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Detrital zircon geochronology and rift-related magmatism : central Mackenzie Mountains, Northwest Territories Leslie, Christopher Dean 2009

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 DETRITAL ZIRCON GEOCHRONOLOGY AND RIFT-RELATED MAGMATISM: CENTRAL MACKENZIE MOUNTAINS, NORTHWEST TERRITORIES   by  Christopher Dean Leslie  B.Sc. University of Alberta, 2006    A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF   MASTER OF SCIENCE   in   The Faculty of Graduate Studies  (Geological Sciences)    THE UNIVERSITY OF BRITISH COLUMBIA (Vancouver)  April 2009      © Christopher Dean Leslie, 2009     ii  ABSTRACT  Neoproterozoic to Cretaceous age strata in the Mackenzie Mountains of the northern Canadian Cordillera record many geological events that affected the western margin of Ancestral North America.  Two of these events that are the focus of this study are; (1) the development of a thick long-lived passive margin sequence of sedimentary rocks; and (2) continental rifting of this passive margin in Ordovician time that was accompanied with alkaline volcanism of the Marmot Formation. Sedimentary units from throughout the Mackenzie Mountains stratigraphy contain ubiquitous detrital zircons with U-Pb ages of 2800 – 2415 Ma and 2080 – 1700 Ma; these reflect a component of sediment derived from basement sources of western Laurentia. More surprising is the abundance of “anomalous” detrital zircons ages (e.g., 1800 – 1000 Ma) in these units that cannot be linked to known local sources.  The most likely source for detrital zircons of these ages are felsic igneous rocks now exposed in southern and eastern North America.  We therefore speculate that in the Neoproterozoic sediments were dispersed from these distal sources across the North American Craton.  Reworking and recycling of these older sediments is evident in most of the younger investigated strata.  In Carboniferous and younger strata there is also a major influence from sources exposed in the Canadian Arctic and possibly along Alaskan margin as evidenced by detrital zircons with Lower Paleozoic ages (e.g., 477 – 392 Ma). Continental extension of the western margin of Laurentia in Early to Middle Cambrian and again in Middle Ordovician time formed a rifted sedimentary basin termed the Misty Creek Embayment.  Alkaline mafic magmatism associated with the Middle  iii Ordovician rifting (460 – 444 Ma) event comprises massive mafic volcanic rocks (MFV and MFX suites), intrusive dykes and sills, volcaniclastic and epiclastic rocks and volcaniclastic filled diatremes (DVI suite). Petrographic studies together with whole rock, mineral chemistry and Nd isotopic studies indicate that the MFV and MFX suites were likely generated by small degrees of partial melting at the base of the subcontinental lithosphere with a minor asthenospheric input.  The DVI parental magmas were also generated by small degrees of partial melting primarily in metasomatized subcontinental lithosphere.                          iv TABLE OF CONTENTS  ABSTRACT…………………………………………………………………………………………….…...ii TABLE OF CONTENTS………………………………………………………………………………......iv LIST OF TABLES………………………………………………………………………………………....vii LIST OF FIGURES………………………………………………………………………………….…....viii ACKNOWLEDGEMENTS…………………………………………………………………………….…..x CO-AUTHORSHIP STATEMENT………………………..………………………………………..……xi CHAPTER 1. INTRODUCTION................................................................................................................ 1 1.1. ORIGINAL DESIGN OF STUDY AND NEW CONTEXT ............................................................... 2 1.2. OBJECTIVE OF CHAPTER TWO..................................................................................................... 3 1.3. OBJECTIVE OF CHAPTER THREE................................................................................................. 4 1.4. CONTENTS OF APPENDICES......................................................................................................... 6 1.5. REFERENCES CITED....................................................................................................................... 7 CHAPTER 2. DETRITAL ZIRCON GEOCHRONOLOGY AND PROVENANCE OF NEOPROTEROZOIC TO CRETACEOUS STRATA OF THE NORTHERN CANADIAN CORDILLERA: INSIGHTS FROM THE CENTRAL MACKENZIE MOUNTAINS, NORTHWEST TERRITORIES ............................................................................................................................................ 8 2.1. INTRODUCTION .............................................................................................................................. 9 2.2. MACKENZIE MOUNTAINS STRATIGRAPHY AND SAMPLING............................................. 12 2.2.1. Mackenzie Mountains Supergroup (MMSG)............................................................................. 13 2.2.2. Windermere Supergroup ........................................................................................................... 15 2.2.3. Early Cambrian Units............................................................................................................... 16 2.2.4. Mackenzie Platform and Selwyn Basin ..................................................................................... 17 2.2.5. Clastic Foredeep....................................................................................................................... 18 2.3. ANALYTICAL METHODS............................................................................................................. 19 2.4. RESULTS ......................................................................................................................................... 21 2.4.1. Katherine Group ....................................................................................................................... 21 2.4.2. Little Dal Group........................................................................................................................ 22 2.4.3. Keele Formation ....................................................................................................................... 22 2.4.4. Backbone Ranges Formation .................................................................................................... 25 2.4.5. Sekwi Formation ....................................................................................................................... 26 2.4.6. Franklin Mountain Formation .................................................................................................. 28 2.4.7. Heritage Trail Formation ......................................................................................................... 28 2.4.8. Clastic Foredeep....................................................................................................................... 29 2.4.9. Summary of Detrital Zircon Results.......................................................................................... 31 2.5. POTENTIAL DETRITAL ZIRCON SOURCES .............................................................................. 31 2.5.1. Archean and Proterozoic Grains .............................................................................................. 33 2.5.2. Paleozoic Grains....................................................................................................................... 37 2.6. INTERPRETATION OF DETRITAL ZIRCON AGES.................................................................... 38 2.6.1. Provenance of the Mackenzie Mountains Supergroup.............................................................. 41 2.6.2. Provenance of the Windermere Supergroup ............................................................................. 42 2.6.3. Provenance of the Backbone Ranges Formation ...................................................................... 44 2.6.4. Provenance of the Sekwi Formation ......................................................................................... 45 2.6.5. Provenance of the Franklin Mountain Formation .................................................................... 47 2.6.6. Provenance of the Tsichu Group............................................................................................... 47 2.6.7. Provenance of the Cretaceous clastic foredeep ........................................................................ 49 2.7. COMPARISONS WITH COEVAL STRATA ALONG THE MIOGEOCLINE .............................. 50 2.8. ANOMALOUS DETRITAL ZIRCON AGES .................................................................................. 58 2.8.1. Eastern North America sources? .............................................................................................. 58  v 2.8.2. Unknown proximal sources?..................................................................................................... 65 2.9. PALEOGEOGRAPHIC IMPLICATIONS ....................................................................................... 66 2.10. CONCLUSIONS............................................................................................................................. 69 2.11. REFERENCES CITED................................................................................................................... 74 CHAPTER 3. THE ORIGIN AND PETROGENESIS OF ORDOVICIAN AGE RIFT-RELATED ALKALINE MAGMATISM OF THE MISTY CREEK EMBAYMENT, CENTRAL MACKENZIE MOUNTAINS, NORTHWEST TERRITORIES..................................................................................... 83 3.1. INTRODUCTION ............................................................................................................................ 84 3.2. GEOLOGICAL SETTING ............................................................................................................... 86 3.3. VOLCANIC OCCURENCES........................................................................................................... 89 3.3.1. Marmot Formation volcanic rocks (MFV)................................................................................ 90 3.3.2. Marmot Formation volcanic xenoliths (MFX) .......................................................................... 91 3.3.3. Diatremes and associated volcaniclastic and intrusive rocks (DVI) ........................................ 92 3.3.4. Petrographic observations........................................................................................................ 92 3.3. SAMPLING METHODS .................................................................................................................. 94 3.4. ANALYTICAL RESULTS............................................................................................................... 94 3.4.1. Mineralogy and Mineral Chemistry.......................................................................................... 94 3.4.1.1. Clinopyroxene .................................................................................................................................... 95 3.4.1.2. Spinel ................................................................................................................................................. 96 3.4.1.3. Phlogopite-biotite ............................................................................................................................... 98 3.4.2. Whole Rock Geochemistry ........................................................................................................ 99 3.4.2.1 Alteration .......................................................................................................................................... 100 3.4.2.2. Major Elements ................................................................................................................................ 103 3.4.2.3. Trace Elements................................................................................................................................. 105 3.4.3. Isotopic and Dating Studies .................................................................................................... 107 3.4.3.1 Neodymium Isotopic Composition.................................................................................................... 108 3.4.3.2. 40Ar/39Ar phlogopite geochronology ................................................................................................ 108 3.5. INTERPRETATION OF MINERAL DATA.................................................................................. 111 3.5.1. Clinopyroxene ......................................................................................................................... 111 3.5.2. Spinel ...................................................................................................................................... 113 3.5.3. Phlogopite-biotite ................................................................................................................... 114 3.6. INTERPRETATION OF GEOCHEMISTRY................................................................................. 116 3.6.1. Crustal Contamination............................................................................................................ 116 3.6.2. Marmot Formation petrogenesis............................................................................................. 118 3.6.2.1. Fractional Crystallization ................................................................................................................. 119 3.6.2.2. Partial Melting.................................................................................................................................. 120 3.7. NATURE OF THE SOURCE ......................................................................................................... 126 3.7.1. Composition of the source....................................................................................................... 126 3.7.2. Depth of source ....................................................................................................................... 129 3.7.2. Source heterogeneities? .......................................................................................................... 130 3.7.3. Problems with source constraints ........................................................................................... 131 3.8. TIMING OF MAGMATISM .......................................................................................................... 132 3.9. COMPARSION WITH VOLCANISM FROM SIMILAR TECTONIC SETTINGS...................... 132 3.9.1. Comparison with rift related mafic volcanic suites................................................................. 132 3.9.2. Comparisons with mafic volcanic rocks in the Northern Cordillera ...................................... 138 3.10. A MODEL FOR THE EVOLUTION OF THE MARMOT FORMATION .................................. 139 3.11. CONCLUSIONS........................................................................................................................... 146 3.12. REFERENCES CITED................................................................................................................. 148 CHAPTER 4. CONCLUSIONS AND SUGGESTIONS FOR FUTURE WORK............................... 157 REFERENCES CITED.......................................................................................................................... 162 APPENDICES .......................................................................................................................................... 163 APPENDIX A. CHAPTER 3 ANALYTICAL METHODS............................................................................... 163 APPENDIX B.  CHAPTER 2 ANALYTICAL DATA..................................................................................... 175 APPENDIX C.  CHAPTER 3 ANALYTICAL DATA..................................................................................... 185  vi APPENDIX D.  MARMOT FORMATION PHOTO DOCUMENTATION .......................................................... 199 APPENDIX E.  VOLCANICLASTIC DIATREMES ....................................................................................... 206      vii LIST OF TABLES   Table 3.1. Geochemical comparisons with East African Rift volcanic rocks.  References are reported in text. ................................................................................................. 137 Table A.1. Mean values and duplicate analyses of standard BAS-1 .............................. 165 Table B.1. U-Pb geochronological data.......................................................................... 176 Table C.1. Microprobe analysis of clinopyroxene from sample 1508............................ 186 Table C.2. Microprobe analysis of spinel from sample 1502C ...................................... 187 Table C.3. Micoprobe analysis of phlogopite from selected DVI and MFV samples.... 188 Table C.4. Major and trace element data from the Marmot Formation and related occureces................................................................................................................. 189 Table C.5. Neodymium data from selected DVI, MFV and MFX occurences .............. 193 Table C.6. 40Ar/39Ar age data for phlogopite from selected DVI and MFV samples..... 194 Table E.1. Locations and descriptions of the Mackenzie Mountain diatremes and associated volcanic rocks........................................................................................ 223                               viii LIST OF FIGURES  Figure 2.1.  Map showing the general geology of the central Mackenzie Mountains with locations of detrital zircon samples........................................................................... 11 Figure 2.2.  Generalized stratigraphy of the central Mackenzie Mountains with approximate stratigraphic location of detrital zircon samples. ................................. 13 Figure 2.3. U-Pb detrital zircon relative age probability distribution curves and histogram plots for the Katherine Group, Little Dal Group, and the Keele Formation. ............ 24 Figure 2.4. U-Pb detrital zircon relative age probability distribution curves and histogram plots for the lower and upper members of the Backbone Ranges Formation and the Sekwi Formation....................................................................................................... 27 Figure 2.5. U-Pb detrital zircon relative age probability distribution curves and histogram plots for the Franklin Mountain Formation, Tsichu Group, and the Cretaceous clastic foredeep..................................................................................................................... 30 Figure 2.6. Tectonic framework of Precambrian basement provinces and Paleozoic granitoids of North America ..................................................................................... 32 Figure 2.7. U-Pb detrital zircon normalized age distribution curves for all sedimentary units investigated from the Mackenzie Mountains. .................................................. 40 Figure 2.8. U-Pb detrital zircon normalized age distributions curves for Windermere Supergroup equivalent strata from along the length of western North America. ..... 53 Figure 2.9. U-Pb detrital zircon normalized age distributions curves for Backbone Ranges Formation equivalent strata from along the length of western North America ........ 56 Figure 2.10. U-Pb detrital zircon normalized age distribution curves for our Mackenzie Mountains compilation compared with detrital zircon data from west Texas.......... 60 Figure 3.1. Map showing Lower Paleozoic paleogeography with location of the Misty Creek Embayment..................................................................................................... 87 Figure 3.2. Southwest to northeast schematic cross-section through the Misty Creek Embayment ............................................................................................................... 89 Figure 3.3. Simplified geology map of the study area showing locations of samples of the Marmot Formation and related intrusive and volcaniclastic rocks. .......................... 90 Figure 3.4. Plots of clinopyroxene chemistry from sample 1508 ..................................... 96 Figure 3.5. Plots of spinel chemisty from sample 1502.................................................... 97 Figure 3.6. Plots of phlogopite chemistry from samples 1494B, F and 1504B................ 99 Figure 3.7. Key compatible elements, compatible element ratios, and incompatible element ratios against the Al2O3/Na2O alteration index of. .................................... 102 Figure 3.8. Chondrite normalized rare earth element plots. ........................................... 104 Figure 3.9. Primitive mantle normalized incompatible element plots.. .......................... 106 Figure 3.10. Petrogenetic and tectonic discrimination diagrams.................................... 107 Figure 3.11. 40Ar-39Ar age spectra for phlogopites from selected samples from the Marmot Formation and related intrusive and volcaniclastic rocks. ........................ 110 Figure 3.12. Potential crustal contaminants. ................................................................... 118 Figure 3.13. Fractional crystalization constraints ........................................................... 120 Figure 3.14. Partial melting constraints. ......................................................................... 121 Figure 3.15. Marmot Formation geochemical models.................................................... 123 Figure 3.16. Mantle composition .................................................................................... 128 Figure 3.17. Geochemistry comparison diagrams with data from the East African Rift.135  ix Figure 3.18. Schematic model cross-section of the western margin of Laurentia through the lithosphere in Early Cambrian and Middle Ordovician. ................................... 145 Figure D.1. Field photographs of the Marmot Formation. ............................................. 201 Figure D.2. Photomicrographs and one field photograph showing the petrographic range of compositions of the MFV and MFX suites. ...................................................... 203 Figure D.3. Photomicrographs and one SEM image of important minerals used in microprobe and dating studies ................................................................................ 205 Figure E.1. Distribution of Mackenzie Mountain diatremes and associated volcanic rocks. Locations of unknown bodies. ................................................................................ 218 Figure E.2. Aerial view looking northeast of the extent of exposure of the Mountain Diatreme.................................................................................................................. 219 Figure E.3. Exposure of the Bits plug............................................................................. 219 Figure E.4. Exposure of the Monica Diatreme ............................................................... 220 Figure E.5. Outcrop image of the Mountain Diatreme central volcaniclastic breccia.... 220 Figure E.6. Photomicrograph of a phlogopite cored juvenile clast from the Mountain Diatreme.................................................................................................................. 221 Figure E.7. Photomicrograph of amoeboid shaped juvenile clast from the Tim diatreme. ................................................................................................................................. 221 Figure E.8. Photomicrograph of the Bits plug ................................................................ 222 Figure E.9. Photomicrograph of a phlogopite porphyritic tuff ....................................... 222            x ACKNOWLEDGEMENTS   Many people have provided various means of intellectual, emotional and financial support whom without I would not be given the opportunity to thank. First and foremost I thank my supervisor Jim Mortensen for not only supplying a never-ending pot of coffee but also for your guidance and your enthusiasm for this project making it unique and very enjoyable. I also thank Kelly Russell and Derek Thorkelson for their insight and for agreeing to be on my thesis committee.  Financial support in part was from a NSERC Discovery Grant held by Jim Mortensen. The Northwest Territories Geoscience Office (NTGO) provided additional financial and logistical support for my research and provided the opportunity to find my true passion in geology.  In particular this research would never have become a reality if it were not for Edith Martel.  She was instrumental in getting the project going and she was kind enough to hire me as a geologist for her Mackenzie Mountains bedrock mapping initiative. Discussions with other geoscientists at the NTGO were also fundamental in directing and making this research fun and exciting. These people include, Hamish Sandeman, Luke Ootes, and Thomas Hadlari.  I therefore express a thank you to the NTGO for everything they have done. I sincerely thank my beautiful girlfriend Karly.  You not only put up with an occasionally less than ideal boyfriend but most importantly you gave me the support I needed over the last few years.  I am also very grateful to my family for their continued understanding and undeniable support. To my colleagues I express a huge thank you for making coming to the department everyday an experience. There are too many colleagues to individually thank, however, insightful discussions with Luke Beranek, Curtis Brett, Tyler Ruks, Kevin Byrne, and Reza Tafti greatly improved this research.  I thank everyone that I have met in the academic community in the past three years.  Whether it was here at UBC, or at a conference, or at the Black Knight Pub in Yellowknife I thank you because I feel that science is not an individual practice but it represents a collaboration of thoughts and ideas you learn from others along the way.                 xi CO-AUTHORSHIP STATEMENT   This thesis comprises two separate chapters (Chapters 2 and 3) each written in manuscript form. Each manuscript has benefited from the help of co-authors and other colleagues who assisted in all aspects of the research.  