidity increases with darker colour short-prismatic to equant-prismatic not obvious rare black-opaques; apatite noted some turbid cores visible ~50 most slightly to moderately fractured (esp. yellow grains) pyrite present 07M-154 zircon M2 mostly yellow; some colourless good stubby-prismatic to short-prismatic not obvious black-opaque and red not obvious <50 most slightly fractured KR-05-08 zircon NM2 colourless to yellow-orange very good short-prismatic with faceted points (elongate-prismatic) not obvious very few some visible <<<50 unfractured zircon is relatively fine- grained but abundant; some apatite KR-05-10 zircon NM2 colourless to very pale yellow (one yellow) very good elongate-prismatic to stubby-prismatic not obvious few; some black- opaque, orange few turbid cores ~50 mostly unfractured; few highly fractured very abundant pyrite; zircon not abundant KR-05-110 zircon M5 colourless to dark yellow some turbid elongate-prismatic to very elongate-prismatic visible in some grains minor round black- brown some turbid cores visible >50 most moderately to highly fractured KR-05-113 zircon M2 colourless to dark yellow good (colourless); turbid (yellow) elongate-prismatic (colourless); stubby- prismatic (yellow) visible in some grains some apatite some turbid cores visible <<50 yellow are slightly to moderately fractured lots of apatite KR-05-126 zircon 0.6-1.8 colourless to pale yellow good stubby-prismatic to elongate-prismatic visible in some grains rare some turbid cores visible <<50 unfractured to slightly fractured KR-05-130 zircon M2 colourless to very pale yellow very good elongate-prismatic some visible few some turbid cores visible <50 some are fractured lots of apatite KR-05-136 zircon M2 colourless to dark yellow good very elongate-prismatic some visible few some turbid cores visible <50 abundant fractures zircon and apatite are abundant; might have picked some apatite KR-05-143 zircon M5 colourless to yellow-orange good; darker grains are cloudy stubby-prismatic to elongate-prismatic visible in some grains some black-opaque not obvious <50 highly fractured to almost unfractured KR-05-148 zircon M2 colourless to pale yellow good; darker grains are cloudy stubby-prismatic to short prismatic? present few black-opaque, moderately abundant apatite not obvious >50 half of grains are highly fractured apatite and zircon are abundant KR-05-164 zircon M2 colourless to very pale yellow good prismatic; some stubby/rounded present black, red-brown, few apatite not obvious <50 moderately fractured abundant coarse-grained apatite 560 Sample Mineral Fraction Code Colour Clarity Morphology Zoning Inclusions Cores % Broken Degree of Fracturing Comments KR-05-175 zircon M2 colourless to dark yellow good, except fractured grains elongate-prismatic (stubby-prismatic and very elongate-prismatic) visible in some grains minor black- opaque, apatite some turbid cores visible >50 darker grains are moderately-highly fractured may have mounted a few apatite KR-05-191 zircon M2 colourless to pale yellow good; darker grains are cloudy stubby-prismatic to elongate-prismatic present few black-opaque, some apatite not obvious <50 half of grains are highly fractured abundant pyrite and apatite KR-05-194 zircon M2 colourless to pale yellow good, except fractured grains prismatic (few rounded) not obvious some, black-opaque and red-orange not obvious >>50 most moderately to highly fractured KR-05-196 zircon M2 colourless to pale yellow (larger grains are darker) good prismatic (few rounded) some visible few, black-opaque, apatite few turbid cores <50 most moderately fractured KR-05-198 zircon NM2 colourless (to yellow) very good, except for inclusions short-prismatic to VERY elongate-prismatic some visible lots VERY small colourless, few black-opaque not obvious <<50 minor fractures very abundant and pristine zircon; lots apatite KR-05-208 zircon M10 colorless to pale yellow good short-prismatic to elongate prismatic (stubby-prismatic) visible in some grains minor very small black-red some turbid cores visible <50 some slightly to moderately fractured KR-05-210 zircon NM2 yellow to pale yellow (rare colourless) good short-prismatic to very elongate-prismatic visible in some grains minor black- opaque; minor apatite some clear, some turbid <50 more fractures in larger grains KR-05-215 zircon N10 colourless to pale brown good stubby to elongate prismatic visible in some grains minor apatite minor turbid <50 most moderately to highly fractured KR-05-22 zircon M2 colourless to yellow good, except fractured grains stubby-prismatic (larger grains look rounded) present abundant apatite some visible <50 highly fractured abundant apatite KR-05-26 zircon M2 colourless to dark yellow good, except fractured grains stubby-prismatic to elongate-prismatic visible in some grains few black-opaque some turbid cores visible <50% some highly fractured, some unfractured very abundant zircon KR-05-32 zircon NM2 colourless to pale yellow good, except fractured grains stubby-prismatic to short/elongate-prismatic not obvious lots black-opaque some turbid cores visible >>50 some highly fractured, most fractured KR-05-43 zircon M2 colourless to yellow very good stubby-prismatic to elongate-prismatic some visible some few turbid cores ? unfractured KR-05-62 zircon M2 colourless to pale yellow very good elongate-prismatic, esp. large grains not obvious some apatite not obvious <<50 moderately fractured KR-05-68 zircon M2 colourless to pale yellow good, some turbid elongate-prismatic (rounded) not obvious ? some darker cloudy cores <<50 mostly unfractured; larger grains more fractured lots of apatite KR-05-76 zircon M10 colourless to dark yellow good elongate-prismatic (some rounded darker grains) present few some turbid cores visible in darker grains <50 relatively unfractured KR-05-97b zircon M2 colourless/very pale yellow to yellow good, except fractured grains prismatic not obvious few black-opaque, apatite not obvious <<50% most highly fractured very abundant pyrite; zircon also abundant but relatively small KR-07-01 zircon M5 pale to dark yellow good; turbidity is rare short-prismatic not obvious minor black- opaque; to fractured to see inclusions not obvious >50 most unfractured to moderately fractured lots of quartz-apaitte- feldspar in separate; zircon not abundant 561 Sample Mineral Fraction Code Colour Clarity Morphology Zoning Inclusions Cores % Broken Degree of Fracturing Comments KR-07-02 zircon M10 colourless to yellow good short-prismatic to stubby-prismatic not obvious some red-brown; rare black-opaque elongate not obvious <50 most unfractured to moderately fractured most very small; not much zircon KR-07-03 zircon M10 colourless to yellow good short-prismatic to long- prismatic not obvious black-opaque and minor red-brown not obvious <50% most moderately to highly fractured KR-07-04 zircon M5 (M2, M10) mostly yellow- brown; rare grey- colourless mostly turbid irregular, corroded; elongate-prismatic to stubby-prismatic to equant-prismatic not obvious not visible due to turbidity not obvious >>50% nd zircon is not abundant in any fraction; generally very small; several with simple twins; largest and most abundant in M5 KR-07-05 zircon M10 colorless < pale yellow-brown good; turbidity increases with darker colour short-prismatic to elongate-prismatic; can be round and wide in 2D not obvious rare not obvious <50 unfractured to highly fractured KR-07-06 zircon M2 colourless to light yellow very good to slightly cloudy (fractures?) short-prismatic to elongate-prismatic not obvious minor black- opaque; apatite not obvious <50 unfractured to slightly fractured; larger grains more fractured minor pyrite KR-07-07 zircon M2 pale to med-yellow good short prismatic some visible minor black- opaque; minor apatite turbid cores abundant >50 most moderately to highly fractured abundant pyrite KR-07-08 zircon M2 colourless to pale yellow very good short-prismatic not obvious minor small black- opaque; apatite few turbid cores <50 unfractured very abundant pyrite KR-07-09 zircon M10 colourless to yellow good to dull/opaque short-prismatic to elongate-prismatic; can be round and wide in 2D present rare black-opaques; apatite noted present >50 unfractured minor pyrite KR-07-10 zircon M10 colourless to yellow good short-prismatic present very few, especially black-opaque not obvious <<50 slightly to moderately fractured relatively large zircon KR-07-11 zircon M10 colourless to med- yellow very good to poor/turbid in darkest grains short-prismatic not obvious rare very small black-opaque rare turbid cores <50 moderately to highly fractured KR-07-12 zircon M10 colourless to pale yellow very good short-prismatic to elongate-prismatic not obvious minor small black- opaque not obvious <<50 unfractured KR-07-13 zircon M5 colourless OR med- yellow very good (colourless); good to slightly turbid (yellow) short-prismatic not obvious rare small black- opaque or red- orange not obvious <<50 unfractured to slightly fractured abundant short to elongate- skinny red-orange crystals (rutile?); zircon is very small and not very abundant KR-07-14 zircon M10 pale yellow (colourless to yellow) very good stubby-prismatic not obvious minor black- opaque/biotite not obvious <50 unfractured (clear) to highly fractured (yellow) relatively large zircon KR-07-15 zircon M5 medium yellow poor (fractured) but transluscent short-prismatic not obvious minor small black- opaque; orange stain/inclusions not obvious >50 highly fractured small sample but zircon is abundant KR-07-16 zircon M10 pale to med-yellow very good to poor (fractured) short-prismatic present very rare not obvious >50 slightly (clear) to highly (yellow- brown) fractured very minor sulphide (pyrite?) 562 Sample Mineral Fraction Code Colour Clarity Morphology Zoning Inclusions Cores % Broken Degree of Fracturing Comments KR-07-17 zircon M5 colourless to pale- med-yellow very good (colourless); good yellow) short-prismatic (stubby and multi-faceted) seen in clear grains very rare turbid cores in darker grains <<50 unfractured to moderately fractured (darker grains) very abundant pyrite (no cubes present; broken or pyrrhotite?) KR-07-18 zircon M10 colourless to PALE- grey/yellow very good short prismatic to short and multi-faceted not obvious very rare not obvious <<50 unfractured abundant short to elongate- skinny red-orange crystals (rutile?) KR-07-19 zircon M5 colourless to pale yellow very good (some turbid) short-prismatic (yellow); elongate-prismatic (colourless) not obvious minor black-opaque +/- biotite; apatite not obvious >50 unfractured; some fractures in darker grains pyrite present KR-07-20 zircon M10 very pale to light- medium yellow very good to slightly cloudy (fractures?) short-prismatic present rare present <50 unfractured to slightly fractured pyrite present; dark red transluscent mineral KR-07-21 zircon M5 colourless to med- yellow (orange- yellow rare) very good (colourless); good, some slightly cloudy (yellow) short-prismatic to elongate-prismatic visible in pale yellow grains minor black- opaque; rare small orange not obvious (rare slightly turbid cores) >50 unfractured clear, to slightly - moderately fractured yellow grains very minor pyrite KR-07-22 zircon M10 light to med-yellow (rare colourless) very good to very turbid short-prismatic to stubby-prismatic visible in light- coloured grains minor small black- opaque not obvious <<50 moderately to highly fractured very abundant small euhedral pyrite KR-07-23 zircon M10 colourless to med- yellow mostly very good; semi-opaque (med-yellow) short-prismatic to stubby-prismatic not obvious rare biotite (in apatite?) not obvious ~50 mostly unfractured mish-mash of material; probably not all zircon mounted (apatite +/- lime- green hexagonal crystal +/- ???) KR-07-24 zircon M10 pale to med-yellow (some colourless) very good but slightly cloudy in darker grains elongate-prismatic visible in some grains very rare visible in some grains >>50 moderately to highly fractured CL-06-33 zircon N5 colourless to light brown good stubby prismatic visible in some grains abundant apatite not visible <<50 unfractured CL-06-34 zircon NM2 or N2 or N5 nd good but slightly cloudy in some grains short to stubby prismatic visible in some grains apatite, rare opaque not visible 50? unfractured to slightly fractured CL-06-35 zircon N2 nd very good (rare brown turbid grains) short to elongate prismatic visible in some grains apatite, rare opaque rare turbid cores 50? unfractured to slightly fractured CL-06-36 zircon N2 nd good short (to stubby) prismatic visible in some grains abundant apatite, some opaque not visible <50? slightly? fractured CL-06-37 zircon N2 nd good (rare turbid brown cores) short to stubby prismatic visible in some grains apatite, opaque not visible <50? unfractured to slightly fractured CL-06-38 zircon N5 colourless to pale yellow/brown very good short to elongate prismatic visible in some some apatite, minor opaque rare dark turbid <50 unfractured to slightly fractured 563 Sample Mineral Fraction Code Colour Clarity Morphology Zoning Inclusions Cores % Broken Degree of Fracturing Comments grains CL-06-39 zircon N5 colourless to pale yellow/brown good (rare turbid brown) short to stubby prismatic not visible abundant apatite, minor opaque rare partially turbid <<50 slightly to moderately fractured CL-06-40 zircon N5 colourless (some pale yellow-brown) good short to stubby prismatic visible in some grains abundant apatite not visible <50 unfractured to slightly fractured 564 B.3 Summary of observed zircon textures from polished grain mounts for samples analyzed by LA–ICP–MS methods. 565 Sample na P a r t i a l l y R e s o r b e d P o r o u s R a d i a l F r a c t u r e s M a g m a t i c ± P a r t i a l l y R e s o r b e d R e s o r b e d I n h e r i t e d U n z o n e d ( B S E ) P o r o u s R a d i a l F r a c t u r e s X E N O C R Y S T S : INCLUSIONSb Summary of Typical Zircon Texturesc RIMS (%) CORES (%) Cassiar Terrane and western Selwyn Basin: KR-07-01 45 27 2 64 2 96 40 56 62 0 (rare ap) thin w-z fract MR with thick BSE-bright p-z overgrowths, on resorbed u-z +/- porous fract IC KR-07-02 61 56 5 51 15 74 2 64 70 2 ap, mnz; 2 mnz grains thin p-z and p-resorbed MR, and resorbed zoned MC mantling resorbed porous IC, hi-fract throughout KR-07-03 54 50 63 54 26 61 7 24 54 5 ap, mnz w-z MR and MC with p-resorbed and a porous overgrowth, and resorbed IC, hi-fract throughout KR-07-04 nd nd nd nd nd nd nd nd nd nd nd nd KR-07-05 44 75 32 39 16 73 9 70 73 1 (ap, mnz) p-z and p-resorbed MR with BSE-bright overgrowths and resorbed w-z MC or u-z/porous IC, hi-fract throughout KR-07-06 54 13 0 20 44 46 7 6 43 0 ap thick w-z MR with variably resorbed w-z to u-z fract MC KR-07-07 51 59 8 53 29 61 22 59 76 0 ap (mnz) w-z MR with resorbed w-z MC or u-z/porous IC, hi-fract throughout KR-07-08 70 16 10 10 70 16 6 6 10 6 ap (mnz) w-z MR with well to p-z MC, abundant p-resorbed throughout KR-07-09 41 41 5 27 32 51 20 27 44 4 (ap, mnz) w-z MR with partially to significantly resorbed MC or u-z/porous IC, hi-fract throughout KR-07-10 48 17 0 19 77 19 0 4 44 0 ap w-z MR with p-resorbed w-z fract MC or very small porous IC KR-07-11 40 35 20 68 25 70 18 60 58 1 (ap, mnz); 1 mnz grain thin w-z MR on u-z to p-z porous resorbed IC that are hi-fract throughout, or on w-z MC KR-07-12 58 36 3 7 28 71 5 0 26 1 ap, mnz w-z MR with partially to significantly resorbed, p-z, fract MC KR-07-13 49 14 2 20 10 37 43 55 65 10 ap very thin w-z MR with large u-z hi-fract and p-resorbed porous IC KR-07-14 61 74 33 34 33 59 0 5 77 0 ap (mnz) w-z MR with thick p-z overgrowths affected by a late resorption event (removing very thin porous rims) with partially to significantly resorbed and fract MC KR-07-15 43 51 28 47 53 40 2 16 81 3 ap mnz w-z p-resorbed (including at least 1 major resorption event) and locally porous MR with significant resorption of zoned MC, hi-fract throughout KR-07-16 45 47 2 22 24 64 7 16 69 0 ap (mnz); 1 mnz grain w-z MR and p-s resorb IC, hi-fract throughout KR-07-17 48 40 0 17 8 46 29 48 27 11 ap (mnz) w-z MR with large u-z and/or resorbed porous IC KR-07-18 55 58 0 25 27 60 11 29 18 7 ap (mnz) w-z MR with partially to significantly resorbed w-z MC to u-z/porous IC KR-07-19 57 19 0 40 16 79 7 23 42 3 ap (mnz) w-z MR with resorbed predominantly w-z MC and lesser u-z porous IC, hi-fract throughout KR-07-20 51 67 4 47 6 90 39 67 41 1 (ap) thin to thick (hi-fract) w-z MR with minor resorb, and hi-fract resorbed u-z and porous IC KR-07-21 46 46 0 50 17 70 15 11 72 1 ap (mnz) thick w-z fract MR with p-resorbed and fract MC KR-07-22 62 74 23 21 18 79 8 27 81 ap (mnz) thin w-z MR with hi-fract resorbed w-z to u-z IC, at least 2 major resorption events KR-07-23 47 13 9 11 4 53 17 23 28 15 ap (mnz); 2 mnz grains very thin w-z MR with large poorly to un-zoned and resorbed rounded IC KR-07-24 58 53 10 57 21 72 9 45 57 0 ap w-z hi-fract MR with zoned resorbed moderate-fract porous MC 566 Sample na P a r t i a l l y R e s o r b e d P o r o u s R a d i a l F r a c t u r e s M a g m a t i c ± P a r t i a l l y R e s o r b e d R e s o r b e d I n h e r i t e d U n z o n e d ( B S E ) P o r o u s R a d i a l F r a c t u r e s X E N O C R Y S T S : INCLUSIONSb Summary of Typical Zircon Texturesc 07M-150 72 76 1 15 17 72 21 6 19 3 ap large grains with w-z thick MR and w-z to u-z MC, displaying 2-3 major resorption events 07M-151 61 74 0 18 62 28 5 2 11 1 ap large grains with w-z thick MR and w-z MC, displaying 2-3 major resorption events 07M-152 66 70 0 6 8 70 35 30 8 2 ap large grains with w-z MR and w-z MC to u-z porous cores, displaying 3+ p-resorbed events (including 1 very late event affecting outermost rims only) 07M-153 58 71 0 19 12 78 28 47 41 1 ap w-z MR and w-z MC frequently mantling u-z porous IC, with 3+ resorption events (including 1 very late event affecting outermost rims only, and 2 affecting inherited cores) 07M-154 59 98 42 66 97 0 3 3 75 0 ap thin to almost no w-z MR and large p-resorbed w-z MC (1 major resorption event); hi-fract throughout; porous grain margins Eastern Selwyn Basin: KR-05-08 45 64 0 0 80 13 20 13 0 0 ap thick w-z MR, with at least 1 resorption