CHAPTER TWO  I completed the written component and the drafting of figures as the senior author. This chapter and the paper will be submitted co-authored by Dr. J.K. Mortensen.  Many aspects of this manuscript benefited from discussions with J.K. Mortensen especially in data interpretation and the discussion section of the paper.  Samples for this study were collected during a regional bedrock-mapping project in the Mackenzie Mountains initiated by the Northwest Territories Geoscience Office during July and August of 2006 and 2007.  Edith Martel, Charlie Roots, and Steve Gordey provided the logistics and time to collect detrital zircon samples.  Edith Martel collected two samples and Darrel Long collected one sample.  I did all the sample preparation and mineral separations as well as the analytical work using laser ablation inductively-coupled plasma mass spectrometry (LA-ICP-MS) at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia (UBC).   xii  CHAPTER THREE   I completed the written component and the drafting of figures as the senior author. This chapter is co-authored by Dr. J.K. Mortensen and Dr. H.A. Sandeman.  Dr. J.K. Mortensen contributed to editing of the text and figures throughout and discussions with him sparked many new ideas that forms part of the discussion section. Samples for this study were collected during a regional bedrock-mapping project in the Mackenzie Mountains initiated by the Northwest Territories Geoscience office during July and August of 2006 and 2007.  Edith Martel, Charlie Roots, and Steve Gordey provided the logistics and time to collect samples for this study. Dr. H.A. Sandeman, through the Northwest Territories Geoscience Office, provided analytical funding for this project.  He also contributed to sample collection in the field and he helped with initial interpretation of whole-rock geochemical data.  I performed all of petrographic analyses as well as the sample preparation and mineral separation for this study.  Whole rock geochemical analyses were done at ALS Chemex Laboratories in North Vancouver, BC.  Thomas Ullrich of the PCIGR at the UBC performed the Ar-Ar age determinations, and Bruno Kieffer of the PCIGR conducted the Nd isotope analyses. Mati Raudsepp assisted with scanning electron microscopy and mineral chemistry at the Electron Microbeam/X-ray Diffraction Facility at UBC  1      CHAPTER 1 Introduction      2 1.1. ORIGINAL DESIGN OF STUDY AND NEW CONTEXT   The original objective for this Masters thesis was to attempt to constrain the nature of the crystalline basement beneath the northern Canadian Cordillera of northeastern British Columbia, eastern Yukon Territory and the southwestern Northwest Territories. The age and composition of this basement (the Nahanni domain) are poorly known because it is nowhere exposed.  The character of the Nahanni domain has been interpreted on the basis of regional potential field (gravity, magnetic) signatures (e.g., Cook and Erdmer 2005) and a single suite of basement xenoliths contained within a small diatreme in the southern Mackenzie Mountains (Jefferson and Parrish 1989; Mortensen and Colpron 1998).  The thesis was initially designed as a three-part study aimed at generating new constraints on the nature of the Nahanni domain basement. First, the ages of inherited zircon cores in Cretaceous granitoids were to be dated, and these ages were be used to provide information on the nature of the material that was melted.  Since it is not possible to determine whether the source rocks for the magmas were part of the Nahanni domain basement or the overlying supracrustal sequence, it was impartial to know the ages of detrital zircons present within the supracrustal column.  To this end, detrital zircons from representative samples from throughout the sedimentary sequence in the study area were to be dated.  Finally, widespread mantle-derived diatremes were to be examined and any contained basement xenoliths were to be investigated as direct samples of underlying basement.   3  Field components of the study were carried out as a part of a regional bedrock- mapping project in the Mackenzie Mountains initiated by the Northwest Territories Geoscience Office, which the author participated in for the entire 2006 and 2007 field seasons.  A total of 14 Cretaceous plutons were sampled for U-Pb dating studies, and nine representative samples from throughout the Neoproterozoic to Cretaceous sedimentary sequence were sampled for detrital zircon dating.  Many of the mafic diatremes and associated intrusive and extrusive rock units were examined and extensively sampled.  Despite a careful examination of all the main mafic diatremes in the study area, not a single basement xenolith was found.  The thesis project was therefore re-designed as two separate standalone projects, one focusing on the detrital zircon geochronology of the Neoproterozoic to Cretaceous sedimentary sequence and the other on the age and petrogenesis of mafic igneous rocks within the central Mackenzie Mountains.  Both of these studies provide constraints on the evolution of the northwestern Laurentian margin, but do not yield any direct new constraints on the nature of the Nahanni domain basement. 1.2. OBJECTIVE OF CHAPTER TWO   U-Pb age determinations from detrital zircons extracted from clastic sedimentary units provide a critical tool for understanding the depositional history and provenance of sediments throughout geologic time (e.g., Gehrels and Ross 1998).  Zircon is a robust mineral that remains relatively undisturbed during physical and chemical weathering and as such it usually retains its initial U-Pb crystallization age.   Intermediate to felsic igneous rocks are the primary sources for detrital zircons, although metamorphic zircons  4 can also be present, together with zircons that have been reworked from older sedimentary units.  Detrital zircon ages can provide evidence for the dominant ages of rock units that were eroded to provide the sediments and in some cases can point to very specific sources.  Strata exposed in the Mackenzie Mountains comprise sediments that were deposited from Neoproterozoic to Cretaceous time, corresponding to approximately 1 billion years of deposition.  These sediments were deposited mainly along the western margin of Ancestral North America, and record the breakup of the supercontinent Rodinia and the establishment of a passive margin that experienced many episodes of tectonic instability.  This chapter documents a detrital zircon dating investigation of sedimentary strata from throughout the Neoproterozoic to Cretaceous sedimentary sequence in the Mackenzie Mountains.  Speculations on the provenance of these sediments are based on these detrital zircon ages.  Sources proximal to the western margin of Laurentia are discussed and considered, together with speculation on more distal sources exposed on the eastern side of North America.  A version of this manuscript will be submitted for publication in a peer-reviewed international scientific journal co-authored by J.K. Mortensen. 1.3. OBJECTIVE OF CHAPTER THREE   Alkaline mafic volcanism is a significant product of continental extension especially when extension rates are sufficient enough to promote continental rifting. Petrographical, geochemical, and isotopical studies on these rift related alkaline igneous rocks can yield considerable information regarding the nature and depth of source and the  5 tectonic evolution of the area through which they were emplaced (e.g., Peircey et al. 2002).  The dynamics of melt generation and its ascent to the earth’s surface profoundly influences the petrologic character of the produced volcanic deposits. The western margin of Ancestral North America throughout its tectonic history has experienced many episodes of tectonic subsidence and uplift, which were direct consequences of continental extension.  The Mackenzie Mountains contains an Early Paleozoic rift basin termed the Misty Creek Embayment (MCE) that developed along the western margin of Laurentia as a result of reactivation of older structures initially formed during breakup of the supercontinent Rodinia.  The MCE records two Early Paleozoic rifting events in Cambrian and Ordovician time, and the latter event was responsible for the generation of widespread mafic alkaline volcanic and associated intrusive rocks. These alkaline mafic rocks comprise the extrusive Marmot Formation and associated intrusive diatremes, dykes and plugs observed along the eastern margin of the MCE. This chapter presents a comprehensive petrographical, geochemical, isotopic and geochronological study from these alkaline igneous rocks.  Speculations on igneous petrogenesis including the nature and depth of the source, the degrees of partial melting, and the affects of magma generation on overlying lithosphere are discussed in this chapter.  A version of this manuscript will be submitted for publication in a peer- reviewed international scientific journal, with J.K. Mortensen and H.A. Sandeman as co- authors.  6   1.4. CONTENTS OF APPENDICES   Analytical methods for Chapter 3 are reported in Appendix A.  Appendix B and Appendix C contains analytical data for Chapter 2 and Chapter 3 respectively. Appendix D is a collection of field and photomicrographs of the Marmot Formation and related intrusive rocks.  This appendix was created in order to present more pictorial documentation of the Marmot Formation than could be included in a paper intended for publication.  Appendix E is comprehensive overview of the mafic diatremes and associated volcaniclastic and intrusive rocks that was prepared for publication as Chapter 4.2 in Martel et al. (in prep; Geology of the central Mackenzie Mountains; Sekwi Mountain (105P), Mount Eduni (106A), and northwest part of Wrigley Lake (95M) map areas, Northwest Territories: NTGO Report xxxx).    7 1.5. REFERENCES CITED   Cook, F.A, and Erdmer, P., 2005, An 1800 km cross section of the lithosphere through the northwestern North American plate: lessons from 4.0 billion years of Earth's history: Canadian Journal of Earth Sciences, v. 42, no. 6, p. 1295-1311.  Gehrels, G., and Ross, G., 1998, Detrital zircon geochronology of Neoproterozoic to Permian miogeoclinal strata in British Columbia and Alberta: Canadian Journal of Earth Sciences, v. 35, p. 1380-1401.  Jefferson C.W. and Parrish, R., 1989, Late Proterozoic stratigraphy, U-Pb zircon ages and rift tectonics, Mackenzie Mountains, northwestern Canada: Canadian Journal of Earth Sciences, v. 26, p. 1784-1801.  Mortensen, J.K., and Colpron, M., 1998, Geochronological and geochemical studies of the Coates Lake Diatreme, southern Mackenzie Mountains, western N.W.T., in Cook, F.A., and Erdmer, P., eds., Slave - Northern Cordillera Lithospheric Evolution (SNORCLE) Transect and Cordilleran Tectonics Workshop Meetings: Lithoprobe Secretariat, The University of British Columbia, Vancouver, B.C., Lithoprobe Report 64, p. 278.  Piercey, S., Mortensen, J., Murphy, D., Paradis, S., and Creaser, R., 2002, Geochemistry and tectonic significance of alkalic mafic magmatism in the Yukon-Tanana terrane, Finlayson Lake region, Yukon: Canadian Journal of Earth Sciences, v. 39, p. 1729-1744.                     8          CHAPTER 2 Detrital zircon geochronology and provenance of Neoproterozoic to Cretaceous strata of the northern Canadian Cordillera: Insights from the central Mackenzie Mountains, Northwest Territories1                        1. A version of this chapter will be submitted for publication. Leslie, C.D., and Mortensen J.K. (2009); Detrital zircon geochronology and provenance of Neoproterozoic to Cretaceous strata of the northern Canadian Cordillera: Insights from the central Mackenzie Mountains, Northwest Territories.    9 2.1. INTRODUCTION   Strata exposed in the central Mackenzie Mountains of the northern Cordilleran orogen in northwestern Canada include sediments deposited from early Neoproterozic through late Paleozoic time (Blusson 1971; Aitken and Cook 1974; Roots and Martel 2008).  Collectively these strata record numerous major geologic events, including the Neoproteozoic breakup of the supercontinent Rodinia, the resultant continental glaciations, and the initiation of a long-lived passive margin built on the western edge of Ancestral North America.  This succession of mainly carbonates and siliciclastic rocks forms the northeastern extent of the Canadian Cordilleran miogeocline (inset fig. 2.1)  Substantial variations in sediment source regions during accumulation of the miogeoclinal strata have been inferred by previous workers, based on results of geochemical, isotopic and detrital zircon studies (e.g., Ross and Bowring 1990; Gehrels et al. 1995, 1996, 1999; Boghossian et al. 1996; Garzione et al. 1997; Rainbird et al. 1997; Gehrels and Ross 1998; Stewart et al. 2001; Ross and Villeneuve 2003; Ross et al. 2005; Lemieux et al. 2007).  Variations in Nd isotopic compositions of sedimentary units within the miogeocline of northern British Columbia, Yukon Territory, and the Northwest Territories (Boghossian et al. 1996; Garzione et al. 1997) have provided an initial framework for interpreting sediment provenance.  Boghossian et al. (1996) and Garzione et al. (1997) concluded that late Neoproterozoic to early Paleozoic clastic strata contain detritus derived primarily from Archean and Proterozoic rocks of the Canadian Shield, whereas mid-Paleozoic to Cretaceous units contain components that may have originated in part from more juvenile sources.  Although the sedimentary provenance  10 history for strata exposed in the Mackenzie Mountains can be partly inferred from these studies, sample suites on which the previous studies were based were mainly from portions of the miogeocline well to the west and south of the Mackenzie Mountains.  In addition, attempting to constrain all possible sources of sediments using Nd isotopic information can be limited due to the nature of bulk sample analysis.   Methods such as U-Pb single grain detrital zircon geochronology can be used to overcome this problem, and to complement results of parallel Nd isotopic studies. In this study we present new U-Pb detrital zircon data from Mackenzie Mountain clastic strata of Early Neoproterozic through Cretaceous age that builds on earlier regional work by Ross and Bowring (1990); Gehrels et al. (1995); Rainbird et al. (1997); Gehrels and Ross (1998) and Ross et al. (2005). We use detrital zircon age information to provide constraints on sediment provenance for units of the miogeocline in the Mackenzie Mountains. We compare our detrital zircon ages with known ages of specific igneous rock units combined with known detrital zircon ages from older sedimentary basins in potential source areas and speculate on potential linkages.  These sources may have been exposed periodically and may not be entirely local to northwestern North America.  We also document how sedimentary sources can change locally within a long- lived (900 M.y.) passive margin.    11  Figure 2.1.  Map showing the general geology of the central Mackenzie Mountains with locations of detrital zircon samples. The distribution of miogeoclinal strata in western Canada is shown in the lower left inset. Geology simplified from Blusson (1971, 1974); Gabrielse et al. (1973); Aitken and Cook (1974).       12   2.2. MACKENZIE MOUNTAINS STRATIGRAPHY AND SAMPLING   Samples of early Neoproterozoic to Early Cretaceous age were collected during 1:250,000 scale regional bedrock mapping of NTS sheets 105P and 106A (fig. 2.1).  This mapping was carried out in collaboration with the Northwest Territories Geoscience Office during July and August of 2006 and 2007.  The entire stratigraphic sequence of the central Mackenzie Mountains is summarized here with emphasis on units that were sampled for detrital zircon U-Pb analysis.  A more detailed discussion of these strata is presented in Roots and Martel (2008).   13  Figure 2.2.  Generalized stratigraphy of the central Mackenzie Mountains (modified from Dewing et al. 2006) with approximate stratigraphic location of detrital zircon samples.  2.2.1. Mackenzie Mountains Supergroup (MMSG)   The Mackenzie Mountains Supergroup (MMSG) in the Mackenzie Mountains (fig. 2.2) comprises strata deposited throughout much of northwestern Laurentia in the  14 late Middle Proterozoic to Early Proterozic time (<1083 Ma and >779 Ma; Heaman et al. 1992, Rainbird et al. 1997).  The minimum age constraint is based on ca. 779 Ma cross cutting dykes and sills (Heaman et al. 1992), whereas the maximum age constraint is previously constrained from a ca. 1083 Ma detrital zircon from the Katherine Group (Rainbird et al. 1997). These units belong to Sequence B strata of Young et al. (1979, 1982) and Rainbird et al. (1996) and comprise shallow water sediments deposited in an intracratonic basin on northwestern supercontinent Rodinia in Neoproterozoic time. The succession consists of 4 main units; from bottom to top these are, unit H1, Tsezotene Formation, Katherine Group, and Little Dal Group (fig. 2.2).  Sandstones and dolostones of unit H1 are conformably overlain by a coarsening upward siliciclastic cycle of the Tsezotene Formation (Gabrielse et al. 1973), which terminates at the base of the Katherine Group (fig. 2.2).  The Katherine Group has been divided into three formations (the lower, middle, and upper) with a further seven internal members (K1-7 of Aitken et al. 1978).  Sample 1 was collected from the top of the upper formation (fig. 2.2).  The carbonate dominated Little Dal group (fig. 2.2) was deposited conformably above the upper formation (member K7) of the Katherine Group and consists of 7 informal units (Turner and Long 2008).  The lowest of these termed the Mudcracked Formation, was sampled for this study (Sample 2; fig. 2.2).  The top of the Little Dal Group consists of flood basalts and stromatolitic dolostones that are in turn unconformably overlain by the Windermere Supergroup (fig. 2.2).   15  2.2.2. Windermere Supergroup   In the northern Canadian Cordilleran Mackenzie Mountains the Windermere Supergroup records the initiation of a complex passive margin in Neoproterozoic time, as a result of continental extension (Ross 1991; MacNaughton et al. 2000).  The basal unit of the Windermere Supergroup (fig. 2.2) in the Mackenzie Mountains is the glaciogenic Rapitan Group, which comprises the Sayunei and the overlying Shezal formations (Jefferson and Parrish 1989).  A separate sequence, termed the Coates Lake Group (fig. 2.2), lies stratigraphically between the overlying Rapitan Group and underlying MMSG.  The exact correlation of the Coates Lake Group is somewhat controversial; however, this sequence was not investigated in this study so it is not discussed further. The Windermere Supergroup in the Mackenzie Mountains consists of two distinct stratigraphic successions.  The lowest succession is a rift related sequence represented by the Rapitan Group (Jefferson and Parrish 1989), which is overlain by an upper post-rift related succession that records the thermal subsidence of the continental margin (Narbonne and Aitken 1995).  Continuous and extensive stratigraphic units of the post-rift deposits comprise fine to coarse grained siliciclastic rocks (Twitya Formation), and overlying carbonates, syndepositional glacial deposits and a cap dolostone of the Keele and Ice Brook formations, and Teepee Dolostone respectively (fig. 2.2).  Sample 3 was collected from the Keele Formation (fig. 2.2). Overlying the “cap dolostone” are four separate formations that collectively record the final stages of thermal subsidence and the progradation of a passive margin (MacNaughton et al. 2000).  They broadly comprise interbedded thick sequences of basinal shales and platformal carbonates, and include,  16 from bottom to top, the Sheepbed, Gametrail, Blueflower, and Risky formations (fig. 2.2).  The upper Risky Formation records the base of the Cambrian, which corresponds to the top of the Windermere Supergroup in the Mackenzie Mountains (Fritz et al. 1991). MacNaughton et al. (1997); however, places the base of the Cambrian units near the base of the Backbone Ranges Formation (see following text).  2.2.3. Early Cambrian Units   Overlying the upper Windermere strata in the Mackenzie Mountains is a succession of Early Cambrian platformal, shelf edge, and deep-water basinal strata that were deposited during a period of renewed crustal extension along the evolving passive margin.  This extension led to the formation of the Selwyn basin as a major deep-water basin that lay immediately west of the Mackenzie Platform (Gordery and Anderson 1993). Collectively these rock units reflect flooding during an Early Cambrian transgression of the passive margin (Dillard et al. 2007 and references therein).  From bottom to top, the Early Cambrian sequence includes; the Backbone Ranges, Vampire and Sekwi formations (fig. 2.2).  The Backbone Ranges Formation contains two members (lower and upper), separated by an unconformity.  The lower member is dominated by fine to coarse-grained nearshore siliciclastics and shallow water carbonates (MacNaughton et al. 2000).  The upper member is more variable and comprises deltaic and shoreface derived sandstones (MacNaughton et al. 2000).  Samples from the base of the lower member (Sample 4) and from the top of the upper member (Sample 5) were collected for this study (fig. 2.2).  Overlying the upper member of the Backbone Ranges  17 Formation is the Vampire Formation, which comprises fine-grained, marine siliciclastic rocks.  Overlying the Vampire Formation are dominantly carbonate rocks with locally interbedded siliciclastic rocks of the Sekwi Formation.  The Sekwi Formation records a shallow to deep water transition with subtidal to shallow subtidal deposits (Dillard et al. 2007) and forms the basal unit of the Mackenzie Platform and Selwyn basin.  Sample 6 is from the top of the Sekwi Formation (fig. 2.2).  2.2.4. Mackenzie Platform and Selwyn Basin   The upper Cambrian to lower Silurian section in the Mackenzie Mountains is represented by strata of the Mackenzie Platform, defined by platformal to shelf transitional carbonates, and their basinal siliciclastic and shale equivalents to the west (Selwyn Basin; Cecile 1982; Gordey and Anderson 1993).  The upper Cambrian to lower Silurian section consists, from bottom to top, of the Franklin Mountain-Rabbitkettle and Whittaker-Mount Kindle formations (fig. 2.2). The Franklin Mountain Formation unconformably overlies the Cambrian Sekwi Formation and generally consists of two internal units.  A basal unit of bioturbated sandstones and dolostones, and an upper unit of thin to thick-bedded grey dolostone are observed in the study area.  A sample from the basal unit was collected for this study (Sample 7; fig. 2.2).  The Franklin Mountain Formation changes gradationally into basinal facies of the Selwyn Basin to the west, where correlative strata are referred to as the Rabbitkettle Formation.  An erosional unconformity defines the top of the Franklin Mountain Formation and the base of the  18 Mount Kindle and coeval Wittaker Formations.  Collectively these strata consist of thin to thick-bedded dolostones, which are locally fossiliferous, and chert bearing.  Two major stratigraphic subdivisions are observed in Devonian strata in the Mackenzie Mountains, one is of Early to Middle Devonian age and the other of Middle to Late Devonian age. Lower to Middle Devonian shallow water carbonates record a transgression onto the western margin of the Ancestral North America (Fritz et al. 1991). This sequence of generally fossiliferous, thin to thick-bedded carbonates includes, from bottom to top, the Delorme, Arnica, Landry, Bear Rock and Hume formations (fig. 2.2). Middle to Upper Devonian strata are dominantly siliciclastic rocks with minor carbonate interbeds. From bottom to top, this sequence comprises the Hare Indian, Canol, and Imperial formations (fig. 2.2).  Overlying these units, which are exposed in the core of a syncline in the southwest corner of map sheet 105P, are siliciclastic rocks of Carboniferous to Permian age that include the Heritage Trail Formation, which is tentatively correlated with the Tsichu Group (Roots and Martel 2008).  Sample 8 is from the base of this unit (fig. 2.2).  2.2.5. Clastic Foredeep   Rock units of Cretaceous age in the central Mackenzie Mountains are dominantly granodioritic intrusions, which crop out in the southwest corner of map sheet 105P. However, a downfaulted block of predominantly shales with interbedded lithic arenites containing Cretaceous age plant fossils (Blusson 1971) occurs in the northeastern part of map sheet 105P (fig. 2.1).  This stratigraphic unit indicates that regional mountain  19 building structures in this area are predominantly post-Cretaceous in age.  A sample was collected from the top of this unit for this study (Sample 9; fig. 2.2).  2.3. ANALYTICAL METHODS   Analytical work was conducted using laser ablation inductively coupled plasma- mass spectrometry (LA-ICP-MS) at the Pacific Center for Isotope and Geochemical Research (PCIGR) at the University of British Columbia.  Zircons were separated from 1- 3 kg samples of fine- to coarse-grained sandstone using conventional crushing and grinding, Wilfley table, heavy liquid, and magnetic separation methods.  A random selection of zircon grains was mounted in an epoxy puck along with several grains of the 337 Ma PL standard zircon (Sláma et al. 2008) and brought to a very high polish.  To avoid selection bias, high quality portions of all grain sizes, free of alteration, inclusions or cores, were selected for analysis.  The surface of the polished mount was washed for 10 minutes with dilute nitric acid and rinsed in ultraclean water prior to analysis. Instrumentation at PCIGR utilizes a New Wave UP-213 laser ablation system and a ThermoFinnigan Element2 single collector, double-focusing, magnetic sector ICP-MS, similar to what is described in Chang et al. (2006).  Line scans rather than spot analyses were employed in order to minimize elemental fractionation during the analyses. Typically 40% laser power and 15 or 25 micron spot size (depending on the size of the zircon grains) was used.  Background measurements were acquired with the laser off for seven seconds, followed by data collection with the laser on for approximately 29 seconds.  The time-integrated signals were analyzed using the GLITTER software  20 package described by Van Achterbergh et al. (2001) and Jackson et al. (2004), which automatically subtracts background measurements, propagates all analytical errors, and calculates isotopic ratios and ages.  Corrections for mass and elemental fractionation were made by bracketing analyses of unknown grains with replicate analyses of the zircon standard.  A typical analytical session consists of alternating analyses of five unknown zircons followed by one analysis of the standard zircon.  We begin and finish each session with four analyses of the standard zircon sample.  Data quality is monitored by doing replicate analysis of a 197 Ma internal zircon standard. Interpretation and plotting of the results employs the Isoplot software of Ludwig (2003).  We measure the U, Pb, and Th signals using time resolved GLITTER software to assess the quality of individual analysis on a grain-by-grain basis.  For detrital zircons with an age older than 1.0 Ga we base our discussions on 207Pb/206Pb ages and we assess the discordance of each analysis using the 206Pb/238U and 207Pb/206Pb age correlation.  Low concentrations of 207Pb hinder the assessment of individual zircon discordance for grains younger than 1.0 Ga; therefore we use 206Pb/238U ages for grains younger that 1 Ga. We exclude zircon analyses that are more than 10% and less than -5% discordant in our interpretations.  All analytical data in the U-Pb study is reported in Appendix B and are shown graphically in relative probability plots with stacked histograms (prepared using the Isoplot 3.0 Excel macro of Ludwig 2003).  These plots (figures 2.3, 2.4, and 2.5) contain both a relative probability curve corresponding to age peaks and a histogram to show the distribution of specific grain populations (e.g., Fedo et al. 2003; Beranek et al. 2006). We use the Detrital Zircon Age Pick Excel macro designed by G.E. Gehrels to identify peaks in our data.  This macro processes data at the 1-sigma uncertainty level and peaks  21 and age groupings are returned at the 2-sigma uncertainty level.  Also, figures 2.7, 2.8, 2.9, and 2.10 were created using the Normalized Age Probability Plot Excel macro created by G.H. Gehrels.  2.4. RESULTS  2.4.1. Katherine Group   This sample (Sample 1; 07CL-1442A; medium grained cross bedded quartz arenite) was collected from near the top of the Katherine Group (K-7 of Aitken et al. 1978).  A total of 45 zircon grains were analyzed of which 33 analyses were acceptable. Results are plotted in Figure 2.3a.  Zircon ages obtained range from 2714 – 1005 Ma with two internal age populations, each of which define several individual peaks in a probability density plot (fig. 2.3a). The main age population contains zircon grains of Mesoproterozoic age (1507 – 1003 Ma, 24 zircons) with defined internal peaks at 1014, 1099, 1156, 1238, 1325, and 1484 Ma.  A subordinate population consists of zircons with ages of 1659 – 1615 Ma (4 grains) with a defined peak at 1641 Ma.  Two Paleoproterozoic zircons yield ages of 1737 and 1972 Ma and two Archean zircons yield ages of 2694 and 2714 Ma, none of which define a population.     22 2.4.2. Little Dal Group   Sample 2 (RAS06-188A; fine to coarse grained, poorly sorted quartz arenite) was collected from the base of the Upper Little Dal Group.  A total of 65 zircon grains were analyzed, of which 59 analyses were acceptable (fig. 2.3b).  Detrital zircon ages obtained from this sample range from 2759 – 1019 Ma with four internal age populations.  The main population comprises Mesoproterozoic zircons with ages from 1378 – 1108 Ma (28 zircons) with a well defined peak at 1118 Ma (12 zircons) and subordinate peaks with ages of 1183, and 1335 Ma.  Two Paleo – Mesoproterozoic populations are defined in our data with ages of 1503 – 1420 Ma and 1613 – 1651 Ma with defined peaks at 1458 Ma (11 zircons) and 1615 Ma respectively. Two groups of zircons with ages 1979 - 1678 Ma (7 zircons) and 2759 – 2329 Ma (6 zircons) do not define statistical populations, however, reflect minor sources.  2.4.3. Keele Formation             Sample 3 (06CL-48A; poorly sorted carbonate cemented lithicarenite) was collected from the middle of the Keele Formation within the “Keele Clastic Wedge” of James et al. (2001).  A total of 68 zircon grains were analyzed, of which 49 analyses were acceptable (fig. 2.3c).  Detrital zircon ages obtained from this sample range from 3439 – 1037 Ma with two main age populations and one subordinate population.  A distinct Mesoproterozoic population with ages 1381 – 1037 Ma (13 zircons) has defined peaks with ages of 1037, 1147, and 1348 Ma.  Another population contains ages 1981 – 1671 Ma with peaks at  23 1695, 1814, 1872 and 1980 Ma.    A subordinate population with ages 2322 – 2220 Ma define a peak at 2317 Ma.  Eight zircons yielded ages between 3439 – 2389 Ma and define peaks in the data at 2410 and 2668 Ma.   24  Figure 2.3. U-Pb detrital zircon relative age probability distribution curves and histogram plots for the Katherine Group (a), Little Dal Group (b), and the Keele Formation (c). Data for samples in this and subsequent figures is shown with the stratigraphically lowest sample on the bottom and samples from higher above.    25  2.4.4. Backbone Ranges Formation   Sample 4 (CL06-47A; fine to coarse grained, poorly sorted feldspathic arenite) was collected from the lower member of the Backbone Ranges Formation.  A total of 52 zircon grains were analyzed, of which 46 analyses were acceptable (fig. 2.4a)  Detrital zircon ages obtained from this sample range from 3070 – 1045 Ma with one main age population.  A major Paleoproterozoic population with ages 2029 – 1753 Ma (33 zircons) define a peak with an age of 1913 Ma.  Two Mesoproterozoic zircons yielded ages of 1045 and 1264 Ma and the oldest zircons gave ages of 3070 - 2084 Ma, all of which do not define a population.  Sample 5 (07CL-1553A; fine grained quartzite) was collected from the top of the upper member of the Backbone Ranges formation.  A total of 64 zircon grains were analyzed, of which 54 analyses were acceptable (fig. 2.4b).  This sample yielded detrital zircon ages between 3047 – 1730 Ma with one main age population and one lesser subordinate population.  A major Paleoproterozoic population, defined by ages 1935 – 1806 Ma (28 zircons), contains a dominant peak with an age of 1828 Ma and a two lesser peaks at 1870 and 1916 Ma.  A subordinate Archean population of zircons with ages 2721 – 2683 Ma (5 zircons) define a peak at 2697 Ma. Two groups of ages 2658 – 1947 Ma and 3047 – 2740 Ma collectively do not individually comprise a statistical population.    26  2.4.5. Sekwi Formation   Sample 6 (CL06-44A; moderately sorted, finely bedded lithic arentie) was collected from top of the Sekwi Formation. A total of 71 zircon grains were analyzed, of which 57 analyses were acceptable (fig. 2.4c)  Detrital zircons analyzed from this sample range in age from 2668 to 999 Ma and define two major age populations.  A dominant population comprise ages 1207 – 999 Ma (14 zircons) with a major peak at 1126 Ma.  A second population of ages, 1745 – 1288 Ma (35 zircons), define a dominant peak with an age of 1434 Ma, and three lesser peaks with ages of 1287, 1643, and 1714 Ma.  Older zircons, 2668 – 1771 Ma, do not collectively form detrital zircon age populations.   27  Figure 2.4. U-Pb detrital zircon relative age probability distribution curves and histogram plots for the lower (a) and upper (b) members of the Backbone Ranges Formation and the Sekwi Formation (c).     28 2.4.6. Franklin Mountain Formation   Sample 7 (EM06-113B; medium grained quartzite) was collected from the base of the lower ‘red beds’ of the Franklin Mountain Formation.  A total of 54 zircon grains were analyzed, of which 42 analyses were acceptable (fig. 2.5a).  Detrital zircon ages obtained from this sample range from 2765 – 1037 Ma with one major age population.  The dominant population contains ages 1500 – 1037 Ma (35 zircons) and defines three internal peaks with ages 1126, 1242 and 1416 Ma.  A subordinate population comprises grains with ages 1969 – 1914 Ma (3 zircons) with a defined peak at 1924 Ma. Four Late Archean zircons (2764 – 2513 Ma) do not define a population.  2.4.7. Heritage Trail Formation   Sample 8 (EM06-70A; fine grained quartz arenite) was collected from the base of the unit 27 of Blusson (1971), which is currently interpreted by Roots and Marte (2008) to correlate with the Heritage Trail Formation of the Carboniferous Tsichu Group (Cecile 2000).  A total of 66 zircon grains were analyzed, of which 55 analyses were acceptable (fig. 2.5b).  Detrital zircons analyzed from this sample range in age from 2872 to 427 Ma with many distinct age populations in the data.  A Silurian population consisting of three zircons with ages 437 – 427 Ma define a peak at 433 Ma.  There is a main population with ages 1129 – 1031 Ma (14 zircons) forming three peaks with ages 1038, 1084 and 1126 Ma. Three lesser subordinate populations with ages 1181 – 1160 Ma (3 zircons),  29 1647 – 1604 Ma (7 zircons), and 1729 – 1703 Ma (5 zircons) contain peaks with ages 1165, 1622, 1641 and 1720 Ma.  Seven Paleoproterozoic zircons, 1970 – 1772 Ma, and five Paleoproterozoic – Late Archean zircons, 2872 – 2367 Ma comprise the oldest zircons in this sample, however, they do not collectively define age populations.  2.4.8. Clastic Foredeep   Sample 9 (07DL-01; medium grained lithic arenite) was collected from the base of map-unit 30 of Blusson (1971).  A total of 74 zircon grains were analyzed, of which 55 analyses were acceptable (fig. 2.5c).  Detrital zircon ages obtained from this sample range in age from 2938 to 392 Ma with two main age populations.  Zircons of Early Ordovician to Early Devonian (477 – 393 Ma, 12 zircons) define a broad age population with two prominent internal peaks at 424 and 442 Ma.  As with Sample 8, there is a nearly continuous spectrum of Proterozoic zircon ages in this sample, with ages from 2011 – 1054 Ma.  A dominant population of zircons with ages 1154 – 1054 Ma (11 zircons) define two peaks at 1068 and 1137 Ma. A subordinate population with ages 1335 – 1254 Ma (5 zircons) yield two peaks with ages 1255 and 1332 Ma.   Zircons with ages 1650 – 1632 Ma (4 zircons) comprise a lesser population with a peak age of 1642 Ma. Paleoproterozoic grains with ages 1886 – 1735 Ma (7 zircons) contain peaks at 1737, 1790, and 1851 Ma.  Paleoproterozoic to Late Archean zircons, 2938 – 1906 Ma, contain a population with ages 2758 – 2723 Ma and define a peak at 2738 Ma.   30  Figure 2.5. U-Pb detrital zircon relative age probability distribution curves and histogram plots for the Franklin Mountain Formation (a), Tsichu Group (b), and the Cretaceous clastic foredeep (c).     31 2.4.9. Summary of Detrital Zircon Results   The following groups of detrital zircon ages are pervasive throughout the strata studied (see fig. 2.7); 1400 – 1000 Ma (exceptions, samples 4 and 5); 1590 – 1400 Ma (exceptions, samples 3, 4 and 5); 1690 – 1600 Ma (exceptions, samples 4 and 5); 2080 – 1700 Ma; and 2800 – 2415 Ma.  Also a group of zircons, 477 – 392 Ma, are restricted to Carboniferous and younger strata (e.g., samples 8 and 9) and Middle Archean zircons (3.4 - 3.0 Ga) were only evident in Neoproterozoic and Cambrian age strata (samples 3, 4 and 5).  2.5. POTENTIAL DETRITAL ZIRCON SOURCES    Compilations of the Precambrian rock units that make up the North American Craton by Hoffman (1989) and Villeneuve et al. (1993) provide a robust geochronological framework for potential basement sources of detrital zircons in the Cordillera (fig. 6).  Comparing this data set with ages obtained from our study can provide insight into possible sources for detrital zircons.  In the following section we summarize potential source ages for detrital zircons in the miogeocline of northwestern Laurentia.  These sources include both ages of known igneous rocks as older sedimentary units with known detrital zircon ages.   32  Figure 2.6. Tectonic framework of Precambrian basement provinces and Paleozoic granitoids of North America (references reported in text; modified from Whitmeyer and Karlstrom 2007). Basement provinces of North America: N, Nahanni; FS, Fort Simpson, H, Hottah; BH, Buffalo Head; TT, Taltson Thelon; R, Rimbey; GB, Great Bear; S, Slave; RP, Rae; HP, Hearn; W, Wyoming; GF, Great Falls tectonic zone; TH, Trans Hudson; SU, Superior; P, Penokean; Mk, Makkovic; T, Torngat; CB, Cumberland batholith; F, Foxe. Precambrian sedimentary basins: TB, Thelon; A, Athabasca; B, Belt Purcell.  33 2.5.1. Archean and Proterozoic Zircon Grains    Archean detrital zircon grains can be directly linked to cratonic sources in the North American Craton to the east (fig. 2.6) such as the 4.0 - 2.5 Ga Slave, Rae, Hearn or Nova provinces (Hoffman 1989).  Early Paleoproterozoic zircons (e.g., 2400 – 2100 Ma) were likely shed from Paleoproterozoic accreted terranes such as the 2324 – 1990 Ma Buffalo Head terrane (Villeneuve et al. 1993) or from the 2440 – 2000 Ma plutons of the adjacent part of the western Rae Province (Bostock and Loveridge, 1988).  Zircons with ages 2000 – 1800 Ma are possibly derived from Paleoproterozoic accretionary and continental orogens comprising the composite Wopmay orogen (e.g., 1845 Ma Fort Simpson arc; 1924 – 1845 Ma Hottah terrane; or 1865 – 1840 Ma Great Bear magmatic zone; Villeneuve et al. 1993) or more distally from the 2000 – 1949 Ma Taltson-Thelon orogen or the 1910 – 1850 Ma Trans-Hudson orogen.  There are no known local felsic igneous sources for zircons with ages 1800 – 1610 Ma in the western North American Craton (fig. 2.6); however, the dominantly mafic ca. 1.71 Ga Bonnet Plume River intrusions presently exposed in the Yukon Territory (Thorkelson et al. 2001) may have been a minor local source for zircons of that age. Distal and widespread felsic igneous sources are evident in southern and eastern North America.  These igneous sources comprise the 1800 – 1600 Ma Yavapai and Mazatzal orogens of southeastern United States (Hoffman 1989), local 1750 – 1600 Ma magmatism (e.g., 1750 – 1680 Ma arc or back-arc magmatism in the southwest Grenville Province and 1710 – 1600 Ma Labradorian magmatism in the northwest Grenville Province) associated with pre-Grenville tectonism (Carr et al. 2000; Davidson, 2008) and the widespread ~1720 and ~1647 Ma granitoid plutonism of the Makkovik orogen of  34 Labrador (Kerr et al. 1992). More proximal sources to the Mackenzie Mountains comprise the western Churchill Province which also contains magmatism that spans 1765 -1750 Ma (Nueltin plutons and coeval Pitz Formation rhyolites; van Breemen et al. 2005; Rainbird and Davis 2007).  The Thelon sedimentary basin (e.g., fig. 2.6) contains abundant 1.8 – 1.7 Ga detrital zircons (Whart Sequence of Rainbird and Davis 2007) and may also be source zircons. Furthermore, the local Paleoproterozoic Wernecke Supergroup exposed in the Yukon Territory also contains abundant 1800 – 1600 Ma detrital zircons (Furlanetto et al. 2009) and may also represent a source for zircons of that age in the Mackenzie Mountains strata deposited following sedimentary recycling.  Zircons with ages 1610 - 1490 Ma fall into the “North American Magmatic Gap” (NAMG; Van Schmus et al. 1993; Ross and Villeneuve 2003); and represents zircon ages with no clearly identified igneous sources in western North America (fig. 2.6).  Possible known sources for zircons of this age are the Gawler Craton of eastern Australia (e.g., Fanning et al. 1998; Blewett et al. 1998) or from the Grenville Province of eastern North America (e.g., Gower and Tucker, 1994; Carr et al. 2000; Davidson 2008).  A source for detrital zircons with ages 1490 - 1400 Ma is not apparent in nearby basement provinces of the western North American Craton (fig. 2.6).  Detrital zircon grains of these ages could be derived from (1) unknown buried basement rocks of that age in northwestern Canada such as the poorly dated Nahanni domain, (2) sources in southern and southeastern United States such as the ~1400 Ma anorogenic plutons (Granite-Rhyolite province) which extend across New Mexico, Arizona, southern California, and northern Sonora (e.g., compilation Figure 2 of Goodge and Vervoort,  35 2006), (3) the Belt Basin of southeastern British Columbia, southwestern Alberta and northern Montana in which detrital zircon grains of this age are abundant in the Belt Supergroup and are inferred to have been derived from local syndepositional magmatism comprising ages ~1470 – 1440 Ma (summarized in Ross and Villeneuve, 2003), (4) Eastern North American sources (summarized from Gower and Tucker, 1994; Carr et al. 2000; Davidson 2008) including the 1520 – 1460 Ma Pinwarian arc related magmatism comprising 1490 and 1472 Ma granitoid plutons; 1499 Ma pegmatites; and 1509 Ma apilitic dykes.  Furthermore, widespread 1480 – 1400 Ma plutonism comprising; 1450 - 1420 Ma arc magmatism; 1480 - 1400 Ma magmatism coinciding with anorogenic magmatism of midcontinental United States (the “Granite-Rhyolite province); widespread 1460 - 1430 Ma granitoid orthogneiss; megacrystic 1460 – 1430 Ma plutons of the Brit domain; and 1450 Ma migmatic orthogneises of the Muskoka domain, or (5) a Mesoproterozoic source to the west of the North American via a cryptic drainage system which was later reworked during the deposition of the lower MMSG. There are two possible sources for zircons with ages 1400 – 950 Ma which collectively comprise what is termed “Grenville” (e.g., Rainbird et al. 1992, 1997; Gehrels and Ross 1998; Ross et al. 2005).  Rainbird et al. (1992, 1997) and Rainbird (2007) suggested that the Katherine Group of the MMSG contain detritus (1250 - 1000 Ma) sourced from the Grenville Province of eastern North America.  Potential zircon igneous sources in the Grenville Province (summarized from Carr et al. 2000; and Davidson 2008) comprise magmatism associated with the 1.29 – 1.19 Ga Elzeverian orogeny, as well as the 1.19 – 1.14 Shawinigan magmatism; 1.17 – 1.15 Ga Frontenac plutons; minor local 1.14 – 1.09 Ga felsic plutons (e.g., ~1100 Ma Hawk granite); mafic  36 and alkalic dykes associated with the Midcontinent Rift and magmatism associated with the 1.08 – 1.02 Ga Ottawan orogeny and the 1.0 – 0.98 Ga Rigolet orogeny.  Rainbird et al. proposed that detrital zircon grains of these ages were shed some 3000 km to the west from the Grenville front via a continental scale braided river system (Rainbird et al. 1992, 1997).  This hypothesis was later supported by Thorkelson et al. (2005) who obtained ages of 1113 -1038 Ma for detrital muscovite in the Tsezotene Formation of the MMSG, which they interpreted to have been also derived from the Grenville Province.  Also, zircon grains of “Grenville” age may have been dispersed from the northern Ellsemerian orogen, exposed on Ellsemere Island of the Canadian Arctic Islands, in which basement of the Pearya terrane also contains ~1000 Ma plutons (Trettin 1991).  It is unlikely that only one of these sources contributed to the ~1050 Ma detrital zircon populations and therefore mixing of these two sources is possible. Gehrels and Ross (1998), Ross et al. (2005), and Lemieux et al. (2007), however, suggest that there may be a more proximal source for 1250 – 1000 Ma zircons recovered from the miogeocline.  These authors suggest that unknown plutons of this age may be buried beneath the Canadian miogeocline or were present outboard as indicated by zircons of this age recovered from strata of the miogeocline. Jefferson and Parrish (1989) and Mortensen and Colpron (1998) reported zircon crystallization ages of 1175 – 1100 Ma from granitic xenoliths recovered from an undated lower Paleozoic diatreme from the southern Mackenzie Mountains.  This granitic clast may provide the only age constraint from the enigmatic Nahanni domain basement that lies beneath much of the Northern Cordillera of northern British Combia, Yukon and western Northwest Territories.  The Nahanni domain was originally defined by Hoffman (1989) as an aeromagnetic low (fig.  37 2.6) that lay to the east and west of the Fort Simpson terrane.  Subsequent work restricted the Nahanni domain to the aeromagnetic low west of the Fort Simspone terrane (Thorkelson et al. 2005 and references therein).  The age and composition of the Nahanni domain is poorly known as it is nowhere exposed and it has not been drilled.   The ages from the granite xenolith however, indicates that at least minor Grenville age magmatism is in the basement of the southern Mackenzie Mountains.  This magmatism is possibly associated with a 1.6 - 0.8 Ga tectonic event in the northern Cordillera inferred from deep penetrating seismic studies (e.g., Cook 1988; Clark and Cook, 1992; Cook and Erdmer, 2005) and the ~1.6, ~1.27 and ~1.15 Ga metamorphic zircons from crustal xenoliths (e.g., Milidagovic 2008) from eastern Yukon.  Also, the ca. 1380 Ma Hart River intermediate sills and the ca. 1270 Ma mafic Bear River dykes (Mackenzie dyke swarm equivalents) presently exposed in Yukon Territory (Thorkelson et al. 2005) are also potential local sources for zircons of these ages.  Furthermore, Ross et al. (2005) suggest that 1.4 - 1.0 Ga detrital zircons from the Hyland Group may have been derived from “intra- Cordilleran Grenville” basement exposed during Windermere rifting.  The Hyland Group is exposed in Yukon and Alaska and is thought to correlate with the Neoproterozoic Windermere Supergroup and Early Cambrian strata of the Mackenzie Mountains (e.g., Rainbird et al. 1996). 2.5.2. Paleozoic Zircon Grains  Silurian and Devonian zircon grains (477 – 392 Ma) are not readily linked to Laurentian igneous sources and have most likely been shed from the northern Canadian Arctic margin or from plutons exposed on Ellesmere Island.  Detrital zircons of these  38 ages are also found in southeastern British Columbia in the Pennsylvanian Spray Lakes Group of Gehrels and Ross (1998); in the Upper Devonian Nation River Formation in eastern Alaska (Gehrels et al. 1999); in Paleozoic strata of the Alexander Terrane in southwestern Yukon (Gehrels et al. 1996) and in Devonian to Triassic strata in central and eastern Yukon (Beranek 2009). Collectively these authors suggest that zircons of these ages are derived from ~390 - 360 Ma plutons (Lane 2007) exposed along the Canadian Alaskan margin and and ~430 Ma plutons (Trettin 1991) in the Canadian Arctic islands.  These sediments were dispersed southward in the middle Paleozoic when clastic sediments from the Ellsemerian orogen blanketed much of northwestern Canada (Gordey 1991).  2.6. INTERPRETATION OF DETRITAL ZIRCON AGES    Paleocurrent data from the Canadian miogeocline suggests that the majority of strata were deposited from sources to the east or northeast (present day coordinates) along the margin of Ancestral North America (e.g., Aitken and McMechan, 1991; Rainbird et al. 1996; Cecile et al. 1997; McNaughton et al. 2000).  Because this study deals with rock units deposited on this margin we can conclude that basement rocks comprising the Canadian Shield were a primary source for the clastic component. However, some sediment recycling and reworking of strata prior to final deposition is inevitable in any passive margin setting (e.g., Gehrels and Ross 1998).  All analysed detrital zircons grains, with the exception of some Paleozoic age zircons, are quite rounded, suggesting substantial physical abrasion possibly by long residence times in river systems or by repeated reworking of sediments.   On the basis of regional tectonics,  39 as inferred from stratigraphic studies, there were profound episodes of uplift and subsidence that affected the miogeocline throughout its depositional history (Bond et al.1983; Bond and Kominz 1984; Cecile et al. 1997).  Furthermore, erosion can be directly associated with uplift and is reflected in the detrital zircon populations.  Possible sources of detrital zircons are outlined in Figure 2.7.  Generalizations that are applicable to all units comprise the Paleoproterozoic and Archean zircons (e.g., 3.4 - 1.8 Ga), of which unique Laurentian sources are readily identifiable and offer a conclusive provenance linkage.  These include the aforementioned composite Wopmay orogen and the Archaean cratonic blocks comprising the Slave, Superior and Churchill provinces (fig. 2.6). Zircons with these ages probably reflect, (1) a common source, which remained a tectonic high, or (2) recycling of sediment prior to deposition.   These age populations are not discussed further except where necessary to distinguish between stratigraphic units.  The following interpretations and discussions therefore focus primarily on younger detrital zircons, because these zircon grains are dominant and unique to the Mackenzie Mountains.   40  Figure 2.7. U-Pb detrital zircon normalized age distribution curves for all sedimentary units investigated from the Mackenzie Mountains.  The colored bands represent unique detrital zircon age groupings that correspond with known ages of igneous rocks from potential sources (listed on top).  References reported in text.   41 2.6.1. Provenance of the Mackenzie Mountains Supergroup  Detrital zircons from Sequence B strata (MMSG in the vicinity of the Mackenzie Mountains; Rainbird et al. 1996) reflect multiple sources, evidenced by multiple age populations (samples 1 and 2).  Furthermore, paleocurrent data indicates a dominance of northwestward transport of Sequence B sediments, implying clastic source regions that lay southeast of the paleo-basin (Rainbird et al. 1996).  Detrital zircons from the same stratigraphic horizon as sample 1, K7 of the Katherine Group, were also dated by Rainbird et al. (1997) using ID-TIMS age determination methods.  