surface in larger grains, and small p-resorbed w-z MC with lesser non-resorbed u-z ± porous IC KR-05-10 37 70 0 3 95 5 19 5 3 0 (ap), (rare mnz) thick w-z MR, with at least 1 but up to 3+ partial resorption surfaces in larger grains, and primarily p- resorbed poorly- to w-z MC (or u-z pristine MC in smaller grains) +/- u-z resorbed porous IC KR-05-22 17 71 0 6 76 29 6 0 6 0 ap thick w-z MR that may have at least 1 resorption surface or leaching, with s- to p-resorbed MC and lesser complexly zoned and s-resorbed cores of unknown origin KR-05-26 44 89 5 7 95 2 18 5 0 0 ap thick w-z MR with 1-2 resorption events (including 1 along very margins of some grains), and p-resorbed poorly- to u-z MC KR-05-32 28 89 4 11 89 14 14 4 0 0 (ap), (rare mnz) thick p-z MR with at least 1 resorption surface, and older p-resorbed poorly- to u-z MC KR-05-43 16 88 0 6 100 0 0 0 0 0 (ap) thick w-z MR that may have at least 1 resorption surface or leaching mantled by a thicker CL-bright outer zone (2nd from grain edge), with pristine to p-resorbed w-z MC KR-05-62 24 33 0 8 79 17 8 8 0 0 ap, (rare mnz) thick w-z MR, with p-resorbed w-z MC and some un- to p-z resorbed cores of unknown origin KR-05-68 49 69 0 16 53 41 37 14 2 4 (rare ap) thick w-z MR frequently demonstrating 2 major growth episodes (evident by 2 repeats of increasingly bright BSE zoning correlative with a resorption surface) with primarily u-z and/or porous+hi-fract, resorbed ± porous IC (some less resorbed and non-porous cores may be magmatic) or older p-resorbed to pristine MC KR-05-76 19 53 5 5 32 68 58 42 21 0 (ap) w-z MR with at least 1 significant resorption surface, and mainly large resorbed porous IC with some older small p-resorbed w-z MC KR-05-97b 58 81 2 5 86 14 5 5 7 0 ap thick w-z MR with at least 1 resorption event, and older p-resorbed poorly/thickly zoned MC KR-05-110 41 83 12 56 76 10 7 2 29 0 (ap), 3 mnz grains thick w-z MR with 1 major resorption surface or 2-3 major growth episodes in larger grains (evident by 2 repeats of increasingly bright BSE zoning and the termination of radiating fractures), and p-resorbed w-z MC or fewer resorbed cores; very hi-fract rims ± MC 567 Sample na P a r t i a l l y R e s o r b e d P o r o u s R a d i a l F r a c t u r e s M a g m a t i c ± P a r t i a l l y R e s o r b e d R e s o r b e d I n h e r i t e d U n z o n e d ( B S E ) P o r o u s R a d i a l F r a c t u r e s X E N O C R Y S T S : INCLUSIONSb Summary of Typical Zircon Texturesc KR-05-113 29 90 3 45 76 28 14 21 38 0 (ap) thin w-z MR with 1-2 significant resorption surfaces, and poorly- to u-z (CL-dark) magmatic? cores; hi-fract rims ± MC KR-05-126 28 89 0 11 100 11 0 21 0 0 (rare ap) small grains with pristine MC-MR and minor resorption surfaces are the most abundant, but thick w-z MR with 1 major resorption surface and 2-3 major growth episodes in larger grains (evident by 2 repeats of increasingly bright BSE zoning and the termination of radiating fractures), with older pristine to p-resorbed w-z MC and fewer resorbed porous IC are also notably present KR-05-130 17 35 0 29 35 53 12 59 12 2 (ap) thick w-z MR with minor resorption near margins, and predominantly resorbed porous IC with older but lesser pristine to p-resorbed w-z MC KR-05-136 24 75 4 29 67 38 17 17 13 0 ap thick w-z MR with at least 2 major growth episodes (evident by 2 repeats of increasingly bright BSE zoning and the termination of radiating fractures, often correlative with a major resorption surface), with older w-z p-resorbed MC; hi-fract rims KR-05-143 52 96 2 31 81 23 40 13 8 0 ap (rare mnz) thick w-z MR with 2-3 major growth episodes in larger grains (evident by 2 repeats of increasingly bright BSE zoning and the termination of radiating fractures, correlative with a major resorption surface), with older un- to w-z p-resorbed MC (CL-bright) and u-z ± porous resorbed IC; hi-fract rims KR-05-148 25 68 0 16 80 12 12 0 0 2 (ap) thick w-z MR with 1 significant resorption event mantled by a relatively thick BSE-dark zone, with pristine to p-resorbed MC KR-05-164 24 75 8 88 17 63 13 ap (rare mnz) w-z bright (BSE) MR with 1-2 major resorption surfaces and large, partially to moderately resorbed, poorly to u-z and BSE-dark MC (apatite inclusions primarily in cores) KR-05-175 29 90 7 45 83 17 17 7 38 1 (ap), (rare mnz) thick w-z MR with at least 1 significant resorption surface and 3+ (4) major growth episodes in larger grains (evident by multiple repeats of increasingly bright BSE zoning and the termination of radiating fractures, with w-z p-resorbed MC and un- to p-z resorbed IC (CL-bright); hi-fract rims + cores KR-05-191 32 44 0 22 97 9 3 3 9 0 (rare ap) thick w-z MR, some with 1 significant resorption surface, and older p-resorbed, poorly to u-z BSE-dark MC; fract rims KR-05-194 54 89 0 17 98 7 13 2 13 0 ap thick poorly- to w-z MR with at least 1 resorption surface, and older p-resorbed well- to p-z MC KR-05-196 52 77 0 92 8 13 13 0 0 ap, (zir) thick p-z non-resorbed MR with pristine to p-resorbed poorly- to u-z MC KR-05-198 48 81 0 2 96 4 52 0 0 0 ap, (zir), (rare mnz) thick w-z MR, typically with 1 major resorption surface, and older large p-resorbed poorly- to u-z (CL-bright) MC KR-05-208 21 100 0 43 86 19 5 14 10 0 (ap) thick w-z MR with at least 1 major resorption surface, and older p-resorbed p-z MC and/or resorbed porous IC; hi-fract rims KR-05-210 28 93 11 54 75 29 7 7 11 0 (ap), (rare zir) thick w-z MR with at least 1 major resorption surface, and older partially to s-resorbed p-z MC; hi-fract rims 568 Sample na P a r t i a l l y R e s o r b e d P o r o u s R a d i a l F r a c t u r e s M a g m a t i c ± P a r t i a l l y R e s o r b e d R e s o r b e d I n h e r i t e d U n z o n e d ( B S E ) P o r o u s R a d i a l F r a c t u r e s X E N O C R Y S T S : INCLUSIONSb Summary of Typical Zircon Texturesc KR-05-215 58 90 19 57 78 26 16 29 29 0 (ap, mnz) w-z MR with at least 1 major resorption surface, and p-resorbed w-z ± porous magmatic to resorbed u-z ± porous (CL-bright) IC; hi-fract rims + cores CL-06-33 32 63 0 9 88 6 0 3 6 0 ap thick w-z MR with older p-resorbed w-z MC CL-06-34 25 72 0 8 100 4 0 4 12 0 ap very small grains with w-z MR and p-resorbed w-z MC and rare but significant resorbed IC; at least 1 significant resorption surface in rim CL-06-35 34 68 3 3 88 9 9 9 3 0 (ap) thick w-z MR with p-resorbed MC; multiple small resorption surfaces and at least 1 significant resorption surface in rim CL-06-36 41 59 2 2 95 2 5 5 5 0 ap thick w-z MR and p-resorbed w-z MC; at least 1 significant resorption surface in rim CL-06-37 21 71 0 0 86 10 0 5 5 0 ap thick w-z MR and p-resorbed w-z MC; at least 1 significant resorption surface in rim CL-06-38 52 44 17 10 35 67 25 25 25 1 ap w-z MR with older p-resorbed MC; abundant porous/u-z and intensely resorbed IC CL-06-39 44 70 2 5 80 18 14 0 18 1 ap thick w-z MR with older p-resorbed MC; at least 1 significant resorption surface in rim CL-06-40 53 53 0 6 87 15 15 0 15 0 ap thick w-z MR with p-resorbed w-z MC; at least 1 significant resorption surface in rim Notes: CL-06 data was previously reported in Rasmussen et al. (2011b). a number of polished zircon grains classified; b ap = apatite; mnz = monazite (or other high Th mineral); brackets indicate a relatively low abundance, and bolding indicates a relatively high abundance; c MR = magmatic rims; MC = magmatic cores (autocrystic and antecrystic not distinguised); IC = inherited cores; w-z = well-zoned; u-z = unzoned; p-z = poorly-zoned; w = well; hi-fract = highly fractured; p-resorbed = partially resorbed. 569 B.4 Representative zircons images. Back-scattered electron (BSE) and cathodoluminescent (CL) images of a representative polished zircon for each sample dated by LA–ICP–MS methods. Samples are labeled by sample ID, name of the magmatic body (italicized if assigned by this study), and—in brackets—the plutonic suite. Plutonic suite abbreviations are: C = Cassiar, A = Anvil, TR = Tay River, Tu = Tungsten, M = Mayo, To = Tombstone, R = Rancheria, BR = Black River, HM = Hess–MacMillan sub-suite, CR = Coal River sub-suite, and ‘e’ = ‘evolved’. Scale bars are in microns, and “N/A” = not available. 570 571 572 573 574 575 576 577 578 579 580 581 582 583 584 585 586 B.5 LA–CP–MS age determinations. Weighted mean plots of 206Pb/238U zircon ages and 207Pb/235U versus 206Pb/238U concordia plots displaying 2-sigma error bars or ellipses (95% confidence interval) for magmatic rim (i.e., crystallization), antecrystic core, and inherited core analyses. Calculated crystallization ages from weighted mean plots are weighted by errors associated with individual analyses (dark grey); rejected analyses not used in the age calculation are pale grey. New names assigned to unnamed plutons and batholiths by this study are italicized. 587 KR-07-01 – Dycer Creek plu W 111 112 113 114 115 116 117 118 20 6 P b/ 23 8 U A ge (M a) 115 116 117 118 119 120 20 6 P b/ 23 8 U A ge (M a) ton: Weighted eighted Mea Mean = 1 Wtd by MSWD Mean Wtd by MSW Mean Cryst n Magmatic 14.8±1.2 [1. data-pt errs o = 6.3, proba = 118.14±0.4 data-pt errs D = 0.31, pro allization Ag Inheritance 0%] 95% co nly, 0 of 7 re bility = 0.000 box hei 2 [0.36%] 2 only, 0 of 5 r bability = 0.8 box hei e Age nf. j. ghts are 2σ σ ej. 7 ghts are 2σ 588 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0172 0.0176 0.0180 0.0184 0.0188 0.112 0.116 0.120 0.124 0.128 0.132 20 6 P b/ 23 8 U 207Pb/235U 110 112 114 116 118 120 data-point error ellipses are 2σ 589 KR-07-02 – Quiet Lake bath W 108 110 112 114 116 118 120 20 6 P b/ 23 8 U A ge (M a) 117 119 121 123 125 20 6 P b/ 23 8 U A ge (M a) olith: Weighted eighted Mea Mean = Wtd b MSW Mean Wtd by MSW Mean Cryst n Magmatic 114.7±1.1 [0 y data-pt errs D = 1.4, pro = 122±15 [1 data-pt errs D = 1.9, pro allization Ag Inheritance .97%] 95% only, 0 of 9 r bability = 0.1 box hei 2%] 95% co only, 0 of 2 r bability = 0.17 box hei e Age conf. ej. 9 ghts are 2σ nf. ej. ghts are 2σ 590 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0165 0.0175 0.0185 0.0195 0.0205 0.10 0.11 0.12 0.13 0.14 0.15 20 6 P b/ 23 8 U 207Pb/235U 108 112 116 120 124 128 data-point error ellipses are 2σ 591 KR-07-03 – Quie 20 6 P b/ 23 8 U A ge (M a) t Lake bath W 92 94 96 98 100 102 20 6 P b/ 23 8 U A ge (M a) 99.0 99.4 99.8 100.2 100.6 101.0 101.4 101.8 102.2 102.6 20 6 P b/ 23 8 U A ge (M a) olith: Weighted eighted Mea Mean = 9 Wtd by MSW Mean = 1 Wtd b MSW Mean Cryst n Magmatic 7.65±0.92 [0 data-pt errs D = 13, prob 00.85±0.57 y data-pt errs D = 1.16, pro allization Ag Inheritance .94%] 95% only, 0 of 12 ability = 0.00 box hei [0.56%] 95% only, 0 of 3 bability = 0.3 box heig e Age conf. rej. 0 ghts are 2σ conf. rej. 1 hts are 2σ 592 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 0.092 0.096 0.100 0.104 0.108 0.112 20 6 P b/ 23 8 U 207Pb/235U 92 94 96 98 100 102 104 data-point error ellipses are 2σ 593 KR-07-05 – Nisutlin batholi 104 106 108 110 112 114 116 118 120 122 20 6 P b/ 23 8 U A ge (M a) 0.016 0.017 0.018 0.019 0.1 20 6 P b/ 23 8 U th: Weighted Concordia Mean = Wtd by MSW 02 0.106 0.1 106 Mean Cryst Plot of Crys 113.2±1.1 [0 data-pt errs o D = 1.9, prob 10 0.114 0.11 207Pb/ 110 114 data-point allization Ag tallization A .97%] 95% c nly, 1 of 13 ability = 0.03 box heig 8 0.122 0.126 235U 118 122 error ellipse e ge onf. rej. 8 hts are 2σ 0.130 0.134 s are 2σ 594 Concordia Plot of Inherited Age 0.0 0.1 0.2 0.3 0.4 0.5 0 2 4 6 8 20 6 P b/ 23 8 U 207Pb/235U 200 600 1000 1400 1800 2200 data-point error ellipses are 2σ 595 KR-0 7-06 – Nisu 20 6 P b/ 23 8 U A (M ) tlin batholi 99 101 103 105 107 109 111 20 6 P b/ 23 8 U A ge (M a) 0.0152 0.0156 0.0160 0.0164 0.0168 0.0172 0.0176 0 20 6 P b/ 23 8 U th: Weighted Concordia Mean = 1 Wtd by MSW .085 0.09 Mean Cryst Plot of Crys 05.13±0.71 data-pt errs D = 1.3, pro 5 0.105 207Pb/ 98 100 102 104 106 10 1 data-poin allization Ag tallization A [0.68%] 95% only, 0 of 14 bability = 0.23 box hei 0.115 0.1 235U 8 10 112 t error ellipse e ge conf. rej. ghts are 2σ 25 0.135 s are 2σ 596 KR-0 7-07 – Young pluton: 99 101 103 105 107 109 111 20 6 P b/ 23 8 U A ge (M a) 0.0158 0.0162 0.0166 0.0170 0.0174 0 20 6 P b/ 23 8 U Weighted Concordia Mean = 1 Wtd by MSW .094 0.098 0. Mean Cryst Plot of Crys 07.20±0.50 data-pt errs D = 1.7, prob 102 0.106 0.1 207Pb/ 102 104 106 1 data-poi allization Ag tallization A [0.47%] 95% only, 0 of 14 ability = 0.04 box hei 10 0.114 0.11 235U 08 110 112 nt error ellips e ge conf. rej. 8 ghts are 2σ 8 0.122 0.126 es are 2σ 597 KR-0 7-08 – St C 20 6 P b/ 23 8 U A (M ) yr pluton: 99 101 103 105 107 109 111 113 115 20 6 P b/ 23 8 U A ge (M a) 0.0160 0.0164 0.0168 0.0172 0.0176 0.0180 0 20 6 P b/ 23 8 U Weighted Concordia Mean = 1 Wtd by MSW .095 0. 1 Mean Cryst Plot of Crys 07.70±0.55 data-pt errs D = 1.2, pro 105 0.1 207Pb/ 04 106 108 110 112 data-poi allization Ag tallization A [0.51%] 95% only, 1 of 16 bability = 0.24 box heig 15 0.125 235U 114 116 nt error ellips e ge conf. rej. hts are 2σ 0.135 es are 2σ 598 KR-07-09 – Fox 20 6 P b/ 23 8 U A (M ) Mountain p 99 101 103 105 107 109 111 113 115 20 6 P b/ 23 8 U A ge (M a) 0.0154 0.0158 0.0162 0.0166 0.0170 0.0174 0.0178 0.0182 0.0186 0 20 6 P b/ 23 8 U luton: Weighted Concordia Mean = Wtd by MSW .08 0.09 Mean Cryst Plot of Crys 107.8±1.1 [ data-pt errs D = 2.3, prob 0.10 0.1 207Pb/ 102 106 1 data-po allization Ag tallization A 1.1%] 95% c only, 0 of 13 ability = 0.00 box hei 1 0.12 235U 10 114 118 int error ellips e ge onf. rej. 6 ghts are 2σ 0.13 0.14 es are 2σ 599 KR-0 7-10 – Big S 20 6 P b/ 23 8 U A (M ) almon bat 97 99 101 103 105 107 109 111 113 115 20 6 P b/ 23 8 U A ge (M a) 0.015 0.016 0.017 0.018 0.0 20 6 P b/ 23 8 U holith: Weighted Concordia Mean = Wtd by MSW 85 0.095 98 Mean Cryst Plot of Crys 106.8±1.3 [ data-pt errs D = 2.9, prob 0.105 0. 207Pb/ 102 106 110 data-poi allization Ag tallization A 1.2%] 95% c only, 0 of 15 ability = 0.00 box hei 115 0.125 235U 114 118 nt error ellips e ge onf. rej. 0 ghts are 2σ 0.135 es are 2σ 600 KR-07-11 – Glen 20 6 P b/ 23 8 U A (M ) 20 6 P b/ 23 8 U A ge (M a) lyon batho W 99 101 103 105 107 109 111 113 20 6 P b/ 23 8 U A ge (M a) 110 112 114 116 118 120 20 6 P b/ 23 8 U A ge (M a) lith: Weighted eighted Mea Mean Wtd by MSW Mean = Wtd by MSW Mean Cryst n Magmatic = 107.72±0.5 data-pt errs D = 1.6, pro 113.3±3.4 [3 data-pt errs D = 2.4, prob allization Ag Inheritance 2 [0.48%] 2 only, 0 of 11 bability = 0.10 box hei .0%] 95% c only, 0 of 3 r ability = 0.08 box hei e Age σ rej. ghts are 2σ onf. ej. 9 ghts are 2σ 601 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0155 0.0165 0.0175 0.0185 0.09 0.10 0.11 0.12 0.13 20 6 P b/ 23 8 U 207Pb/235U 102 106 110 114 118 data-point error ellipses are 2σ 602 KR-0 7-12 – Tinti 20 6 P b/ 23 8 U A (M ) na pluton: 45 46 47 48 49 50 51 52 20 6 P b/ 23 8 U A ge (M a) 0.0069 0.0071 0.0073 0.0075 0.0077 0.0079 0.0081 0 20 6 P b/ 23 8 U Weighted Concordia Mean = 4 Wtd by MSW .036 0.040 Mean Cryst Plot of Crys 9.06±0.41 [0 data-pt errs o D = 3.3, prob 0.044 0.048 207Pb/ 46 48 5 data-poi allization Ag tallization A .83%] 95% nly, 1 of 15 ability = 0.00 box hei 0.052 0.05 235U 0 52 nt error ellips e ge conf. rej. 0 ghts are 2σ 6 0.060 es are 2σ 603 KR-0 7-13 – Squa 20 6 P b/ 23 8 U A (M ) re Lake pl 106 108 110 112 114 116 20 6 P b/ 23 8 U A ge (M a) 0.0166 0.0170 0.0174 0.0178 0.0182 0 20 6 P b/ 23 8 U uton: Weighted Concordia Mean = Wtd by MSW .09 0.10 Mean Cryst Plot of Crys 111.7±1.1 [0 data-pt errs D = 2.6, prob 0.11 0.1 207Pb/ 108 110 112 114 116 data-po allization Ag tallization A .98%] 95% c only, 0 of 7 r ability = 0.01 box heig 2 0.13 235U int error ellips e ge onf. ej. 6 hts are 2σ 0.14 0.15 es are 2σ 604 KR-0 7-14 – Blac 20 6 P b/ 23 8 U A (M ) k River plu 43 44 45 46 47 48 49 50 20 6 P b/ 23 8 U A ge (M a) 0.0067 0.0069 0.0071 0.0073 0.0075 0.0077 0.0079 0 20 6 P b/ 23 8 U ton: Weighted Concordia Mean = 4 Wtd by MSW .038 0.04 Mean Cryst Plot of Crys 7.09±0.34 [0 data-pt errs o D = 1.4, prob 2 0.046 207Pb/ 44 45 46 47 data-poi allization Ag tallization A .72%] 95% nly, 1 of 16 ability = 0.15 box heig 0.050 235U 48 49 50 nt error ellips e ge conf. rej. hts are 2σ 0.054 es are 2σ 605 KR-0 7-15 – Meis 20 6 P b/ 23 8 U A (M ) ter Lake pl 77 78 79 80 81 20 6 P b/ 23 8 U A ge (M a) 0.0119 0.0121 0.0123 0.0125 0.0127 0 20 6 P b/ 23 8 U uton: Weighted Concordia Mean = 7 Wtd by MSW .077 0.079 77 7 Mean Cryst Plot of Crys 9.18±0.43 [0 data-pt errs o D = 6.3, prob 0.081 207Pb/ 8 79 80 data-poi allization Ag tallization A .55%] 95% nly, 0 of 13 ability = 0.00 box hei 0.083 0. 235U 81 nt error ellips e ge conf. rej. 0 ghts are 2σ 085 0.087 es are 2σ 606 KR-0 7-16 – Cass 20 6 P b/ 23 8 U A (M ) iar batholi 106 108 110 112 114 116 20 6 P b/ 23 8 U A ge (M a) 0.0164 0.0168 0.0172 0.0176 0.0180 0 20 6 P b/ 23 8 U th: Weighted Concordia Mean = 1 Wtd by MSW .106 0.110 106 108 Mean Cryst Plot of Crys 11.49±0.48 data-pt errs D = 1.9, prob 0.114 0 207Pb/ 110 112 1 data-poi allization Ag tallization A [0.43%] 95% only, 1 of 13 ability = 0.03 box hei .118 0.122 235U 14 116 nt error ellips e ge conf. rej. 8 ghts are 2σ 0.126 es are 2σ 607 KR-0 7-17 – Cass 20 6 P b/ 23 8 U A (M ) iar batholi 105 107 109 111 113 115 20 6 P b/ 23 8 U A ge (M a) 0.0162 0.0166 0.0170 0.0174 0.0178 0.0182 0 20 6 P b/ 23 8 U th: Weighted Concordia Mean = 1 Wtd by MSW .104 0.108 104 106 Mean Cryst Plot of Crys 11.09±0.57 data-pt errs D = 2.7, prob 0.112 0.1 207Pb/ 108 110 112 data-poi allization Ag tallization A [0.51%] 95% only, 1 of 13 ability = 0.00 box hei 16 0.120 0 235U 114 116 nt error ellips e ge conf. rej. 2 ghts are 2σ .124 0.128 es are 2σ 608 KR-0 7-18 – Cass 20 6 P b/ 23 8 U A (M ) iar batholi 107 109 111 113 115 117 119 121 20 6 P b/ 23 8 U A ge (M a) 0.0164 0.0168 0.0172 0.0176 0.0180 0.0184 0.0188 0.0192 0 20 6 P b/ 23 8 U th: Weighted Concordia Mean = 1 Wtd by MSW .095 0. Mean Cryst Plot of Crys 14.30±0.89 data-pt errs D = 1.5, pro 105 0.1 207Pb/ 108 112 data-poi allization Ag tallization A [0.77%] 95% only, 0 of 14 bability = 0.12 box hei 15 0.12 235U 116 120 nt error ellips e ge conf. rej. ghts are 2σ 5 0.135 es are 2σ 609 KR-0 7-19 – Mark 20 6 P b/ 23 8 U A (M ) er Lake ba 99 101 103 105 107 109 111 113 115 117 20 6 P b/ 23 8 U A ge (M a) 0.0158 0.0162 0.0166 0.0170 0.0174 0.0178 0.0182 0.0186 0 20 6 P b/ 23 8 U tholith: Weighted Concordia Mean = 1 Wtd by MSW .085 0.095 Mean Cryst Plot of Crys 08.41±0.47 data-pt errs D = 0.73, pro 0.105 0.1 207Pb/ 104 108 112 data-poi allization Ag tallization A [0.44%] 95% only, 1 of 16 bability = 0.7 box hei 15 0.125 235U 116 nt error ellips e ge conf. rej. 5 ghts are 2σ 0.135 0.145 es are 2σ 610 KR-0 7- 20 – Cab 20 6 P b/ 23 8 U A (M ) in Creek pl 106 107 108 109 110 111 112 113 20 6 P b/ 23 8 U A ge (M a) 0.0165 0.0167 0.0169 0.0171 0.0173 0.0175 0.0177 0 20 6 P b/ 23 8 U uton: Weighted Concordia Mean = 1 Wtd by MSW .104 0.108 0 106 107 10 Mean Cryst Plot of Crys 09.75±0.50 data-pt errs D = 2.4, prob .112 0.116 207Pb/ 8 109 110 111 112 113 data-poin allization Ag tallization A [0.45%] 95% only, 0 of 15 ability = 0.00 box hei 0.120 0.124 235U t error ellipse e ge conf. rej. 3 ghts are 2σ 0.128 0.132 s are 2σ 611 KR-0 7- 21 – Alle 20 6 P b/ 23 8 U A (M ) n Stock plu 75 76 77 78 79 20 6 P b/ 23 8 U A ge (M a) 0.0116 0.0118 0.0120 0.0122 0.0124 0 20 6 P b/ 23 8 U ton: Weighted Concordia Mean = 7 Wtd by MSW .074 0.076 75 Mean Cryst Plot of Crys 7.22±0.39 [0 data-pt errs o D = 4.8, prob 0.078 207Pb/ 76 77 data-poi allization Ag tallization A .51%] 95% nly, 0 of 14 ability = 0.00 box heig 0.080 0. 