These authors suggest that 1.3 - 1.0 Ga age zircons may have originated from the Grenville Province and were transported to the west for up to 3000 km via a continental scale braided river system. Furthermore, a local source for zircons of these ages that may have been exposed synchronously with deposition of Sequence B strata is not readily identifiable in western North America (see above). MMSG samples analyzed in this study contain significant 1379 – 1005 Ma, 1480 – 1400 Ma and 1679 – 1610 Ma detrital zircon age components (fig. 2.3a,b).  These ages are significant owing to the absence of known local igneous sources. Distal sources are readily identifiable in eastern North America including those exposed in the composite Grenville Province, as well as the Granite-Rhyolite province of the U.S. mid-continent, Makkovik orogen and the Yavapai and Mazatzal orogens.  Also, zircons with ages 1980 – 1885 Ma may have been dispersed from multiple sources such as the Hottah terrane or more distally from the Taltson-Thelon and/or the Trans-Hudson orogens.  The oldest zircons from MMSG samples (e.g., 2759 – 2586 Ma) are likely derived from the Slave Province or Hearne Province.  Collectively these age groupings may reflect erosion of  42 detritus from distal sources and were in turn transported to the western margin of ancestral North America in Neoproterozoic time.  In order to incorporate eastern North America detritus (e.g., 1.3 – 1.0, 1.48 – 1.4, and 1.68 – 1.61 Ga) river systems draining the easternmost Grenville Province are required, as was suggested by Rainbird et al. (1997) and Rainbird (2007).  We therefore interpret the Katherine Group and Little Dal Group of the MMSG to contain sediments shed from eastern Laurentian cratonic sources, and to possibly include sources as far east as the Grenville Province. 2.6.2. Provenance of the Windermere Supergroup  We can rule out significant reworking of older sediments in MMSG into Windermere Supergroup strata because zircon grains with ages 1660 – 1400 Ma are absent and zircons with ages 2000 – 1700 Ma and 2400 – 2200 Ma zircons are abundant (fig. 2.3).  Underlying MMSG strata (e.g., samples 1 and 2) contain a significant 1660 – 1400 Ma component and a pronounced depletion in 2000 – 1700 Ma and 2400 – 2200 Ma zircons (fig. 2.3).  Therefore, in Windermere time MMSG strata were not exposed, thus “Grenville” age zircons in Sample 3 may reflect continued sourcing from the Grenville Province.  The absence of zircons with ages 1660 – 1400 Ma in Sample 3 is significant because eastern North America contains multiple sources with these ages (e.g., fig. 2.6; Grenville Province, Makkovik, Yavapai and Mazatzal orogens). If we consider an eastern North American source for the 1300 – 1000 Ma zircons in this sample then we would expect these older 1660 – 1400 Ma zircons to be present (e.g., MMSG strata); however, this is not the case.  This may be due to rivers draining local eastern North American sources, primarily ‘Grenvillian,’ in the latest Neoproterozoic.  Therefore, the abundant  43 1800 – 1660 Ma zircons in the Keele Formation may be primarily derived from anomalous eastern regions such as the Makkovik orogen and the ~1750 Ma Nueltin plutons exposed in the Churchill Province or were reworked from intracratonic basis of Laurentia such as the Thelon Basin. Because we suggest that direct eastern North America sources dispersed sediment until Windermere time, it is therefore possible that older sources (e.g., the Mesoproterozoic Granite-Rhyolite province) were not exposed at this time or were not sampled by these more northerly river systems. In addition the abundance of 2400 – 2200 Ma detrital zircons is significant as this is the only unit to contain a high proportion of detrital zircon grains derived from the Buffalo Head terrane. This suggests that basement rocks of the Buffalo Head terrane were also exposed in Windermere time.  Although not a favored model in this study, Ross et al. (2005) suggest that rifting of the western Laurentian margin, in Late Neoproterozoic - Early Cambrian time, may have led to tectonic uplift and exposure of intra-Cordilleran “Grenville” age basement rocks.  This hypothesis was proposed on the basis of relatively abundant 1.4 - 1.0 Ga detrital zircons recovered from the late Neoproterozoic – Early Cambrian Hyland Group of Yukon and Alaska (Ross et al. 2005).  This hypothetical, now buried, basement could potentially be a source for the 1300 – 1000 Ma zircons observed in our age equivalent Windermere Supergroup sample (Sample 3).  We cannot entirely preclude this scenario because little is known about the nature of the basement beneath the Mackenzie Mountains.  The anomalous ca. 1175 – 1100 Ma granite xenoliths from a lower Paleozoic diatreme (e.g., Jefferson and Parrish, 1989; Mortensen and Colpron, 1998) and the  44 inferred ~1.6, ~1.27 and ~1.15 Ga metamorphic events (e.g., Milidragovic 2008) may provide the only clue into nature of the basement beneath the Mackenzie Mountains. 2.6.3. Provenance of the Backbone Ranges Formation   Renewed thermal subsidence resulted in westward basin thickening and deposition of thick shelf to terrestrial strata in latest Neoproterozoic to Early Cambrian time.  This subsidence prompted a widespread transgression onto the continent and formed the sub-Cambrian unconformity (e.g., Fritz et al. 1991; Narbonne and Aitken 1995; MacNaughton et al. 1997).  This unconformity beveled the top of the Windermere strata to the west and down section into the Katherine Group strata to the east as exposed in the Mackenzie Arch.  This erosional event is not reflected in the Backbone Ranges Formation samples (e.g., Sample 4, base of lower member and Sample 5, top of upper member) as the detrital zircon ages for these samples are markedly different from zircon populations observed in Windermere and MMSG strata.  Both samples from the Backbone Ranges Formation, collected 600 m apart stratigraphically, contain a distinct population of Paleoproterozoic (2087 – 1730 Ma) zircons, which form 71% of zircons analyzed from these samples (fig. 2.4).  This likely reflects a local source with little influence from older strata (e.g., two zircons at 1045 and 1264 Ma; Sample 4; fig 2.4a). The lower member of the Backbone Ranges Formation (e.g., Sample 4) contains a prominent 1913 Ma peak whereas the upper member contains a younger 1828 Ma peak (fig. 2.4a,b).  The older ~1910 Ma zircons of the lower member may have originated from the 1910 – 1850 Ma Trans-Hudson orogen followed by a shift to younger proximal (e.g., Fort Simpson arc) sources for the upper member.  However, we consider a Trans-  45 Hudson orogen source unlikely as Paleoproterozoic detrital zircons are dominant and Archean (e.g., Slave Province) detrital zircons are minor in these strata.  A westerly migrating river system with abundant Tran-Hudson zircon grains should also contain abundant Slave Province zircon grains as it would be likely to incorporate Slave Province detritus.  Therefore, the composite 2000 – 1800 Ma Wopmay orogen (fig. 2.6) is a logical candidate for zircon sourcing for the prominent Paleoproterozoic age zircon grains.  The difference between lower and upper members may reflect a westerly shift in sources. The lower member may represent a time when the 1924 – 1845 Ma Hottah terrane was exposed and shed detritus to the west.  The upper member in turn may record a westerly shift in source and represent exposure of the ca. 1845 Ma Fort Simpson arc.  This westerly migration of sources in Cambrian time may be in response to continued subsidence of the margin coeval with uplift of inboard basement terranes such as the Fort Simpson arc.  This local Wopmay orogen source with minor Slave Province input interpretation is plausible owing to the absence of “Grenville” age zircons in the Backbone Ranges Formation.  As previously mentioned there is no known “Grenville” age magmatism within the Wopmay orogen.  Therefore in latest Neoproterozoic to Early Cambrian time, sediments were derived from local sources with minimal reworking and interaction with older Windermere Supergroup or MMSG strata. 2.6.4. Provenance of the Sekwi Formation  In the Lower Cambrian an erosional phase began which corresponds the formation of the Selwyn basin and the initial rifting of the Misty Creek Embayment on the northeastern margin of the Selwyn basin (Cecile et al. 1997).  The tectonic nature of  46 this rifting event is believed to represent a lower plate extensional setting on which strata of the Mackenzie Platform and Selwyn basin were deposited (Cecile et al. 1997).   Older strata, possibly including basement rock, may have been exposed during these periods of block faulting and extension.  This is evident in detrital zircon populations from the Sekwi Formation (Sample 6), which are very different from those in the underlying Backbone Ranges Formation units (e.g., Samples 4 and 5).  The Backbone Ranges Formation strata contain almost no zircons with Mesoproterozoic ages (with the exception of two zircons from the lower member) and a dominant population with Paleoproterozoic ages.  The Sekwi Formation contains abundant Mesoproterozoic zircons with minor amounts of Paleoproterozoic zircons. This difference suggests that tectonism associated with rifting directly affected sediment dispersal patterns to the extent that underlying strata (e.g., Backbone Ranges Formation) were not incorporated into sediments comprising the Sekwi Formation. Detrital zircon populations from the Sekwi Formation (fig. 2.4c) are more similar to the age populations observed in the MMSG strata (e.g., Samples 1 and 2, fig. 2.3).  These similarities comprise abundant ~ 1130, ~ 1450, ~ 1640, ~ 1720, and ~2700 Ma detrital zircon populations and minimal 1980 – 1870 Ma detrital zircons (fig. 2.7).  Similarities such as these likely reflect recycling of MMSG sediments exposed in the Mackenzie Arch along the western margin. Unlike MMSG strata; however, the Sekwi Formation, contains a dominance of ~1450 Ma zircons which may also reflect continued sourcing from more distal 1480 – 1400 Ma sources exposed in the southeastern United States and Labrador (see above), with minor recycling from older strata.   47 2.6.5. Provenance of the Franklin Mountain Formation   The base of the Franklin Mountains Formation (Sample 7) rests on a regionally persistent unconformity, which suggests probable reworking of older strata.  There is a similarity of some detrital zircon age populations of this sample (fig. 2.5a) with those from our sample from the underlying Sekwi Formation and from the MMSG (fig. 2.7). However, the lack of 1910 - 1500 Ma age zircons with and a minor 2000 – 1900 Ma detrital zircon component from this sample is significant because grains of these ages are abundant in the Sekwi Formation (fig. 2.4c) and the MMSG (fig. 2.3a,b) strata and the Backbone Ranges Formation strata (fig. 2.4a,b), respectively.  Significant reworking and erosion of sediments below an unconformity should incorporate all detrital zircon populations from the underlying strata.  Furthermore, this may suggest that a Mesoproterozoic age (1500 - 1000 Ma) detrital zircon source, similar to that sourcing the Sekwi Formation, was also exposed in Early Ordovician time such as older MMSG strata with some primary grain sorting eliminating 1910 – 1500 Ma zircons. This bimodal age distribution may be derived from a local and now buried source, although we do not favor this model. This local source would have a maximum age limit of 1500 Ma owing to the absence of older zircons4 in this unit. 2.6.6. Provenance of the Tsichu Group  Lower to Middle Devonian age strata record a transgression onto the western margin of the Ancestral North America defined by widespread deposition of shallow water carbonates (Fritz et al. 1991).  Middle to Upper Devonian strata that blanketed much of northwestern North America are dominantly siliciclastic rocks with minor  48 carbonate interbeds, and record an influx of clastic detritus derived from the Innuitian and Ellsmerian orogenies and eastward from block faulted uplifted parts of the Selwyn basin (Gordey 1991; Trettin 1991).  An additional source may include the Caledonian orogen, now exposed on eastern Greenland and Scotland, which may have shed detritus to the southwest across northern Laurentia (e.g., McNicoll et al. 1995). Carboniferous time marks a period in which stable shelf conditions returned to the miogeocline.  It is likely that components of the 1100 – 1000 Ma and ~430 Ma detrital zircon populations (fig. 2.5b) from our Tsichu Group sample (Sample 8) were sourced from Canadian Arctic Islands (e.g., Innuitian orogen), either directly from southward draining systems or reworked from older Devonian strata.  Proterozoic zircons from this sample contain similar age peaks to those in underlying strata (fig. 2.7); therefore it is suggested that a component of recycling of older units contributed to these strata.  Zircons with ages 1100 – 1000 Ma in these younger strata may have an additional distinct source as they are particularly dominant in the Tsichu Group and the “clastic fordeep” as evidenced by peaks at 1038, 1084, and 1068 Ma (fig. 2.5b,c).  The Pearya Terrane on western Ellsemere Island contains 1100 – 1000 Ma felsic plutons (fig. 2.6; Trettin, 1991), and owing to the abundance of detrital zircon grains of this age in the Tsichu Group, the Pearya Terrane is viable source for some of the 1100 – 1000 Ma detrital zircons. The Alexander Terrane of western Alaska and Yukon contains abundant Devonian and Silurian plutons throughout its southern extent.  These plutons are believed to have sourced detrital zircons of this age, recovered from mid Paleozoic strata on the Alexander Terrane (Gehrels et al. 1996).  However, the Alexander Terrane is not thought to have  49 docked with western North America until Late Triassic time (Butler et al. 1997), thus making it an unlikely source for detrital zircons in miogeoclinal strata. Additionally, it is unlikely that the northern Innuition and Ellsemerian orogens (fig. 2.6) are the only sources for these sediments as sources for our Proterozoic and Archean grains are not present in the northern Canadian Arctic Islands. We suggest that reworking of older northern Canadian Cordillera sedimentary units (e.g., MMSG and Sekwi Formation) with mixing of northern Arctic Island sources carrying young Paleozoic (430 Ma) and 1100 – 1000 Ma zircons provided much of the detritus for the Tsichu Group strata. 2.6.7. Provenance of the Cretaceous clastic foredeep   The nature of the Cretaceous “Clastic Foredeep” in the Mackenzie Mountains (our Sample 9) is enigmatic because it is unclear whether these strata represent remnants of a much larger easterly migrating prodelta deposit or an isolated, fault controlled, intermontane basin (Roots and Martel, 2008).  The Proterozoic and Archean detrital zircon age populations from this sample are very similar to those observed from the Tsichu Group (fig. 2.5b) and therefore likely reflect reworking and recycling of these sediments.  Young zircons from this sample (fig. 2.5c) range in age from 477 – 392 Ma, which may reflect a much more dominant Early Paleozoic source than was indicated for the ca. 430 Ma zircons from the Tsichu Group.  Detritus may have additionally been shed from the Canadian and Alaskan margin ca. 390 – 360 Ma felsic plutons (fig. 2.6; Lane 2007) that were incorporated during potential mid-Cretaceous southward sedimentation.   50 2.7. COMPARISONS WITH COEVAL STRATA ALONG THE MIOGEOCLINE   Neoproterozoic sedimentary strata in the Mackenzie Mountains comprise the Katherine and Little Dal groups (part of Sequence B strata of Young et al. (1979, 1982) and Rainbird et al. (1996)) and form the western extent of sediments deposited from the proposed Neoproterozoic river systems of Rainbird et al. (1997) and Rainbird (2007). Detrital zircons from the K7 member of the Katherine group were initially studied by Rainbird et al. (1997) using ID-TIMS methods.  This is the same stratigraphic horizon that comprises Sample 1.  Our data is very similar to ages reported by Rainbird et al. (1997), with a dominant Grenville Province and a broadly eastern North American age signature present in our Katherine Group sample (fig. 2.3a).  Rifting of western Laurentia and deposition basal synrift coarse clastic rocks and diamictites of the lower Windermere Supergroup in the northern miogeocline began at ca. 755 Ma (Windgate and Giddings 2000), and was followed by erosion and deposition of the upper Windermere passive margin strata at ca. 570 Ma (Colpron et al. 2002).  The nature of Windermere sedimentation along the margin comprises passive margin sedimentation in the north coeval with initial rifting in the south.  This contrast is shown by comparing detrital zircon ages (fig. 2.8) from Windermere Supergroup and equivalents in the Hyland Group in Yukon (Ross et al. 2005), Horsethief Creek Group in southern Canadian Cordillera (Gehrels and Ross 1998), the Caddy Canyon Formation in Idaho (Stewart et al. 2001), and the Mutual Formation in Utah (Stewart et al. 2001) with our sample 3 (Keele Formation), representing northern Cordillera Windermere strata. Ross and Bowring (1990) also reported detrital zircon ages from Windermere strata from  51 throughout the Canadian Cordillera; however their ages were based on multi-zircon fractions and including only a small number of samples.  Although we cannot plot their data for direct comparison purposes, we can use some of their interpretations regarding zircon provenance.  It is apparent that local tectonism primarily controlled dispersal of detrital zircons in late Neoproterozoic time. Our sample (Sample 3) contains zircons with ages 1381 – 1037 Ma which may represent an eastern North American “Grenville” source. This signature is absent from Windermere strata of British Columbia (Ross and Bowring, 1990; Gehrels and Ross, 1998) but is present in Caddy Canyon and Mutual formations of the western United States (e.g., fig. 8).  In western United States, “Grenville” age zircons (1300 – 950 Ma) are interpreted to have been dominantly derived from local sources (e.g., Pikes Peak Granite; Stewart et al. 2001).  Additionally, Paleoproterozoic zircons (e.g., 2000 – 1700 Ma) are dominant in the Windermere of southern British Columbia, but are absent from the western United States, but are present in moderate proportions in the Mackenzie Mountains. These differences may reflect pronounced differences in source or a response to the different tectonic regimes affecting the northern and the southern parts of the Cordillera.  Thermal uplift in the Canadian Cordillera in northeastern BC as a response to initial rifting may have exposed ancient bedrock (e.g., Peace River Arch) yielding a dominant Paleoproterozoic detrital zircon signature, while renewed subsidence in the north led to passive margin sedimentation with a unique northern Windermere Supergroup detrital zircon signature.  Alternatively, bedrock along the length of the margin may have been exposed at this time and the nature of the bedrock to the north may comprise Mesoproterozoic (e.g., 1.4 – 1.0 Ga) sources (e.g., “intra-Cordilleran Grenville age” basement of Ross et al. 2005).  The detrital zircon  52 signature in the Keele Formation, therefore, likely reflects reworking and recycling of early Windermere Supergroup passive margin strata.  This model for Windermere Supergroup deposition is plausible in view of the spatial distibution of basement provinces of varying ages along the western margin of Laurentia (e.g., heterogeneous north to south, fig. 2.6) and the very limited constraints on the nature of the basement beneath the Mackenzie Mountains.   53  Figure 2.8. U-Pb detrital zircon normalized age distributions curves for Windermere Supergroup equivalent strata from along the length of western North America (the lowest plot represents southern miogeocline strata whereas the top plot represents northern miogeocline strata).  Hyland Group data from Ross et al. (2005); Horsethief Creek Group data from Gehrels and Ross (1998); Caddy Canyon Quartzite and Mutual Formation data from Stewart et al. (2001).  54   Gehrels and Ross (1998), Gehrels et al. (1999), and Stewart et al. (2001) interpret Early Cambrian strata of the miogeocline to consist of detritus derived from a complex source comprising several basement provinces exposed in the western Canadian Shield, particularly Paleoproterozoic terranes (e.g., Fort Simpson arc, Ksituan, Chinchaga and Kiskatinaw terranes).  This interpretation was based on abundant Paleoproterozoic age zircons recovered from the Adams Argillite of eastern Alaska, Atan Group of northeastern British Columbia, the Hamill Group of southeastern British Columbia and the Osgood Formation of Nevada.  These strata are time equivalent and collectively all contain a similar zircon signature dominated by Paleoproterozoic grains (fig. 2.9).  Our data from the Backbone Ranges Formation (samples 4 and 5), however, are dominated by 2060 – 1800 Ma zircons, whereas equivalent units of the southern Cordillera (e.g., Hamill and Osgood groups) are dominated by 1860 – 1730 Ma zircons.  The Osgood Group also contains younger ~1100 and 1450 Ma grains that were likely derived from local Laurentian sources (e.g., ~1100 Ma Pikes Granite and the Mesoproterozoic Granite- Rhyolite province). The Atan Group and Adams Argillite contain Paleoproterozic zircons of similar age to the Backbone Ranges Formation suggesting a similar source for zircons of that age in these strata (fig. 2.9).  However, the Adams Argillite also contains a “Grenville” age signature, possibly derived from recycling of older MMSG strata or from local “Grenville” age plutons that are presently not exposed.  Our lower Backbone Ranges Formation sample (Sample 4) contains two Mesoproterozoic zircons (dated at 1045 and 1264 Ma), which may be from a similar source as zircons of the same age in the Adams Argillite. Therefore, in Early Cambrian time, sediments with a similar  55 Paleoproterozoic source, likely the composite Wopmay orogen with minor sediment recycling, accumulated along the miogeocline north of the Atan Group.  South of this area Early Cambrian units contain detrital zircons with different source regions, possibly comprising basement rock exposed in the Peace River Arch.     56  Figure 2.9. U-Pb detrital zircon normalized age distributions curves for Backbone Ranges Formation equivalent strata from along the length of western North America (the lowest plot represents southern miogeocline strata whereas the top plot represents northern miogeocline strata).  Adams Argillite data from Gehrels et al. (1999); Atan Group and Hamil Group data from Gehrels and Ross (1998); Osgood Mountains Quartzite data from Stewart et al. (2001).  57   Ordovician strata of the miogeocline have been suggested by Gehrels et al. (1995) and Gehrels and Ross (1998) to be anomalous because they show a unique detrital zircon age signauture.  These authors infer that the Peace River Arch was the source for much of the detritus contained within Ordovician strata that are exposed in British Columbia and further south.  Our sample 7, from the Late Cambrian - Ordovician Franklin Mountain Formation, however, contains a very minor component of Paleoproterozoic zircon grains and no zircons with ages 1915 – 1500 Ma, where the former are considered to be typical of the Peace River Arch. It is thus unlikely that the proximal Wopmay orogen or the Peace River Arch was a dominant source for the detritus comprising the Franklin Mountain Formation.  This sample is thus suggested to be unlike age equivalent strata to the south.  Middle Paleozoic strata that are broadly equivalent to the Tsichu Group (our sample 8) have been studied in the Pennsylvanian Spray Lakes Group of southeastern BC (Gehrels and Ross 1998); the Upper Devonian Nation River Formations in eastern Alaska (Gehrels et al. 1999); the Paleozoic Alexander Terrane in southwestern Yukon (Gehrels et al. 1996); and Devonian to Triassic strata of central and eastern Yukon (Beranek 2009).  Collectively all these sedimentary units contain abundant ~430 Ma detrital zircons and are suggested to have been sourced from igneous rocks exposed in the Canadian Arctic.  Therefore, clastic sediments derived from the Ellsemerian orogen is a dominant contributor to sedimentary units from Devonian to Cretaceous age throughout much of the Cordilleran miogeocline in western Canada and Alaska.    58 2.8. ANOMALOUS DETRITAL ZIRCON AGES  The Grenville Province of eastern Laurentia has been a suggested source for 1.25– 1.0 Ga zircons recovered from MMSG strata (Rainbird et al. 1992, 1997).  The Grenville Province also contains an older Meso - Paleoproterozoic tectonic history that spanned 1.80 – 1.25 Ga recorded by major ca. ~1.7, ~1.5, and ~1.45 Ga episodes of magmatism (Gower and Tucker, 1994; Carr et al. 2000; Davidson, 2008).  The ~1.7, ~1.5 and ~1.45 Ga magmatic events in the Grenvile Province outlined above are therefore also potential sources for detrital zircons of this age, and may have been transported along with 1.25 - 1.0 Ga zircons in the postulated Neoproterozoic river systems of Rainbird et al. (1992, 1997) and Rainbird (2007).  These river systems might be expected to have incorporated additional detritus proximal to the Grenville Province, such as sources associated with the Makkovik orogen, the Mesoproterozoic Granite-Rhyolite province of mid-continental United States, Yavapai and Mazatzal orogens, and the Nueltin plutons of the Churchill Province (fig. 2.6).  Collectively we refer to these potential sources as ‘eastern North American.’  2.8.1. Eastern North America sources?   There is abundant evidence in our data for anomalous detrital zircon ages (ages with no obvious western North American sources, e.g., 1800 - 950 Ma) that must have been dispersed from eastern North America. Eastern North American detrital zircons make up a large percentage of all grains analyzed (e.g., 91% of Sample 1 - Katherine Group, 81% of Sample 2 - Little Dal Group, 59% of Sample 3 - Keele Fm., 88% of  59 Sample 6 - Sekwi Fm., 83% of Sample 7 - Franklin Mt. Fm., 75 % of Sample 8 - Tsichu Group, and 49% of Sample 9 - the Cretaceous clastic foredeep). Therefore, understanding the provenance of these grains is critical.  