235U 78 79 8 nt error ellips e ge conf. rej. 0 hts are 2σ 082 0.084 0 es are 2σ 612 KR-0 7- 22 – Cas 20 6 P b/ 23 8 U A (M ) siar bathol 99 101 103 105 107 109 111 113 115 117 119 20 6 P b/ 23 8 U A ge (M a) 0.0155 0.0165 0.0175 0.0185 0 20 6 P b/ 23 8 U ith: Weighted Concordia Mean = 1 Wtd by MSW .10 102 10 Mean Cryst Plot of Crys 13.80±0.79 data-pt errs D = 2.5, prob 0.11 207Pb/ 6 110 114 data-po allization Ag tallization A [0.70%] 95% only, 1 of 11 ability = 0.00 box hei 0.12 235U 118 int error ellips e ge conf. rej. 8 ghts are 2σ 0.13 es are 2σ 613 KR-0 7- 23 – Gra 20 6 P b/ 23 8 U A (M ) vel Creek p 77 79 81 83 85 87 20 6 P b/ 23 8 U A ge (M a) 0.0120 0.0124 0.0128 0.0132 0.0136 0 20 6 P b/ 23 8 U luton: Weighted Concordia Mean = 8 Wtd by MSW .072 0.076 Mean Cryst Plot of Crys 1.92±0.84 [ data-pt errs o D = 5.4, prob 0.080 0.08 207Pb/ 78 80 82 data-poi allization Ag tallization A 1.0%] 95% c nly, 0 of 10 r ability = 0.00 box heig 4 0.088 0 235U 84 86 88 nt error ellips e ge onf. ej. 0 hts are 2σ .092 0.096 es are 2σ 614 KR-0 7- 24 – Cas 20 6 P b/ 23 8 U A (M ) siar bathol 110 111 112 113 114 115 116 117 20 6 P b/ 23 8 U A ge (M a) 0.0171 0.0173 0.0175 0.0177 0.0179 0.0181 0.0183 0.0185 0 20 6 P b/ 23 8 U ith: Weighted Concordia Mean = 1 Wtd by MSW .110 0.11 110 Mean Cryst Plot of Crys 13.74±0.55 [ data-pt errs D = 3.9, prob 4 0.118 207Pb/ 112 114 data-poi allization Ag tallization A 0.49%] 95% only, 0 of 16 ability = 0.00 box hei 0.122 235U 116 118 nt error ellips e ge conf. rej. 0 ghts are 2σ 0.126 es are 2σ 615 07M-150 – Carol 20 6 P b/ 23 8 U A (M ) yn batholit W 88 90 92 94 96 98 100 102 20 6 P b/ 23 8 U A ge (M a) 97 99 101 103 105 107 20 6 P b/ 23 8 U A ge (M a) h: Weighted eighted Mea Mean = Wtd by Mean = Wtd by MSW Mean Cryst n Magmatic 95.7±1.4 [1 data-pt errs 102.0±2.2 [2 data-pt errs D = 0.16, pro allization Ag Inheritance .4%] 95% co only, 0 of 11 box heig .2%] 95% c only, 0 of 2 r bability = 0.6 box heig e Age nf. rej. hts are 2σ onf. ej. 9 hts are 2σ 616 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.013 0.014 0.015 0.016 0.017 0.018 0.07 0.09 0.11 0.13 0.15 20 6 P b/ 23 8 U 207Pb/235U 86 90 94 98 102 106 110 114 118 data-point error ellipses are 2σ 617 07M-151 – South 20 6 P b/ 23 8 U A ge (M a) Fork (volc W 92 94 96 98 100 102 20 6 P b/ 23 8 U A ge (M a) 95 97 99 101 103 105 107 20 6 P b/ 23 8 U A ge (M a) anic): Weighted eighted Mea Mean = 9 Wtd by MSW Mean = Wtd by MSW Mean Cryst n Magmatic 7.22±0.81 [0 data-pt errs D = 0.53, pro 102.5±1.2 [ data-pt errs D = 0.57, pro allization Ag Inheritance .83%] 95% only, 0 of 10 bability = 0.8 box heig 1.1%] 95% c only, 0 of 6 r bability = 0.7 box heig e Age conf. rej. 5 hts are 2σ onf. ej. 3 hts are 2σ 618 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 0.0166 0.0170 0.08 0.09 0.10 0.11 0.12 20 6 P b/ 23 8 U 207Pb/235U 94 98 102 106 data-point error ellipses are 2σ 619 07M- 152 – Buttle Creek plu 99 101 103 105 107 109 20 6 P b/ 23 8 U A ge (M a) 0.0152 0.0156 0.0160 0.0164 0.0168 0.0172 0 20 6 P b/ 23 8 U ton: Weighted Concordia Mean = 1 Wtd by MSW .094 0.098 0 98 Mean Cryst Plot of Crys 03.99±0.53 data-pt errs D = 0.76, pro .102 0.106 207Pb/ 100 102 104 1 data-poi allization Ag tallization A [0.51%] 95% only, 0 of 16 bability = 0.7 box hei 0.110 0.114 235U 06 108 nt error ellips e ge conf. rej. 3 ghts are 2σ 0.118 0.122 es are 2σ 620 07M- 153 – Glenlyon bathol 99 101 103 105 107 109 111 113 115 20 6 P b/ 23 8 U A ge (M a) 0.0158 0.0162 0.0166 0.0170 0.0174 0.0178 0.0182 0 20 6 P b/ 23 8 U ith: Weighted Concordia Mean = Wtd by MSW .098 0.102 0 102 1 Mean Cryst Plot of Crys 108.5±1.1 [ data-pt errs D = 2.8, prob .106 0.110 0. 207Pb/ 04 106 108 11 data-poi allization Ag tallization A 1.0%] 95% c only, 0 of 14 ability = 0.00 box heig 114 0.118 0. 235U 0 112 114 116 nt error ellips e ge onf. rej. 1 hts are 2σ 122 0.126 es are 2σ 621 07M-154 – Little Salmon pluton: Weighted Mean Crystallization Age Concordia Plot of Crystallization Age 88 90 92 94 96 98 100 102 104 box heights are 2σ 20 6 P b/ 23 8 U A ge (M a) 90 92 94 96 98 100 102 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 0.075 0.085 0.095 0.105 0.115 207Pb/235U data-point error ellipses are 2σ 20 6 P b/ 23 8 U 622 KR-05-08 – Jorgensen plut 91 93 95 97 99 101 20 6 P b/ 23 8 U A ge (M a) 0.0142 0.0146 0.0150 0.0154 0.0158 0 20 6 P b/ 23 8 U on: Weighted Concordia Mean = Wtd b MSW .085 0. 92 9 Mean Cryst Plot of Crys 95.90±0.77 y data-pt errs D = 2.1, pro 095 0.1 207Pb/ 4 96 98 100 102 data-poi allization Ag tallization A [0.80%] 95% only, 0 of 12 bability = 0.0 box hei 05 0.11 235U nt error ellips e ge conf. rej. 17 ghts are 2σ 5 0.125 es are 2σ 623 Concordia Plot of Inherited Ages 0.00 0.02 0.04 0.06 0.08 0.0 0.2 0.4 0.6 0.8 20 6 P b/ 23 8 U 207Pb/235U 50 150 250 350 450 550 data-point error ellipses are 2σ 624 KR-05-10 – Powers pluton: 94 96 98 100 102 104 106 108 20 6 P b/ 23 8 U A ge (M a) 0.0144 0.0148 0.0152 0.0156 0.0160 0.0164 0.0168 20 6 P b/ 23 8 U Weighted Concordia Mean = 9 Wtd by MSW 0.08 0.09 Mean Cryst Plot of Crys 8.89±0.68 [0 data-pt errs D = 1.3, pro 0.10 207Pb/ 96 100 10 data-po allization Ag tallization A .69%] 95% only, 1 of 15 bability = 0.18 box heig 0.11 0 235U 4 108 int error ellips e ge conf. rej. hts are 2σ .12 0.13 es are 2σ 625 480 490 500 510 520 530 20 6 P b/ 23 8 U A ge (M a) 0.00 0.02 0.04 0.06 0.08 0.10 0. 20 6 P b/ 23 8 U 5 Weighte Concord Mea Wtd by MSW 0 0.2 0 150 d Mean Inh ia Plot of Inh n = 496.7±7 data-pt errs D = 1.9, pro 0.4 207Pb/2 250 350 data-po erited Age erited Ages .7 [1.6%] 2σ only, 0 of 2 r bability = 0.17 box heig 0.6 35U 450 int error ellip ej. hts are 2σ 0.8 550 ses are 2σ 626 KR-0 5-22 – Roy pluton: 88 92 96 100 104 108 20 6 P b/ 23 8 U A ge (M a) 0.0135 0.0145 0.0155 0.0165 20 6 P b/ 23 8 U Weighted Concordia Mean = Wtd b MSW 0.05 0.07 Mean Cryst Plot of Crys 96.06±0.93 y data-pt errs D = 0.73, pr 0.09 207P 90 94 98 10 data-po allization Ag tallization A [0.97%] 95% only, 0 of 16 obability = 0. box hei 0.11 0 b/235U 2 106 int error ellips e ge conf. rej. 75 ghts are 2σ .13 0.15 es are 2σ 627 KR-0 5-26 – Fish pluton: 90 92 94 96 98 100 20 6 P b/ 23 8 U A ge (M a) 0.0140 0.0144 0.0148 0.0152 0.0156 0.0160 0 20 6 P b/ 23 8 U Weighted Concordia Mean = Wtd b MSW .085 0.09 90 92 94 96 Mean Cryst Plot of Crys 96.11±0.60 y data-pt errs D = 0.85, pr 5 0.105 207Pb 98 100 102 data-poi allization Ag tallization A [0.63%] 95% only, 0 of 10 obability = 0. box hei 0.115 0. /235U nt error ellips e ge conf. rej. 57 ghts are 2σ 125 0.135 es are 2σ 628 KR-05-32 – Ivo pluton: 90 92 94 96 98 100 102 20 6 P b/ 23 8 U A ge (M a) 97 99 101 103 105 107 20 6 P b/ 23 8 U A ge (M a) Weighted Weighte Mean = 9 Wtd by MSW Mean = 10 Wtd by MSW Mean Cryst d Mean Inh 6.04±0.95 [0 data-pt errs D = 1.5, pro 3.19±0.86 [ data-pt errs D = 1.03, pro allization Ag erited Ages .99%] 95% only, 0 of 12 bability = 0.13 box hei 0.84%] 95% only, 0 of 7 re bability = 0.4 box heig e conf. rej. ghts are 2σ conf. j. 0 hts are 2σ 629 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0135 0.0145 0.0155 0.0165 0.0175 0.07 0.09 0.11 0.13 0.15 20 6 P b/ 23 8 U 207Pb/235U 90 94 98 102 106 110 data-point error ellipses are 2σ 630 KR-0 5-43 – Park pluton: 88 90 92 94 96 98 100 102 104 106 20 6 P b/ 23 8 U A ge (M a) 0.0135 0.0145 0.0155 0.0165 20 6 P b/ 23 8 U Weighted Concordia Mean = Wtd by MSW 0.07 0.08 Mean Cryst Plot of Crys 96.1±1.2 [1 data-pt errs D = 2.0, prob 0.09 0.1 207Pb/ 90 94 98 data-po allization Ag tallization A .3%] 95% co only, 0 of 16 ability = 0.01 box hei 0 0.11 235U 102 106 int error ellips e ge nf. rej. 4 ghts are 2σ 0.12 0.13 es are 2σ 631 KR-0 5-62 – Mount Christie 84 88 92 96 100 104 20 6 P b/ 23 8 U A ge (M a) 0.013 0.014 0.015 0.016 0 20 6 P b/ 23 8 U pluton: Weighted Concordia Mean = Wtd by MSW .06 0.08 8 Mean Cryst Plot of Crys 95.1±1.2 [1 data-pt errs D = 0.85, pro 0.10 207Pb/ 6 90 94 98 102 data-po allization Ag tallization A .3%] 95% co only, 0 of 12 bability = 0.5 box heig 0.12 235U int error ellips e ge nf. rej. 9 hts are 2σ 0.14 es are 2σ 632 KR-05-68 – Mount Wilson p W 92 94 96 98 100 102 104 20 6 P b/ 23 8 U A ge (M a) 99 101 103 105 107 20 6 P b/ 23 8 U A ge (M a) luton: Weighted eighted Mea Mean = 9 Wtd by MSW Mea Wtd by MSW Mean Cryst n Magmatic 7.74±0.75 [0 data-pt errs D = 0.61, pro n = 103.1±1. data-pt errs D = 0.47, pro allization Ag Inheritance .77%] 95% only, 0 of 12 bability = 0.8 box hei 8 [1.7%] 2σ only, 0 of 4 r bability = 0.7 box hei e Age conf. rej. 2 ghts are 2σ ej. 0 ghts are 2σ 633 Concordia Plot of Crystallization and Magmatic Inheritance Ages Concordia Plot of Inheritance Ages 0.014 0.015 0.016 0.017 0.07 0.08 0.09 0.10 0.11 0.12 0.13 20 6 P b/ 23 8 U 207Pb/235U 90 94 98 102 106 110 data-point error ellipses are 2σ 0.0 0.1 0.2 0.3 0.4 0 1 2 3 4 5 6 20 6 P b/ 23 8 U 207Pb/235U 200 600 1000 1400 1800 data-point error ellipses are 2σ 634 KR-05-76 – Pelly River plut W 89 91 93 95 97 99 101 103 105 107 20 6 P b/ 23 8 U A ge (M a) 96 98 100 102 104 106 108 110 20 6 P b/ 23 8 U A ge (M a) on: Weighted eighted Mea Mean Wtd by MSW Mean = Wtd b MSW Mean Cryst n Magmatic = 98.0±1.1 [ data-pt errs D = 1.3, pro 104.2±1.4 y data-pt errs D = 0.117, p allization Ag Inheritance 1.2%] 95% c only, 0 of 15 bability = 0.2 box hei [1.3%] 95% c only, 0 of 6 robability = 0 box hei e Age onf. rej. 1 ghts are 2σ onf. rej. .99 ghts are 2σ 635 Concordia Pl W 0.0135 0.0145 0.0155 0.0165 0.0175 20 6 P b/ 23 8 U 1750 1770 1790 1810 1830 20 7 P b/ 23 5 U A ge (M a) ot of Crysta eighted Mea 0.05 0.07 Wt MS llization and n Magmatic 0.09 207Pb/ 90 94 98 102 10 data-po Mean = 1792 d by data-pt e WD = 0.0066 Magmatic I Inheritance 0.11 0.13 235U 6 110 int error ellips ±39 [2.2%] rrs only, 0 of , probability box hei nheritance A Age 0.15 es are 2σ 2σ 2 rej. = 0.94 ghts are 1σ ges 636 Concordia Plot of Inheritance Ages 0.0 0.1 0.2 0.3 0.4 0 2 4 6 20 6 P b/ 23 8 U 207Pb/235U 200 600 1000 1400 1800 data-point error ellipses are 2σ 637 KR-05-97b – O'Grady batho W 87 89 91 93 95 97 20 6 P b/ 23 8 U A ge (M a) 93 94 95 96 97 98 99 100 20 6 P b/ 23 8 U A ge (M a) lith: Weighted eighted Mea Mean = Wtd by MSW Mean Wtd by MSW Mean Cryst n Magmatic 92.84±0.63 [ data-pt errs D = 1.7, prob = 97±15 [15 data-pt errs D = 4.3, prob allization Ag Inheritance 0.68%] 95% only, 1 of 13 ability = 0.07 box hei %] 95% con only, 0 of 2 r ability = 0.03 box hei e Age conf. rej. 2 ghts are 2σ f. ej. 7 ghts are 2σ 638 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0134 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0.085 0.095 0.105 0.115 20 6 P b/ 23 8 U 207Pb/235U 88 92 96 100 data-point error ellipses are 2σ 639 KR-0 5-110 – Cac pluton: 91 93 95 97 99 101 20 6 P b/ 23 8 U A ge (M a) 0.0140 0.0144 0.0148 0.0152 0.0156 0.0160 0 20 6 P b/ 23 8 U Weighted Concordia Mean = 9 Wtd by MSW .055 0.0 Mean Cryst Plot of Crys 7.37±0.96 [0 data-pt errs D = 1.7, prob 65 0.075 207Pb/ data-poi allization Ag tallization A .99%] 95% only, 0 of 12 ability = 0.06 box heig 0.085 235U nt error ellips e ge conf. rej. 5 hts are 2σ 0.095 90 91 92 93 94 95 96 es are 2σ 640 KR-05-126 – (plug SW of O W 86 90 94 98 102 106 20 6 P b/ 23 8 U A ge (M a) 95 97 99 101 103 105 107 109 20 6 P b/ 23 8 U A ge (M a) 'Grady): Weighted eighted Mea Mean Wtd b MSW Mean = Wtd b MSW Mean Cryst n Magmatic = 95.9±1.4 y data-pt errs D = 0.49, p 101.7±2.5 y data-pt errs D = 0.31, pr allization Ag Inheritance [1.5%] 95% c only, 0 of 12 robability = 0 box hei [2.4%] 95% c only, 0 of 3 obability = 0. box hei e Age onf. rej. .91 ghts are 2σ onf. rej. 74 ghts are 2σ 641 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.013 0.014 0.015 0.016 0.017 0.05 0.07 0.09 0.11 0.13 0.15 20 6 P b/ 23 8 U 207Pb/235U 86 90 94 98 102 106 110 data-point error ellipses are 2σ 642 KR-05-130 – Central Nahan W 88 90 92 94 96 98 100 102 104 20 6 P b/ 23 8 U A ge (M a) 97 99 101 103 105 20 6 P b/ 23 8 U A ge (M a) ni pluton: Weighted eighted Mea Mean = 9 Wtd by MSW Mean = Wtd by MSW Mean Cryst n Magmatic 7.12±0.76 [0 data-pt errs D = 0.90, pro 101.3±1.5 [ data-pt errs D = 0.14, pro allization Ag Inheritance .79%] 95% only, 0 of 15 bability = 0.5 box hei 1.5%] 95% c only, 0 of 4 r bability = 0.9 box hei e Age conf. rej. 6 ghts are 2σ onf. ej. 4 ghts are 2σ 643 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.013 0.014 0.015 0.016 0.017 0.018 0.07 0.09 0.11 0.13 0.15 20 6 P b/ 23 8 U 207Pb/235U 86 90 94 98 102 106 110 114 data-point error ellipses are 2σ 644 KR-05-136 – North Nahann W 90 92 94 96 98 100 102 104 20 6 P b/ 23 8 U A ge (M a) 96 98 100 102 104 106 20 6 P b/ 23 8 U A ge (M a) i pluton: Weighted eighted Mea Mean = 9 Wtd by MSW Mean = Wtd by MSW Mean Cryst n Magmatic 6.87±0.73 [0 data-pt errs D = 1.6, prob 101.8±1.3 [ data-pt errs D = 0.67, pro allization Ag Inheritance .75%] 95% only, 0 of 18 ability = 0.04 box heig 1.3%] 95% c only, 0 of 4 r bability = 0.5 box heig e Age conf. rej. 9 hts are 2σ onf. ej. 7 hts are 2σ 645 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0135 0.0145 0.0155 0.0165 0.0175 0.075 0.085 0.095 0.105 0.115 0.125 0.135 20 6 P b/ 23 8 U 207Pb/235U 90 94 98 102 106 110 114 data-point error ellipses are 2σ 646 KR-05-143 – Circular Stock W 84 88 92 96 100 104 108 20 6 P b/ 23 8 U A ge (M a) 92 96 100 104 108 112 20 6 P b/ 23 8 U A ge (M a) pluton: Weighted eighted Mea Mean = Wtd b MSW Mean = Wtd b MSW Mean Cryst n Magmatic 96.86±0.61 y data-pt errs D = 0.74, pr 103.0±1.6 y data-pt errs D = 1.9, pro allization Ag Inheritance [0.63%] 95% only, 1 of 18 obability = 0. box heig [1.5%] 95% c only, 0 of 5 bability = 0.1 box heig e Age conf. rej. 75 hts are 2σ onf. rej. 2 hts are 2σ 647 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0125 0.0135 0.0145 0.0155 0.0165 0.0175 0.04 0.06 0.08 0.10 0.12 0.14 0.16 20 6 P b/ 23 8 U 207Pb/235U 84 88 92 96 100 104 108 112 data-point error ellipses are 2σ 648 KR-05-148 – Rifle Range pl 87 89 91 93 95 97 99 101 103 20 6 P b/ 23 8 U A ge (M a) 0.0134 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 0.0166 0 20 6 P b/ 23 8 U uton: Weighted Concordia Mean = Wtd by MSW .08 0.09 88 9 Mean Cryst Plot of Crys 96.3±1.0 [1 data-pt errs D = 1.8, prob 0.10 207Pb/ 2 96 100 10 data-po allization Ag tallization A .1%] 95% co only, 0 of 16 ability = 0.02 box hei 0.11 0 235U 4 int error ellips e ge nf. rej. 8 ghts are 2σ .12 0.13 es are 2σ 649 KR-05-164 – Hole-in-the-Wa 87 89 91 93 95 97 99 101 20 6 P b/ 23 8 U A ge (M a) 0.0134 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 20 6 P b/ 23 8 U ll batholith Weighted Concordia Mean = Wtd by MSW 0.08 0 88 : Mean Cryst Plot of Crys 95.3±1.1 [1 data-pt errs D = 4.2, prob .09 0.1 207Pb/ 92 96 data-po allization Ag tallization A .1%] 95% co only, 0 of 15 ability = 0.00 box hei 0 0.1 235U 100 int error ellips e ge nf. rej. 0 ghts are 2σ 1 0.12 es are 2σ 650 KR-05-175 – Lened pluton: 94 95 96 97 98 99 100 101 20 6 P b/ 23 8 U A ge (M a) 0.0145 0.0147 0.0149 0.0151 0.0153 0.0155 0.0157 0.0159 0 20 6 P b/ 23 8 U Weighted Concordia Mean = 9 Wtd by MSW .088 0.092 Mean Cryst Plot of Crys 7.57±0.78 [0 data-pt errs D = 3.0, prob 0.096 0.1 207Pb/ 94 96 9 data-poi allization Ag tallization A .79%] 95% only, 0 of 11 ability = 0.00 box hei 00 0.104 0 235U 8 100 102 nt error ellips e ge conf. rej. 1 ghts are 2σ .108 0.112 es are 2σ 651 Concordia Plot of Inherited Age 0.0 0.1 0.2 0.3 0.4 0 2 4 6 20 6 P b/ 23 8 U 207Pb/235U 200 600 1000 1400 1800 data-point error ellipses are 2σ 652 KR-05-191 – Marion pluton 88 92 96 100 104 108 20 6 P b/ 23 8 U A ge (M a) 0.013 0.014 0.015 0.016 0.017 0.018 0 20 6 P b/ 23 8 U : Weighted Concordia Mean Wtd by MSW .05 0.07 Mean Cryst Plot of Crys = 96.7±1.4 [ data-pt errs D = 0.93, pr 0.09 0 207Pb/ 86 90 94 98 102 10 1 data-po allization Ag tallization A 1.4%] 95% c only, 0 of 10 obability = 0.5 box heig .11 0.13 235U 6 10 114 int error ellips e ge onf. rej. 0 hts are 2σ 0.15 es are 2σ 653 KR-05-194 – Coal River bat W 88 90 92 94 96 98 20 6 P b/ 23 8 U A ge (M a) 92 93 94 95 96 97 98 99 100 20 6 P b/ 23 8 U A ge (M a) holith: Weighted eighted Mea Mean = 9 Wtd by MSWD Mean = 9 Wtd by MSW Mean Cryst n Magmatic 2.76±0.93 [ data-pt errs o = 2.4, prob 5.79±0.76 [0 data-pt errs D = 0.90, pro allization Ag Inheritance 1.0%] 95% c nly, 0 of 10 r ability = 0.01 box hei .79%] 95% only, 0 of 5 r bability = 0.4 box hei e Age onf. ej. 1 ghts are 2σ conf. ej. 6 ghts are 2σ 654 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0136 0.0140 0.0144 0.0148 0.0152 0.0156 0.0160 0.082 0.086 0.090 0.094 0.098 0.102 0.106 0.110 20 6 P b/ 23 8 U 207Pb/235U 88 90 92 94 96 98 100 102 data-point error ellipses are 2σ 655 KR-0 5-196 – Coal River bat 92 94 96 98 100 102 104 106 20 6 P b/ 23 8 U A ge (M a) 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 0.0166 0 20 6 P b/ 23 8 U holith: Weighted Concordia Mea Wtd MS .085 0.095 94 9 Mean Cryst Plot of Crys n = 98.4±1.0 by data-pt er WD = 1.9, p 0.105 0 207Pb/ 8 102 