These zircons either accumulated via direct sourcing from the east or were derived from sediment reworking of older strata (e.g., MMSG) and subsequent deposition.  In a compilation (fig. 2.10) of all our anomalous 1800 – 950 Ma detrital zircons from all sedimentary units there are defined peaks that comprise ages 1125, 1245, 1324, 1430, 1479, 1642, 1725, and 1776 Ma.  Local igneous sources for these ages are unknown in western North America; however, possible igneous sources are widespread in the Grenville Province, Makkovik orogen, the Mesoproterozoic Granite-Rhyolite province and the Yavapai and Mazatzal orogens.  Neoproterozoic river systems may have originated from areally restricted eastern North American sources (e.g., Rainbird et al. 1992, 1997; and Rainbird 2007); therefore, it cannot be assumed that all river systems would have carried the same populations of eastern North American grains.  Additionally, deposition of sediments with a unique detrital zircon age signature may have been locally controlled and in turn reworking of these sediments would be expected to produce an equally unique age signature.  Peaks in our 1.8 – 1.0 Ga detrital zircon compilation (see above) are compared with known sources in eastern North America (ages of sources are from Kerr et al. 1992; Gower and Tucker, 1994; Carr et al. 2000; Davidson, 2008).  Examples of these comparisons are as follows; 1125 Ma peak and the ca. 1130 Ma Marcy anorthosite and anorthosite-mangerite-charnockite granites; 1245 Ma peak and magmatism associated with the ca. 1.29 – 1.19 Elzeverian orogeny such as the 1.25 – 1.22 Ga TTG suite and ~1.25 plutonism associated with the Composite Arc Belt; 1324, 1430, and 1479 Ma peaks  60 and the ca. 1.5 – 1.3 Ga mid-continent Granite-Rhyolite province of mid-continent United States and associated 1.52 – 1.30 Ga magmatism observed in Grenville Province (see above); 1642, 1725 and 1778 Ma peaks and widespread ~1647 and ~1720 Ma plutonism of the Makkovik orogen or the 1.8 – 1.65 Yavapai-Mazatzal orogens.  These sources represent candidates for detrital zircon populations from Mackenzie Mountains strata; however, they are certainly not the only possible sources.   Figure 2.10. U-Pb detrital zircon normalized age distribution curves for our Mackenzie Mountains compilation compared with detrital zircon data from west Texas reported by Gleason et al. (2007).   61  To test the hypothesis of a possible eastern North American source connection with the Mackenzie Mountains strata we compare our Mackenzie Mountains compilation with detrital zircon data from sedimentary units of varying ages from eastern North America with well constrained provenance (fig. 2.10).  These include; data from Pennsylvanian sandstones from the Marathon basin of west Texas (Gleason et al. 2007) and from Paleozoic foreland basin deposits of the Appalachian basin of eastern United States (Gray and Zeitler 1997; Becker et al. 2005); Lower Cambrian to Devonian age sedimentary rocks deposited in foreland basins associated with the Taconian and Acadian orogeneis of New England (McLennan et al. 2001); and from modern day rivers draining the Appalachian Mountains (Eriksson et al. 2003).  Collectively all authors suggest that eastern North American sources were the primary source for detritus for all ages of strata. This is evidenced by a large percentage of Grenville 1.25 - 1.0 Ga zircons recovered from these sedimentary units (e.g., 60% and 71% from Mclennan et al. (2001) and Eriksson et al. (2003), respectively). Detrital zircon ages reported by Gleason et al. (2007) from Pennsylvanian sandstones (Haymond Formation) is interesting because the data indicate a largely Laurentian source for Proterozoic detrital zircons, in particular a dominance of sources exposed in eastern North America (e.g., 83%, excluding young Paleozoic Gondwana derived grains).  These authors infer that detrital zircons yielding ages 1300 – 900 Ma were derived from plutonic sources in the Grenville Province; ~1400 Ma zircons were derived from the mid-continent Granite-Rhyolite province; and zircons with ages ~1750 – 1600 to have been dispersed from the Yavapai-Mazatzal orogens.  Owing to the close proximity of these samples to their inferred sources these inferences appear to be reasonable.  These data, therefore, offer a reference for eastern North American derived  62 detritus because a large range of eastern North American sources is reflected in the data (e.g., not only Grenvillian). It is assumed that sediments were initially shed from these sources in the Proterozoic and were reworked into the Haymond Formation in the Pennslyvannian. Characteristic detrital zircon peaks for this reference (a compilation of two samples from Haymond Formation; fig. 2.10; Gleason et al. 2007) comprise ages 1123, 1242, 1330, 1435, 1638, and 1706 Ma which are remarkably similar to peaks defined in a compilation of all data from the Mackenzie Mountains strata (e.g., 1125, 1245, 1324, 1430, 1479, 1642, 1725, and 1776 Ma).  These similarities suggest that anomalous detrital zircon ages from our study were initially dispersed from eastern North America.  Our data, however, contains a higher proportion of ~1330 Ma grains, as well as a 1479 Ma peak, a 1725 Ma peak, and a 1776 Ma peak, which are not present in the Gleason et al. (2007) reference.  These minor differences likely reflect from river systems sampling additional sources north of the Haymond Formation paleo-basin.  For example, detrital zircon data presented by Becker et al. (2005) from Lower Pennsylvanian strata deposited in the Appalachian basin during the Alleghanian orogeny (assembly of the supercontinent Pangea), comprises sediments deposited north of the Haymond Formation of Gleason et al. (2007).  As with the study of Gleason et al. (2007), proximal eastern North American sources for detrital zircons are interpreted for these strata.  A compilation of 1800 – 950 Ma detrital zircon data reported from Becker et al. (2006) (not shown) yield 24 statistical peaks (possibly owing to the size of the data set).  Some characteristic peaks from this data are 1132, 1239, 1324, 1425, and 1649 Ma, which again are also characteristic of the Mackenzie Mountains strata. Also of interest is the fact that the Becker et al. (2006) data contain peaks with ages 1490 Ma, 1727 Ma and  63 1750 Ma, which are present in the Mackenzie Mountains compilation but were not observed in the study of Gleason et al. (2007).  Furthermore, modern rivers draining the Appalachian Mountains were investigated for their detrital zircon signatures (Eriksson et al. 2003) and a compilation of these data (not shown) contain peaks at 1729 and 1761 Ma, which again are present in the Mackenzie Mountains compilation.  Finally, Stewart et al. (2001) analyzed detrital zircons recovered from Cambrian units exposed in the southwest United States.  These authors infer that zircons with ages 1.9 - 1.6 Ga have been dispersed from the Yavapai, Mazatzal and Mojave provinces and that grains with ages 1.45 - 1.4 Ga have been derived from the widespread Mesoproterozoic Granite- Rhyolite province. A compilation of their data (not shown) shows dominant peaks in these age ranges (e.g., 1.45 - 1.4 and 1.6 - 1.9) at 1428, 1732, and 1775 Ma.  These peak ages are statistically identical to peaks that appear in Mackenzie Mountains compilation (e.g., 1430, 1725, and 1776 Ma). These similarities are unique for the northern Cordillera as detrital zircons of these ages are believed to not be found north of Sonora and Nevada (Gehrels et al. 1995).  Zircons with ages 1.6 - 1.4 Ga are present in all but three samples analyzed in the present study.  This span of ages comprises anomalous magmatism in the mid continental United States (e.g., Granite-Rhyolite province) and the North American magmatic gap (NAMG, Ross and Villeneuve 2003).  The NAMG comprises magmatism with ages unknown in western North America. NAMG zircons have been suggested to have been initially dispersed from widespread plutons of this age exposed in Australia (Gehrels and Stewart 1998; Ross and Villeneuve 2003). The Belt-Purcell Supergroup of the Belt basin, of southeast British Columbia and northwestern United States, contains abundant 1.6 –  64 1.4 Ga zircons (e.g., Ross and Villeneuve, 2003) and owing to its relative proximity to the MMSG it may have acted as a potential source. Ross and Villeneuve (2003) speculate that detrital zircons with ages 1480-1400 Ma from the Belt-Purcell Supergroup were dispersed from local syndepositional magmatism of that age. Therefore, owing to the proximity of the Belt basin to the Mackenzie Mountains, it is a possible source for detrital zircons of these ages (1.6 - 1.4 Ga) in the Mackenzie Mountains strata.  To test this hypothesis we compare our Mackenzie Mountains compilation with data presented by Ross and Villeneuve (2003).  The Belt-Purcell Supergroup compilation (not shown) contains dominant detrital zircon peaks between 1.6 – 1.4 Ga with ages at 1445, 1505, and 1605 Ma.  Although some Mackenzie Mountains strata contains zircons with these ages (Table B1), we consider it unlikely that the Belt-Purcell Supergroup was a dominant source of these zircons to the north.  The youngest detrital zircon reported from the Belt Purcell Supergroup is ca. 1377 Ma (Ross and Villeneuve 2003), which is significantly older than potential Grenville Province sources. Therefore if the Belt Purcell Supergroup was the only source for ~1.45 Ga detrital zircons then we should also see dominant 1505 and 1605 Ma peaks in Mackenzie Mountains strata; however, this not the case. Furthermore, in order for a significant quantity of clastic material to reach the paleo- MMSG basin from the Belt Purcell basin a significant amount of longshore transport of detritus would be required.  There is no evidence for northwards longshore transport of sediment to the north in Neoproterozoic time (e.g., time of deposition of MMSG strata) in the stratigraphic record (Aitken and McMechan 1991). Additionally, sediments comprising MMSG strata are believed to be dominantly derived from the east (Rainbird et al. 1996). It is, however, possible that these 1.45 Ga detrital zircon grains were  65 entrained in river systems carrying eastern North American zircons or that there is a possible northern Laurentian connection with Australia and these detrital zircons were, therefore, alternatively shed from Australia.  2.8.2. Unknown proximal sources?  We suggest that the anomalous detrital zircons recovered from the Mackenzie Mountains strata to have been dispersed primarily from eastern North American sources. This is based on the argument that local sources of similar ages with a significant felsic (zircon producing) component (e.g., ~75% of all detrital zircons from units bearing anomalous ages) are nowhere observed in western North America.   We cannot preclude the possibility that proximal sources of these ages may be present but not yet recognized (e.g., Ross et al. 2005; Milidragovic 2008).  However, for these sources to completely reconcile our data they would have to represent contemporaneous magmatism with magmatic events spanning 1.8 – 1.0 billion years observed in eastern North America, and that the Rodinia SWEAT (southwest U.S.-East Antarctic) reconstruction of Moores (1991) be applied to explain the 1.3-1.0 Ga zircons.  Moores (1991) speculate that the Grenville orogen may extend around the coast of East Antarctica into India and Australia when these continental blocks were juxtaposed along western Laurentia.  This reconstruction therefore places Grenville age magmatism and metamorphism along the western margin of Laurentia in Neoproterozoic time. Alternatively, Milidragovic (2008), on the basis of inferred metamorphic zircon crystallization ages (e.g., 1.6, 1.27, and 1.15) apply the “missing-link” hypothesis of Li et al. (1995; 2008) to explain these ages.  This  66 model speculates that oblique collision of the Yangtze Craton with Cathayasia explains the range of Grenville age zircons along the western margin of Laurentia.  These reconstructions, however, cannot explain the abundant older 1.8 - 1.4 Ga anomalous detrital zircons we identified in Mackenzie Mountains strata. Any uplift of the Cordilleran margin could potentially have exposed basement rocks.  Similarly extension related block faulting may expose ancient bedrock.  This interpretation was suggested by Ross et al. (2005) to explain a large component of 1.4 - 1.0 Ga detrital zircons recovered from the Neoproterozoic to Cambrian age Hyland Group of Yukon and Alaska (fig. 2.8).  These authors suggest Windermere rifting and extension caused exposure of an “intra-Cordilleran” Grenville age basement rock which in turn sourced the Hyland Group. If ~1100 Ma igneous rock units (e.g., Jefferson and Parrish 1989; Mortensen and Colpron 1998) and the 1.15 Ga metamorphic events (Milidragovic 2008) are widespread in the basement of the northern Cordillera these rocks may have been exposed periodically throughout the evolution of the miogeocline, and therefore may have dispersed sediments to the west (for a detailed overview of Grenville age tectonism along the length of the northern Cordillera see Milidragovic 2008).  This scenario, however, does not explain the abundant anomalous ~1.43, ~1.47, ~1.64, ~1.72, and 1.77 Ga detrital zircons observed in the Mackenzie Mountains strata.  2.9. PALEOGEOGRAPHIC IMPLICATIONS   Numerous authors have proposed that different continental blocks were juxtaposed against the western margin of Laurentia during the assembly of the Rodinian  67 supercontinent between ~1.0 and ~0.75 Ga.  These include Australia (Hoffman 1991; Dalziel 1991; Moores 1991), Siberia (Sears and Price 2000, 2003), and South China (Li et al. 1995, 2008 and Milidragovic 2008).  On the basis of reconstructions presented by these authors a suitable candidate for a western Laurentian neighbor in the Neoproterozoic is Australia and a northern neighbor may have been the Siberia continent. Australia lying adjacent to western Laurentia in the Neoproterozoic was proposed by Ross and Villeneuve (2003) on the basis of abundant NAMG detrital zircons and eastward directed paleocurrent data in the Belt basin. Australia has abundant 1.61 - 1.49 Ga sources (e.g., Gawler Craton, Blewett et al. 1998) and during assembly of Rodina sediments may have been dispersed to the east from these Australian sources to the paleo- Belt basin.  A northern Cordilleran connection with Australia; however, is not favored on the basis of detrital zircon age data. For example, strata interpreted by Ross and Villeneuve (2003) from the Belt-Purcell Supergroup to contain a large Australian influence are dominated by NAMG zircons (e.g., Revett Formation of the Belt-Purcell Supergroup). Also, Gehrels et al. (1996) report evidence that the Alexander Terrane, of western Yukon and Alaska, was once linked with Australia in the middle Paleozoic.  This was suggested on the basis of abundant NAMG zircons recovered from Paleozoic strata of the Alexander Terrane.  Our MMSG samples, on the other hand contain only a minor NAMG zircon population.  If Australia was juxtaposed against northwestern North America in the Neoproterozoic, one would expect a significant Australian signature (similar to Belt Purcell Supergroup and Alexander Terrane).  Instead MMSG data shows a largely Laurentian signature.  Additionally, in order for NAMG zircons of the MMSG to be directly derived from Austrailia, a crytic pre-MMSG river system would have to  68 have existed on the Rodinian supercontinent shedding detritus eastward.  These sediments would then would have be reworked in westward flowing fluvial systems to reconcile paleocurrent data of Rainbird et al. (1996).  This scenario cannot be entirely precluded as there is little information regarding the nature of the base of the MMSG, which is nowhere exposed. The western margin of northern Larurentia from Neoproterozic time onwards experienced many episodes of tectonic instability marked by periods of extension and subsequent thermal subsidence (Bond and Kominz 1984; Bond et al. 1983; Cecile et al. 1997). As suggested by Ross et al. (2005), a “intra-Cordilleran Grenville” age basement province may have existed along the western margin of Laurentia and was exposed during continental extension and deposition of Windermere Supergroup equivalent strata (e.g., Hyland Group).  This scenario, although not favored in this study, would imply that the amalgamation of supercontinent Rodina between 1.3 – 1.0 Ga is not only in eastern North America but also on the western margin of Laurentia (e.g., Rodinia reconstruction of Moores 1991, see above).  As suggested above, an eastern North American source is viable owing to anomalous detrital zircon age populations of which proximal sources are nowhere observed.  Transporting sediments across Laurentia in Neoproterozoic time was first suggested by Young (1979), Rainbird et al. (1992, 1997) and Rainbird (2007) to explain data from Sequence B strata.  Large scale river systems analogous to those draining the modern day Andes were suggested by these authors.  Additionally, the lack of vegetation and imposing mountain belts in Neoproterozic time on Laurentia supports the presence of migrating large scale braided river systems (Rainbird et al. 1992, 1997; and Rainbird  69 2007).  Our new data from Sequence B (e.g., MMSG) and younger units support these initial hypotheses.  We suggest that eastern North America, including the Grenville Province, was a dominant source for much of the Mackenzie Mountains sedimentary units either directly or by reworking of older units.  This scenario would imply that large braided sand sheets originating from along the length of the Grenville Province were active until Early Cambrian time.  On the basis of Ar-Ar cooling ages and the occurrence of a sub-Cambrian unconformity resting on Proterozoic metamorphic rocks of the Grenville Province, the Grenville Province was characterized by significant topographic relief for ~400 m.y. (Cosca et al. 1991). We can, therefore, conclude that direct Grenville Province sourcing of sediments that were ultimately deposited on the western margin of Laurentia may have persisted until late Neoproterozoic time, coeval with rifting and the deposition of the upper passive margin related Windermere Supergroup strata.  This is evidenced, as suggested above, by a distinct zircon signature observed in Sample 3 (Keele Formation) which shows little evidence exists for recycling of older MMSG strata.  Therefore, following deposition of the Windermere Supergroup (e.g., Cambrian strata), eastern North American age zircons (fig. 2.7) were likely derived from recycling of older strata or possibly from a local exposed and now buried source (e.g., Ross et al. 2005).  2.10. CONCLUSIONS    The temporal evolution of a long-lived passive margin can be very complex, especially for one that experienced many episodes of tectonic instability.  The entire  70 Cordilleran miogeocline is thought to have evolved synchronously from the Paleozoic onwards.  However, this detailed study from the Mackenzie Mountains has shown that local variations are likely to exist during sediment accumulation and deposition in a passive margin setting.  A study of this breadth in the Mackenzie Mountains is possible because of the relatively intact complete succession of strata from Neoproterozoic age to Cretaceous age.  Conclusions from our U-Pb dating study of detrital zircons in the Mackenzie Mountains strata are as follows:  1.  Strata deposited from Neoproterozoic to Cretaceous time in the northern miogeocline contain a clear eastern derivation from Archean and Proterozoic sources exposed in the Canadian Shield.  This is evidenced by pervasive 2080 – 1700 Ma and 2800 – 2415 Ma age groupings in all investigated strata (fig. 2.7).  Local sources in the western Canadian Shield such as the Wopmay orogen and Slave Province are potential candidates for much of the detritus.  2.  There is a major component (e.g., ~75% of all detrital zircons from units bearing anomalous ages) of detrital zircon grains with ages 1.8 – 1.0 Ga (fig. 2.7) that cannot be linked to local northwestern Laurentian sources. All anomalous detrital zircons with ages 1.80 - 1.4 Ga are significant and unique to the Mackenzie Mountains. Zircons of these ages are typically confined to the southern United States component of the miogeocline (exposed in Sonora and Nevada). Although considered unlikely, it is conceivable that local unknown and now unexposed “intra-Cordilleran Grenville” basement may have contributed detritus to the Mackenzie Mountains strata, if such basement ever existed and was exposed.  Additionally, the NAMG (1.61 – 1.49 Ga) zircons that have been identified  71 in our sample suite may have a connection with Australia, since there is little evidence for reworking from the Belt basin. However, further evidence to support a SWEAT reconstruction (Moores 1991) or a “missing link” model favored by Milidragovic (2008) is needed in the northern Cordillera to justify this configuration.  3. “Grenville” age zircon grains (1.4 –1.0 Ga) are pervasive in all stratigraphic units that were investigated (fig. 2.7), except the Backbone Ranges Formation, and can be readily linked to two sources.  These are; (1) Grenville Province of eastern North America for all strata; (2) 1.1 – 1.0 Ga plutons exposed in the Pearya Terrane of western Ellsemere Island (fig. 2.6) for our Carboniferous and younger strata. Owing to strong rounding of “Grenville” age detrital zircons from the MMSG we consider it unlikely that they could have been derived directly from a local source.  Furthermore, sediments shed from proximal “Grenville” sources would be deposited in local basins would be relatively immature deposits.  Conversely, MMSG clastic strata are quite mature (Aitken and McMechan 1991) and the detrital zircons appear to be quite rounded.  4.  Windermere Supergroup strata of the Mackenzie Mountains contain a similar “Grenville” age (and eastern North American) signature; however, it does not result from recycling of older MMSG strata.  It is suggested that 1.4 – 1.0 Ga zircons may have been dispersed from the Grenville orogen until latest Neoproterozoic.  Rifting of Laurentia in the late Neoproterozoic may have exposed local “intra-Cordilleran Grenville” age basement; however, detrital zircons from the Keele Formation (Sample 3) are quite  72 rounded, suggesting substantial physical weathering and long distance transport, possibly in eastern North America derived river systems.  5.  The sub-Cambrian unconformity is not manifested in the detrital zircon signature obtained from overlying younger strata as there is little indication for recycling of sediments.  There is a distinct shift in source in Early Cambrian strata as evidenced by a disappearance of “Grenville” ages and a new dominance of Wopmay orogen Paleoproterozoic zircons as evidenced in the Backbone Ranges Formation.  Middle Cambrian time marks a substantial shift to younger eastern North American detrital zircons with pronounced depletion of Paleoproterozoic Wopmay orogen detrital zircons as evidenced in the Sekwi Formation detrital zircon signature.  This may reflect erosion and reworking of MMSG strata; however, there is an anomalous abundance of 1490 – 1400 Ma detrital zircons, possibly indicating primary fluvial grain size sorting, that we cannot presently account for.  Late Cambrian to Ordovician strata (our Sample 7) are dominated by a “Grenville” age signature but lack other eastern North American (e.g., 1800 – 1500 Ma) zircon grains typical of underlying strata.  6.  Carboniferous and younger strata contain a significant ~430 Ma zircon component. We propose detrital zircons of these ages from the Carboniferous Heritage Trail Formation were either originally deposited in the Upper Devonian and subsequently reworked and redeposited, or they were also derived directly from the same source as the Upper Devonian strata.  It is likely that this source is located in the Canadian Arctic, and may include the mid-Paleozoic plutons presently exposed on Ellesmere Island. Our  73 sample from the Cretaceous clastic foredeep contains a very similar zircon age population as is seen in the Carboniferous Heritiage Trail Formation.  This similarity may reflect significant recycling of strata in the Cretaceous or may indicate that the source region sourcing the Tsichu Group was also exposed during mid-Cretaceous time with additional input from sources further to the west.  7.  We have shown that provenance correlations along the length of the miogeocline should be made with caution as substantial local variation is inevitable.  8. We confirm the initial speculation of Rainbird et al. (1992, 1997) and Rainbird (2007) and propose that the majority of anomalous detrital zircon grains (e.g., 1800 – 950 Ma) contained in strata from northwest North America were initially dispersed from eastern North American sources.    74 2.11. 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INTRODUCTION    Robust geochemical and geochronological studies coupled with field and petrographic investigations of alkaline mafic volcanic rocks can yield considerable information regarding both the igneous petrogenesis and the tectonic evolution of the area through which they occur  (e.g., Downes et al. 1995; Furman 1995; Jung and Masberg 1998; Furman and Graham 1999; Jung and Hoerners 2000; Rogers et al. 2000; Spath et al. 2001; Macdonald et al. 2001 Piercey et al. 2002; Kalmarkar et al. 2005).  Constraining the nature of the mantle source for the magmas has a profound influence on the dynamics of melt generation and magma transport to the surface. It is generally believed that large volumes of melts generated during crustal extension are generated by asthenospheric melting  (e.g., McKenzie and Bickle 1988); however, some workers now suggest that melting of an enriched subcontinental lithosphere with an asthenospheric component may also be responsible for significant alkaline melt generation in some extensional settings (e.g., Jung and Hoernes 2000; Spath et al. 2001).  The stratigraphic association between crustal extension and alkaline magmatism is generally well established in Paleozoic rocks of the Canadian Cordillera (e.g., Goodfellow et al. 1995; Cecile et al. 1997); however, the dynamics of melt generation and its effects on localized rifting has yet to be resolved.  The Misty Creek Embayment (MCE) of central Mackenzie Mountains, Northwest Territories is a lower Paleozoic rift basin developed as a result of at least two discrete rifting events, one in Lower to Middle Cambrian time and another in late Early Ordovician to Middle Ordovician time (Cecile  85 1978; 1982, Cecile et al. 1997). The MCE contains widespread occurrences of alkaline mafic volcanic rocks, termed the Marmot Formation, that were emplaced over a short period of time during the second rifting event.  Intrusive rocks, including dykes and sills as well as a number of intrusive diatremes are associated with the Marmot Formation. The Marmot Formation has received limited petrographic work in this part, largely because of the inaccessibility of the region. The Mountain Diatreme which is one of the largest diatremes associated with the Marmot Formation was discovered in 1973 by Dynasty Explorations Ltd., and became a focus for evaluating the diamond potential of the area.  Regional exploration in the vicinity of the MCE by Archer Cathro, Ltd., in the early 1990’s led to the discovery of many additional diatremes and documentation of associated volcanic and volcaniclastic occurrences (e.g., Copland 1994). Volcano- stratigraphic studies and reconnaissance level petrographic and geochemical investigations of volcanic rocks in the Misty Creek Embayment were carried out by Goodfellow et al. (1980) and Cecile (1982). Results of additional petrographic and geochemical studies specifically focusing on the Mountain diatreme are given by Oldershaw (1977), Goodfellow et al. (1995) and Godwin and Price (1986). Initial age constraints for the Mountain diatreme and associated volcanic rocks were reported by McArthur et al. (1980) and Godwin and Price (1986). A regional synthesis of the nature and geochemistry of Paleozoic volcanic rocks of the miogeocline in northwestern Canada is presented by Goodfellow et al. (1995).  In this contribution we provide geochronological, whole rock and mineral geochemical, isotopic and petrographic constraints on the Marmot Formation and associated intrusive, volcaniclastic and epiclastic occurrences.  These new data are used  86 to constrain the petrogenesis of the Ordovician age alkaline mafic magmatism and to speculate on the nature of the source of the magmas.  We also provide constraints on the Early Paleozoic tectonic evolution of the MCE and more broadly the Canadian Cordilleran miogeocline.  3.2. GEOLOGICAL SETTING   The Early Paleozoic evolution of the western Canadian Cordillera was marked by the development of a passive margin built on the western edge of ancestral North America as a result of supercontinent Rodinia breakup in Neoproterozoic time (Devlin and Bond 1988; Fritz et al. 1991; Gordey and Anderson 1993; Cecile et al. 1997). Variations in stratigraphic thicknesses and anomalous sedimentary facies, combined with the presence of local accumulations of alkaline mafic volcanic rocks suggests renewed extension of the passive margin occurred episodically throughout much of the Paleozoic (Fritz et al. 1991; Goodfellow et al. 1995).  During this time the northern miogeocline shoaled eastward which accompanied the development of expansive carbonate platforms (e.g., Mackenzie, Redstone and Ogilvie platforms).  These platform areas pass to the west to embayments, basins and troughs (fig. 3.1; Fritz et al. 1991).   87  Figure 3.1. Map showing Lower Paleozoic paleogeography with location of the Misty Creek Embayment (modified from Cecile et al. 1997).  Red box represents the bounding contraints of Figure 3.3.   To the south and west of the Mackenzie platform the strata thins to basinal facies siliciclastic rocks deposited in what is termed the Selwyn basin (fig. 3.1; Gordey and Anderson 1993). An irregular rifted basin termed the Misty Creek Embayment (Fig 3.1; Cecile 1978; 1982) occurs along the margin between the Mackenzie platform and the Selwyn basin.  The 100 x 150 km MCE represents a lower Paleozoic basin developed as  88 a result of three superimposed stages of rifting.  The first rifting event coincides with the initial breakup of the supercontinent Rodinia in the latest Neoproterozoic.  This rifting event was marked with the emplacement of voluminous mafic igneous rocks (e.g., the Tsezotene Sills; Ootes et al. 2008) into the Neoproterozoic Mackenzie Mountains Supergroup.  The second rifting event coincides with the development the lower Paleozoic passive margin and marks a time of substantial subsidence and infill of basinal facies strata (Cambrian rifting of Figure 3.2).  This rifting event formed the initial geometry of the MCE in Lower to Middle Cambrian time. The third rifting event was superimposed on the second event and culminated in Upper Cambrian to Lower Ordovician time (Cecile 1978, 1982; Cecile et al. 1997).  As with the second rifting event, considerable subsidence is recorded in the strata; however, this rifting event is also characterized by voluminous submarine alkaline mafic magmatism (Ordovician rifting of Figure 3.2).  These volcanic deposits, which comprise the Marmot Formation, are interstratified with upper Lower and Middle Ordovician basinal shale facies in the center of the MCE.    89  Figure 3.2. Southwest to northeast schematic cross-section through the Misty Creek Embayment showing two classic rift ‘steers-head’ profiles (modified from Cecile et al. 1997).  3.3. VOLCANIC OCCURENCES   Mafic alkaline magmatism of Ordovician age within the MCE can be divided into three distinct suites that are spatially and lithologically restricted. The Marmot Formation comprises two internal suites that we herein term ‘Marmot Formation volcanic rocks’ (MFV) and ‘Marmot Formation volcanic xenoliths’ (MFX). The third suite is restricted to the eastern margin of the MCE and comprises what we herein term ‘diatremes and associated volcaniclastic and intrusive rocks’ (DVI).  All suites collectively belong to the Group 1 volcanic rocks of Goodfellow et al. (1995) and contain well-preserved primary submarine depositional features.  Previous speculations on the correlation of the Marmot Formation (MFV) with the DVI suite have been made by Cecile (1982) and Goodfellow et al. (1995) on the basis of limited geochemical data.  We have confirmed this correlation and herein use the term Marmot Formation to include all three suites.  90   Figure 3.3. Simplified geology map of the study area showing locations of samples of the Marmot Formation and related intrusive and volcaniclastic rocks. Geology from Blusson (1974).  Red dashed line denotes the location for the cross section presented in Figure 3.2.  3.3.1. Marmot Formation volcanic rocks (MFV)  The first suite comprises the Marmot Formation, as described by Cecile (1982), and is mainly exposed in the center of the MCE (fig. 3.3).  The bulk of the MFV suite comprises volcanic and volcaniclastic and epiclastic material that is interstratified with Ordovician siliclastic and carbonate rocks (Cecile 1982).  At the top of the Marmot Formation type section Cecile (1982), however, identified a two-hole crinoid osicle fossil  91 in a limestone bed that is interstratified with volcanic rock and tuffs. The presence of a two-holed crinoid suggests an upper Lower to early Middle Devonian age constraint for the cessation of volcanism. It is possible that volcanism in the MCE persisted to this time; however, the bulk of MCE volcanism is Ordovician in age (Cecile 1982; Goodfellow et al. 1995) and is therefore the focus of this study. The dominant MFV occurrences comprise texturally coherent thick accumulations of mafic volcanic rocks (e.g., sample 1508; fig. D.1b) that overlie basinal facies carbonates of the Middle Ordovician Duo Lakes Formation.  This suite also contains brecciated coherent material such as those deposits interpreted to reflect volcanic vents (e.g., samples 1502 and 1504; fig. D.1a). Additionally, the MFV suite also includes hyaloclastites, brecciated pillowed basalts, amygdaloidal basalts, and stratabound lapilli tuffs (e.g., fig. D.1.f, g, h).  3.3.2. Marmot Formation volcanic xenoliths (MFX)  The second Marmot Formation suite comprises mafic volcanic xenoliths sampled from some MFV occurrence (e.g., sample1505) and is what we herein term ‘MFX.’ These xenoliths range from gravel to boulder size and are all strongly rounded (e.g., fig. D.1c). All xenoliths are petrograhically similar, albeit less altered to the MFV samples and do not contain an identifiable intrusive texture (fig. D.2e, f), thus they likely represent shallow level volcanic material.     92 3.3.3. Diatremes and associated volcaniclastic and intrusive rocks (DVI)  Lithic volcaniclastic breccia filled diatremes (e.g., Mountain Diatreme) and associated volcaniclastic, epiclasitic and intrusive rocks that are restricted to the eastern margin of the MCE comprise our third suite (DVI).  For a comprehensive overview of these deposits see Appendix E.  3.3.4. Petrographic observations  All Marmot Formation occurrences are varibly altered primarily by carbonate ± quartz, chlorite, sericite, and serpentine and in some cases all primary volcanic textures and mineralogy are overprinted (e.g., fig. D.2b,d and fig. E.8). In most cases this alteration is too strong to petrographically classify these rocks on the basis of mineralogy, therefore, the following petrographic observations are primarily focused on samples with an identifiable mineralogy.  Mafic volcanic rocks comprising the MFV suite typically contain euhedral clinopyroxene macrocrysts as well as clinopyroxene and minor olivine microphenocrysts set in a plagioclase, clinopyroxene, analcite (?) and glass bearing groundmass (e.g., fig. D.2a). These samples are of liquid composition with no cumulate textures. Clinopyroxene macrocrysts in the majority of these samples show concentric zoning, classic hour-glass zoning and simple core and rim zoning (e.g., sample 1508; fig. D.3b, c).  Variably resorbed and sieve textured clinopyroxene comprise individual macrocrysts and cores of macrocyrsts with a euhedral rims are also evident in most samples (e.g., fig. D.3a, b).  The question of whether these macrocrysts are phenocrysts or xenocrysts will be addressed in a later section.  There are also intergranular mafic  93 volcanic rocks that make up a large portion of the MFV suite (e.g., fig. D.2d). Petrographically the majority of these mafic volcanic occurrences are tephrites.  Mafic volcanic rocks of the MFV suite also include local volcaniclastic equivalents which consist of lapilli sized juvenile clasts displaying the typical MFV mineral assemblage, with a carbonate and lithic clast dominated interclast matrix (e.g., fig. D.2c).  Spinel, clinopyroxene and local phlogopite comprise identifiable fresh mineral phases in these deposits.  The occurrence of phlogopite will be discussed in a later section.  Many of the MFV occurrences are pervasively altered to chlorite, sericite and carbonate with clinopyroxene ± olivine macrocrysts pseudomorphed by carbonate making the majority of these deposits are difficult to petrographically classify. Xenoliths of the MFX suite contain glomeroporphyritic textures (fig. D.2f) that comprises clinopyroxene phenocrysts set in a plagioclase, clinopyroxene, analcite (?) and glass bearing groundmass.  The MFX suite is for the most part less carbonate altered than the MFV suite (e.g., fig. D.3e, f). Detailed petrography of the DVI suite is presented in Appendix E.  The majority of the diatreme occurrences of the DVI suite are pervasively altered to carbonate and chlorite and contain a high proportion of country rock xenoliths with affinity to the host Ordovician Franklin Mountain Formation.  In some cases, crystal shapes are preserved in juvenile magmaclasts and therefore speculations on the compositions of primary magmas can be made (e.g., melilitites and alnoite; see Appendix E).      94  3.3. SAMPLING METHODS  The Marmot Formation was examined and sampled in 2007 at seventeen different localities throughout the MCE.  This work was done during regional 1:250,000 scale bedrock mapping in collaboration with the Northwest Territories Geoscience Office. Samples were submitted to ALS Chemex laboratories in North Vancouver, B.C., for whole-rock geochemical analysis.  Ar-Ar age determinations and Nd isotopic analyses were performed at the Pacific Center for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia.  Electron microprobe analyses were conducted at the Electron Microbeam / X-Ray Diffraction Facility (EM-XDF) at the University of British Columbia.  3.4. ANALYTICAL RESULTS  3.4.1. Mineralogy and Mineral Chemistry  The mineralogy and mineral chemistry of several key mineral phases in the Marmot Formation were examined in order to shed light on the magma genesis.  Mineral chemistry for clinopyroxene is presented in Table C.1, spinel chemistry is presented in Table C.2, and phlogopite-biotite chemistry is presented in Table C.3.  Analytical procedures are reported in Appendix A.    95  3.4.1.1. Clinopyroxene   Fresh clinopyroxene is typical of MFV samples, especially those comprising the massive CPX phyric mafic volcanic flows exposed near the center of the MCE (e.g., sample 1508).  Mineral chemistry studies were concentrated on CPX crystals from these occurrences as they offer distinct insight into two stages of CPX growth (cores and rims). In most MFV occurrences (e.g., sample 1508) CPX grains are typically the most abundant and freshest primary mineral; therefore microprobe analyses on these grains was used to offer constraints on magmatic evolution.  All CPX cores and rims analyzed from the massive mafic volcanic rocks (sample 1508) are classified as diopsides (fig. 3.4d). Cores contain variable TiO2 (0.83-1.94 wt.%) and Cr2O3 (0-0.91 wt %), whereas these contents in rims are 2.14-3.32 wt.% and 0-0.16 wt.% respectively (table C.1). Al2O3 contents in cores are low (2.8-5.03 wt.%), whereas rims yield contents higher Al2O3 contents (4.83-7.44 wt.%). Mg numbers (100*Mg/(Fe+2 + Mg)) for cores (79.9-86.8) are higher than Mg numbers from the rims (72.3-79.0).  In a plot of Al2O3 vs. SiO2 (fig. 3.4a) used to constrain crystallizing magma alkalinity, cores are characterized by chemical compositions reflecting alkaline crystallization conditions (high SiO2, low Al2O3) whereas rims reflect peralkaline crystallization conditions (low SiO2, high Al2O3).   96  Figure 3.4. Plots of clinopyroxene chemistry from sample 1508. (A) SiO2 vs. Al2O3 with alkalinity fields from Nisbet and Pearce (1977). (B) Al3+ vs. Ti4+ showing Al/Ti ratios. The 2.6 line denotes pressure of crystallization of ~30kbars (from Lloyd 1981). (C) Al-Cr-Na plot with clinopyroxene equilibrium fields from Morris et al. (2002).  (D) Ternary classification of pyroxenes.  3.4.1.2. Spinel   Spinel grains are abundant in all samples examined from the Marmot Formation. Typically they are small euhedral grains, contained within typical groundmass assemblages (e.g., samples 1494 and 1498). Interestingly, spinel grains are also observed as inclusions in carbonate pseudomorphs after olivine macrocrysts (e.g., samples 1475 and 1502, fig. D.3f).  We analyzed spinel grains from a heavy mineral separate of sample  97 1502 which likely includes grains from both the groundmass and from within olivine inclusions.  These grains have variable MgO (15.13-18.22 wt.%), Al2O3 (18.65-38.32 wt.%), TiO2 (0.57-1.18), and chromium numbers (100*Cr/(Cr+Al)) ranging from 31.99 to 63.10.  A plot of our analyses together with spinel compositions from previous studies (Copland 1994) (fig. 3.5a) indicates that the spinels are volcanic related and are not of peridotitic affinity, and fall near the OIB and within the MORB field of Kamenetsky et al. (2001).  In a plot of chromium number against magnesium number (from Kamenetsky et al. 2001) our spinel analyses form an array of increasing Mg number with decreasing Cr number, suggesting these grains were in equilibrium with olivine with a forsterite content between 89-90% (fig. 3.5b).   Figure 3.5. Plots of spinel chemisty from sample 1502. (A) TiO2 vs. Al2O3 with spinel crystallization discrimination fields from Kamenetsky et al. (2001).  (B) Cr-number vs. Mg-number (Cr/Cr+Al vs. Mg/Mg+Fe) with black lines representing the forsterite content of coexisting olivine (from Kamenetsky et al. 2001).  Historic data from early mineral assessment work (e.g., Copland 1994)    98 3.4.1.3. Phlogopite-biotite   Mica grains were analyzed from two occurrences (samples 1494B and 1494F from the DVI Mountain Diatreme, and sample 1504, a phlogopite and CPX phyric fragmental basalt belonging to the MFV suite).  In the Mountain Diatreme samples, juvenile clasts are typically cored by large phlogopite macrocrysts (up to 2 cm in width) and small grains are also present in the groundmass assemblages of both the altered juvenile clasts and the interclast matrix.  Sample 1504 contains large (up to 0.5cm) mica grains as macrocrysts as well as small grains interspersed throughout the groundmass. All macrocryst grains are remarkably fresh with no internal alteration or zoning; however, all grains are typically subhedral and show resorbed grain boundaries (fig. D.3d, f). We analyzed macrocrystic and groundmass phlogopite from all three samples.  In a conventional mica Mg-Al-Fe plot (fig. 3.6a), all our grains are classified as phlogopites (macrocrysts in both samples) and biotites (groundmass grains in both samples).  Macrocrysts from the Mountain Diatreme (e.g., DVI samples 494 B,F) are aluminous (15.92-16.6 wt.% Al2O3) with high TiO2 (4.06-4.7 wt.%), variable FeO (6.44- 7.50 wt.%), low Cr2O3 (0-0.19 wt.%) and Mg# ranging from 82.7 to 84.6.  Groundmass grains from the Mountain Diatreme are geochemically distinct from their macrocryst counterparts.  These groundmass grains are aluminum poor (14.4-14.8 wt.% Al2O3) with low TiO2 (2.42-3.17 wt.%), rich in FeO (15.86-16.99 wt.%), are low in Cr2O3 (0-0.045 wt.%) with Mg# ranging from 58.1-62.7 (fig. 3.6b).  Macrocryst mica data from MFV sample 1504 indicate highly variable Al2O3 (11.90-15.90 wt.%) and Cr2O3 (0-0.26 wt.%) contents, high TiO2 (3.94-6.03 wt.%), and variable Mg# varying between 55.9 and 78.9.  99 Groundmass micas from this sample are titanium poor (0.02-2.64 wt.% TiO2) with high Cr2O3 (0-0.4 wt.%) contents and Mg# ranging from 59.2 to 64.2 (fig. 3.6b). Our phlogopite data fall near the mafic volcanic field and are interpreted to represent non- peridotite affinity    Figure 3.6. (A) Ternary classification of micas from samples 1494B, F and 1504B (B) Si4+ vs. Mg-number (Mg/Mg+Fe) plot of phlogopite chemistry. Dashed black line denotes the garnet spinel transition and fields from Arai (1984).   3.4.2. Whole Rock Geochemistry   Whole rock geochemical analysis were determined for a representative suite of samples from the Marmot Formation. Lithogeochemical data is presented in Table C.4 and analytical procedures employed in the study are reported in Appendix A.   100 3.4.2.1 Alteration   An attempt was made to select the least altered samples for lithogeochemical analyses.  On the basis of field and petrographic observations; however, it is inferred that low-grade seawater and carbonate alteration has affected the majority of these samples to some extent.  Petrographic studies indicate that the typical alteration minerals apparent in most Marmot Formation samples are clay minerals, carbonate, chlorite, serpentine, and minor quartz.  Also, because the volcanic occurrences are hosted within regional carbonate strata, carbonate derived fluids and secondary volatiles were fundamental in these alteration processes.  Therefore, major elements (e.g., SiO2, Na2O, K2O and CaO) and LILE (large ion lithophile elements; e.g., Cs, Rb, Ba, and Sr) concentrations in these samples are suspect owing to the mobility of these elements during low grade alteration (e.g., Maclean, 1990; Piercey et al. 2002).  Furthermore, the high LOI (loss on ignition) values from these rocks, ranging up to 34 wt. %, support the likelihood of elemental mobility.  However, major elements Al2O3, TiO2, and P2O5 and HFSE, REE and Th were likely relatively immobile during low grade alteration processes (e.g., Pearce and Cann 1973; Winchester and Floyd 1977; Wood 1980, Jenner 1996; and Piercey et al. 2002). To test that these elements remained immobile we have plotted key compatible elements and compatible and incompatible element ratios used in this chapter against the Al2O3/Na2O alteration index (fig. 3.7) of Spitz and Darling (1987).  It is clear that for the most part there is little to no correlation between these key elements and the alteration index indicating that these elements were likely immobile during the aforementioned conditions.  There is, however, possible correlations of some ratios with the alteration  101 index within the DVI suite (fig. 3.7f, g, h), indicating that these elements were likely mobile and are suspect in the DVI suite.  The lack of correlation within the MFV and MFX suite permits us to use these elements only for discussion on the MFV and MFX suite.   We base our interpretations of primary igneous geochemistry of the sample suite primarily on immobile elements to avoid element mobility and erroneous interpretations. We use Zr as our fractionation monitor for volcanic evolution, because Zr is usually relatively immobile during alteration and its concentration will likely reflect primary volcanic fractionation trends (e.g., Piercey et al. 2002).  Elements that do not define a linear trend with Zr are interpreted to have been mobile and to reflect non-volcanic processes.  102  Figure 3.7. Key compatible elements, compatible element ratios, and incompatible element ratios used throughout this chapter against the Al2O3/Na2O alteration index of Spitz and Darling (1978).  103 3.4.2.2. Major Elements   As expected on the basis of field observations, LOI values are high for the DVI samples (e.g., 8.91 – 34.0 wt. %) and likely reflect eruption through a carbonate platform, resulting in carbonate alteration, secondary volatiles, and unavoidable carbonate lithic fragments.  There is a limited range in some major element contents within the DVI suite (Al2O3, 3.66–10.1 wt. %; TiO2, 0.73-2.95 wt. %), suggesting little mobility of these elements.  Some major elements on the other hand show a substantial compositional range (e.g., MgO, 4.95-23.54 wt. %; CaO, 5.45-41.71 wt.%), which likely reflects element mobility and alteration processes.  Owing to the high LOI values, and Al2O3/Na2O ratios >25 (see Piercey et al. 2004; table C.4), we do not utilize major elements for characterizing the DVI petrogenesis.  Furthermore, there is a wide range in compositions of Marmot Formation samples in a total alkalis vs. silicate (TAS) diagram (fig. 10a) suggesting lithologic classifications such as basaltic andesites to picro-basalts. On the basis of petrography we can preclude this diagram and the use of major elements for lithologic classification of the Marmot Formation.  MFV and MFX samples have quite similar major element contents (table. C.4) and therefore will be reported together. LOI values for these samples are somewhat lower than the DVI samples (1.46-13.55 wt. %) possibly reflecting eruption through basinal strata (e.g., shales) rather than carbonates, as well as lower degrees of seawater alteration and minimal country rock xenoliths. Like the DVI samples these samples exhibit a narrow range of some elements (Al2O3, 10.28-19.18 wt. %, TiO2, 1.53-4.32 wt. %), but display less variable mobile major elements contents (e.g., MgO, 0.860-13.9 wt. %, CaO,  104 4.11-17.79 wt. %).  Normative mineralogy is not reported in this paper due to the nature of major elements during alteration.  However, we suggest that TiO2 remained relatively immobile during alteration, as evidenced by its narrow range of contents among all samples. We therefore use TiO2 as petrogenetic indicator (e.g., fig. 3.10b).   Figure 3.8. Chondrite normalized rare earth element plots. Chondrite values from Sun and McDonough (1989).     105 3.4.2.3. Trace Elements  Marmot Formation volcanic rocks are characterized with basaltic affinity (Zr/TiO2 = 50.49-208.22) and are moderate to extremely alkalic (Nb/Y = 2.57–7.44). They can therefore be geochemically classified as alkali basalts to foidites (fig. 3.10b; Peace 1996).  Rare earth element (REE) normalized patterns (fig. 3.8) for all Marmot Formation samples are quite similar with enrichment in light REE (LREE) and depletion in the heavy REE (HREE).  The DVI samples exhibit a greater LREE enrichment than the MFV and MFX samples (e.g., average La/Yb; 79.4 and 29.6 respectively).  Two samples (1501D and 1505E) exhibit a negative Eu anomaly, whereas, one sample (1500A4) is characterized with a positive Eu anomaly.  Primitive mantle normalized diagrams are again quite similar (fig. 3.9).  All but one sample (1496B) exhibits a weak to strong positive Nb anomaly.  DVI samples exhibit a pronounced negative Zr anomaly, whereas the MFX and MFV samples are characterized with mainly positive Zr anomalies.  Ti anomalies are typically weakly to strongly negative in the DVI samples, weakly positive in the MFV samples and absent in the MFX samples.  The Th-Zr-Nb plot (fig. 3.10d) of Wood (1980) confirms that all samples are characteristic of an OIB source. All samples fall within or near the alkaline array on a Ti-V plot (fig. 3.9c) of Shervais (1982) with the exception of sample 1501B, owing to the low TiO2 content of this sample.  106  Figure 3.9. Primitive mantle normalized incompatible element plots. Primitive mantle values from Sun and McDonough (1989).  107  Figure 3.10. (A) Total alkalis vs. silicate (TAS) diagram for volcanic rocks after Le Bas et al. (1986). (B) Zr/TiO2 vs. Nb/Y diagram (Pearce 1996). (C) Ti vs. V discrimination diagram of Shervais (1982). ARC, arc-related basalt; BABB, back- arc-basin basalt; BON, boninite; IAT, island-arc tholeiite; LOTI, low-Ti tholeiite; MORB, mid-ocean-ridge basalt; OFB, ocean-floor basalt; OIB, ocean-island basalt. (D) Th-Zr-Nb tectonic discrimination diagram of Wood (1980).  