106 data-poi allization Ag tallization A [1.0%] 95% rs only, 0 of robability = 0 box hei .115 0.125 235U nt error ellips e ge conf. 14 rej. .025 ghts are 2σ 0.135 es are 2σ 656 KR-05-198 – Hole-in-the-Wa W 84 86 88 90 92 94 96 98 100 20 6 P b/ 23 8 U A ge (M a) 87 89 91 93 95 97 99 101 103 105 20 6 P b/ 23 8 U A ge (M a) ll batholith Weighted eighted Mea Mean = Wtd by MSW Mean = Wtd by MSW : Mean Cryst n Magmatic 91.3±2.1 [2 data-pt errs D = 2.0, prob 96.73±0.97 data-pt errs D = 0.62, pro allization Ag Inheritance .3%] 95% co only, 0 of 5 r ability = 0.09 box hei [1.0%] 95% c only, 0 of 10 bability = 0.7 box heig e Age nf. ej. 4 ghts are 2σ onf. rej. 8 hts are 2σ 657 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0125 0.0135 0.0145 0.0155 0.0165 0.07 0.09 0.11 0.13 0.15 0.17 20 6 P b/ 23 8 U 207Pb/235U 84 88 92 96 100 104 108 data-point error ellipses are 2σ 658 KR-05-208 – Little Hyland p 90 92 94 96 98 100 102 20 6 P b/ 23 8 U A ge (M a) 0.0140 0.0144 0.0148 0.0152 0.0156 0.0160 0.0164 20 6 P b/ 23 8 U luton: Weighted Concordia Mean = 9 Wtd by MSW 0.08 0.0 90 Mean Cryst Plot of Crys 6.80±0.60 [0 data-pt errs D = 1.2, pro 9 0.10 207Pb/ 92 94 96 98 100 1 data-po allization Ag tallization A .62%] 95% only, 0 of 17 bability = 0.25 box hei 0.11 235U 02 104 int error ellips e ge conf. rej. ghts are 2σ 0.12 es are 2σ 659 KR-05-210 – Nahanni Rang 93 95 97 99 101 20 6 P b/ 23 8 U A ge (M a) 0.0142 0.0146 0.0150 0.0154 0.0158 0.0162 20 6 P b/ 23 8 U e Road plu Weighted Concordia Mean = 9 Wtd by MSW 0.08 0 ton: Mean Cryst Plot of Crys 7.19±0.49 [0 data-pt errs D = 0.75, pro .09 0.1 207Pb/ 92 94 96 98 1 data-po allization Ag tallization A .50%] 95% only, 0 of 15 bability = 0.7 box hei 0 0.1 235U 00 102 104 int error ellips e ge conf. rej. 3 ghts are 2σ 1 0.12 es are 2σ 660 CL-06-33 – Dechen’La pluto W 90 91 92 93 94 95 96 20 6 P b/ 23 8 U A ge (M a) 92 94 96 98 100 20 6 P b/ 23 8 U A ge (M a) n: Weighted eighted Mea Mean = 9 Wtd by MSWD Mean = 9 Wtd by MSW Mean Cryst n Magmatic 3.68±0.52 [0 data-pt errs = 1.9, prob 6.39±0.72 [0 data-pt errs D = 5.8, prob allization Ag Inheritance .56%] 95% c only, 1 of 9 re ability = 0.070 box hei .74%] 95% only, 0 of 15 ability = 0.00 box hei e Age onf. j. ghts are 2σ conf. rej. 0 ghts are 2σ 661 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0.088 0.092 0.096 0.100 0.104 0.108 20 6 P b/ 23 8 U 207Pb/235U 90 92 94 96 98 100 data-point error ellipses are 2σ 662 CL-06-34 – Natla pluton: 89 90 91 92 93 94 95 20 6 P b/ 23 8 U A ge (M a) 0.0139 0.0141 0.0143 0.0145 0.0147 0.0149 0 20 6 P b/ 23 8 U Weighted Concordia Mean = 9 Wtd by MSWD .090 0.09 90 91 92 9 Mean Cryst Plot of Crys 2.29±0.58 [0 data-pt errs o = 5.4, prob 4 0.098 207Pb/ 3 94 95 data-poi allization Ag tallization A .63%] 95% c nly, 0 of 12 r ability = 0.000 box hei 0.102 0. 235U nt error ellips e ge onf. ej. ghts are 2σ 106 0.110 es are 2σ 663 Concordia Plot of Inheritance Ages 0.0 0.1 0.2 0.3 0.4 0 2 4 6 20 6 P b/ 23 8 U 207Pb/235U 200 600 1000 1400 1800 data-point error ellipses are 2σ 664 CL-06-35 – Logan pluton: 90 92 94 96 98 100 20 6 P b/ 23 8 U A ge (M a) 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0 20 6 P b/ 23 8 U Weighted Concordia Mean = Wtd by MSW .086 0.090 0 90 9 Mean Cryst Plot of Crys 95.6±1.2 [1 data-pt errs D = 7.9, prob .094 0.098 0. 207Pb/ 2 94 96 98 1 data-poi allization Ag tallization A .3%] 95% co only, 0 of 11 ability = 0.00 box hei 102 0.106 0. 235U 00 102 nt error ellips e ge nf. rej. 0 ghts are 2σ 110 0.114 es are 2σ 665 CL-06-36 – Mount Christie 93 94 95 96 97 98 99 20 6 P b/ 23 8 U A ge (M a) 0.0145 0.0147 0.0149 0.0151 0.0153 0.0155 0 20 6 P b/ 23 8 U pluton: Weighted Concordia Mean = 9 Wtd by MSW .093 0.095 0 93 94 Mean Cryst Plot of Crys 6.08±0.51 [0 data-pt errs o D = 2.9, prob .097 0.099 0. 207Pb/ 95 96 97 9 data-poi allization Ag tallization A .53%] 95% c nly, 0 of 12 r ability = 0.00 box heig 101 0.103 0. 235U 8 99 nt error ellips e ge onf. ej. 1 hts are 2σ 105 0.107 es are 2σ 666 CL-06-37 – Mount Christie 93 94 95 96 97 98 99 100 20 6 P b/ 23 8 U A ge (M a) 0.0144 0.0146 0.0148 0.0150 0.0152 0.0154 0.0156 0.0158 0 20 6 P b/ 23 8 U pluton: Weighted Concordia Mean = 9 Wtd by MSW .088 0.092 0 Mean Cryst Plot of Crys 6.72±0.63 [0 data-pt errs D = 2.7, prob .096 0.100 207Pb/ 94 96 98 10 data-poi allization Ag tallization A .65%] 95% only, 0 of 13 ability = 0.00 box hei 0.104 0.108 235U 0 nt error ellips e ge conf. rej. 1 ghts are 2σ 0.112 0.116 es are 2σ 667 CL-06-38 – Ross River plut W 89 90 91 92 93 94 95 20 6 P b/ 23 8 U A ge (M a) 92 93 94 95 96 97 98 99 20 6 P b/ 23 8 U A ge (M a) on: Weighted eighted Mea Mean = 9 Wtd by MSWD Mean = 9 Wtd by MSW Mean Cryst n Magmatic 2.99±0.92 [0 data-pt errs = 4.3, prob 5.25±0.49 [0 data-pt errs D = 1.8, prob allization Ag Inheritance .99%] 95% c only, 0 of 7 re ability = 0.000 box heig .52%] 95% only, 1 of 13 ability = 0.05 box heig e Age onf. j. hts are 2σ conf. rej. 2 hts are 2σ 668 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0138 0.0142 0.0146 0.0150 0.0154 0.0158 0.084 0.088 0.092 0.096 0.100 0.104 0.108 0.112 20 6 P b/ 23 8 U 207Pb/235U 90 92 94 96 98 100 data-point error ellipses are 2σ 669 CL-06-39 – Keele River plut W 89 90 91 92 93 94 95 96 97 20 6 P b/ 23 8 U A ge (M a) 92 94 96 98 100 102 20 6 P b/ 23 8 U A ge (M a) on: Weighted eighted Mea Mean = 9 Wtd by MSW Mean = 9 Wtd by MSW Mean Cryst n Magmatic 3.22±0.39 [0 data-pt errs o D = 1.4, prob 6.29±0.87 [0 data-pt errs D = 0.47, pro allization Ag Inheritance .42%] 95% c nly, 0 of 16 r ability = 0.12 box hei .90%] 95% only, 0 of 5 r bability = 0.7 box hei e Age onf. ej. ghts are 2σ conf. ej. 5 ghts are 2σ 670 Concordia Plot of Crystallization and Magmatic Inheritance Ages 0.0139 0.0141 0.0143 0.0145 0.0147 0.0149 0.0151 0.0153 0.086 0.090 0.094 0.098 0.102 0.106 0.110 20 6 P b/ 23 8 U 207Pb/235U 90 92 94 96 98 data-point error ellipses are 2σ 671 CL-06-40 – Mile 222 pluton 92 93 94 95 96 97 98 99 20 6 P b/ 23 8 U A ge (M a) 0.0144 0.0146 0.0148 0.0150 0.0152 0.0154 0.0156 0 20 6 P b/ 23 8 U : Weighted Concordia Mean = 9 Wtd by MSW .090 0.094 Mean Cryst Plot of Crys 6.58±0.56 [0 data-pt errs D = 1.7, prob 0.098 0 207Pb/ 94 96 98 data-poi allization Ag tallization A .58%] 95% only, 0 of 11 ability = 0.08 box hei .102 0.106 235U 100 nt error ellips e ge conf. rej. 2 ghts are 2σ 0.110 es are 2σ 672 B.6 LA–CP–MS geochronological data. Sample conditions and results (age determinations, isotopic ratios, and isotope counts/second) for individual analyses used to calculate final weighted mean ages. (available for download at: http://hdl.handle.net/2429/44385) 673 B.7 ID-TIMS age determinations. Pb207/U235 versus Pb206/U238 concordia plots displaying 2-sigma error ellipses (95% confidence interval) for samples analyzed by ID–TIMS methods. Black error ellipses represent zircon data and grey error ellipses represent monazite data (unless noted otherwise). 674 ARG-33 – Tsa Da Glisza pluton: Concordia Plot of Crystallization Age FL-46 – Square Lake pluton: Concordia Plot of Crystallization Age 108 110 112 114 116 0.0168 0.0170 0.0172 0.0174 0.0176 0.0178 0.0180 0.0182 0.112 0.114 0.116 0.118 0.120 0.122 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 110 111 112 113 114 0.0172 0.0174 0.0176 0.0178 0.0180 0.114 0.115 0.116 0.117 0.118 0.119 0.120 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 675 KG-31 – Tsa Da Glisza pluton: Concordia Plot of Crystallization Age KM-07 – Mink pluton: Concordia Plot of Crystallization Age 108 110 112 114 116 0.0168 0.0170 0.0172 0.0174 0.0176 0.0178 0.0180 0.0182 0.112 0.114 0.116 0.118 0.120 0.122 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 108 110 112 114 116 0.0168 0.0170 0.0172 0.0174 0.0176 0.0178 0.0180 0.0182 0.112 0.114 0.116 0.118 0.120 0.122 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 676 MC-CL-01 – dyke (Clear Lake): Concordia Plot of Crystallization Age MP-18 – Money pluton: Concordia Plot of Crystallization Age 105 106 107 108 109 110 0.0164 0.0166 0.0168 0.0170 0.0172 0.1085 0.1095 0.1105 0.1115 0.1125 0.1135 0.1145 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 108 110 112 114 116 118 0.0168 0.0172 0.0176 0.0180 0.0184 0.112 0.114 0.116 0.118 0.120 0.122 0.124 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 677 03M-203 – felsic dyke, Cantung: Concordia Plot of Crystallization Age 03M-204 – aplite, Cantung: Concordia Plot of Crystallization Age 97 97.4 97.8 98.2 98.6 99 0.0151 0.0152 0.0153 0.0154 0.0155 0.098 0.100 0.102 0.104 0.106 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 90 110 130 150 170 190 0.012 0.016 0.020 0.024 0.028 0.04 0.08 0.12 0.16 0.20 0.24 0.28 0.32 0.36 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 678 03M-207 – Mine Stock-apical, Cantung: Concordia Plot of Crystallization Age 03M-214 – Shelf Lake pluton: Concordia Plot of Crystallization Age 96.4 96.8 97.2 97.6 98 98.4 98.8 0.01505 0.01515 0.01525 0.01535 0.01545 0.0985 0.0995 0.1005 0.1015 0.1025 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 93 94 95 96 97 98 99 100 0.0145 0.0147 0.0149 0.0151 0.0153 0.0155 0.0157 0.095 0.097 0.099 0.101 0.103 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 679 03M-215 - Mount Appler pluton: Concordia Plot of Crystallization Age 03M-216 - Faille pluton: Concordia Plot of Crystallization Age 93 94 95 96 97 98 99 100 0.0145 0.0147 0.0149 0.0151 0.0153 0.0155 0.0157 0.095 0.097 0.099 0.101 0.103 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 95 96 97 98 99 0.0148 0.0150 0.0152 0.0154 0.098 0.099 0.100 0.101 0.102 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 680 03M-217 - Mount Sir James McBrien pluton: Concordia Plot of Crystallization Age 03M-218 - Hole-in-the-Wall batholith (north): Concordia Plot of Crystallization Age 95 96 97 98 99 0.0148 0.0150 0.0152 0.0154 0.098 0.099 0.100 0.101 0.102 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 96.8 97.2 97.6 98 98.4 0.01505 0.01515 0.01525 0.01535 0.0996 0.1000 0.1004 0.1008 0.1012 0.1016 0.1020 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 681 03M-219 - Pyramid Mountain pluton: Concordia Plot of Crystallization Age 03M-220 - Ragged pluton: Concordia Plot of Crystallization Age 92 93 94 95 96 97 98 0.0143 0.0145 0.0147 0.0149 0.0151 0.0153 0.0155 0.094 0.096 0.098 0.100 0.102 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 92 93 94 95 96 97 98 0.0143 0.0145 0.0147 0.0149 0.0151 0.0153 0.0155 0.094 0.096 0.098 0.100 0.102 2 0 6 P b /2 3 8 U 207Pb/235U data-point error ellipses are 2s 682 B.8 ID–TIMS geochronological data. Sample conditions and results (isotopic ratios and age determinations) for individual analyses used to calculate final concordia ages in Appendix B.7. (available for download at: http://hdl.handle.net/2429/44385) 683 B.9 Summary of Zircon Inheritance. 684 Summary of Zircon Inheritance Overview of Zircon Inheritance Textures: New data on zircon inheritance detected in the course of this study (as well as inferred zircon inheritance ages from Heffernan, 2004, Murphy et al., 2001b, and Gordey, in press) provide preliminary information on the age(s) of potential melt sources in the crust for the various plutonic suites. Back-scattered electron (BSE) and cathodoluminescent (CL) images using the scanning electron microscope (SEM) have been utilized to describe zircon textures from all of the samples dated using LA–ICP–MS methods (Table 2.3; Appendix B.3). Polished cross-sections through zircons typically display well-zoned (or oscillatory-zoned) magmatic rims of variable thicknesses mantling some combination of unzoned to poorly (coarse and diffuse) or well (thin and sharp) zoned oscillatory magmatic cores. Zircons with magmatic cores of the same age as that of that magmatic rims are referred to as “autocrystic” and are interpreted to have grown entirely within the specific pulse of magma that it is found (Miller et al., 2007). In some samples, however, a significant population of magmatically zoned cores were found to be several m.y. older than the rims (Table 2.2, Chapter 2; Appendices B.5 and B.6). These older magmatic cores are similar to the description of “antecrystic” cores described by Miller et al. (2007) for zircons initiated grown in an earlier magma pulse but were scavenged by a later magma pulse. This can be envisioned to occur when magma mixing occurs, or by the assimilation of slightly older magmatic rocks. It is significant, however, that these antecrystic zircons are overgrown by what may be relatively thick (i.e., several 10’s μm) magmatic rims precipitated from the later host magma. This implies that zircon antecrysts may have experienced a relatively lengthy period of interaction with the younger magma, allowing thick magmatic rims to precipitate. Antecrystic cores were commonly, but not exclusively or ubiquitously, associated with textural evidence for a significant resorption event, resulting in the truncation of older oscillatory-zoned cores against the resorption surface (e.g., Appendices B.3 and B.4). Thus, in some cases the antecrystic zircons were in disequilibrium with the younger magma prior to the precipitation of the younger magmatic rims. The presence of antecrystic cores, therefore, may have implications regarding the type of magmatic processes experienced by a particular intrusion. 685 Old inherited cores are also present in zircons from nearly all of the samples examined by BSE and CL methods on the SEM (Appendix B.3). These are typically rounded (resorbed), and less well zoned (unzoned to coarsely zoned). The age and proportion of old inherited cores are thought to provide information on the melt source(s) to the magmas (e.g., Miller et al., 2007); however, these cores were not analyzed systematically in this study. Therefore, concordant ages of inheritance reported in Table 2.2 (Chapter 2; Appendices B.5 and B.6) cannot be considered representative of the magma protolith(s). Several of the LA–ICP–MS inherited ages reported in Table 2.2 (Chapter 2; Appendices B.5 and B.6) also only comprise one concordant zircon analysis, which limits their statistical significance. Additionally, upper intercept ages calculated on the basis of discordant ID–TIMS analyses in zircon or monazite (even if pinned by a concordant analysis given the age of crystallization) must still be regarded with caution. These upper intercept ages only provide an average age of the inherited component for the fraction of zircon grains that was analyzed. If the zircon fraction contains grains with widely variable inheritance, upper intercept ages will be a statistically useless average age of all ages of the inheritance. Xenocrystic grains, which lack magmatic rims, are incorporated into a magma from wall rocks during its ascent and emplacement (e.g., Miller et al., 2007). As such, these grains provide no real information on the age or composition of the melt source(s). Ages from xenocrystic zircons that were inadvertently dated, however, are provided (Table 2.2, Chapter 2; Appendices B.5 and B.6). Zircon Textures and Inheritance in the Plutonic Suites: A summary of basic zircon textures, based on details in Appendix B.3, as well the ages of zircon inheritance detected in magmatic and inherited cores (including xenocrysts) for each plutonic suite follows. Representative zircon images for each sample dated by LA–ICP–MS methods are provided in Appendix B.4. Well-developed oscillatory zoning and inclusions of apatite and monazite are typical of magmatic zircon of the Cassiar plutonic suite, and resorbed poorly to unzoned inherited cores or unzoned xenocrysts lacking magmatic rims are nearly ubiquitous (Appendix B.3). Old inherited zircon was observed in nearly all of the Cassiar suite samples, but only dated in the Nisutlin 686 batholith (Fig. 2.3a; Table 2.3). Four analyses of three xenocrysts in the Quiet Lake batholith (KR-07-02) produced ages of ~360 Ma, whereas an older inherited zircon core of 2.0 Ga was detected in the Nisutlin batholith (KR-07-05). Also notable is that the Quiet Lake and Glenlyon batholiths, the Fox Mountain pluton, and the Dycer Creek pluton (immediately north of the Quiet Lake batholith) also contain poorly to well-zoned, partially resorbed (BSE and CL images suggest 1+ major resorption events) cores that are interpreted to be magmatic, but that are 3–5 m.y. older than the interpreted crystallization ages for the same intrusions as determined from magmatically zoned zircon rims. Only one Anvil suite sample was dated by LA–ICP–MS methods, but it did not produce any inherited ages (KR-07-13). The zircons in this sample of the Square Lake pluton were large, with resorbed and porous, highly fractured inherited cores (of unknown age), mantled by thin, oscillatory-zoned magmatic rims (Appendix B.3). Three upper intercept ages for the Money pluton, determined from discordant ID–TIMS data, are 760 Ma, 1.66 Ga, and 2.36 Ga (Table 2.2, Chapter 2; Appendices B.7 and B.8). None of the Hyland plutonic suite samples were dated by LA–ICP–MS by this study and there is no information on zircon textures on an ages of inherited zircon cores. Two upper intercept ages of 1.87 Ma were obtained from a sample of the Mount Billings batholith (sample 99SH-022; Heffernan, 2004), based on discordant ID–TIMS analyses pinned by concordant fractions. Zircons recovered from two samples of the Seagull plutonic suite and dated by LA–ICP–MS methods are typically highly fractured and have resorbed and well-developed oscillatory zoning in magmatic rims with porous margins or overgrowths. Slightly older magmatic or antecrystic cores have similar textures as the magmatic rims. Resorbed inherited cores are present in some of the plutons, but were not dated. The Little Salmon plug (sample 07M-154) does not have any apparent inheritance, but the sample collected from the Quiet Lake batholith (sample KR-07-03) contains Middle Triassic (234 Ma) xenocrysts, as well as antecrystic cores dated at ~101 Ma (Table 2.2, Chapter 2; Appendices B.5 and B.6). Zircons recovered from the two Tay River and South Fork samples in the western Selwyn Basin are characterized by a relatively large size, thick and oscillatory-zoned magmatic rims displaying evidence for 2–3 