3.4.3. Isotopic and Dating Studies   Neodymium isotopic compositions and 40Ar/39Ar age determinations were made on a suite of samples from the Marmot Formation in order to further characterize the suite and to place age constraints on this magmatism.  Neodymium compositions are presented in Table C.5, and 40Ar/39Ar geochronological data are presented in Table C.6. Analytical procedures are reported in Appendix A.  108  3.4.3.1 Neodymium Isotopic Composition   The Nd isotopic compositions were determined for a total of 13 Marmot Formation samples. Initial !Nd values were calculated for 455Ma (average eruption age, see below), and range from +2.7 to +8.7, corresponding to initial 144Nd/143Nd values of 0.51220 – 0.51251.  The DVI samples show the lowest !Nd values (+2.7-+3.2) whereas the MFV and MFX samples show higher values of +3.9 to +8.7.  The latter samples approach the depleted mantle evolution curve on an !Nd vs. time plot (fig. 3.12b). There is an intermediate cluster of analyses (6 of 13 total analyses) that comprise an average !Nd of +6.0. Our Marmot Formation samples yield values that are unlike any potential contaminants (fig. 3.12b), and therefore likely reflect primary mantle source isotopic values. Depleted mantle model ages (tDM) range from 455 Ma (sample 1508B) to 870 Ma (sample 1498B), with MFV and MFX samples yielding model ages between 620 and 781 Ma.  3.4.3.2. 40Ar/39Ar phlogopite geochronology   We use 40Ar/39Ar geochronological methods on small phlogopite macrocrysts from the Marmot Formation to constrain the timing of volcanism.  An important question is whether these phlogopite grains represent mantle derived xenocrysts or crystallized from the evolving magma and are therefore phenocrystic.  Mantle derived xenocrystic phlogopite might not record the age of volcanism whereas comagmatic phlogopite would.  109 On the basis of petrography and mineral chemistry, phlogopite grains appear to be phenocrystic as there are no xenocrystic cores and grains are remarkably unaltered. There are, however, minor rims on a few macrocrysts which are likely result of alteration with the evolved melt.  This is suggested on the basis of grain boundary relationships with the groundmass as observed in back scattered electron (BSE) images, (fig. D.3d). These rims (or alteration fronts; light grey in BSE images, fig. D.3d), are also observed along grain fractures and along mica cleavage planes that have been separated, thus suggesting late chemical alteration of exposed grain edges.  The argon closure temperature in phlogopite is low (~480oC, Giletti 1974); therefore the argon isotopic system in xenocrystic phlogopite would be reset by the magmatic temperatures and thus even xenoocyrsts would likely record the age of magmatism  Eight phlogopite or biotite grains were selected from eight separate Marmot Formation occurrences from locations throughout the MCE were dated using 40Ar/39Ar step-heating methods. Not all Marmot Formation deposits contain mica, thus a bias towards mica bearing occurrences is unavoidable. Micas were picked from hand samples, the smallest grain size attainable were selected in most cases as these produce the most robust ages (e.g., degas inherited argon more readily and produce reliable emplacement ages; Kelley and Wartho 2000; Downes et al. 2006).  A hindrance in many 40Ar/39Ar geochronological investigations of kimberlites is the appearance of excess argon (e.g., when the concentration of 40Ar is greater than that produced by argon decay).  Usually saddle shaped profiles in conventional age vs. cumulative 39Ar percent diagrams are an indication of excess argon (for a review see Kelley 2002) and only an estimation of emplacement age can be interpreted (e.g., Kaneoka and Aoki 1978).  110  Figure 3.11. 40Ar-39Ar age spectra for phlogopites from selected samples from the Marmot Formation and related intrusive and volcaniclastic rocks.   111 Ages from all eight samples are based on statistically valid plateaus (fig. 3.11). The age data appear to define three distinct clusters, one is at 459.6 ± 2.4 Ma to 460.6 ± 2.6 Ma (defined by four ages), another is at 455 ± 2.5 Ma and 451.9 ± 2.6 Ma (defined by two ages), and a third at 444.4 ± 2.6 and 444.4 ± 2.5 Ma (defined by two ages).  Excess argon is not readily identifiable in our data, thus these results are interpreted to represent ages of igneous intrusion of eruption.  Implications of these ages are discussed in a later section.  3.5. INTERPRETATION OF MINERAL DATA  3.5.1. Clinopyroxene   Clinopyroxene in mafic igneous rocks has been shown by previous authors to provide geochemical clues into the mantle source and magmatic evolution of basalts (e.g., Nisbet and Pearce 1977; Liotard et al. 1988; O’Brien et al. 1988; Dobosi et al. 1991; Bindi et al. 1999).  This is in part due to the wide range of chemical substitutions into the crystal structure of CPX and its overall resilience to late stage alteration. The observed zoning in CPX grains from massive mafic volcanic rocks (e.g., MFV sample 1508) likely represents typical CPX fractionation as evidenced by higher contents of Mg, Si, and Cr in the core and lesser contents of Fe, Ti and Al than in the rims. Petrographic observations show that this transition from core to rim is sharp (fig. E.1), possibly suggesting there was a rapid change in pressure and temperature conditions during magmatic evolution.  Furthermore, Cr has a high CPX/melt distribution coefficient  112 (Dodosi et al. 1991) and will therefore be enriched in CPX during early stages of crystallization (e.g., cores) at higher pressures.  Additionally, Ti accommodation in CPX is favored at low pressure and temperature conditions; thus Ti contents should be higher in rims of crystallizing CPX (Dobosi et al. 1991).  Many grains display slightly resorbed cores surrounded by euhedral rims.  The resorbed nature of these cores may be a result of unstable conditions related to changing pressure and temperature conditions.  This scenario was also suggested for CPX grains with similar textures from alkali basalts in Austria by Dobosi et al. (1991).  A simple CPX crystallization evolution is needed for the massive mafic volcanic rocks owing to the remarkable linearity in the data on variation diagrams (fig. 3.4a,b). Any deviation from this trend would record different crystallizing magma conditions or could reflect magma mixing. However, these grains appear to record simple CPX fractionation with two stages of crystallization at unique pressure and temperature conditions from the same evolving alkaline to peralkaline magma (fig. 3.4a). Based on consistency of mineral chemistry from sample 1508, as well as the euhedral shape of rims and the significant lack of other phenocrystic phases within this sample, we interpret that the host basalt reflects magma from which the CPX crystallized, and that none of the CPX grains are xenocrystic in origin. Additional study is required to confirm this interpretation since CPX bearing nodules have been observed in some samples.  A plot of Al/Ti vs. Al+Ti has been used (e.g., Lloyd 1981) to constrain depths of CPX crystallization (fig. 3.4b).  This is based on experimental work (e.g., Edgar et al. 1976) in which these authors grew synthetic CPX crystals (as well as olivine, phlogopite and minor ilmenite) at different temperatures and pressures from an ultra-potassic starting material from the East African rift.  The synthetic CPXs that crystallized during  113 experiments of Edgar et al. (1976) are remarkably similar to CPXs from the Marmot Formation (e.g., Ca/(Ca+Mg) > 0.5 and low TiO2 contents (1.02-1.31 wt. %, similar to our cores).  These authors suggest that Al/Ti ratio from CPXs of 2.6 marks a pressure of crystallization at around 30 kbars and temperature at around 1250oC, and that there is a positive correlation between the Al/Ti ratio and the pressure and temperature conditions during crystallization.  Our results from the massive mafic volcanic rocks (sample 1508) yield low average Al/Ti ratios near the 2.6 mark (rim; 3.10, core; 4.87). The higher Al/Ti ratio of cores compared to rims reflects a higher pressure during crystallization for the cores than that of the rims.  However, both values are near the 2.6 mark (fig. 3.4b); therefore, CPX fractionation may have occured at pressures around 30 kbars. Furthermore, on a ternary Cr-Al-Na plot, Morris et al. (2002) distinguished between CPX compositions from mantle peridotite xenocrysts and those of non-peridotite affinity.  Our data fall in the crust equilibrated field (fig. 3.4c), suggesting that these CPX crystals crystallized from a melt and were not in equilibrium with peridotite and are not xenocrystic.  3.5.2. Spinel  As with CPX, the composition of spinels in volcanic rocks can provide constraints on the melt chemistry from which they crystallized.  Our spinel compositional data coupled with spinel compositions previously reported from the area (e.g., early assessment work by Copland 1994), indicates that all spinels from the Marmot Formation are of volcanic origin (e.g. fig. 3.5a; TiO2-Al2O3 plot of Kamenetsky et al. 2001) and are  114 typical of spinels derived from an enriched source (e.g., OIB).  Additionally, our spinel results suggest that spinels and crystallizing magma coexisted with primitive olivine with forsterite content ranging from 89 to 90 % (fig. 3.5b).  Although equivocal, owing to the lack of fresh olivine, such olivine compositions (Fo>88) is typical of primitive melts (Frey et al. 1977).  3.5.3. Phlogopite-biotite  As noted by Arai (1984), aluminum contents of phlogopite from ultramafic rocks is controlled fundamentally by the equilibrium pressure and as such can be used as a geobarometer.  The average Al2O3 contents of groundmass phlogopite from sample 1494 is 14.62 wt.% which is similar to Al2O3 contents from the alteration rims on macrocrystic phlogopite (14.97 wt.%), suggesting that groundmass phlogopite was in pressure equilibrium with these rims.  Other than Al2O3 the chemistry of the rims is unlike groundmass assemblages (e.g., lower Mg #, higher TiO2 contents), suggesting that formation of the rims was not coevel with groundmass assemblages.  The rounding of individual grain boundaries may reflect a late stage chemical phenomenon in the highly volatile melt in which they were transported to the surface.  Closer inspection of these grain boundaries in BSE images (fig. D.3d) show that for the most part, rounding appears to be chemically destructive (e.g., erosion of grain edges giving the grain margins a jagged and irregular appearance).  There may also be an earlier phase of mechanical rounding as evidenced by some grains.  115 Utilizing phlogopite geochemical data from phlogopite bearing peridotite xenoliths has been used to provide constraints on the formation of phlogopite magacyrysts observed in kimberlites (e.g., Dawson and Smith 1975).  These methods provide a valuable tool for identifying the origin of Marmot Formation phlogopite macrocyrsts used in 40Ar/39Ar dating methods (fig. 3.11).   Dawson and Smith (1975) attribute low TiO2, Cr2O3, and MgO contents of micas to reflect magmatic crystallization and not from peridotic sources.  Canil and Scarfe (1989) report phlogopite geochemistry from grains recovered from peridotic xenoliths entrained in Pliocene to Pleistocene age basalts located eastern British Columbia.  These data represent mantle derived phlogopite geochemistry and comprise high Cr2O3 (~ 1.25 wt.%), Al2O3 (~ 15.68 wt.%) and TiO2 (~ 5.68 wt. %) contents with Mg#s varying between 81.0 and 89.8.  The highest Cr2O3 phlogopite content from the Marmot Formation is 0.4 wt.% and the average of all analyses is 0.07 wt.%, which are significantly lower that the minimum value reported from the peridotite xenoliths of Canil and Scarfe (1989). Additionally, our TiO2 and Al2O3 contents (~3.26 wt.% and ~14.74 wt.%) are lower than those values reported by Canil and Scarfe (1989).  There is, however, a Ti rich phlogopite from sample 1504 and probably reflects the high TiO2 contents of the host rock (table C.4).  Moreover, our phlogopite data, from all suites, are characterized with low Mg# (average 70.2) and are low in SiO2 contents (average 36.5 wt.%).  These observations indicate that they were not derived from garnet peridotite which is characterized by high Mg# (84 – 95) and high SiO2 (fig. 3.6b, fields from Arai 1984).  Furthermore, our low Mg#’s are additionally unlike mica derived from spinel peridotite but are most similar to phlogopite derived from mafic volcanic rocks (e.g., fig. 3.6b; lower Mg#). Al2O3 contents from phlogopite in  116 spinel peridotites and from rims of kimberlitic phlogopite are generally higher than those in garnet peridotite (Arai 1984) owing to a lower pressure of crystallization.  Marmot Formation phlogopite from the Mountain Diatreme (e.g., 1494 B and F) on average have high Al2O3 contents (~ 15.23 wt. %) compared to garnet peridotites (~13.32 wt.%, data from garnet lherzolite reported in Dawson and Smith 1975; Delaney et al. 1980). Therefore, macrocrystic phlogopite from the Marmot Formation did not crystallize in equilibrium with garnet peridotite (the inferred location of melt genesis).  We suggest that the phlogopite grains from the Marmot Formation crystallized at shallower depths during the evolution of the magma and are not xenocrystic.  3.6. INTERPRETATION OF GEOCHEMISTRY  3.6.1. Crustal Contamination   Crustal contamination is commonly expected for mantle-derived magmas that erupt through continental lithosphere (Wendlandt et al. 1995) and may limit interpretations on magmatic evolution.  Therefore evaluating the nature and extent of contamination, if any, is critical.  Crustal contamination is expected when ascending magma interacts with and assimilates melted country rock.  None of our Marmot Formation samples contain xenoliths of unknown sources (e.g., basement rock xenoliths or mantle inclusions are completely absent), suggesting little interaction with lithospheric contaminants early in the magma ascent.  Additionally, the very low average Th/Nb and Zr/Nb ratios (0.063-0.145 and 0.767-4.39 respectively) for the Marmot Formation are inconsistent with typical bulk continental crust values (0.32 and 9.0 respectively; values  117 from Taylor and McLennan 1985), again suggesting minimal contamination.  Similarly, all of our samples have higher average Nb/Ta and Zr/Hf ratios, (19.5 and 39.5 respectively) compared to typical bulk crustal values of 11 and 33.3 respectively (bulk continental crust data Taylor and McLennan 1985).  Furthermore, high Ti/Yb ratios is considered to be strong evidence against crustal influence (Wendlandt et al. 1995 and references therein).  Marmot Formation rocks show average Ti/Yb ratios of 8805 which is somewhat higher than continental crust values of ~2455 (Taylor and McLennan 1985) but are more similar to average OIB values of ~7140 (OIB data from Sun and McDonough 1989). The majority of the Marmot Formation samples are characterized by positive Nb (with the exception of two DVI samples) in extended incompatible element plots (fig. 3.9).  Positive Nb anomalies are typical of OIB magmas (e.g., fig. 3.12a) and are not characteristic of continental crust (fig. 3.12a). Finally, positive !Nd values for all Marmot Formation samples analyzed suggests that continental crust was not a contributing component of the source.  We conclude, therefore, that the Marmot Formation was not significantly contaminated by continental lithosphere during its evolution.        118  Figure 3.12. (A) Primitive mantle normalized incompatible element plots showing the envelop of data from all studied samples of the Marmot Formation and average compositions of OIB, N-MORB (data from Sun and McDonough 1989) and continental crust (data from Taylor and Mclennan 1985). (B) !Nd vs. time plot with Nd evolution curves of depleted mantle, continental crust and CHUR.  Also plotted are potential crustal contaminants (data from; 1, Garizone et al. (1997); 2, Dudas and Lustwerk (1995); 3, Rainbird et al. (1997); 4, Villeneuve et al. (1991). Symbols as in Figure 3.7.   3.6.2. Marmot Formation petrogenesis     There are minor geochemical variations amongst the Marmot Formation samples analyzed in this study. For example, Al2O3/TiO2 values range from 2.92-5.18 for the DVI samples, and 2.81-8.19 for the MFV and MFX samples.  These values are significantly lower than typical primitive mantle ratios (~22) or normal mid-ocean-ridge basalts (N- MORB ~9.4) but are more similar to typical ocean island basalts (OIB ~5; Sun and McDonough 1989; Karmalkar et al. 2005). As suggested by Jung and Masberg (1998), this lack of variation implies that the geochemical composition of all samples is largely controlled by mantle processes or mantle compositions, and that fractional crystallization in high-level magma chambers (e.g., plagioclase fractionation, see below) and/or country  119 rock contamination were not important.  3.6.2.1. Fractional Crystallization   Clinopyroxene is the most dominant phenocryst phase in some samples (e.g., MFV and MFX samples), suggesting that it was an important constituent in the evolution of the melt.  Mineral chemistry of clinopyroxene macrocrysts indicates two stages of crystal growth at different pressures and temperatures, and suggests simple clinopyroxene fractionation (see above).  Additional evidence for clinopyroxene fractionation is indicated by a negative slope in a plot of Cr vs. Zr (fig. 3.13b).  This plot suggests that spinel was also a fractionating phase.  The ultramafic nature of the Marmot Formation and the presence of olivine macrocrysts pseudomorphed by carbonate in DVI samples also indicates olivine fractionation.  The negative correlation depicted in a plot of of Ni vs. Zr (fig. 3.13a) is consistent with olivine fractional crystallization.  The absence of prominent negative Eu anomalies in REE plots (fig. 3.8) suggests that low-pressure plagioclase fractionation has not taken place.  These features together show that an early crystal fractionation process involving clinopyroxene, spinel and olivine largely controlled the petrogenesis and evolution of the Marmot Formation.   120  Figure 3.13. (A) Ni vs. Zr plot and (B) Cr vs. Zr plot showing the incompatible nature of these elements. Inverse relationships between Zr and Ni and Cr are consistent with olivine, and clinopyroxene + spinel fractionation respectively. Symbols as in Figure 3.7.  3.6.2.2. Partial Melting   Trace element ratios with similar degrees of incompatibility on the numerator and denominator, such as Hf/Zr, La/Sm, Nb/Th, and Yb/Tb can be used qualitatively to investigate degrees of partial melting amongst a single igneous suite (e.g., Jung and Masberg 1988).  These ratios are plotted together with primitive mantle (Sun and McDonough 1989) in binary diagrams in Figures 3.14 a,b,c,and d. It is evident in these plots that there is a positive correlation amongst all of the data with the exception of Figure 3.14d in which there are two arrays.  The Yb vs. Tb plot (fig. 3.14d) shows an array defined by the MFV and MFX samples and another array defined by the DVI samples. This difference in the Tb/Yb likely reflects the mobility of this ratio in the DVI samples during post emplacement alteration (e.g., fig. 3.7).  Constant ratios (e.g., fig. 3.14 a,b,and c) indicate an initial similar degree of partial melting with a similar post melting  121 magma evolution.  On the basis of these constraints, a similar initial degree of partial melting is indicated for the genesis of the parental magmas for all suites.  Figure 3.14. (A) Hf vs. Zr, (B) Nb vs. Th, (C) La vs. Sm and (D) Yb vs. Tb variation plots (primitive mantle values from Sun and McDonough 1989). Symbols as in Figure 3.7.  We have identified six samples from the Marmot Formation (e.g., MFV samples 1502 A,C, 1503A, MFX samples 1505E, and DVI samples 1497A, 1498B) that meet the requirements of Frey et al. (1978) and Bartels et al. (1991) for identification of the most primitive samples (e.g., Mg # > 68, Ni contents 90-670 ppm, and Cr contents > 500 ppm).  Additionally, spinel chemistry from sample 1502C (fig. 3.5b) indicates that primitive olivine (Fo = 88-89) was in equilibrium with a primitive melt (e.g., Fo>88  122 typical of primitive melts; Frey et al. 1978); therefore sample 1502C is likely primitive in nature.  Geochemical modeling for the genesis of the Marmot Formation is based on the primitive samples identified above. A simplistic model for single stage batch melting is assumed owing to the primitive nature of the samples, and the lack of evidence for crystal fractionation within these primitive samples.  A modeling equation (e.g., CL/CO = 1/(Di(1-F)+F); where CL is the concentration of trace element in sample; CO is the concentration of trace element in source; Di is bulk distribution coefficient; and F is the weight fraction of melt produced ) for batch melting of Shaw (1970) is applied to our most primitive samples.  On the basis of modeling highly incompatible elements (e.g., Di approaches 0) the equation reduces to CL/CO = 1/F, and we can rearrange for F = CO/CL (e.g., Franz et al. 1999).  We assume average incompatible element concentrations (e.g., Ce = 1.78 ppm; Ta = 0.041 ppm; Nb = 0.658-0.713 ppm; and Th = 0.085 ppm) for mantle composition (e.g., Frey et al. 1978; Sun and McDonough 1989).  Green (1973) proved experimentally that highly alkaline rocks (e.g., nephelinites to melilitites) could be produced by low degrees of melting in the presence of H2O and garnet from an average mantle composition (e.g., primitive mantle). Therefore, the need to infer a unique mantle composition, owing to lack of mantle xenoliths, for the genesis of the Marmot Formation is not required. Low degrees of partial melting (0.71-1.5%) are required to produce the MFV and MFX samples and similar low degrees of partial melting (0.5-0.9%) for the DVI samples.  Simplistic batch melting for the genesis of our primitive samples is favored; however, these processes do not account for inevitable in situ fractionation of magmas prior to ascent (e.g., crystal fractionation leading to development of the non-  123 primitive Marmot Formation samples).  The REE ratios Ce/Y and Zr/Nb are believed to be insensitive to moderate degrees of crystal fractionation (e.g., olivine and clinopyroxene; Hardarson and Fitton 1991).  These ratios can, therefore, be used to model source constraints and degrees of melting, especially if crystal fractionation is inferred (fig. 3.13).  Fractional melting curves, calculated for various mantle compositions (Hardarson and Fitton 1991) are plotted on a Ce/Y vs. Zr/Nb diagram (fig. 3.15a).  Our least fractionated samples (e.g., the most primitive) from the MFV and MFX roughly follow the GP (primitive garnet lherzolite) melting curve, above the SP (primitive spinel lherzolite) curve corresponding to <1% melting.  The primitive DVI samples are not plotted on this diagram owing to the possible mobility of the Ce/Y ratio in the DVI suite (e.g., fig. 3.7h). This diagram suggests that the derivation of the primary magma of the Marmot Formation was from small degrees of melt from a garnet bearing mantle source  Figure 3.15. (A) Ce/Y vs. Zr/Nb plot for our MFV and MFX samples. The continuous lines are melting curves calculated by Hardarson and Fitton (1991) for mantle compositions; GD, depleted garnet lherzolite; GP, primitive garnet lherzolite; SD, depleted spinel lherzolite; SP, primitive spinel lherzolite. Numbers on curves refer to percentages of melt. (B) Tb/Yb vs. La/Yb plot for our MFV and MFX samples. Continous lines are for melting of fertile lherzolite mantle calculated by Macdonald et al. (2001). Numbers on curves represent modal amounts of garnet. OIB field from Macdonald et al. (2001).  Symbols as in Figure 3.7.   124 3.6.3. Mantle enrichment  Enriched garnet bearing peridotite is believed to be the dominant source for a majority of alkaline volcanics with an OIB type source affinity (e.g., Furman 1995; Civetta et al. 1998; Jung and Hoernes 2000; Rogers et al. 2000; Macdonald et al. 2001; Karmalkar et al. 2005; Song et al. 2008).  The nature and cause of this enrichment, however, is still debatable.  Enrichment may result from recycling of oceanic crust or pelagic and/or terrigenous sediments (e.g., Hofmann and White 1982; Weaver 1991) or may reflect metasomatic processes within the upper mantle (Halliday et al. 1995; Niu and O’Hara 2003). Crustal contamination of the mantle via sediment assimilation is commonly attributed to the enriched nature of OIB-like rocks (e.g., HIMU, Weaver 1991).  Weaver (1991) suggests that the ancient subducted oceanic lithosphere imparts the overriding mantle wedge with an Nb and Ta rich residue.  This residue likely explains the pronounced enrichment of these elements in OIB type rocks (e.g., HIMU, EMI, EMII). As such this process should be reflected in field observations, variations in incompatible elements, incompatible element ratios, primitive mantle normalized plots, and Nd isotopic composition.  However, as stated above, all of these means of assessing crustal contamination of the mantle source indicate little influence of terrigenous or pelagic sources; therefore mantle source enrichment of the Marmot Formation was probably not related to sediment assimilation. The presence of macrocrystic phlogopite may indicate a history of mantle metasomatism early in the evolution of the Marmot Formation.  Mantle metasomatism  125 typically leads to the modification of mantle peridotite and precipitation of incompatible element enriched minerals such as amphibole and phlogopite (e.g., Canil and Scarfe 1989; Halliday et al. 1995; Spath et al. 2001).  Metasomatic processes are generally interpreted to enrich the peridotitic mantle in incompatible elements, probably by infiltration of enriched fluids of either carbonititic or silicate composition.  High La/Yb, Nb/La and Nb/Ta ratios coupled with very high Zr/Hf and very low Ti/Eu ratios are characteristics of carbonatite metasomatism (Green 1995; Kalmarkar et al. 2005; and references therein).  