significant resorption events around oscillatory-zoned antecrystic cores 687 (Appendix B.3). Apatite inclusions are particularly abundant in these suites. Resorbed inherited cores and xenocrysts are also abundant in the Tay River sample. Only magmatic inherited cores were analysed, however, and produced a pronounced at of ~102 Ma in both samples (although the plutonic sample also has zircon cores as old as 107 Ma; Table 2.2, Chapter 2; Appendices B.5 and B.6). One Tay River suite sample analyzed by ID–TIMS methods (sample A-098; Pigage 2004) produced upper intercept ages for three discordant fractions of: 1.57, 1.61, and 1.86 Ga. Several upper intercept ages reported by Gordey (in press) include: 1.18, 2.13, 2.24, and 2.47 Ga for the Tay River suite; and, 1.18, 2.13, 2.47, 2.60, and 2.70 for the South Fork calderas. In the eastern Selwyn Basin, Tay River suite plutons also contain relatively large zircons with thick oscillatory-zoned magmatic rims around poorly to unzoned resorbed magmatic cores with no detectable difference in ages from cores to rims. Zircons from at least two samples (samples KR-05-32 and KR-05-191), however, have antecrystic cores of ~103–104 Ma (~7 m.y. older than the age of the rims; Table 2.2, Chapter 2; Appendices B.5 and B.6). Magmatic portions of individual grains also commonly display evidence for 1–2 resorption events. Zircons from one Tay River suite pluton analyzed by Heffernan (2004; sample 98-HAS-06) yielded an upper intercept age of 1.83 Ga. Heffernan (2004) also reported upper intercept ages of ~1.8–1.9 Ga (sample 98-HAS-02) and a poorly constrained 2.1 Ga (sample 99-SH-001) for ‘evolved’ phases. As the only xenocrystic inheritance age estimates for the Tay River suite and its ‘evolved’ examples are from upper intercepts for discordant analyses, interpretation of these must be tentative. Zircons from at least two small quartz monzodiorite porphyry bodies of the Coal River sub-suite in the southeast Selwyn Basin are texturally similar to those described for the Tay River suite plutons (although antecrystic cores are typically well-zoned), but also have magmatically zoned and internally resorbed inherited cores and (or) xenocysts with widely variable concordant analyses dated at circa 340 and 470 Ma (from one grain; sample KR-05-08), and at circa 450, 494, and 508 Ma (from one grain; sample KR-05-10; Table 2.2, Chapter 2; Appendices B.5 and B.6). The Powers plug, a nearby felsic plug of the Coal River sub-suite sampled by Heffernan (2004; sample 98-HAS-12) produced two discordant zircon analyses that project to upper intercept ages of ~1.8–1.9 Ga and ~2.25 Ga. Tungsten suite zircons typically have thick, well-zoned magmatic rims overgrowing partially to strongly resorbed, and poorly to well-zoned magmatic cores. Resorbed ± porous inherited zircon cores are also typical, especially in the samples from the southeast Selwyn Basin (i.e., >19% of 688 zircons examined for the southeast vs. 10–17% of zircons examined for the northeast; Appendix B.3). Magmatic rims and cores also typically display 1–3 (or more) recognizable resorption– growth events, and zircon grains are generally highly fractured. Apatite inclusions are not as abundant within zircon as in other plutonic suites, but monazite inclusions are typically present. Magmatic inheritance in several Tungsten suite plutons ranges from 103 to 100 Ma, and inherited cores returned a LA–ICP–MS age of 1.79 ± 14 (northeast Selwyn Basin; Table 2.2, Chapter 2; Appendices B.5 and B.6). Calculated upper intercept ages from ID–TIMS data for discordant fractions from an aplitic dyke at the Cantung mine in the southeast Selwyn Basin produced ages of: 1.82, 2.18, and 2.44 Ga (Table 2.2, Chapter 2; Appendices B.7 and B.8). Zircons from the Mayo plutonic suite samples typically have thick, oscillatory-zoned magmatic rims with well-zoned antecrystic cores ranging in age from 101–102 Ma, 103–104 Ma (in Mayo- age examples of the Hess–MacMillan sub-suite) and 107 Ma (Table 2.2, Chapter 2; Appendices B.5 and B.6). Up to two major resorption–growth episodes are also commonly observed in the Mayo suite zircons (Appendix B.3). Resorbed and porous inherited cores are present in some samples (the Mount Wilson and Pelly River plutons of the Hess–MacMillan sub-suite) where inherited cores returned LA–ICP–MS ages of 1.77 and 1.79 Ga (Table 2.2, Chapter 2; Appendices B.5 and B.6). A magmatically zoned xenocryst also returned an age of 1.59 Ga (Table 2.2, Chapter 2; Appendices B.5 and B.6) Zircons from Tombstone suite intrusions in the northwest Selwyn Basin are typically large, anhedral, and have thick, poorly zoned magmatic rims with rare inherited cores (Hart et al., 2005). Conversely, in the northeast to southeast Selwyn Basin Tombstone suite zircons are also large, but have mainly euhedral grain boundaries and thick well-zoned magmatic rims (the thin oscillatory zoning in these rims was poorly delineated BSE images, but is evident in CL images). The magmatic rims also display evidence for at least one major resorption event, and older partially resorbed well-zoned antecrystic cores are apparent in many of the zircons (Appendix B.3). Strongly resorbed anhedral antecrystic cores from many of the Tombstone-age plutons in the northeast and southeast Selwyn Basin are 95–96 Ma. Old inherited cores range from abundant to absent between samples of different intrusive phases. Ages of inherited cores obtained by LA–ICP–MS methods include 1.8 Ga, and 1.9 Ga (Table 2.2, Chapter 2; Appendices B.5 and B.6). A single, but concordant 125 Ma analysis from an apparently magmatic core of a 689 Tombstone-age sample from the Hess–MacMillan sub-suite was also obtained (Table 2.2, Chapter 2; Appendices B.5 and B.6). Zircons from the Rancheria samples examined here have variably resorbed, oscillatory-zoned magmatic rims developed on rounded resorbed and poorly zoned inherited cores (Gravel Creek Stock), or oscillatory-zoned and partially to significantly resorbed antecrystic cores lacking evidence for older inheritance (Meister Lake and Allen Stock plutons; Appendix B.3). None of the inherited cores in the Gravel Creek stock were dated. Abundant apatite and rarer monazite inclusions are typical of these zircons. Zircons from the Black River plutonic suite have well-developed, partially resorbed and (or) porous oscillatory magmatic rims on variably resorbed and highly fractured antecrystic cores. This suite does not appear to contain any zircon inheritance (Appendix B.3). Apatite and monazite inclusions were also present in Black River suite zircons 690 Appendix C – Supporting Materials for Reported 40Ar/39Ar Ages 691 C.1 40Ar/39Ar plateau plots for plutonic and mafic dyke samples from the eastern Selwyn Basin. a) Mafic dykes samples. b) Plutonic samples. 692 a) Mafic dykes samples. KR-05-77 – mafic dyke, Pelly River pluton: 0 20 40 60 80 100 120 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.71±0.68 Ma (2s, including J-error of .5%) MSWD = 1.12, probability=0.35 Includes 91.9% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.0008 0.15 0.16 0.17 0.18 0.19 0.20 39Ar/40Ar 36Ar 40Ar Age = 96.00±0.84 Ma Initial 40Ar/36Ar =284±20 MSWD = 1.12 data-point error ellipses are 2s 693 KR-05-157 – mafic dyke, Roy pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.52±0.69 Ma (2s, including J-error of .5%) MSWD = 1.2, probability=0.30 Includes 99.12% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.0008 0.0010 0.0012 0.0014 0.11 0.13 0.15 0.17 0.19 0.21 39Ar/40Ar 36Ar 40Ar Age = 96.56±0.76 Ma Initial 40Ar/36Ar =293±15 MSWD = 1.5 data-point error ellipses are 2s 694 KR-05-158 – mafic dyke, Roy pluton: 0 20 40 60 80 100 120 140 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 97.61±0.69 Ma (2s, including J-error of .5%) MSWD = 0.88, probability=0.49 Includes 94.1% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0004 0.0008 0.0012 0.0016 0.0020 0.0024 0.07 0.09 0.11 0.13 0.15 0.17 0.19 0.21 39Ar/40Ar 36Ar 40Ar Age = 97.61±0.77 Ma Initial 40Ar/36Ar =295±18 MSWD = 1.11 data-point error ellipses are 2s 695 KR-05-204 – mafic dyke, Cantung pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 90.95±0.77 Ma (2s, including J-error of .5%) MSWD = 0.28, probability=0.89 Includes 63.4% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.0008 0.0010 0.15 0.16 0.17 0.18 0.19 0.20 0.21 39Ar/40Ar 36Ar 40Ar Age = 90.4±1.5 Ma Initial 40Ar/36Ar =316±49 MSWD = 0.13 data-point error ellipses are 2s This is effectively a two-point isochron - of dubious value. 696 0.000 0.001 0.002 0.003 0.004 0.00 0.04 0.08 0.12 0.16 0.20 0.24 39Ar/40Ar 36Ar 40Ar Age = 91.3±1.3 Ma Initial 40Ar/36Ar =302±15 MSWD = 2.6 data-point error ellipses are 2s 697 KR-05-213 – mafic dyke: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.72±0.76 Ma (2s, including J-error of .5%) MSWD = 1.8, probability=0.098 Includes 99.99% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.000 0.001 0.002 0.003 0.04 0.08 0.12 0.16 0.20 0.24 39Ar/40Ar 36Ar 40Ar Age = 96.6±1.4 Ma Initial 40Ar/36Ar =301±50 MSWD = 2.1 data-point error ellipses are 2s 698 b) Plutonic samples. KR-05-22 – Roy pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.17±0.51 Ma (2s, including J-error of .5%) MSWD = 0.73, probability=0.67 Includes 71.9% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0001 0.0002 0.0003 0.0004 0.176 0.180 0.184 0.188 0.192 0.196 0.200 0.204 39Ar/40Ar 36Ar 40Ar Age = 96.27±0.55 Ma Initial 40Ar/36Ar =285±19 MSWD = 0.70 data-point error ellipses are 2s 699 KR-05-26 – Fish pluton: 0 20 40 60 80 100 120 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 97.20±0.57 Ma (2s, including J-error of .5%) MSWD = 1.9, probability=0.10 Includes 59.3% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s Interpret with care; slightly hump-shaped 0.0002 0.0004 0.0006 0.0008 0.0010 0.0012 0.125 0.135 0.145 0.155 0.165 0.175 39Ar/40Ar 36Ar 40Ar Age = 97.6±1.7 Ma Initial 40Ar/36Ar =288±23 MSWD = 2.0 data-point error ellipses are 2s 700 KR-05-32 – Ivo: 0 20 40 60 80 100 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 64.98±0.43 Ma (2s, including J-error of .5%) MSWD = 2.1, probability=0.053 Includes 83.3% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0004 0.0008 0.0012 0.0016 0.0020 0.10 0.14 0.18 0.22 0.26 0.30 39Ar/40Ar 36Ar 40Ar Age = 64.93±0.83 Ma Initial 40Ar/36Ar =297±21 MSWD = 2.5 data-point error ellipses are 2s 701 KR-05-43 – Park pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.95±0.51 Ma (2s, including J-error of .5%) MSWD = 1.4, probability=0.19 Includes 66.5% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.155 0.165 0.175 0.185 0.195 0.205 39Ar/40Ar 36Ar 40Ar Age = 97.02±0.52 Ma Initial 40Ar/36Ar =291.0±8.5 MSWD = 1.5 data-point error ellipses are 2s 702 KR-05-62 – Mount Christie pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 94.52±0.53 Ma (2s, including J-error of .5%) MSWD = 1.2, probability=0.30 Includes 72.2% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00002 0.00006 0.00010 0.00014 0.00018 0.00022 0.00026 0.00030 0.00034 0.182 0.186 0.190 0.194 0.198 39Ar/40Ar 36Ar 40Ar Age = 95.02±0.81 Ma Initial 40Ar/36Ar =265±36 MSWD = 0.98 data-point error ellipses are 2s 703 KR-05-68 – Mount Wilson pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.37±0.53 Ma (2s, including J-error of .5%) MSWD = 1.19, probability=0.31 Includes 71.1% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00000 0.00002 0.00004 0.00006 0.00008 0.193 0.194 0.195 0.196 0.197 39Ar/40Ar 36Ar 40Ar Age = 97.07±0.76 Ma Initial 40Ar/36Ar =86±170 MSWD = 0.92 data-point error ellipses are 2s 704 0.000 0.001 0.002 0.003 0.004 0.00 0.04 0.08 0.12 0.16 0.20 0.24 39Ar/40Ar 36Ar 40Ar Age = 96.13±0.87 Ma Initial 40Ar/36Ar =273±17 MSWD = 9.8 data-point error ellipses are 2s 705 KR-05-76 – Pelly River pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 97.09±0.52 Ma (2s, including J-error of .5%) MSWD = 0.83, probability=0.59 Includes 78.9% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0001 0.0002 0.0003 0.0004 0.0005 0.165 0.175 0.185 0.195 39Ar/40Ar 36Ar 40Ar Age = 97.12±0.55 Ma Initial 40Ar/36Ar =293±15 MSWD = 0.93 data-point error ellipses are 2s 706 KR-05-110 – Cac pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 94.60±0.52 Ma (2s, including J-error of .5%) MSWD = 0.50, probability=0.81 Includes 69.8% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0001 0.0002 0.0003 0.182 0.186 0.190 0.194 0.198 0.202 39Ar/40Ar 36Ar 40Ar Age = 94.66±0.63 Ma Initial 40Ar/36Ar =289±35 MSWD = 0.61 data-point error ellipses are 2s 707 KR-05-113 – South Nahanni pluton: 0 20 40 60 80 100 120 140 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.82±0.51 Ma (2s, including J-error of .5%) MSWD = 0.49, probability=0.82 Includes 56.9% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00000 0.00002 0.00004 0.00006 0.00008 0.00010 0.00012 0.191 0.193 0.195 0.197 0.199 39Ar/40Ar 36Ar 40Ar Age = 95.67±0.63 Ma Initial 40Ar/36Ar =318±84 MSWD = 0.57 data-point error ellipses are 2s 708 KR-05-130 – Central Nahanni pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.97±0.50 Ma (2s, including J-error of .5%) MSWD = 0.89, probability=0.53 Includes 78.5% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0001 0.0002 0.0003 0.0004 0.0005 0.0006 0.16 0.17 0.18 0.19 0.20 39Ar/40Ar 36Ar 40Ar Age = 95.91±0.53 Ma Initial 40Ar/36Ar =297±11 MSWD = 1.09 data-point error ellipses are 2s 709 KR-05-136 – North Nahanni pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.77±0.51 Ma (2s, including J-error of .5%) MSWD = 0.50, probability=0.88 Includes 83.4% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00000 0.00004 0.00008 0.00012 0.00016 0.00020 0.184 0.186 0.188 0.190 0.192 0.194 0.196 0.198 39Ar/40Ar 36Ar 40Ar Age = 96.81±0.52 Ma Initial 40Ar/36Ar =282±36 MSWD = 0.51 data-point error ellipses are 2s 710 KR-05-143 – Circular Stock pluton: 0 40 80 120 160 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 97.57±0.53 Ma (2s, including J-error of .5%) MSWD = 1.5, probability=0.15 Includes 70.7% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.0008 0.15 0.16 0.17 0.18 0.19 0.20 39Ar/40Ar 36Ar 40Ar Age = 97.60±0.67 Ma Initial 40Ar/36Ar =294±18 MSWD = 1.9 data-point error ellipses are 2s 711 KR-05-148 – Rifle Range pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.60±0.52 Ma (2s, including J-error of .5%) MSWD = 0.99, probability=0.44 Includes 77% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00000 0.00004 0.00008 0.00012 0.00016 0.186 0.188 0.190 0.192 0.194 0.196 0.198 39Ar/40Ar 36Ar 40Ar Age = 96.56±0.59 Ma Initial 40Ar/36Ar =298±42 MSWD = 1.2 data-point error ellipses are 2s 712 KR-05-164 – Hole-in-the-Wall batholith: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.21±0.53 Ma (2s, including J-error of .5%) MSWD = 1.11, probability=0.35 Includes 44.8% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0001 0.0002 0.0003 0.182 0.186 0.190 0.194 0.198 0.202 39Ar/40Ar 36Ar 40Ar Age = 95.16±0.67 Ma Initial 40Ar/36Ar =296±32 MSWD = 1.4 data-point error ellipses are 2s 713 KR-05-175 – Lened pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 94.78±0.62 Ma (2s, including J-error of .5%) MSWD = 1.2, probability=0.29 Includes 39% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.000 0.001 0.002 0.003 0.00 0.04 0.08 0.12 0.16 0.20 0.24 39Ar/40Ar 36Ar 40Ar Age = 94.79±0.61 Ma Initial 40Ar/36Ar =294.0±5.9 MSWD = 1.4 data-point error ellipses are 2s 714 KR-05-191 – Marion pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.45±0.51 Ma (2s, including J-error of .5%) MSWD = 1.4, probability=0.18 Includes 95.9% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.16 0.17 0.18 0.19 0.20 39Ar/40Ar 36Ar 40Ar Age = 95.44±0.58 Ma Initial 40Ar/36Ar =294±24 MSWD = 1.7 data-point error ellipses are 2s 715 KR-05-194 – Coal River batholith: 0 20 40 60 80 100 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 67.03±0.39 Ma (2s, including J-error of .5%) MSWD = 2.2, probability=0.056 Includes 93.4% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0004 0.0008 0.0012 0.0016 0.16 0.18 0.20 0.22 0.24 0.26 0.28 0.30 39Ar/40Ar 36Ar 40Ar Age = 67.12±0.58 Ma Initial 40Ar/36Ar =288±18 MSWD = 2.0 data-point error ellipses are 2s 716 KR-05-196 – Coal River batholith: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 94.92±0.51 Ma (2s, including J-error of .5%) MSWD = 1.01, probability=0.44 Includes 100% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.000 0.001 0.002 0.003 0.004 0.00 0.04 0.08 0.12 0.16 0.20 0.24 39Ar/40Ar 36Ar 40Ar Age = 94.84±0.50 Ma Initial 40Ar/36Ar =297.1±4.0 MSWD = 1.11 data-point error ellipses are 2s 717 KR-05-198 – Hole-in-the-Wall batholith: 0 40 80 120 160 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 92.51±0.52 Ma (2s, including J-error of .5%) MSWD = 0.45, probability=0.64 Includes 35.3% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00002 0.00006 0.00010 0.00014 0.00018 0.00022 0.190 0.192 0.194 0.196 0.198 0.200 0.202 0.204 39Ar/40Ar 36Ar 40Ar Age = 92.72±0.86 Ma Initial 40Ar/36Ar =401±87 MSWD = 0.58 data-point error ellipses are 2s 718 KR-05-208 – Little Hyland pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 89.28±0.51 Ma (2s, including J-error of .5%) MSWD = 0.69, probability=0.66 Includes 79.4% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00000 0.00004 0.00008 0.00012 0.00016 0.00020 0.202 0.204 0.206 0.208 0.210 0.212 0.214 39Ar/40Ar 36Ar 40Ar Age = 89.74±0.84 Ma Initial 40Ar/36Ar =245±77 MSWD = 0.59 data-point error ellipses are 2s 719 KR-05-210 – Nahanni Range Road pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 86.49±0.50 Ma (2s, including J-error of .5%) MSWD = 1.6, probability=0.13 Includes 69.5% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.0008 0.0010 0.0012 0.0014 0.12 0.14 0.16 0.18 0.20 0.22 39Ar/40Ar 36Ar 40Ar Age = 86.23±0.59 Ma Initial 40Ar/36Ar =307±16 MSWD = 1.5 data-point error ellipses are 2s 720 03M-207 – Mine Stock-apical, Cantung pluton: 50 60 70 80 90 100 110 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.08±0.43 Ma (2s, including J-error of .4%) MSWD = 1.5, probability=0.097 Includes 94.4% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00005 0.00015 0.00025 0.00035 0.058 0.059 0.060 0.061 0.062 0.063 0.064 39Ar/40Ar 36Ar 40Ar Age = 95.65±0.94 Ma Initial 40Ar/36Ar =267±40 MSWD = 1.5 data-point error ellipses are 2s 721 03M-216 – Faille pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 96.42±0.52 Ma (2s, including J-error of .4%) MSWD = 1.5, probability=0.16 Includes 89.2% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.00005 0.00015 0.00025 0.00035 0.00045 0.00055 0.00065 0.053 0.055 0.057 0.059 0.061 0.063 39Ar/40Ar 36Ar 40Ar Age = 96.9±2.0 Ma Initial 40Ar/36Ar =277±77 MSWD = 1.8 data-point error ellipses are 2s 722 03M-217 – Mount Sir James McBrien (Cirque) pluton: 0 50 100 150 0 20 40 60 80 100 A g e ( M a ) Cumulative 39Ar Percent Plateau age = 95.33±0.46 Ma (2s, including J-error of .4%) MSWD = 1.2, probability=0.27 Includes 70.8% of the 39Ar Plateau steps are filled, rejected steps are open box heights are 2s 0.0000 0.0002 0.0004 0.0006 0.0008 0.051 0.053 0.055 0.057 0.059 0.061 0.063 0.065 39Ar/40Ar 36Ar 40Ar Age = 96.0±1.0 Ma Initial 40Ar/36Ar =261±43 MSWD = 1.16 data-point error ellipses are 2s 723 C.2 40Ar/39Ar analysis conditions and individual analyses used for final plateau age calculations. a) Mafic dyke samples. b) Pluton samples. (available for download at: http://hdl.handle.net/2429/44385) 724 Appendix D – Geochemical Datasets 725 D.1 ALS Chemex major and trace element geochemical data and relevant calculated geochemical parameters. a) Geochemical sample descriptions. b) Major elements (wt.%). c) Rare earth elements (ppm). d) Trace elements (ppm). e) Calculated geochemical parameters used in text. 726 a) Geochemical sample descriptions. Plutonic Suite (sub-suite or region) Pluton Sample Lithology a,b Early to 'mid'-Cretaceous magmatism (in the Selwyn Basin) Anvil (YTT inlier) Square Lake KR-07-13 m-gr qtz-phyric Q>/=P>/=K bt-ms-tourmaline (leuco)granite South Fork South Fork 07M-151 welded dark grey pl-qtz-bt-(hbl) lapilli tuff w/aphanitic to microcrystalline groundmass Tay River (west) Carolyn/Orchay 07M-150 pl-hbl-qtz-phyric (chl-ep ± ms-cal) diorite-granodiorite porphyry Tay River (east) Faille 03M-216 f-m-gr mc-crystic K>P=Q bt>ms monzogranite Tay River (east) Mt. Sir James MacBrien 03M-217 m-c-gr mc-phyric K>P>Q bt>hbl monzogranite Tay River (east) Hole-in-the-Wall (north) 03M-218 m-c-gr mc-phyric K>Q=P bt>hbl monzogranite Tay River (east) Marion KR-05-191 (m)-c-gr mc-phyric (sparse) Q>/=K=P bt monzogranite Tay River (east) Coal River KR-05-196 m-gr equigr K=P>Q bt>hbl monzogranite Tay River (east) Roy KR-05-22 c-gr equigr to pl-phyric P>K>Q bt monzogranite Tay River (east) Fish KR-05-26 f-m-gr pl-phyric P>Q=K bt monzogranite Tay River (east) Roy KR-05-157 v f-gr pl-qtz-phyric quartz diorite dyke Tay River (east) Roy KR-05-158 v f-gr pl-qtz-phyric quartz diorite dyke Tay River (Coal River) Jorgensen KR-05-08 f-gr pl-qtz-hbl-phyric qtz monzodiorite porphyry Tay River (Coal River) Powers KR-05-10 f-gr pl-(hbl)-phyric qtz monzodiorite porphyry Tay River (Coal River) Park KR-05-43 f-m-gr equigr P>>Q>K hbl>bt granodiorite Tay River ('evolved') Shelf Lake 03M-214 f-m-gr mc-crystic K>P>Q bt>ms monzogranite Tay River ('evolved') Mt. Appler 03M-215 f-m-gr mc-crystic K>P>Q bt>ms monzogranite Tay River ('evolved') Pyramid Mtn 03M-219 m-gr mc-phyric bt>>ms K=Q>P monzogranite Tay River ('evolved') Hole-in-the-Wall KR-05-164 f-m-gr mc-crystic K>P>Q bt monzogranite Tay River ('evolved') Ivo KR-05-32 f-gr equigr K>/=Q>P bt>ms grt leucogranite Tungsten (southeast) Mine Stock (apical) 03M-207 f-m-gr (mc-phyric) K=P=Q bt>>ms monzogranite Tungsten (southeast) Ragged 03M-220 f-gr mc-phyric (sparse) bt monzogranite Tungsten (southeast) Circular Stock KR-05-143 m-gr mc-phyric K>P>Q bt monzogranite Tungsten (southeast) Rifle Range KR-05-148 c-gr mc-phyric (sparse) Q>/=K>P bt monzogranite Tungsten (southeast) West Tuna KR-05-207 mylonitic c-gr mc-phyric bt monzogranite Tungsten (southeast) Little Hyland KR-05-208 strongly foliated (mc-augen) f-m-gr equigr Q=K>P bt>>ms monzogranite Tungsten (southeast) Nahanni Range KR-05-210 m-c-gr equigr Q=K>P bt>>ms monzogranite Tungsten (southeast) Mine Stock KR-05-215 f-m-gr mc-phyric (small and sparse) K=Q>P bt>>>ms monzogranite Tungsten (southeast) Mine Stock KR-05-204 altered foliated f-gr bt-phyric kersantite (lamprophyric) dyke Tungsten (southeast) Cantung KR-05-212 v f-gr leucocratic qtz/feldspathoid-phyric? monzonitic porphyry dyke Tungsten (southeast) Cantung KR-05-213 altered v f-gr bt-phyric kersantite (lamprophyric) dyke Tungsten (northeast) Cac KR-05-110 m-gr mc-phyric K=P=Q bt monzogranite Tungsten (northeast) Lened KR-05-175 f-m-gr mc-crystic (pl sub-phyric) K>/=Q=P bt monzogranite Mayo (northeast) Logan CL06-35A f-gr equigr P>K>Q bt>hbl>>cpx qtz monzodiorite Mayo (northeast) Mt. Christie CL06-36A (f)-m-gr mc-phyric (crowded) K>P>Q hbl>bt>>cpx qtz monzonite Mayo (northeast) Christie Pass CL06-37A f-m-gr mc-phyric (near-crowded) K>>P=Q hbl>>bt>>cpx qtz monzonite Mayo (northeast) Keele River CL06-39A f-gr mc-crystic hbl>bt >>cpx K>P>/=Q qtz monzonite Mayo (northeast) Mile 222 CL06-40A f-m-gr mc-phyric hbl>bt>>cpx K=P>Q qtz monzonite Mayo (northeast) South Nahanni KR-05-113 m-c-gr mc-crystic (pl-qtz sub-phyric) K=Q=P bt>hbl>>cpx monzogranite Mayo (northeast) (plug SW of O'Grady) KR-05-126 altered qtz-mc-pl-bt-phyric monzogranitic leuco-porphyry Mayo (northeast) Central Nahanni KR-05-130 m-gr equigr bt granodiorite Mayo (northeast) North Nahanni KR-05-136 f-gr equigr P=K>Q bt>hbl monzogranite Mayo (northeast) Cirque Lake KR-05-47b moderately foliated m-gr pl-bt-qtz-mc-phyric monzogranitic porphyry Mayo (northeast) Rockslide Mtn KR-05-59 m-gr equigr (sparse mc ph-xsts) bt monzogranite Mayo (northeast) Mt. Christie KR-05-62 m-gr mc-phyric K>P>Q hbl>>bt qtz monzonite Mayo (northeast) Gun KR-05-71 m-gr bt monzogranite Mayo (northeast) South Nahanni KR-05-112 alt'd v f-gr equigr mafic dyke (lamprophyre?) Mayo (Hess–MacMillan) Mount Wilson KR-05-68 (foliated) f-gr mc-crystic Q=K=P bt monzogranite Mayo (Hess–MacMillan) Pelly River KR-05-76 foliated f-gr equigr Q=K=P bt qtz diorite Mayo (Hess–MacMillan) Pelly River KR-05-77 altered bt-phyric kersantite (lamprophyric) dyke Tombstone (northeast) Dechen’La CL06-33A m-gr mc-crystic hbl-phyric K>P>/=Q hbl>bt>>cpx monzogranite Tombstone (northeast) Natla CL06-34A f-m-gr mc-crystic K>P>Q bt monzogranite Tombstone (northeast) O'Grady KR-05-97b f-m-gr mc-phyric (crowded) K>>P>>Q hbl>cpx>bt tur-bearing qtz monzonite 727 Plutonic Suite (sub-suite or region) Pluton Sample Lithology a,b Tombstone (Hess– MacMillan) Ross River CL06-38A f-m-gr mc-phyric (scattered) Q=P>K bt (leuco)granite Tombstone (southeast) Coal River KR-05-194 m-gr mc-phyric K=P>/=Q bt monzogranite Tombstone (southeast) Hole-in-the-Wall KR-05-198 c-gr mc-phyric (sparse) Q>>P>K bt monzogranite Early to 'mid'-Cretaceous magmatism (in the Cassiar Terrane) Cassiar (muscovite-bearing) Dycer Creek KR-07-01 f-m-gr mc-phyric K=P=Q bt>ms monzogranite Cassiar (muscovite-bearing) Quiet Lake KR-07-02 m-gr equigr (mc-phyric) Q>K=P bt>ms monzogranite Cassiar (muscovite-bearing) Cassiar KR-07-16 m-(c)-gr mc-phyric Q>K>P bt>>ms (leuco)granite Cassiar (muscovite-bearing) Cassiar KR-07-17 f-(m)-gr equigr (mc-phyric) Q=K>P bt>ms monzogranite Cassiar (muscovite-bearing) Cassiar KR-07-18 f-(m)-gr equigr (mc-bt-phyric) Q=K>P bt>ms monzogranite Cassiar (muscovite-bearing) Cassiar KR-07-22 v f-gr mylonitic mc-augen K=Q>P bt-ms monzogranite Cassiar (biotite-bearing) Buttle Creek 07M-152 foliated m-gr equigr Q>P>K bt granodiorite Cassiar (biotite-bearing) Glenlyon 07M-153 f-m-gr equigr P=Q>K bt granodiorite Cassiar (biotite-bearing) Nisutlin KR-07-05 f-m-gr foliated equigr Q>/=K>P bt leucogranite Cassiar (biotite-bearing) Young KR-07-07 f-gr equigr (mc-phyric) K=Q=P bt monzogranite Cassiar (biotite-bearing) St. Cyr KR-07-08 f-(m)-g equigr P=Q>K bt granodiorite Cassiar (biotite-bearing) Fox Mountain KR-07-09 f-gr qtz-fsp-bt (leuco)granitic porphyry Cassiar (biotite-bearing) Big Salmon KR-07-10 f-m-gr (mc-phyric) K=Q>/=P bt monzogranite Cassiar (biotite-bearing) Glenlyon KR-07-11 f-gr equigr P=K=Q bt leucogranite Cassiar (biotite-bearing) Marker Lake KR-07-19 c-gr mc-crystic K=Q-P bt granodiorite Cassiar (biotite-bearing) Cabin Creek KR-07-20 m-c-gr equigr Q=K-P bt granodiorite Cassiar (biotite-bearing) Cassiar KR-07-24 m-gr equigr Q=K>P bt monzogranite/granodiorite Cassiar (hornblende- bearing) Nisutlin KR-07-06 c-gr mc-phyric Q>K>P bt>hbl monzogranite Seagull Little Salmon 07M-154 (m)-c-gr equigr P=Q>K bt monzogranite Seagull Quiet Lake KR-07-03 m-gr equigr K>/=Q>P bt>ms (leuco)granite Seagull? Nisutlin KR-07-04 f-gr equigr ms>bt Q>/=P>K leucogranite Late Cretaceous magmatism (in the Cassiar Terrane) Rancheria Allen Stock KR-07-21 c-gr mc-crystic K>Q>P bt monzogranite Rancheria Meister Lake KR-07-15 m-gr equigr K=Q>P ms>bt (leuco)granite Rancheria Gravel Creek KR-07-23 c-gr equigr (sub-mc-phyric) Q=K=P bt>ms monzogranite Latest Cretaceous to Eocene magmatism (in the Cassiar Terrane) Black River Tintina KR-07-12 f-m-gr mc-crystic qtz-phyric Q>K>P bt (sub-phyric) (leuco)granite sub- porphyry Black River Black River KR-07-14 m-gr mc-phyric bt-(ms) Q=K>>P monzogranite Notes: a mineral abbreviations are after Kretz (1983) or P = plagioclase, K = K-feldspar, and Q = quartz; b textural abbreviations are: f = fine, m = medium, c = coarse, gr = grained, equigr = equigranular, -crystic = megacrystic. 728 b) Major elements (wt.%). Sample SiO2 TiO2 Al2O3 FeT Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI TOTAL Early to 'mid'-Cretaceous magmatism (in the Selwyn Basin) KR-07-13 71.74 0.16 15.63 1.11 0.04 0.96 0.02 0.54 1.18 3.36 4.92 0.07 1.03 99.82 07M-151 68.07 0.48 15.18 3.70 0.42 2.95 0.06 1.21 2.89 2.32 3.28 0.09 2.18 99.62 07M-150 64.05 0.62 16.26 4.63 0.71 3.52 0.08 2.00 4.30 2.58 2.77 0.11 2.38 99.90 03M-216 67.58 0.43 15.45 3.49 0.53 2.65 0.06 0.92 2.53 2.37 4.86 0.18 0.63 98.68 03M-217 67.84 0.45 15.61 3.44 0.83 2.34 0.06 1.27 3.09 2.50 4.54 0.18 0.40 99.56 03M-218 68.19 0.44 15.49 3.26 0.88 2.14 0.05 1.19 2.96 2.55 4.82 0.15 0.39 99.63 KR-05-191 70.61 0.33 14.83 2.35 0.20 1.93 0.03 0.76 2.10 3.03 4.45 0.10 0.94 99.65 KR-05-196 68.27 0.49 15.14 3.41 0.77 2.37 0.05 1.26 2.84 2.89 4.27 0.13 0.57 99.46 KR-05-22 69.95 0.43 14.75 2.30 0.81 1.34 0.03 0.84 2.53 3.28 3.91 0.12 0.32 98.56 KR-05-26 72.82 0.21 14.01 2.04 0.55 1.34 0.07 0.32 1.38 3.27 4.60 0.06 0.51 99.35 KR-05-157 54.05 0.84 16.75 7.21 0.86 5.70 0.12 4.98 6.90 2.16 3.08 0.45 2.79 99.51 KR-05-158 52.71 0.81 15.74 7.41 1.04 5.72 0.11 6.37 7.25 1.99 2.71 0.34 3.27 98.92 KR-05-08 59.64 0.53 16.17 3.99 0.35 3.27 0.09 2.39 4.17 3.48 2.24 0.12 5.27 98.27 KR-05-10 62.15 0.71 16.89 4.63 1.13 3.14 0.06 2.16 4.26 4.02 2.17 0.13 2.89 100.25 KR-05-43 63.69 0.56 16.60 4.47 1.54 2.63 0.08 2.05 4.61 3.37 2.52 0.12 0.55 98.74 03M-214 70.72 0.27 15.30 2.00 0.06 1.75 0.03 0.73 1.80 2.86 4.67 0.20 0.87 99.55 03M-215 70.97 0.39 14.99 2.38 0.18 1.98 0.03 0.56 1.65 2.66 5.17 0.13 0.33 99.34 03M-219 71.98 0.36 14.89 2.14 0.23 1.72 0.03 0.61 1.47 2.63 5.00 0.17 0.50 99.82 KR-05-164 70.33 0.23 15.00 1.86 0.16 1.53 0.04 0.35 1.09 2.91 6.22 0.09 0.31 98.55 KR-05-32 74.77 0.21 13.16 1.28 0.14 1.02 0.03 0.25 1.09 2.43 4.95 0.03 0.37 98.62 03M-207 70.50 0.35 15.60 2.75 0.08 2.39 0.04 0.68 2.40 2.82 3.93 0.13 0.41 99.71 03M-220 66.98 0.47 16.03 3.39 0.49 2.60 0.05 1.30 3.18 2.74 3.56 0.13 1.62 99.57 KR-05-143 69.73 0.34 15.46 2.49 0.20 2.06 0.04 0.69 2.31 3.05 4.31 0.11 0.50 99.13 KR-05-148 69.22 0.26 15.72 2.11 0.11 1.80 0.03 0.54 1.63 2.77 6.21 0.10 0.41 99.14 KR-05-207 71.23 0.23 15.16 2.21 0.22 1.79 0.06 0.56 1.90 3.45 4.00 0.08 0.56 99.49 KR-05-208 72.25 0.20 14.61 2.15 0.01 1.92 0.10 0.45 1.64 3.33 3.89 0.10 0.45 99.22 KR-05-210 73.27 0.17 14.46 1.75 0.18 1.41 0.07 0.40 1.19 3.19 3.95 0.07 0.71 99.26 KR-05-215 71.53 0.26 15.04 2.20 0.05 1.93 0.04 0.53 1.93 3.26 4.26 0.10 0.43 99.66 KR-05-204 52.18 1.03 17.23 6.19 0.28 5.31 0.19 5.33 7.26 1.93 5.56 0.61 1.45 99.23 KR-05-212 52.92 0.85 15.60 6.10 0.68 4.87 0.05 7.48 8.59 1.57 3.77 0.36 1.61 99.32 KR-05-213 52.85 0.85 15.45 7.79 1.15 5.96 0.12 7.86 7.67 2.21 3.24 0.36 1.20 99.84 KR-05-110 67.50 0.42 15.44 3.08 0.22 2.57 0.06 0.91 2.47 2.56 4.77 0.18 0.88 98.41 KR-05-175 69.63 0.43 15.35 2.86 0.00 2.57 0.06 0.89 2.31 2.75 4.62 0.16 0.60 99.78 CL06-35A 61.54 0.80 17.28 5.74 0.63 4.59 0.09 2.74 5.22 2.35 3.21 0.12 0.56 99.75 CL06-36A 63.04 0.64 14.72 5.07 1.37 3.32 0.09 2.96 4.13 1.96 5.91 0.25 0.93 99.89 CL06-37A 63.36 0.60 14.72 4.77 0.62 3.73 0.08 2.89 4.28 2.02 5.80 0.24 0.80 99.75 CL06-39A 63.69 0.61 15.66 4.58 0.46 3.70 0.08 2.86 3.67 2.08 5.23 0.22 0.96 99.85 CL06-40A 61.89 0.65 15.55 5.23 0.89 3.90 0.09 3.34 4.54 1.97 5.38 0.26 0.62 99.73 KR-05-113 69.93 0.39 14.80 2.74 0.31 2.18 0.04 1.18 2.51 2.59 4.74 0.16 0.34 99.54 KR-05-126 71.14 0.11 15.05 0.73 0.02 0.64 0.03 0.13 1.49 3.17 4.63 0.05 1.96 98.54 KR-05-130 65.08 0.57 16.25 4.14 0.44 3.32 0.06 2.26 4.10 2.76 3.40 0.12 0.35 99.24 KR-05-136 66.75 0.47 15.97 3.67 0.54 2.81 0.05 1.69 3.47 2.61 3.79 0.09 0.47 99.14 KR-05-47b 68.18 0.42 15.62 3.51 0.45 2.75 0.06 1.04 2.88 2.36 4.41 0.12 0.93 99.67 KR-05-59 68.93 0.51 14.98 3.00 0.15 2.56 0.05 1.27 3.37 2.40 3.60 0.09 0.70 99.00 KR-05-62 61.34 0.74 14.37 5.06 0.36 4.22 0.09 3.64 4.68 1.93 5.64 0.29 0.72 98.69 KR-05-71 66.73 0.62 15.84 3.25 0.18 2.76 0.04 1.50 3.70 2.32 4.00 0.10 0.59 98.81 KR-05-112 51.38 0.78 16.14 5.92 0.50 4.87 0.07 9.02 6.74 1.96 2.89 0.12 3.94 99.32 KR-05-68 68.11 0.52 15.47 3.05 0.20 2.56 0.04 1.38 3.39 2.20 3.88 0.09 0.84 99.06 KR-05-76 56.81 1.04 16.61 9.84 7.20 2.37 0.11 3.12 6.61 4.65 0.37 0.27 0.42 99.94 KR-05-77 50.72 0.93 11.88 6.64 1.06 5.01 0.12 7.97 6.74 1.13 4.01 0.41 7.79 98.60 CL06-33A 66.26 0.47 15.34 3.79 0.65 2.82 0.07 2.17 3.55 2.43 4.76 0.17 0.61 99.79 CL06-34A 69.14 0.33 15.13 2.91 0.35 2.30 0.05 1.07 2.15 2.34 5.46 0.18 0.99 99.85 KR-05-97b 63.35 0.57 15.09 4.37 1.23 2.82 0.07 1.99 3.81 2.15 6.44 0.27 0.43 98.73 CL06-38A 69.11 0.32 15.04 2.60 0.33 2.04 0.04 0.86 2.37 2.70 4.32 0.08 0.81 98.32 KR-05-194 70.19 0.54 14.94 2.91 0.34 2.31 0.05 0.71 1.42 2.80 5.26 0.19 0.57 99.63 KR-05-198 68.34 0.49 15.10 3.70 0.27 3.08 0.06 0.75 2.01 2.76 4.95 0.13 0.61 98.99 729 Sample SiO2 TiO2 Al2O3 FeT Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI TOTAL Early to 'mid'-Cretaceous magmatism (in the Cassiar Terrane) KR-07-01 70.28 0.27 15.22 1.97 0.47 1.35 0.04 0.65 1.60 3.16 4.30 0.13 0.79 98.56 KR-07-02 72.00 0.19 15.42 1.58 0.30 1.15 0.03 0.47 1.46 3.36 4.26 0.14 0.93 99.97 KR-07-16 73.11 0.24 14.53 1.67 0.60 0.96 0.04 0.56 1.46 3.35 4.06 0.11 0.46 99.74 KR-07-17 73.00 0.23 14.76 1.74 0.39 1.21 0.03 0.51 1.39 3.01 4.58 0.12 0.49 99.99 KR-07-18 71.87 0.28 13.91 1.60 0.53 0.96 0.03 0.50 0.66 2.36 5.96 0.17 0.95 98.40 KR-07-22 73.02 0.15 14.62 1.52 0.09 1.28 0.03 0.35 1.05 3.51 4.39 0.13 1.04 99.95 07M-152 68.66 0.43 16.39 3.48 0.62 2.57 0.07 0.90 3.49 3.69 1.78 0.15 0.73 99.86 07M-153 68.28 0.37 16.31 3.17 0.32 2.56 0.06 1.11 3.16 3.10 3.37 0.15 0.55 99.79 KR-07-05 73.89 0.07 14.98 1.15 0.15 0.90 0.02 0.30 1.29 2.97 4.34 0.08 0.70 99.90 KR-07-07 69.30 0.37 15.41 2.80 0.38 2.17 0.07 0.96 2.27 2.86 4.39 0.16 0.95 99.66 KR-07-08 63.61 0.61 16.88 4.84 1.41 3.08 0.09 1.98 4.12 3.04 2.91 0.15 1.44 99.79 KR-07-09 71.02 0.29 15.16 2.46 0.26 1.98 0.05 0.77 2.14 3.11 4.03 0.12 0.63 99.90 KR-07-10 72.31 0.38 14.12 2.54 1.04 1.35 0.06 0.85 1.76 3.00 4.26 0.14 0.36 99.89 KR-07-11 71.78 0.14 15.82 1.19 -0.01 1.08 0.02 0.43 1.73 3.37 4.61 0.08 0.60 99.89 KR-07-19 66.70 0.66 15.87 4.23 1.54 2.42 0.08 1.45 2.96 3.56 3.14 0.28 0.68 99.81 KR-07-20 72.17 0.19 15.30 1.76 0.62 1.02 0.03 0.56 2.41 3.23 3.42 0.08 0.40 99.82 KR-07-24 70.26 0.29 15.67 2.22 0.65 1.41 0.03 0.78 2.25 3.49 3.80 0.13 0.55 99.64 KR-07-06 69.59 0.42 15.27 2.66 0.95 1.54 0.06 0.87 2.04 3.25 4.61 0.18 0.76 99.90 07M-154 73.91 0.25 13.70 1.87 0.44 1.28 0.03 0.22 0.80 3.60 4.94 0.05 0.34 99.77 KR-07-03 76.00 0.03 13.83 0.77 0.06 0.64 0.03 0.14 0.67 3.49 4.60 0.03 0.29 99.91 KR-07-04 76.64 0.01 13.73 0.42 0.00 0.38 0.04 0.04 0.23 4.29 4.09 0.01 0.32 99.83 Late Cretaceous magmatism (in the Cassiar Terrane) KR-07-21 70.42 0.36 14.71 2.33 0.90 1.28 0.04 0.64 1.52 3.20 4.86 0.17 0.49 98.91 KR-07-15 74.10 0.08 14.75 1.18 0.40 0.70 0.03 0.22 0.35 3.27 4.88 0.11 0.84 99.86 KR-07-23 73.13 0.16 14.45 1.91 0.84 0.96 0.02 0.25 0.57 2.89 4.50 0.14 1.10 99.17 Latest Cretaceous to Eocene magmatism (in the Cassiar Terrane) KR-07-12 73.10 0.22 14.25 1.95 0.24 1.54 0.04 0.33 0.97 3.02 5.25 0.12 0.59 99.90 KR-07-14 73.79 0.21 13.72 1.98 0.63 1.21 0.04 0.32 0.96 2.87 5.08 0.10 0.53 99.63 730 c) Rare earth elements (ppm). Sample Ba Co Cr Cs Cu Ga Hf Mo Nb Ni Pb Rb Sn Sr Ta Th Tl U V W Y Zn Zr Early to 'mid'-Cretaceous magmatism (in the Selwyn Basin) KR-07-13 403 1.4 20 21.1 26.3 7.0 11 10 52 347.0 10 169 1.50 19 1.7 4.62 9 2 8.3 54 282 07M-151 1120 5.8 20 5.1 5 18.6 8.7 0 13 0 22 131.0 3 240 0.90 16 0.5 3.38 48 2 22.6 64 367 07M-150 741 9.1 20 3.4 7 17.9 5.0 3 11 20 116.5 3 320 0.80 13 0.6 2.70 84 4 21.6 62 193 03M-216 1285 4.8 6.7 19.0 14.0 21 6 31 183.5 5 425 1.70 33 0.6 4.20 54 1 24.2 62 576 03M-217 1250 6.6 6.3 18.0 14.0 20 7 25 175.5 4 458 1.50 31 0.6 5.80 61 2 16.4 47 547 03M-218 1155 6.0 4.7 17.0 11.0 18 8 21 187.0 3 417 1.70 33 0.5 9.10 60 4 16.8 35 424 KR-05-191 856 3.1 20 7.3 14.0 8.0 22 18 191.0 4 283 2.00 28 6.00 30 5 18.1 37 274 KR-05-196 1015 5.7 20 8.4 15.0 7.0 19 9 165.0 3 322 1.70 26 5.20 49 7 18.8 34 277 KR-05-22 658 4.1 30 4.5 17.0 9.0 2 22 9 14 176.5 3 335 1.90 25 7.10 36 3 21.0 25 315 KR-05-26 389 2.0 20 4.9 18.0 9.0 28 20 217.0 3 169 2.00 34 4.40 17 6 32.4 36 281 KR-05-157 1200 23.3 80 8.6 26 16.0 6.0 2 19 42 13 118.5 2 663 1.20 14 3.60 186 4 22.3 101 215 KR-05-158 1180 27.5 300 6.1 24 15.0 5.0 15 77 5 104.0 2 558 0.90 13 3.40 182 4 21.8 97 180 KR-05-08 1050 13.0 70 2.0 16 19.0 5.0 2 14 49 11 111.5 3 454 1.00 10 2.40 87 3 18.6 114 191 KR-05-10 775 10.2 30 2.0 5 20.0 6.0 2 14 14 9 74.4 2 720 1.10 9 2.30 106 3 16.7 49 196 KR-05-43 750 10.6 30 2.6 5 19.0 7.0 18 9 16 91.6 2 463 1.60 16 2.20 82 2 16.0 57 240 03M-214 574 3.5 14.7 17.0 9.0 17 6 43 212.0 6 253 2.50 16 0.7 5.30 32 2 15.6 26 349 03M-215 508 3.6 10.4 20.0 11.0 22 6 34 239.0 7 166 2.40 25 0.7 11.60 25 2 17.5 52 453 03M-219 358 3.7 14.2 20.0 9.0 21 7 35 293.0 9 147 2.50 21 0.8 6.80 24 2 11.9 57 354 KR-05-164 930 2.4 20 15.6 17.0 7.0 18 39 313.0 10 113 2.80 18 0.6 13.70 12 8 25.6 48 242 KR-05-32 229 1.2 20 3.3 14.0 10.0 23 26 175.0 2 86 1.50 37 3.40 6 3 21.4 24 337 03M-207 666 3.3 7.1 19.0 7.0 12 5 39 173.0 3 314 1.20 19 0.5 9.20 25 1 17.2 40 288 03M-220 804 5.9 6.2 20.0 8.0 12 8 35 150.0 5 351 1.20 19 0.5 4.00 51 1 17.9 40 301 KR-05-143 710 3.4 20 8.5 17.0 7.0 11 35 192.0 6 215 1.60 20 8.90 26 2 21.2 37 260 KR-05-148 1080 2.7 20 7.8 16.0 8.0 18 37 255.0 3 339 1.60 24 0.7 7.10 20 3 21.2 28 276 KR-05-207 372 2.1 10 10.2 13.0 6.0 9 27 147.5 5 134 1.20 12 6.10 16 5 14.6 32 199 KR-05-208 387 2.7 20 20.9 17.0 6.0 13 32 262.0 10 128 3.50 14 0.5 7.20 21 6 22.6 50 241 KR-05-210 247 2.2 10 15.8 15.0 7.0 11 33 211.0 6 91 2.20 13 0.5 5.50 13 7 18.4 32 250 KR-05-215 584 2.8 20 9.8 17.0 7.0 11 41 203.0 3 201 1.30 18 0.5 8.80 23 6 20.9 33 268 KR-05-204 2160 15.2 160 17.7 173 17.0 6.0 24 13 5 309.0 101 502 1.60 18 0.9 7.70 163 8 23.4 225 200 KR-05-212 3400 29.6 370 3.7 54 15.0 5.0 17 109 104.0 7 585 1.20 17 5.00 203 9 20.9 32 169 KR-05-213 1430 31.5 390 7.4 28 16.0 5.0 17 122 5 104.0 6 547 1.20 16 4.80 201 5 20.3 62 173 KR-05-110 845 4.0 20 8.5 16.0 8.0 16 38 205.0 4 283 2.10 25 0.5 10.40 35 4 21.2 43 272 KR-05-175 963 4.5 30 8.6 17.0 9.0 17 43 218.0 8 315 2.10 30 0.6 4.80 41 4 21.9 46 295 CL06-35A 579 12.8 40 8.8 9 19.3 8.5 11 13 30 117.5 3 291 0.90 15 0.5 2.02 93 3 22.3 74 350 CL06-36A 1075 13.6 100 14.6 25 17.5 8.8 2 18 24 47 271.0 5 469 1.50 56 1.5 10.10 104 8 23.4 69 350 CL06-37A 1030 12.3 100 21.7 18 17.2 8.9 2 18 25 49 276.0 6 455 1.70 52 1.5 14.40 96 10 23.6 64 346 CL06-39A 1145 11.0 90 12.5 10 18.1 7.5 16 22 38 223.0 4 432 1.40 40 1.0 6.40 84 10 21.1 62 303 731 Sample Ba Co Cr Cs Cu Ga Hf Mo Nb Ni Pb Rb Sn Sr Ta Th Tl U V W Y Zn Zr CL06-40A 1330 15.1 120 12.1 20 18.9 9.0 18 26 44 251.0 5 510 1.40 43 1.3 7.66 114 4 24.7 69 354 KR-05-113 681 5.4 50 10.8 16.0 8.0 16 8 31 228.0 2 335 1.60 37 7.90 43 3 19.0 27 274 KR-05-126 258 0.6 20 12.0 22.0 5.0 12 40 240.0 11 161 1.40 6 0.6 3.40 5 10.6 39 163 KR-05-130 1020 10.0 60 10.6 8 18.0 8.0 14 16 30 147.5 4 364 1.30 18 5.40 68 2 16.0 51 259 KR-05-136 644 8.9 40 12.4 5 18.0 8.0 13 14 37 163.0 2 286 1.40 24 7.40 61 3 19.6 53 270 KR-05-47b 1040 6.1 30 8.1 6 20.0 10.0 16 6 35 204.0 4 409 1.30 30 0.6 6.80 48 4 17.4 56 344 KR-05-59 634 7.1 40 7.7 6 19.0 8.0 10 7 33 162.5 4 253 1.20 15 1.90 41 3 20.8 52 293 KR-05-62 1240 17.3 150 14.6 27 18.0 9.0 2 18 25 41 272.0 7 502 1.50 48 1.0 12.40 130 9 24.5 59 309 KR-05-71 721 8.9 40 7.3 21.0 9.0 11 8 33 192.5 4 271 1.10 17 0.5 3.70 38 6 15.8 54 313 KR-05-112 2770 28.5 380 5.0 5 18.0 5.0 10 104 6 94.8 3 392 0.70 16 5.70 154 4 16.6 44 157 KR-05-68 524 7.3 40 13.0 12 20.0 9.0 9 12 27 184.5 4 242 1.00 20 4.30 34 4 17.8 47 326 KR-05-76 673 7.2 50 9.0 22.0 10.0 10 15 33 189.0 4 292 1.20 17 0.5 5.30 35 52 14.6 50 346 KR-05-77 1430 33.5 700 10.3 37 16.0 9.0 20 136 14 198.0 4 387 1.50 39 0.7 8.40 162 9 19.1 61 290 CL06-33A 937 9.0 80 9.1 5 17.4 8.5 0 13 16 36 210.0 3 446 1.00 24 1.1 7.34 71 2 17.2 54 339 CL06-34A 611 4.6 20 22.2 0 17.4 10.3 0 14 10 30 266.0 7 305 1.50 30 1.4 4.90 48 3 16.3 61 424 KR-05-97b 1160 9.8 50 30.0 17 17.0 8.0 18 10 48 346.0 6 526 1.80 52 1.2 14.50 100 11 20.8 66 269 CL06-38A 444 4.5 30 8.5 5 18.7 9.1 0 12 10 37 206.0 2 175 1.30 21 0.8 5.16 21 3 16.4 42 379 KR-05-194 524 3.9 20 12.1 17.0 10.0 3 31 23 237.0 7 173 2.70 34 0.5 5.60 25 5 19.4 67 366 KR-05-198 883 5.0 20 11.4 19.0 10.0 30 29 224.0 8 145 2.60 19 4.40 37 10 33.6 91 355 Early to 'mid'-Cretaceous magmatism (in the Cassiar Terrane) KR-07-01 1090 3.1 10 9.6 9 19.4 9.1 12 9 28 205.0 4 319 1.40 23 1.0 6.63 32 2 12.7 58 383 KR-07-02 862 2.2 20 23.1 10 21.4 7.8 14 8 27 265.0 7 226 2.90 17 1.3 3.09 20 3 10.1 56 328 KR-07-16 1195 3.1 20 7.6 19.2 7.3 20 8 31 195.0 4 383 1.70 23 1.0 7.91 35 4 26.3 50 340 KR-07-17 850 2.2 10 3.3 18.9 9.0 18 23 215.0 4 223 1.40 20 0.9 2.78 19 8 12.0 44 386 KR-07-18 780 1.8 10 5.4 24 21.2 10.0 11 32 302.0 4 200 0.60 46 1.5 4.56 16 3 12.4 133 410 KR-07-22 942 1.5 20 3.8 0 16.6 7.0 13 51 143.0 6 288 1.10 46 0.9 7.11 16 4 20.1 39 271 07M-152 525 4.3 10 12.6 0 20.8 8.6 18 13 152.0 8 332 1.40 11 0.6 3.47 35 8 15.7 74 359 07M-153 1005 4.8 10 4.1 0 19.4 6.8 13 21 139.0 3 359 1.00 13 0.5 2.02 45 4 15.6 71 280 KR-07-05 679 0.9 20 4.1 18.3 6.7 13 7 38 168.5 2 252 1.20 13 0.7 4.30 0 3 11.4 38 278 KR-07-07 833 4.0 20 9.5 16.9 7.3 21 8 27 219.0 4 358 2.30 22 1.0 8.41 39 11 22.4 51 302 KR-07-08 737 8.9 20 5.2 7 17.0 6.2 13 10 14 120.5 3 303 0.90 15 0.5 2.87 91 17 23.8 57 258 KR-07-09 751 3.3 10 2.8 15.7 6.9 14 8 23 155.0 3 254 1.20 19 0.7 3.89 33 2 18.0 38 284 KR-07-10 645 4.1 20 6.3 15.8 8.2 21 8 24 243.0 3 238 1.50 32 1.4 6.32 38 6 17.8 44 346 KR-07-11 837 2.5 20 7.7 19.8 8.2 11 8 56 213.0 6 312 1.40 10 1.0 3.01 15 3 11.3 36 361 KR-07-19 1335 6.8 10 4.0 21.0 10.0 19 17 129.0 4 485 1.20 22 0.6 4.13 74 2 25.3 93 449 KR-07-20 1630 3.0 20 2.6 18.3 7.7 8 25 119.0 2 850 0.70 14 0.5 1.84 24 3 6.0 34 319 KR-07-24 1025 4.0 20 2.8 20.1 7.2 11 29 123.5 4 540 1.00 15 0.5 3.35 39 3 9.4 60 299 KR-07-06 1115 3.6 20 3.4 0 16.5 9.6 53 8 21 167.0 3 470 3.50 53 0.9 7.86 42 3 26.8 48 398 07M-154 400 1.7 10 13.5 25.9 11.1 54 0 23 245.0 10 72 4.70 25 1.0 5.18 9 6 36.2 74 417 732 Sample Ba Co Cr Cs Cu Ga Hf Mo Nb Ni Pb Rb Sn Sr Ta Th Tl U V W Y Zn Zr KR-07-03 170 0.6 10 5.0 17.3 7.7 20 0 31 247.0 5 53 2.10 19 1.2 5.50 0 2 26.7 32 264 KR-07-04 11 0.5 10 55.0 22.1 8.2 96 7 45 912.0 4 3 12.60 15 3.1 29.40 0 14 26.7 16 257 Late Cretaceous magmatism (in the Cassiar Terrane) KR-07-21 1065 3.5 20 4.7 20.6 9.7 70 0 16 228.0 4 531 3.50 36 0.7 12.50 27 3 19.2 35 402 KR-07-15 349 1.3 10 11.7 20.4 7.4 43 7 28 325.0 20 102 7.00 24 1.3 6.16 5 11 18.2 40 307 KR-07-23 302 1.5 10 6.0 27.9 9.1 41 0 26 322.0 5 95 4.00 22 1.2 8.48 7 4 16.5 50 385 Latest Cretaceous to Eocene magmatism (in the Cassiar Terrane) KR-07-12 519 2.5 20 26.5 28.2 12.4 48 8 38 493.0 13 146 6.70 54 1.9 11.80 14 13 54.5 47 500 KR-07-14 209 2.2 20 6.4 18.5 9.0 27 7 32 295.0 8 76 3.00 46 1.1 6.07 9 6 39.1 40 366 733 d) Trace elements (ppm). Sample La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Early to 'mid'-Cretaceous magmatism (in the Selwyn Basin) KR-07-13 30.8 59.3 6.5 23 4.42 0.52 3.46 0.39 1.7 0.28 0.8 0.11 0.60 0.09 07M-151 43.3 85.3 8.9 32 6.03 1.15 5.61 0.76 4.4 0.83 2.6 0.35 2.40 0.35 07M-150 33.2 64.9 6.8 25 4.96 1.05 4.60 0.69 4.0 0.79 2.4 0.33 2.27 0.35 03M-216 85.3 153.5 16.5 56 9.00 1.70 7.80 1.00 4.4 0.90 2.5 0.40 2.30 0.40 03M-217 60.6 108.5 11.5 39 6.20 1.30 5.30 0.70 3.0 0.60 1.7 0.20 1.50 0.30 03M-218 54.1 97.6 10.3 35 5.70 1.30 5.00 0.60 3.0 0.60 1.7 0.30 1.60 0.30 KR-05-191 50.9 92.8 9.5 34 5.70 1.30 4.60 0.60 3.5 0.70 2.0 0.30 2.10 0.30 KR-05-196 66.6 120.0 12.0 42 6.40 1.20 5.30 0.70 3.7 0.70 2.1 0.30 2.00 0.30 KR-05-22 57.9 102.0 10.8 35 5.60 1.10 4.90 0.70 3.5 0.70 2.1 0.30 2.20 0.30 KR-05-26 64.8 118.5 12.9 44 7.90 1.10 7.20 1.00 5.8 1.10 3.1 0.50 3.20 0.50 KR-05-157 50.4 97.8 10.8 44 8.00 2.00 6.40 0.90 4.6 0.90 2.5 0.30 2.20 0.30 KR-05-158 45.1 88.7 10.0 41 7.20 1.80 6.10 0.90 4.7 0.80 2.4 0.30 2.20 0.30 KR-05-08 30.4 54.6 6.0 21 3.80 1.00 3.50 0.50 2.7 0.60 1.6 0.20 1.60 0.20 KR-05-10 26.6 49.7 5.7 20 3.80 1.10 3.50 0.50 2.9 0.60 1.7 0.30 1.60 0.30 KR-05-43 38.3 64.3 6.8 22 3.70 1.00 3.60 0.50 2.7 0.60 1.7 0.30 1.70 0.30 03M-214 31.5 60.9 7.0 24 4.50 0.90 3.90 0.50 2.7 0.50 1.6 0.20 1.40 0.20 03M-215 47.6 93.9 10.3 36 6.60 1.00 5.60 0.70 3.3 0.60 1.6 0.30 1.50 0.30 03M-219 38.4 77.3 8.8 31 5.60 0.80 4.40 0.50 2.3 0.40 1.2 0.20 1.00 0.20 KR-05-164 30.5 61.3 6.8 26 5.00 0.90 4.40 0.70 4.7 0.90 2.9 0.40 3.00 0.40 KR-05-32 48.5 94.7 10.7 37 6.80 0.60 6.10 0.80 4.0 0.80 2.1 0.30 1.80 0.30 03M-207 46.2 88.0 10.0 34 6.30 1.20 5.20 0.70 3.1 0.60 1.7 0.20 1.50 0.20 03M-220 43.6 79.7 9.1 32 5.50 1.20 4.80 0.60 3.2 0.60 1.8 0.30 1.70 0.30 KR-05-143 38.0 75.0 8.3 32 6.10 1.10 4.80 0.70 4.0 0.70 2.1 0.30 2.00 0.30 KR-05-148 45.6 89.5 9.7 37 6.50 1.50 5.00 0.80 4.1 0.80 2.2 0.30 2.10 0.30 KR-05-207 26.2 50.2 5.3 20 3.80 0.70 3.10 0.50 2.8 0.50 1.5 0.20 1.40 0.20 KR-05-208 27.6 53.9 5.6 21 4.50 0.70 4.10 0.70 4.2 0.80 2.4 0.30 2.30 0.30 KR-05-210 24.3 48.0 5.0 19 4.00 0.40 3.40 0.60 3.5 0.70 2.0 0.30 2.10 0.30 KR-05-215 40.6 79.3 8.4 31 5.50 1.00 5.00 0.70 4.0 0.70 2.2 0.30 2.00 0.20 KR-05-204 59.1 112.0 12.4 48 8.30 2.00 6.80 0.90 4.9 0.90 2.6 0.30 2.20 0.30 KR-05-212 61.7 117.5 12.6 49 8.50 2.40 7.10 0.90 4.6 0.80 2.3 0.30 1.90 0.30 KR-05-213 57.1 111.0 12.0 46 8.30 2.20 6.90 0.90 4.4 0.80 2.1 0.30 1.80 0.30 KR-05-110 51.4 98.5 10.8 41 6.80 1.40 5.40 0.80 4.1 0.80 2.4 0.30 2.20 0.30 KR-05-175 58.3 111.0 11.8 44 7.40 1.40 5.90 0.80 4.5 0.80 2.4 0.30 2.20 0.30 CL06-35A 38.0 75.0 8.5 32 5.78 1.23 5.46 0.74 4.0 0.82 2.4 0.34 2.08 0.33 CL06-36A 68.0 133.0 15.0 56 9.91 1.81 8.21 0.94 4.5 0.82 2.5 0.34 2.08 0.33 CL06-37A 85.1 162.0 17.9 64 10.70 1.80 8.78 0.98 4.6 0.85 2.5 0.35 2.13 0.34 CL06-39A 75.4 143.0 16.0 57 9.15 1.57 7.46 0.86 4.1 0.75 2.2 0.29 1.79 0.28 CL06-40A 76.9 148.5 16.8 62 10.50 1.91 8.69 1.00 4.8 0.88 2.6 0.35 2.13 0.34 KR-05-113 49.3 98.0 10.6 41 7.20 1.20 5.40 0.70 3.8 0.70 2.0 0.30 1.80 0.30 KR-05-126 10.4 20.0 2.3 9 2.40 0.60 2.30 0.40 2.1 0.30 0.9 0.10 0.70 0.10 KR-05-130 41.9 78.9 8.4 32 5.40 1.20 4.30 0.60 3.2 0.60 1.7 0.20 1.60 0.20 KR-05-136 40.2 75.4 7.9 30 5.10 1.00 4.10 0.60 3.8 0.70 2.1 0.30 2.00 0.30 KR-05-47b 60.4 111.5 12.8 45 7.50 1.40 5.80 0.70 3.3 0.60 1.6 0.20 1.50 0.20 KR-05-59 35.3 69.4 7.8 28 5.30 1.00 4.50 0.60 3.7 0.70 2.1 0.30 1.90 0.30 KR-05-62 79.7 152.5 17.4 61 10.50 1.80 8.20 1.00 4.8 0.90 2.4 0.30 2.20 0.30 KR-05-71 37.8 74.2 8.3 29 4.90 1.10 4.20 0.50 2.9 0.60 1.6 0.20 1.40 0.20 KR-05-112 44.3 83.4 9.0 35 5.70 1.60 4.70 0.60 3.5 0.60 1.8 0.20 1.60 0.20 KR-05-68 43.7 85.2 9.8 35 6.10 1.00 4.70 0.60 3.3 0.60 1.8 0.20 1.50 0.20 KR-05-76 40.8 80.2 9.0 32 5.40 1.10 4.20 0.50 2.6 0.50 1.4 0.20 1.30 0.20 KR-05-77 45.8 101.0 13.0 51 9.30 1.60 6.50 0.80 3.6 0.70 2.0 0.30 1.70 0.20 CL06-33A 44.6 86.9 10.0 38 6.69 1.38 5.55 0.66 3.3 0.62 1.8 0.24 1.55 0.25 CL06-34A 51.9 106.0 11.7 43 7.50 1.37 6.29 0.72 3.3 0.58 1.6 0.22 1.29 0.20 KR-05-97b 62.4 128.0 14.4 58 9.90 1.90 6.90 0.90 4.3 0.80 2.2 0.30 1.90 0.30 CL06-38A 42.1 81.1 9.0 33 5.96 0.89 5.19 0.63 3.1 0.55 1.7 0.24 1.39 0.21 KR-05-194 58.4 108.5 11.7 43 7.20 1.20 5.90 0.80 4.1 0.70 2.0 0.30 1.80 0.20 KR-05-198 57.8 104.5 12.3 48 8.80 1.40 7.60 1.20 6.7 1.30 3.7 0.50 3.20 0.50 734 Sample La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Early to 'mid'-Cretaceous magmatism (in the Cassiar Terrane) KR-07-01 46.3 85.7 9.2 32 5.38 0.81 4.54 0.54 2.5 0.43 1.2 0.17 0.99 0.15 KR-07-02 33.4 62.8 6.7 23 4.13 0.56 3.58 0.44 2.0 0.34 0.9 0.13 0.77 0.12 KR-07-16 44.1 78.6 8.4 29 5.43 0.90 4.66 0.70 4.0 0.82 2.6 0.39 2.50 0.37 KR-07-17 41.1 79.5 8.3 30 5.50 0.71 4.78 0.56 2.5 0.42 1.2 0.15 1.02 0.15 KR-07-18 70.3 145.5 15.4 54 9.83 0.78 7.40 0.77 2.9 0.44 1.2 0.14 0.93 0.13 KR-07-22 41.8 78.9 8.2 28 5.98 1.00 5.37 0.74 4.0 0.70 2.1 0.28 1.81 0.25 07M-152 36.3 70.0 7.4 27 5.40 1.19 4.63 0.62 3.1 0.56 1.6 0.21 1.38 0.20 07M-153 34.4 67.8 7.2 27 4.94 1.11 4.45 0.60 3.1 0.55 1.7 0.22 1.44 0.20 KR-07-05 24.3 46.6 5.2 19 3.75 0.62 3.41 0.47 2.2 0.38 1.0 0.15 0.87 0.13 KR-07-07 49.2 89.8 9.7 35 5.88 1.24 5.37 0.70 3.9 0.78 2.4 0.36 2.24 0.36 KR-07-08 40.8 77.1 8.5 31 5.44 1.12 5.25 0.73 4.1 0.84 2.5 0.38 2.37 0.37 KR-07-09 43.8 82.3 8.9 31 5.45 0.97 5.02 0.63 3.2 0.63 1.8 0.28 1.69 0.26 KR-07-10 34.7 66.8 7.6 27 4.99 0.83 4.33 0.56 3.0 0.60 1.9 0.29 1.90 0.32 KR-07-11 20.3 38.4 4.2 15 3.14 1.03 2.85 0.41 2.2 0.39 1.1 0.16 0.99 0.16 KR-07-19 62.5 114.0 12.0 43 7.69 1.56 6.88 0.86 4.6 0.85 2.6 0.34 2.36 0.34 KR-07-20 35.5 62.4 6.1 21 3.43 1.02 2.86 0.29 1.2 0.20 0.6 0.08 0.55 0.07 KR-07-24 42.0 75.3 7.6 26 4.53 1.02 3.72 0.45 2.0 0.33 1.0 0.12 0.82 0.12 KR-07-06 96.1 166.5 17.8 60 9.12 1.59 8.10 0.97 4.8 0.92 2.9 0.41 2.61 0.41 07M-154 48.9 99.3 10.6 39 8.29 0.54 8.07 1.25 7.1 1.35 3.9 0.54 3.53 0.48 KR-07-03 18.4 39.1 4.1 14 3.54 0.32 3.60 0.66 4.7 0.92 3.2 0.47 3.59 0.49 KR-07-04 17.4 42.7 4.8 15 3.19 0.03 2.88 0.44 2.6 0.56 2.0 0.40 3.38 0.64 Late Cretaceous magmatism (in the Cassiar Terrane) KR-07-21 82.5 147.0 14.3 46 7.83 1.42 6.92 0.87 4.2 0.76 2.1 0.28 1.83 0.27 KR-07-15 34.6 64.3 6.5 21 3.93 0.54 3.75 0.58 3.2 0.58 1.7 0.26 1.54 0.23 KR-07-23 55.2 106.0 10.6 36 6.62 0.55 5.80 0.72 3.4 0.52 1.5 0.18 1.07 0.14 Latest Cretaceous to Eocene magmatism (in the Cassiar Terrane) KR-07-12 86.9 169.5 18.2 63 11.60 1.18 10.70 1.59 9.2 1.84 5.6 0.90 5.56 0.85 KR-07-14 58.9 118.5 13.1 45 8.54 0.63 8.07 1.15 6.6 1.31 4.0 0.62 4.00 0.60 735 e) Calculated geochemical parameters used in text. Sample A/CNK A/NK Fe2O3/FeO MALI crn (wt.%) 10000* Ga/Al Rb/Sr SmCN EuCN GdCN Eu/Eu* LaCN/YbCN Early to 'mid'-Cretaceous magmatism (in the Selwyn Basin) KR-07-13 1.13 1.26 0.04 7.10 2.73 3.18 2.05 27.89 7.97 15.84 0.38 36.82 07M-151 1.26 1.81 0.14 2.71 2.57 2.32 0.55 38.41 18.83 26.30 0.59 12.94 07M-150 1.21 2.04 0.20 1.05 1.32 2.08 0.36 31.42 17.10 21.38 0.66 10.49 03M-216 1.11 1.41 0.20 4.70 1.89 2.32 0.43 57.82 28.31 36.96 0.61 26.60 03M-217 1.09 1.47 0.35 3.95 1.16 2.18 0.38 39.52 21.41 24.79 0.68 28.98 03M-218 1.05 1.39 0.41 4.41 0.83 2.07 0.45 36.25 21.41 23.33 0.74 24.25 KR-05-191 1.07 1.32 0.10 5.38 1.29 1.78 0.67 36.25 21.41 21.38 0.77 17.39 KR-05-196 1.06 1.41 0.33 4.32 0.73 1.87 0.51 40.83 19.69 24.79 0.62 23.89 KR-05-22 1.07 1.37 0.60 4.66 0.68 2.18 0.53 35.60 17.97 22.84 0.63 18.88 KR-05-26 1.04 1.19 0.41 6.49 1.21 2.43 1.29 50.63 17.97 34.04 0.43 14.53 KR-05-157 1.05 2.13 0.15 -1.66 0.00 1.81 0.18 51.29 33.48 30.14 0.85 16.43 KR-05-158 0.99 2.24 0.18 -2.55 0.00 1.80 0.19 46.06 30.03 28.68 0.83 14.70 KR-05-08 1.18 1.93 0.11 1.55 0.52 2.22 0.25 23.84 16.24 16.03 0.83 13.63 KR-05-10 1.17 1.87 0.36 1.93 0.29 2.24 0.10 23.84 17.97 16.03 0.92 11.93 KR-05-43 1.14 1.91 0.59 1.28 0.08 2.16 0.20 23.18 16.24 16.52 0.83 16.16 03M-214 1.16 1.35 0.03 5.73 2.64 2.10 0.84 28.41 14.52 17.98 0.64 16.14 03M-215 1.09 1.27 0.09 6.18 2.23 2.52 1.44 42.14 16.24 26.25 0.49 22.76 03M-219 1.14 1.29 0.13 6.16 2.81 2.54 1.99 35.60 12.79 20.41 0.47 27.54 KR-05-164 1.00 1.08 0.10 8.04 1.57 2.14 2.77 31.68 14.52 20.41 0.57 7.29 KR-05-32 1.05 1.18 0.14 6.29 1.84 2.01 2.04 43.44 9.34 28.68 0.26 19.33 03M-207 1.20 1.54 0.04 4.35 2.52 2.30 0.55 40.18 19.69 24.30 0.63 22.09 03M-220 1.20 1.70 0.19 3.12 2.04 2.36 0.43 34.95 19.69 22.36 0.70 18.40 KR-05-143 1.11 1.40 0.10 5.05 1.70 2.08 0.89 38.87 17.97 22.36 0.61 13.63 KR-05-148 1.01 1.15 0.06 7.35 1.52 1.92 0.75 41.48 24.86 23.33 0.80 15.58 KR-05-207 1.13 1.36 0.12 5.55 1.82 1.62 1.10 23.84 11.07 14.09 0.60 13.42 KR-05-208 1.15 1.36 0.01 5.58 2.11 2.20 2.05 28.41 11.07 18.95 0.48 8.61 KR-05-210 1.20 1.36 0.13 5.95 2.89 1.96 2.33 25.14 5.90 15.55 0.30 8.30 KR-05-215 1.11 1.34 0.03 5.59 1.68 2.14 1.01 34.95 16.24 23.33 0.57 14.56 KR-05-204 0.88 1.51 0.05 0.23 0.00 1.86 0.62 53.25 33.48 32.09 0.81 19.27 KR-05-212 0.83 1.92 0.14 -3.25 0.00 1.82 0.18 54.56 40.38 33.55 0.94 23.29 KR-05-213 0.88 1.89 0.19 -2.22 0.00 1.96 0.19 53.25 36.93 32.58 0.89 22.75 KR-05-110 1.11 1.39 0.08 4.86 1.86 1.96 0.72 43.44 23.14 25.28 0.70 16.76 KR-05-175 1.11 1.38 0.00 5.06 1.82 2.09 0.69 47.37 23.14 27.71 0.64 19.01 CL06-35A 1.15 2.07 0.14 0.34 0.63 2.11 0.40 36.78 20.21 25.57 0.66 13.10 CL06-36A 0.87 1.22 0.41 3.74 0.00 2.25 0.58 63.77 30.21 38.95 0.61 23.45 CL06-37A 0.86 1.23 0.17 3.54 0.00 2.21 0.61 68.93 30.03 41.73 0.56 28.66 CL06-39A 1.01 1.41 0.12 3.64 0.21 2.18 0.52 58.80 26.07 35.30 0.57 30.21 CL06-40A 0.93 1.39 0.23 2.81 0.00 2.30 0.49 67.63 31.93 41.29 0.60 25.90 KR-05-113 1.05 1.34 0.14 4.82 1.08 2.04 0.68 46.06 19.69 25.28 0.58 19.65 KR-05-126 1.11 1.29 0.03 6.31 2.17 2.76 1.50 14.69 9.34 10.19 0.76 10.66 KR-05-130 1.13 1.77 0.13 2.06 0.68 2.09 0.41 34.29 19.69 19.92 0.75 18.78 KR-05-136 1.14 1.66 0.19 2.93 1.36 2.13 0.57 32.33 16.24 18.95 0.66 14.42 KR-05-47b 1.13 1.53 0.16 3.89 1.81 2.42 0.50 48.02 23.14 27.22 0.64 28.88 KR-05-59 1.12 1.66 0.06 2.63 1.10 2.40 0.64 33.64 16.24 20.90 0.61 13.33 KR-05-62 0.84 1.24 0.09 2.89 0.00 2.37 0.54 67.63 30.03 38.90 0.59 25.99 KR-05-71 1.11 1.66 0.06 2.62 1.07 2.51 0.71 31.03 17.97 19.44 0.73 19.37 KR-05-112 1.00 2.22 0.10 -1.89 0.00 2.11 0.24 36.25 26.59 21.87 0.94 19.86 KR-05-68 1.14 1.69 0.08 2.69 1.59 2.44 0.76 38.87 16.24 21.87 0.56 20.90 KR-05-76 1.08 2.34 3.04 -1.59 0.00 2.50 0.65 34.29 17.97 19.44 0.70 22.51 KR-05-77 0.75 1.51 0.21 -1.60 0.00 2.54 0.51 59.78 26.59 30.63 0.62 19.32 CL06-33A 1.01 1.41 0.23 3.64 0.00 2.14 0.47 42.73 22.79 26.01 0.68 20.64 CL06-34A 1.06 1.28 0.15 5.65 1.76 2.17 0.87 48.02 22.62 29.61 0.60 28.86 KR-05-97b 0.86 1.15 0.44 4.78 0.00 2.13 0.66 63.71 31.76 32.58 0.70 23.56 CL06-38A 1.11 1.42 0.16 4.65 1.72 2.35 1.18 37.95 14.34 24.26 0.47 21.73 KR-05-194 1.10 1.23 0.15 6.64 2.41 2.15 1.37 46.06 19.69 27.71 0.55 23.27 736 Sample A/CNK A/NK Fe2O3/FeO MALI crn (wt.%) 10000* Ga/Al Rb/Sr SmCN EuCN GdCN Eu/Eu* LaCN/YbCN KR-05-198 1.08 1.30 0.09 5.70 1.71 2.38 1.55 56.52 23.14 35.98 0.51 12.96 Early to 'mid'-Cretaceous magmatism (in the Cassiar Terrane) KR-07-01 1.17 1.36 0.35 5.86 2.57 2.41 0.64 34.16 12.97 21.09 0.48 33.55 KR-07-02 1.19 1.35 0.26 6.16 2.80 2.62 1.17 25.99 8.66 16.42 0.42 31.11 KR-07-16 1.14 1.31 0.63 5.95 2.01 2.50 0.51 34.49 14.52 21.68 0.53 12.65 KR-07-17 1.14 1.29 0.32 6.20 2.46 2.42 0.96 34.95 11.24 22.26 0.40 28.90 KR-07-18 1.06 1.10 0.55 7.66 2.64 2.88 1.51 63.25 12.45 35.01 0.26 54.22 KR-07-22 1.13 1.24 0.07 6.85 2.32 2.15 0.50 38.08 16.24 25.13 0.52 16.57 07M-152 1.33 2.06 0.24 1.98 2.29 2.40 0.46 34.29 19.52 21.53 0.72 18.87 07M-153 1.21 1.69 0.12 3.31 1.99 2.25 0.39 31.29 18.14 20.65 0.71 17.14 KR-07-05 1.20 1.37 0.16 6.02 3.11 2.31 0.67 23.51 9.69 15.59 0.51 20.03 KR-07-07 1.14 1.41 0.18 4.98 2.05 2.07 0.61 37.43 20.38 25.13 0.66 15.75 KR-07-08 1.21 1.91 0.46 1.83 1.45 1.90 0.40 34.56 18.31 24.55 0.63 12.35 KR-07-09 1.14 1.42 0.13 5.00 1.94 1.96 0.61 34.62 15.72 23.43 0.55 18.59 KR-07-10 1.10 1.30 0.77 5.50 1.59 2.11 1.02 31.61 13.31 20.07 0.53 13.10 KR-07-11 1.12 1.32 6.25 2.17 2.37 0.68 19.52 16.76 12.87 1.06 14.71 KR-07-19 1.20 1.60 0.63 3.74 1.64 2.50 0.27 49.26 25.90 32.48 0.65 19.00 KR-07-20 1.18 1.55 0.61 4.24 1.76 2.26 0.14 21.42 16.59 12.92 1.00 46.30 KR-07-24 1.16 1.44 0.46 5.04 1.83 2.42 0.23 28.61 16.59 17.10 0.75 36.74 KR-07-06 1.08 1.30 0.61 5.82 1.44 2.04 0.36 58.61 26.41 38.42 0.56 26.41 07M-154 1.00 1.07 0.35 7.74 1.03 3.57 3.43 53.18 8.31 38.27 0.18 9.94 KR-07-03 1.07 1.14 0.09 7.42 1.93 2.36 4.64 22.14 4.52 16.52 0.24 3.68 KR-07-04 1.08 1.10 -0.01 8.15 1.85 3.04 285.00 19.85 13.01 3.69 Late Cretaceous magmatism (in the Cassiar Terrane) KR-07-21 1.07 1.21 0.71 6.54 1.61 2.65 0.43 50.18 23.48 32.67 0.58 32.34 KR-07-15 1.19 1.21 0.57 7.80 3.65 2.61 3.19 24.69 8.31 17.25 0.40 16.12 KR-07-23 1.26 1.30 0.88 6.82 4.07 3.65 3.39 42.27 8.48 27.22 0.25 37.00 Latest Cretaceous to Eocene magmatism (in the Cassiar Terrane) KR-07-12 1.06 1.14 0.15 7.30 2.02 3.74 3.39 74.82 19.34 51.07 0.31 11.21 KR-07-14 1.05 1.14 0.52 6.99 1.96 2.55 3.90 54.82 9.86 38.27 0.22 10.56 737 D.2 Accuracy analysis for 2006–2007 samples analyzed at ALS Chemex. (available for download at: http://hdl.handle.net/2429/44385) 738 D.3 Comparison between trace elements analyzed by inductively coupled plasma – mass spectrometry (ICP–MS) methods at ALS Chemex and trace element data determined by high resolution ICP–MS methods at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) for 2005 samples. (available for download at: http://hdl.handle.net/2429/44385) 739 D.4 Sliding normalization to the high-K calc-alkaline Yenchichi–Telabit Series (after Liégeois et al., 1998). (available for download at: http://hdl.handle.net/2429/44385) 740 Appendix E – Supporting Materials for Zircon Compositions 741 E.1 Zircon images and analysis spots. 742 196_a: 196_b: 196_c: 196_d: 196_e: 196_f: 196_g: Actual TE and HF analysis locations Requested TE and HF analyses Zones Inh (J3?) Inh (L4) Inh Inh S3 S 2 3 -S 1 8 S22-S23 S10 S 1 8 -S 1 9 J 3 -J 4 S19 -S2 0 S24 -S2 5 S1 9-S 20 S2 4-S 25 S 4 Inh (L 5? ) S4 S3 S18 J3 L5 L2-L3 S21-S22 J4 S22-S23 TE Hf TE Hf Hf Hf Hf TE TE Hf 196-Hf-1 196-TE-1 196-TE-2 196-TE-3 196-Hf-3 196-Hf-2 196-Hf-4 196-Hf-5 probe analysis for Hf 743 196_j: 196_k: 196_l: 196_m: 196_o: 196_h: 196_i: Actual TE and HF analysis locations Requested TE and HF analyses Zones In h ( J 3 ? ) S 2 2 S11-S12 S18 Inh Inh Inh(P1?) S14-S19 L4-S4 S3? S24-J4 S5? S19-S24 S12-S17 S1-S2 S3 TE TE TE Hf Hf TE 196-TE-5 196-TE-4 196-Hf-6 196-TE-6 196-HF-7 196-TE-7 note: 'i' and 'j' are very close together 744 196_p: 196_q: 196_r: 196_s: 196_t: 196_u: 196_v: Actual TE and HF analysis locations Requested TE and HF analyses Zones Inh Inh S22 S1 7-S 18 , S2 2-S 23 S13 Inh (S 13 ?) S 2 3 -J 3 S 1 9 -S 2 0 , S 2 4 -S 2 5 S 2 -S 3 L 3 -S 3 S1-S2 S10 S 2 4 S3TE TE TE 196-TE-8 196-TE-9 196-TE-10 745 196_w: 196_x: 196_y: 196_z: 196_aa: 196_bb: Actual TE and HF analysis locations Requested TE and HF analyses Zones ? J4 S13? S13 S24-J4 S10-P2 J3-J4 S17-S 22 746 08_a: 08_b: 08_c: 08_d: 08_e: 08_f: 08_g: Actual TE and HF analysis locations Requested TE and HF analyses Zones L2-S2 L3-S3 L3-S3 S14-S19 diss? S 2 -S 7 S 1 -S 6 S1-S6,S2-S7 S1-S6,S2-S7 S1-S6 S3-S8 S2-S7 S1-S6,S2-S7 S1-S6 S1-S6 S2-S7 Q 3 Inh Inh Inh ? ( S1 9-S 14 ) Inh? (J4?) L3 S19- S20 diss? Inh L3? TE TE TE Hf TE TE TE TE Hf Hf 08-TE-1 08-TE-2 08-Hf-1 08-TE-3 08-Hf-2 08-TE-5 08-TE-6 08-TE-7 08-Hf-3 08-TE-4 probe analysis for Hf 747 08_h: 08_i: 08_j: 08_k: 08_l: 08_m: 08_n: S5 S2 S2 P1-P2 S2-S3 Inh S14-S19 S2-S7 G1 G 1 -P 1 P1 P1-P2 S3 Actual TE and HF analysis locations Requested TE and HF analyses Zones S 1 7 -S 2 2 S2-S3, S6-S7 TE Hf TE TE Hf Hf 08-TE-8 08-Hf-4 08-TE-10 08-Hf-608-Hf-5 08-TE-9 748 08_o: 08_p: 08_q: 08_r: 08_s: 08_t: 08_u: S6 S20-P4 S6-S7 P3 Inh S25? L4-L5 S4-S9 ? L2 S2 S21-S22 S2-S3, S7-S8 S19-S24 S2-S7 S2-S3, S7-S8 J4 Actual TE and HF analysis locations Requested TE and HF analyses Zones TE TE08-TE-11 08-TE-12 749 08_w: 08_x: 08_y: S18-S23 S18-S23 G1-P1 S3 Actual TE and HF analysis locations Requested TE and HF analyses Zones 750 143_a: 143_b: 143_c: 143_d: 143_e: 143_f: 143_g: Actual TE and HF analysis locations Requested TE and HF analyses Zones S 5 -P 1 In h L3? S2 L4-L5 L3 P2-P3S4-S5 P1 S11-S12 D? L 4 Inh In h (P 1) Inh (P3) Q2-Q3 Inh? Inh L 4 -S 4 S 8 -S 1 3 S 5 G1 Hf TE Hf TE Hf TE TE Hf143-Hf-1 143-Hf-2 143-TE-2 143-Hf-3 143-TE-3 143-TE-4 143-Hf-4 143-TE-1 probe analysis for Hf 751 143_h: 143_i: 143_j: 143_k: 143_l: 143_m: 143_n: possible xenocryst Inh ? ( P1 ) Inh (P2) P2 S9-S10 S 9- P 2 P 2- P 3 P 2- P 3? S1 9- S2 4 In h In h ( L3 -S 3-S 8) S1 3-S 8? S8 -S 9 S4 -S 9 S1 8- S2 3 Actual TE and HF analysis locations Requested TE and HF analyses Zones Hf TE TE TE 143-TE-5 143-TE-6 143-Hf-5 143-TE-7 note: distance between 'i' and 'j' is fairly long 752 143_o: 143_p: 143_q: 143_r: 143_s: 143_t: 143_u: S9 -S1 0 L 4 S 9 -S 1 4 S10 L3-S3 S1 2- S7 -S 2 P2 -P 3 S 6 -S 2 -L 3 L 3 In h (Q 2 -Q 3 ) S 4 -S 5 In h ( Q 3 2 -Q 3 ) Inh (S13-S14) Q ? S4-S5-L5 Inh (S3-S8) Inh Actual TE and HF analysis locations Requested TE and HF analyses Zones TE HfTE TE TE TE 30 micron spot Hf TE TE 143-TE-8 143-TE-9 143-TE-11 143-TE-12 143-TE-13 143-TE-14 143-Hf-7 143-TE-10 143-Hf-6 753 143_v: 143_w: 143_x: 143_y: 143_z: S1 9- S2 4 S1 3 L4 S2 4 Q1? S21-S22-S 23 In h ? Inh? (S6) Inh L1-S1 S 5 -L 5 xenocryst? Actual TE and HF analysis locations Requested TE and HF analyses Zones 754 S3-L3 Inh? P1-G 1 L 4 In h ? Inh? Inh? L4 L2-S 2L3-S 3-S 8 G1 S7-S12? ? monazite monazite 110_a: 110_b: 110_c: 110_d: 110_e: 110_f: 110_g: Actual TE and HF analysis locations Requested TE and HF analyses Zones TE TE Hf 110-TE-1 110-TE-2 110-Hf-1 probe analysis for Hf 755 ?L4-L5 S4-S5 ? S3 S5-P1 S4-S9 S4-S5 ? L3 S9 S6-S7, S11-S12 110_h: 110_i: 110_j: 110_k: 110_l: 110_m: 110_n: Actual TE and HF analysis locations Requested TE and HF analyses Zones L3 TE TE Hf110-Hf-2 110-TE-4 110-TE-5 756 Actual TE and HF analysis locations Requested TE and HF analyses Zones S3-S8 L3 monazite 110_o: 110_p: 757 175_a: 175_b: 175_c: 175_d: 175_e: 175_f: 175_g: Actual TE and HF analysis locations Requested TE and HF analyses Zones Inh (P2-P1) S3-S4 Inh (S19-S20?) ? S 7 -S 1 2 P1 Inh? S2 L4-S4 S13 G1 G1 L3-L4 L 2 S4 S3 Inh (L1-L2, S1-S2?) TE TE TE TE TE TE Hf Hf 175-TE-1 175-TE-2 175-TE-3 probe analysis for Hf 175-TE-4 175-Hf-2 175-TE-5 175-TE-6 175-Hf-1 758 175_h: 175_i: 175_j: 175_k: 175_l: 175_n: S 7 -S 8 S3Inh? Inh S9 ? Inh? S5-P1? S8-S9 P1 S3 L5 Actual TE and HF analysis locations Requested TE and HF analyses Zones 759 Actual TE and HF analysis locations Requested TE and HF analyses Zones 36_1_a: 36_1_b: 36_1_c: 36_1_d: 36_1_e: 36_1_g: 36_1_h: S 2 4 -J 4 S 2 0 -S 2 5 S 2 2 S 2 3 In h In h ? L 4 -L 5 , S 4 -S 5 S 2 4 -J 4 ? S 2 4 S 1 7 -S 2 2 S 1 8 S23-J3 S 1 8 S4 -S 5 In h S 1 8 ? Hf TE Hf 36_1-Hf-1 probe analysis for Hf 36_1-Hf-2 36_1-TE-1 760 36_1_i: 36_1_j: 36_1_k: 36_1_l: 36_1_m: 36_1_n: 36_1_o: S 2 -S 3 S2 2- S2 3 S23-S24 S5-P1 J2-J4 S 2 2 -S 2 3 P2-P3 P5 S2 4 S2 2-S 23 G2 Inh ? S2/Inh? Actual TE and HF analysis locations Requested TE and HF analyses Zones TE TE 36_1-TE-2 36_1-TE-3 761 36_1_q: 36_1_r: 36_1_s: 36_1_t: 36_1_u: 36_1_v: 36_1_w: P2-P3 S 2 3 -S 2 4 S 2 3 -S 2 4 S 2 1 S 1 9 -S 2 4 S 1 8 -S 2 3 S 8 S 1 7 -S 2 2 S 4 J 4 S19-S24 S 2 4 S 1 8 S 2 5S 2 3 P1-P4 ? d is s ? Actual TE and HF analysis locations Requested TE and HF analyses Zones TE TE 36_1-TE-4 36_1-TE-5 762 36_1_x: 36_1_y: 36_1_z: 36_1_aa: 36_1_bb: S 5 -P 1 P1 In h ? S5 -P 1 S 1 8 -S 2 3 S18-S23 S 1 7 -S 2 2 S 1 8 -S 2 3 S2 1- S2 2 S2 4- J4 (In h? ) S 2 4 In h ? S 2 1 Actual TE and HF analysis locations Requested TE and HF analyses Zones Hf36_1-Hf-3 763 68_a: 68_b: 68_d: 68_e: 68_f: 68_g: 68_c: Inh S 3 -L 3 S 5 J 4 -J 5 ? S2 P1 L2 S3-L3 L5-G1 Q1-Q2 L2 S2 S 1 2 -S 1 7 S 1 6 -S 1 7 L 3 J4 Inh S2-L2 Actual TE and Hf analysis locations Requested TE and Hf analyses Zones TE TE Hf TE Hf 68-TE-1 68-TE-2 68-Hf-1 probe analysis for Hf 68-TE-3 68-Hf-2 764 68_h: 68_i: 68_j: 68_k: 68_l: 68_m: 68_n: L3 L5-G1 S24 P1-G1 S2 S1-S6 S2-L2 L1-L2 P1-G1S3-S4 G1 S1 S21 Inh L 3 -S 3 P1-P2 S3 P1 L2-S2 S4 S3-S8 S16-S21 Actual TE and HF analysis locations Requested TE and HF analyses Zones TE Hf TE TE Hf Hf TE TE 68-TE-4 68-TE-5 68-TE-6 68-TE-7 68-Hf-4 68-Hf-5 68-TE-8 68-Hf-3 765 68_o: 68_p: 68_q: 68_r: 68_s: 68_t: 68_u: Inh S 5 I n h S 3 S 2 -S 3 , L 2 -L 3 G1 L1 L1 -S1 Inh (L 1) Inh (L4) Inh Inh (G2) S 3 L5 L4-S4 S3 S2 S 2 -S 3 L 5 L1 -L2 Inh (S1) Actual TE and HF analysis locations Requested TE and HF analyses Zones TE Hf 68-Hf-6 68-TE-9 766 68_v: 68_w: 68_x: 68_y: 68_z: Inh? (L2-S2) L2 S24-J4 S 2 1 -J 1 S 2 S25? L 5 L 2 L 2 -L 4 S24 S2 4-J 4 G 2 L 3 Actual TE and HF analysis locations Requested TE and HF analyses Zones TE TE 68-TE-10 68-TE-11 767 76_a: 76_b: 76_c: 76_d: 76_e: In h S 5 -S 1 0S 4 -S 5 G 1 Inh In h ( S 3 ) G 1 L4S5-P1 S3 S2 Inh? (Q2?) Actual TE and HF analysis locations Requested TE and HF analyses Zones S 4 Inh ? ( Q2 -S6 ? TE TE TE Hf TE 76-TE-1 76-Hf-1 76-TE-3 76-TE-4 76-TE-2 probe analysis for Hf note: very large gap between 'a' and 'b' grains note: one cruddy grain between 'b' and 'c' grains 768 Actual TE and HF analysis locations Requested TE and HF analyses Zones 97b_a: 97b_b: 97b_c: 97b_d: 97b_e: 97b_f: 97b_h: S 19-S 24 S 19-S 24 Inh (L2?) L5-G 1 Inh diss? L1 S 2 0 -S 2 5 S 2 0 -S 2 5 S 2 3 S 1 8 S 8 -S 9 S 2 3 -J 3S 2 4 -J 4 S 2 4 -J 4 S 1 8 -S 2 3S1 0 S 1 7 J 4 S 1 9 -S 2 4 S 2 5 -J 5 S 2 4S 2 3 S 1 7 -S 1 2 G 2 S 2 5 -J 5 L 3 -S 3 S 5 S 2 3 S 1 8 -S 2 3 S 1 8 -S 2 3 TE TE TE TE TE TE TE TE Hf TE TE TE TE Hf Hf Hf Hf Hf 97b-Hf-1 97b-TE-1 97b-TE-2 97b-TE-3 97b-TE-4 97b-TE-5 97b-Hf-4 97b-TE-7 97b-TE-8 97b-TE-9 97b-Hf-5 97b-Hf-6 97b-TE-12 97b-TE-11 97b-TE-10 97b-TE-6 97b-Hf-2 97b-Hf-3 probe analysis for Hf note: cruddy grain between 'f' and 'h' note: cruddy grain between 'h' and 'i' 769 97b_j: 97b_k: ? 97b_l: 97b_m: 97b_o: 97b_p: S 2 3 S 2 4 S 2 4 S 2 4 -S 2 5 S 1 2 S 1 9 -S 2 0 ? S 1 8 ? S 24 S 24 S25-J5 S 2 5 S 1 8 S 1 3 J 5 S 1 9 -S 2 4 In h J 1 -J 2 J3 S17-S22 97b_i: S 2 5 -J 5 S 2 4 S 2 1 -S 2 2 S 2 0 -S 2 5 S 2 3 Actual TE and HF analysis locations Requested TE and HF analyses Zones Hf Hf Hf TE TE TE TE TE 97b-TE-13 97b-Hf-8 97b-Hf-7 97b-TE-14 97b-TE-15 note: cruddy grain between 'm' and 'n' 97b-TE-17 97b-TE-16 97b-Hf-9 770 Actual TE and HF analysis locations Requested TE and HF analyses Zones 97b_t: 97b_u: 97b_v: 97b_x: 97b_y: S 1 8 -S 2 3 S 1 4 -S 1 9 S 2 4 -J 4 J 3 J 4 S 2 4 J 4 J 4 S 2 2 S 2 4 -J 4 J5 S22 P1-P2 97b_q: 97b_r: S 1 3 -S 1 5 J 2 S 2 2 -J 2 S 2 5 S 1 9 -S 2 4 S 2 4 -J 4 S 1 9 -S 2 4 S 1 7 J 3 TE TE TE 97b-TE-18 97b-TE-19 note: several cruddy grains between 'q' and 'r' 771 97b_z: J 4 J 4 J 2 J 3 J 3 J 3 J 3 Actual TE and HF analysis locations Requested TE and HF analyses Zones 772 E.2 Complete zircon trace element analyses. (available for download at: http://hdl.handle.net/2429/44385) 773 E.3 Complete zircon Lu–Hf isotope analyses. (available for download at: http://hdl.handle.net/2429/44385) 774 E.4 Details for zircon Lu–Hf isotope calculations. 775 Details for zircon Lu–Hf isotope calculations* Included data files: In the provided Summary Excel files, the “RESULTS” sheet summarizes the relevant data for each analysis that might be included in a paper. The “supp. data” sheet includes extra data or intermediate calculation values that might be included as an Appendix or online supplementary data table. The “PL stats” tab summarizes the data for the Plesovice standard for the given day. Plots and calculations for each analysis can be viewed in each individual analysis’s worksheet in the MCTOOL Excel files. The MCTOOL files also contain a “MCTOOL-Hf” tab that lists information on the instrument parameters used and Neptune multicollector’s cup configuration. If the client requests, the raw instrument output files (*.EXP) for each analyses can be provided. The relevant data from the files are already imported into the MCTOOL Excel files. Hafnium isotope calculations (600 laser shots): Before proceeding with the Hf isotopic calculations, all raw signals are blank corrected using the mean of ~20-25 pts recorded as the gas blank baseline before ablation. A user-selected interval of ~50 data pts (~50 seconds of data since the integration time is ~1 data point per second) covering the sample transient peak (~60 seconds of ablation) is used for calculation of the Hf ratio with mass bias correction by exponential law. A mass bias factor β is calculated for Yb and Hf. The measured (meas) signals are the blank- corrected signals from the instrument, while the corrected values (corr) are the expected natural values for these isotopic ratios. ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ ⎥⎥ ⎥⎥ ⎥ ⎦ ⎤ ⎢⎢ ⎢⎢ ⎢ ⎣ ⎡ ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ = )(171 )(173 171 173 171 173 ln ln )( Yb Yb meas corr M M Yb Yb Yb Yb Ybβ ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ ⎥⎥ ⎥⎥ ⎥ ⎦ ⎤ ⎢⎢ ⎢⎢ ⎢ ⎣ ⎡ ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ = )(177 )(179 177 179 177 179 ln ln )( Hf Hf meas corr M M Hf Hf Hf Hf Hfβ (Eq. 1 and 2) Where: corrected value for 179Hf/177Hf = 0.7325 (Patchett and Tatsumoto, Contrib Mineral Petrol, 1980 75, 263-267) corrected values for 173Yb/171Yb = 1.1301 and 176Yb/171Yb = 0.7938 (Segal et al, J. Anal. At. Spectrom., 2003, 18, 1217–1223) M = atomic mass of the isotope 776 The signal measured for the 176 mass is a combination of 176Hf, 176Yb and 176Lu. In order to determine 176Hf, the interferences from Yb and Lu must be subtracted. The amount of Yb and Lu contributing to the 176 signal is calculated by: )( )(176 )(173 173 176 173176 Yb Yb Yb corr measmeas M M Yb YbYbYb β ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×= (Eq. 3) )( )(176 )(175 175 176 175176 Yb Lu Lu corr measmeas M M Lu LuLuLu β ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×= (Eq. 4) measmeasmeasmeas LuYbLuYbHfHf 176176176176 )( −−++= (Eq. 5) Where: corrected value for 176Lu/175Lu = 0.2656 (Chu et al, J. Anal. At. Spectrom., 2002, 17, 1567–1574) A mass bias factor cannot be calculated for Lu since it only has two isotopes, so β(Yb) is used to approximate β(Lu) in Eq. 4. This approximation is assumed to be acceptable since (1) Lu is chemically similar to Yb and (2) the contribution of 176Lu to the total 176 signal is relatively small. Finally, the 176Hf/177Hf is corrected for mass bias as follows: )( )(177 )(176 177 176 177 176 Hf Hf Hf meascorr M M Hf Hf Hf Hf β ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛=⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ (Eq. 6) This calculation is done for each data point of an individual analysis. The final mass bias and interference corrected 176Hf/177Hf reported for the analysis is the mean of the 50 points in the user-selected interval. Outliers are removed by 2SD rejection. The error for the analysis is reported as: n SDSE ×= 22 (Eq. 7) Where: SE = standard error of the mean SD = standard deviation n = number of data points in selected interval For 176Yb/177Hf and 176Lu/177Hf, the mass bias correction is approximated using the mass bias factor for Hf, β(Hf): 777 Rebecca Lam, April 2009 Memorial University of Newfoundland *from Rebecca Lam, April 2009 (Memorial University of Newfoundland) )( )(177 )(176 177 176 177 176 Hf Hf Yb meascorr M M Hf Yb Hf Yb β ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛=⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ (Eq. 8) )( )(177 )(176 177 176 177 176 Hf Hf Lu meascorr M M Hf Lu Hf Lu β ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛=⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ (Eq. 9) No outlier rejection is done for these ratios since Yb and Lu are not necessarily homogenous in a zircon. To show that the mass bias correction is working, the mass bias corrected 178Hf/177Hf is calculated as follows: )( )(177 )(178 177 178 177 178 Hf Hf Hf meascorr M M Hf Hf Hf Hf β ⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛×⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛=⎟⎟⎠ ⎞ ⎜⎜⎝ ⎛ (Eq. 10) Reported literature values for 178Hf/177Hf typically range from 1.4671 to 1.4674. There doesn’t appear to be a universally accepted value, but for comparison, refer to Table 6 of Thirlwall and Anczkiewicz, International Journal of Mass Spectrometry, 235 (2004) 59–81. The Plesovice zircon is analyzed as a quality control standard. Data on this zircon can be found in Slama et al, Chemical Geology, 249 (2008) 1-35. Their reported value for 176Hf/177Hf is 0.282482 ± 0.000013 (2SD). 778 Appendix F – Supporting Materials for Reported Magmatic Fluid Compositions 779 F.1 Complete compositions, molar calculations, and activity calculations for all apatites used in interpretations and figures in Chapter 6. (available for download at: http://hdl.handle.net/2429/44385) 780 F.2 Complete compositions, molar calculations, and activity calculations for all biotites used in interpretations and figures in Chapter 6. (available for download at: http://hdl.handle.net/2429/44385) 781 F.3 Complete compositions, molar calculations, and activity calculations for all apatites used in interpretations and figures in Chapter 7. (available for download at: http://hdl.handle.net/2429/44385) 782 F.4 Complete compositions, molar calculations, and activity calculations for all biotites used in interpretations and figures in Chapter 7. (available for download at: http://hdl.handle.net/2429/44385) 783 Appendix G – Data Compilations 784 G.1 Geochronological data compilation for Cretaceous to Eocene magmatic rocks. (available for download at: http://hdl.handle.net/2429/44385) 785 G.2 Geochemical data compilation for Cretaceous to Recent magmatic rocks. (available for download at: http://hdl.handle.net/2429/44385) 786 G.3 Isotopic data compilation for Cretaceous to Recent magmatic rocks. (available for download at: http://hdl.handle.net/2429/44385) 787 G.4 Isotopic data compilation for northwest Laurentian basement rocks, and accreted pericratonic and oceanic terrane rocks. (available for download at: http://hdl.handle.net/2429/44385) 788
Thesis/Dissertation
eng
Geological Sciences
Vancouver : University of British Columbia Library
Attribution-NonCommercial-NoDerivs 3.0 Unported
Graduate
The timing, composition, and petrogenesis of syn- to post-accretionary magmatism in the northern Cordilleran miogeocline, eastern Yukon and southwestern Northwest Territories
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