Our samples, in particular the MFV and MFX samples, are characterized by average La/Yb, Nb/La, Nb/Ta, Zr/Hf, and Ti/Eu ratios of 29.6, 1.84, 17.6, 38.4 and 6868 respectively. The DVI samples also show high average La/Yb, Nb/La, Nb/Ta and Zr/Hf ratios of 72.7, 1.2, 20.0 and 42.5 respectively and a low average Ti/Eu ratio of 4080.  These ratios, indicate enrichment with respect to typical primitive mantle (La/Yb, 1.47; Nb/La, 1.007; Nb/Ta, 17.4 Zr/Hf, 36.2; and Ti/Eu, 7619, values from Sun and McDonough 1989) probably via carbonitic metasomatism.  The strikingly lower Ti/Eu ratios of the DVI samples may indicate a greater degree of carbonititic metasomatism and may also explain the dominance of phlogopite macrocrysts observed in the DVI occurrences (e.g., Mountain Diatreme). Therefore, variations in incompatible element ratios are probably a response to varying affects of metasomatic processes within the source. Determining when mantle enrichment induced by metasomatism took place is difficult owing to the lack of mantle xenoliths.  However, depleted mantle model (tDM) ages from all but one sample (1508D) yield ages between 620 and 870 Ma which may record Neoproterozoic mantle enrichment (e.g., Piercey et al. 2006).  Widespread  126 occurrences of large igneous provinces emplaced along the margin of western Laurentia (e.g., ca. 780 Ma Gunbarrel event, Harlan et al. (2003); Ootes et al. (2008) and ca. 720 Ma Franklin event, Heaman et al. (1992)) accompanied the Neoproterozoic breakup of supercontinent Rodinia.  It was speculated by Piercey et al. (2006) that magmatism associated with these large igneous provinces was responsible for the fertilization of the subcontinental mantle beneath the northern Canadian Cordillera.  Subsequent reactivation of this margin (e.g., MCE) and tapping of the fertilized subcontinental mantle may have yielded the initial OIB like mantle characteristics observed in the Marmot Formation.  3.7. NATURE OF THE SOURCE  Because of the absence of lower crustal and mantle derived xenoliths in the Marmot Formation inferences concerning the mineralogical and geochemical composition of the source are speculative.  The following discussion is primarily based on trace element geochemistry.  3.7.1. Composition of the source   The MFV and MFX samples are petrographically similar but collectively are different than the DVI samples.  Utilizing incompatible elements with similar bulk partition coefficients in a binary plot, we can assess the potential of different sources for the genesis of the three suites.  For example, plots of incompatible element ratios all exhibit a positive correlation (fig. 14a,b,c and d) away from primitive mantle suggesting  127 that the derivation of the parental Marmot Formation was from a primitive mantle source. On the basis of these relationships, it is suggested that the DVI samples were derived from a similar source as the MFV and MFX rocks.  An OIB-like source for the Marmot Formation is confirmed by the Th-Zr-Nb plot (fig. 3.10d) in which all suites fall in the OIB field. The Nd isotopic nature of individual OIB mantle reservoirs (e.g., HIMU, EM1, and EM2, of Zindler and Hart 1986) at 455 Ma is not currently constrained; however, values are most broadly similar to an OIB endmember HIMU-like (!Nd ~+6).  The source of the DVI samples (!Nd +2.7 – +3.2) on the basis of Nd isotopes; however, suggests the source was more enriched to that which generated the MFV and MFX suites.  The tight cluster of analyses (!Nd ~ +6.0) likely reflects the true isotopic signature of the source and reflects melting of a HIMU type source. On the basis of incompatible elements all samples are geochemically similar to OIB endmember HIMU. Incompatible element ratios characteristic of HIMU comprise Zr/Nb (3.2-5.0), La/Nb (0.66-0.77), Th/Nb (0.078-0.101), and Th/La (0.107-0.133) (Weaver 1991).  These ratios from our data (average of all three suites) comprise Zr/Nb (2.68), La/Nb (0.71), Th/Nb (0.09), and Th/La (0.14).  Additionally, our data plot in a tight cluster close to HIMU in an Nb/Y vs. Zr/Y diagram (fig. 3.16; fields after Condie 2005).  The minor discrepancies between Nd isotopic compositions and incompatible element ratios (e.g., our !Nd values not confined to HIMU values) are difficult to interpret in terms of identification of the source owing to the conclusive evidence outlined by geochemistry.  Moreover, as suggested by Wendlandt et al. (1995), the isotopic composition of primary basaltic melts reflects not only the composition of the source but also the degree of partial melting.   128  Figure 3.16. Nb/Y vs. Zr/Y plot showing mantle compositional components and fields for basalts from various tectonic settings (compositions and fields from Condie 2005). Abbreviations: UC, upper continental crust; PM, primitive mantle; DM, shallow depleted mantle; HIMU, high mu (U/Pb) source; EM1 and EM2, enriched mantle sources; ARC, arc- related basalts; NMORB, normal ocean ridge basalt; OIB, oceanic island basalt; DEP, deep depleted mantle; EN, enriched component; REC, recycled component.  Symbols as in Figure 3.7.  The source composition was garnetiferous as evidenced from substantial enrichments of LREE over HREE (average LaN/YN = 23.1).  Yitrium is preferentially incorporated into garnet and left in the residuum, thus melts derived from a dgarnet bearing source should be depleted in Y with respect to LREE (e.g., high LaN/YN). The Ce/Y vs. Zr/Nb diagram (see above) also indicates that the source was predominantly garnet bearing.  In a La/Yb vs. Tb/Yb plot (Macdonald et al. 2001) our least fractionated samples from the MFV and MFX suites fall between curves corresponding to 2 to 4 modal percent garnet in the source (fig. 3.15b).  The DVI primitive samples were not plotted on this diagram owing to the mobility of the Tb/Yb ratio in the DVI suite (e.g., fig. 3.7f). These curves were derived from geochemical modeling of fertile lherzolite mantle (Macdonald et al. 2001). On the basis of these constraints the source that was melted to generate the primary magma for the Marmot formation was likely a garnet bearing lherzolite.  129 The presence of phlogopite as a macrocrystic, probably phenocrystic, phase in some of the Marmot Formation occurrences (e.g., DVI) suggests that there was a hydrous phase, probably either phlogopite or amphibole, in the source.  Identifying which of these phases was present can be done utilizing major element data, especially those elements indicative of phlogopite and amphibole with characteristic partition coefficients such as K, Rb, Ba, and Sr (e.g., Furman 2007; and references therein).  The likelihood of element mobility of these incompatible elements was high during low-grade alteration that most of our rocks have experienced (see above).  Utilizing these methods is therefore likely to produce erroneous results and was not attempted.  Similarly, highly incompatible element ratios such as Zr/Hf and Nb/Ta indicate derivation from a metasomatically altered source (see above; Green 1995).  The partition coefficients of these incompatible elements (e.g., Nb and Ta) in amphibole and phlogopite are similar in alkaline melts (e.g., Green et al. 1993; Adam and Green 2006) and as such these elements should behave in a similar manner. Therefore distinguishing between amphibole and phlogopite on the basis of incompatible elements is also unreliable.  We can, however, infer that the metasomatic enrichment of the source for the Marmot Formation was likely to involve residual phlogopite because phlogopite is a dominant phenocryst/macrocryst phase present in many occurrences of the Marmot Formation.  3.7.2. Depth of source  The presence of garnet in the source provides a pressure constraint on the depth of melt extraction and limits the source for the genesis of the Marmot Formation to depths at  130 or below the garnet-spinel transition, corresponding to depths of ~ 90 km.  Additionally, the presence of residual phlogopite bearing peridotite requires pressures greater than ~25 kbar with a narrow temperature interval (~1000oC) at which phlogopite can coexist with melt above the peridotite solidus (Green 1970; Lloyd et al. 1985; Wallace and Green 1988; Furman 2007 and references therein).  This phlogopite pressure constraint correlates to depths of at least 70 km.  Therefore, phlogopite and garnet can coexist at depths greater than ~90 km and at temperatures of ~1000oC. Less than 10% partial melting under these conditions would be expected to generate melilitites and highly alkaline basalts (Green 1970).  3.7.2. Source heterogeneities?  Neodymium isotopic compositions from the Marmot Formation suggest that the nature of the source was spatially somewhat isotopically heterogeneous as indicated by a range in !Nd values (+2.7 - +8.7).  However, the variations in isotopic compositions are minor; thus on a large scale all represent derivations from a source approaching depleted mantle reservoir compositions.  On the basis of geochemistry (see above), the depth of the source for all suites appears to have been approximately the same therefore source heterogeneity may be laterally distributed rather than layered.  This would imply that, as suggested above, the effects of metasomatism are great along the margins of the basins where the DVI occurrences are concentrated (e.g., eastern MCE) whereas farther west in the central and western MCE the effects of this metasomatism were not as pronounced. Mantle heterogeneities can exist on small (e.g., 10 meter) and large (>1000 meter) scales  131 (Zindler and Hart 1986); thus the possibility of slight source heterogeneity beneath the MCE is reasonable.  3.7.3. Problems with source constraints  We interpret garnet to have been present in the residuum; however, as mentioned above, no mantle xenoliths have been recovered from the Marmot Formation and our ability to rigorously test this interpretation is therefore limited. The majority of mantle xenoliths studied from western North America are spinel peridotites (e.g., Roden et al. 1984; Francis 1987; Canil and Scarfe 1989; Francis and Lunden 1990; Shi et al. 1998) that were entrained in recent volcanic erruptions (e.g., <10 Ma) and are therefore ~450 m.y. younger than Marmot Formation volcanism. Milidragovic et al. (2005) and Milidragovic (2008), however, report the presence of garnet peridotite xenoliths in Early Cambrian lamprophyres (e.g., Quartet Mountain lamprophyres) located ~200 km northwest of the MCE.  These xenoliths confirm the presence of a garnet bearing upper mantle at least locally beneath the western margin of North America in Early Paleozoic time.  The Archean Slave Province to the east of the MCE contains kimberlites that were roughly coeval with emplacement of the Marmot Formation (e.g., 460 Ma Drybones Bay Kimberlite, Carbno and Canil 2002; Heaman 2003).  The Drybones Bay Kimberlite contains abundant garnet peridotite xenoliths.  Although it is interpreted to represent sampling from greater depths below thick mantle lithosphere (~160 km; Carbno and Canil 2002), it also suggests that in Middle Ordovician time the mantle beneath North America was at least in part garnetiferous.  132 3.8. TIMING OF MAGMATISM   Our 40Ar/39Ar indicates that Marmot Formation volcanism spanned ~16 m.y., beginning at ~460 Ma and culminating at ~444 Ma, corresponding to Middle to Late Ordovician time (fig. 3.10). Furthermore, diatreme bodies (e.g., DVI samples) are observed to cross cut the Franklin Mountain Formation, suggesting a maximum age of Llandeilian (Middle Ordovician, Cecile 1982), which is consistent with our phlogopite 40Ar/39Ar age data.  Cecile (1982) identified a two-holed crinoid oscicle fossil within interbedded limestones with volcanic rocks and tuffs at the top of the Marmot Formation type section (see section 17 of Cecile 1982). The presence of a two-holed crinoid suggests an upper Lower to early Middle Devonian age constraint for the cessation of MCE volcanism.  As suggested earlier, the bulk of MCE volcanism is Ordovician in age and therefore this age of 444 Ma represents the time of cessation of this Ordovician magmatism. The age range of 444 – 460 Ma for the major period of alkaline volcanism within the MCE suggests tapping of a relatively large, homogeneous, and long lived mantle source.  3.9. COMPARSION WITH VOLCANISM FROM SIMILAR TECTONIC SETTINGS  3.9.1. Comparison with rift related mafic volcanic suites   The composition of OIB is highly variable and as such several isotopically distinct end-member mantle sources (Zindler and Hart 1986) have been defined to  133 explain this variability (e.g., HIMU, EM1, and EM2).  Collectively, however, all end- member compositions represent the derivation from enriched and in some cases heterogeneous mantle sources (Weaver 1991). OIB-like rocks are typically alkaline in character and enriched in LILE (K, Rb, Cs, Ba, Pb and Sr) and HFSE (high field strength elements; Th, U, Ce, Zr, Hf, Nb, Ta, and Ti). The geochemistry of the Marmot Formation is largely consistent with an OIB type source (e.g., fig. 3.10c and fig. 3.16); however, this does not require that the Marmot Formation erupted into an ocean island setting from a mantle plume. Continental rift settings such as the East African Rift contain voluminous alkaline rocks that are consistent with OIB type geochemistry (e.g., Furman 1995; Furman and Graham 1999; Rogers et al. 2000; Spath et al. 2001; Macdonald et al. 2001) and whose genesis is related to substantial mantle upwelling below a divergent margin. In the following sections we compare the tectonic setting and volcanic geochemistry from well-studied locations of the East African Rift and the Rhine Graben of Germany with our data from the Marmot Formation to constrain a petrogenetic and tectonic model for the evolution of the MCE and the Marmot Formation.  As a first order comparison, we concentrate on the East African rift system as this tectonic entity is well studied. The general tectonic model for the East African rift involves continental extension concentrated along weakened zones within Proterozoic basement mobile belts (e.g., Kenya rift of Macdonald et al. 2001).  This extension began with lithospheric thinning (pre rift stage) and was followed by the development of graben structures accompanied with coeval mafic volcanism (rift stage).  The lithospheric thickness below these active rifts is significantly thinner than that of distal cratonic basement (e.g., 20-35 km crust thickness beneath Kenya rift, Macdonald et al. 2001).  134 The tectonic nature of the MCE is believed to have involved a simple graben structure as a result of crustal extension, with coeval mafic volcanism of the Marmot Formation. Significant differences between these two areas include the dramatically different scale of crustal extension and the nature and the pre-rift thickness of underlying crust.  Thinning of the lithospheric crust below the MCE occurred ~300 m.y. prior to formation of the MCE during breakup of supercontinent Rodinia (Colpron et al. 2002). Reactivation of these older crustal structures led to the formation of the MCE. Lithospheric thickness in Ordovician time in western North America is not currently known; however, a modern day passive margin analogy such as the Norwegian passive margin comprises 26 - 35 km crust and a total lithospheric thickness of ~125 km (Sheck-Wenderoth and Maystrenko 2008). Therefore it is reasonable to expect that the crust beneath the MCE would have at most been ~30 km thick and the vertical distance to the asthenosphere-lithosphere transition may have been ~125 km prior to rifting.  Following rifting the depth to this transition likely decreased due to upwelling of the asthenosphere imparting a thinner lithosphere. Geochemical evidence (discussed above) indicates that ascending magmas did not interact significantly with the lithosphere and therefore may represent pristine mantle melts, similar to those observed in the East African Rift, albeit having been overprinted by post emplacement alteration.   135  Figure 3.17. Comparison diagrams with data from the East African Rift (Kenya Rift data from Rogers et al. 2000; Spath et al. 2001; Macdonald et al. 2001, and West Rift data from Furman 1995; Furman and Graham 1999). (A) Primitive mantle normalized incompatible element plots showing average MFV, MFX and DVI suites with average compositions from basalts from the Kenya Rift and from the West Rift. (B) Ce/Y vs. Zr/Nb plot with data from the Kenya rift and the West Rift. Black line denotes the upper limit of Figure 3.14A. Symbols as in Figure 3.7.  Geochemical evidence for this relationship is clearly evident in extended incompatible element profiles and plots of incompatible elements (e.g., fig. 17a and 17b respectively).  Data shown for samples from the East African Rift are separated into alkaline rocks from the Western Rift (Furman 1995; Furman and Graham 1999) and from the Kenya Rift (Rogers et al. 2000; Spath et al. 2001; Macdonald et al. 2001). Incompatible element ratios are used for comparison purposes between Marmot Formation rocks and those from East African Rift (e.g., table 3.1).  It is evident that rocks of the Marmot Formation are geochemically similar to those observed in the East African Rift (table 3.1).  In Figure 3.17a we plot average composition of all of our Marmot Formation suites together with data from the Kenya Rift (Rogers et al. 2000; Spath et al. 2001; Macdonald et al. 2001) and the Western Rift (Furman 1995; Furman and Graham 1999) in an extended incompatible element plot.  There are marked similarities in these  136 plots as evidenced by the enrichment of HFSE relative to LILE (steep profile), positive Nb and Ti (except DVI) anomalies, and negative Zr anomalies (except MFV which exhibits a weak positive anomaly).  Other similarities included the relative enrichment of Nb, Ce, La, and Nd (LILE) of Western Rift and DVI samples compared to Kenya Rift and MFV and MFX samples of the Marmot Formation. On the basis of trace element modeling, Spath et al. (2001) suggested that the Chyulu Hills lavas in the Kenya Rift were generated by 1-3% partial melting of a hydrous amphibole-bearing source.  These constraints are similar to the melting and source constraints that we have inferred for the Marmot Formation; however, we cannot constrain the role of amphibole or phlogopite in the Mamot Formation source due to the mobility of Rb, Ba and K (see above). Furthermore, in a plot of incompatible element ratios Ce/Y against Zr/Nb the Marmot Formation MFX and MFV samples data fall into the data field defined by rocks from the Kenya Rift and are, for the most part, not unlike rocks from the Western Rift, possibly suggesting a similar source and similar low degrees of partial melting (fig. 3.17b).           137 Table 3.1. Geochemical comparisons with East African Rift volcanic rocks. References are reported in text.    Ti/Zr Zr/Y Zr/Nb Nb/Y Th/La Yb/Tb Ce/Y  *DVI 62.07 11.54 1.90 6.35 0.12 1.30 10.26 *MFV and *MFX 62.30 12.49 3.00 4.27 0.14 1.85 4.67  Kenya Rift 71.71 7.90 3.83 2.06 0.15 - 3.42 Western Rift 65.50 6.70 2.23 3.01 0.09 2.26 5.40  Primitive Mantle 116.00 2.46 15.70 0.16 0.12 4.56 0.39 N-MORB 110.00 2.64 31.80 0.08 0.05 5.02 0.27 OIB 55.00 9.66 5.83 1.66 0.11 2.05 2.75 HIMU - - 4.10 - 0.10 - -  BCC 54.00 5.00 9.09 0.55 0.22 3.66 1.65 UCC 15.80 8.64 7.60 1.14 0.36 3.44 2.90 LCC 85.60 3.68 11.70 0.32 0.10 3.73 1.21  * denotes average of suite   For the most part the East African Rift system evolved past the rift stage to a stage of oceanic crust formation with voluminous basalt generation.  The Rhine Graben in Germany, however, represents a failed rift that did not evolve to full ocean crust.  It is characterized by minor alkaline volcanism and offers a compelling analogy to the MCE. The Rhine Graben developed in the northern Alpine foreland as a result of reactivation of preexisting structures and was broadly coeval with the collisional phases of the Alpine and Pyrenean orogenies (Schumacher 2002).  Although this tectonic scenario is different than that proposed for rifting of the MCE and more broadly along the eastern margin of the Selwyn basin (e.g., Cecile et al. 1997), the geometry of the Rhine Graben rift and petrology of the associated volcanic rocks is very similar to those of the MCE.  The Rhine Graben is somewhat narrower and longer than the MCE (average width of 36 km a  138 length of ~300 km, vs. ~100 x 150 km for the MCE).  Rift related volcanism in the Rhine Graben spanned ~65 m.y. from 80 to 13 Ma (Prodehl et al. 1995) and magmatism is characterized by diatremes, dykes, and local flows.  These volcanic rocks are concentrated along the margins of the graben where pervasive crustal scale faults permit ascent of mantle derived magmas (Prodehl et al. 1995).  These volcanic rocks, in particular the Tertiary age deposits, petrographically comprise alkali-basalts, and it is interpreted that primary magmas were olivine melilitites and olivine nephelinites (Prodehl et al. 1995).  These primary magmas are thought to have been derived from low degrees of partial melting of a garnet lherzolitic mantle at depths of ~100 km (Prodehl et al. 1995).  These petrological considerations of Rhine Graben magmatism are very similar to what is envisaged for the Marmot Formation. Based on the similarities discussed above, we interpret the Marmot Formation to represent rift related magmatism derived from small degrees of partial melting of variably enriched, metasomatized OIB-like source.  3.9.2. Comparisons with mafic volcanic rocks in the Northern Cordillera   Lower Paleozoic alkaline volcanic rocks are widespread throughout much of the northern Canadian Cordilleran miogeocline and some of these units have been tentatively correlated with the Marmot Formation (Goodfellow et al. 1995). These include the Niddery, Itsi Lakes, and MacMillan volcanic rocks of eastern Yukon and the Vulcan volcanic rocks which are located south of the MCE of western North West Territories (Goodfellow et al. 1995).  Collectively these deposits have been interpreted to represent melting of a metasomatized subcontinental lithosphere with local asthenospheric input  139 (Goodfellow et al. 1995).  These authors suggest that observed enrichments in Nb and REE concentrations are an indication of metasomatism in the subcontinental lithosphere and infer that the genesis of these volcanic rocks were related to an episodically subsiding passive margin. Additionally, the Ordovician to Silurian Menzie Creek volcanic rocks of the Selwyn basin (e.g., Pigage 2004) may be also correlatives of the Marmot Formation, although poor age constraints on the Menzie Creek Formation makes this correlation uncertain.  Lamprophyre dykes (Quartet Mountain suite) documented by Milidragovic (2008) are located ~200 km to the northwest of the MCE and are suggested to be ~530 Ma which is somewhat older than the Marmot Formation.  All these Lower Paleozoic volcanic rocks, however, are broadly geochemically similar to the Marmot Formation. Diatremes that have also been identified in British Columbia, where they intrude Cambrian to Devonian miogeocline strata (Pell 1986), may be in part also correlative with the Marmot Formation.  Cecile et al. (1997) divide volcanic occurrences in the northern Cordilleran miogeocline into three periods of volcanic activity, with Cambrian, Ordovician, and Devonian ages.  Strictly speaking, comparisons of the Marmot Formation should only be made with other Ordovician volcanic (see Goodfellow et al. (1995) for a detailed correlation chart).  3.10. A MODEL FOR THE EVOLUTION OF THE MARMOT FORMATION   Petrological and petrochemical evidence presented here collectively provide an internally consistent model for the evolution of the Marmot Formation.  Rifting along the western margin of Laurentia in Early Paleozoic time reflected reactivation of lithospheric  140 structures that were formed during initial breakup of the supercontinent Rodinia in Neoproterozoic time.  Normal and associated transfer faults that have been identified in the basement underlying the Neoproterozoic Mackenzie Mountains Supergroup strata (Turner and Long 2008) may be structures that were reactivated in the Early Paleozoic. The formation of the MCE defines a typical ‘steers head’ rift basin, with a center graben structure developed on the lower plate segment of a typical rifted continental passive margin setting (Northern Lower Plate Zone of Cecile et al. 1997).  The presence of discrete volcaniclastic melilitic diatreme bodies (see Appendix E) along the eastern margin of the basin reflects the crustal scale normal fault systems along this margin (e.g., analogous to the Rhine Graben, Prodehl et al. 1995).  The volcanic deposits near the center of the basis comprise lava flows and volcanic vents, indicating less explosive or volatile charged volcanism.  Collectively all deposits have textural features consistent with submarine eruptions (e.g., hyaloclastites, pillow lavas, sedimentary layering and sorting).  Intracontinental mafic magmas can be generated by melting of either the continental lithosphere, the asthenospheric mantle, or, in some cases, both (e.g., Jung and Hoernes 2000, and references therein).  The derivation of the majority of these magmas requires an OIB-like enriched mantle source, as has been suggested for the Marmot Formation.  We speculate that the enrichment of the Marmot Formation mantle source may have occurred through infiltration of incompatible element and volatile rich metasomatic melts (e.g., carbonititic) leading to the precipitation of phlogopite (or amphibole).  The generation of these primitive magmas from the continental lithosphere requires very high continental stretching rates during extension (McKenzie and Bickle,  141 1988).  As suggested above the lithospheric thickness prior to rifting of the MCE may have been approximately ~125 km.  Following rifting and initial generation of the Marmot Formation magma it is inevitable that asthenospheric upwelling and continental extension would have thinned the overlying lithosphere.  The volume of magma generated is minor and as such it is doubtful that the lithosphere thinned greatly, possibly to ~90 km.  Additionally, the relatively local graben structure that defines the MCE suggests that excessive rates of stretching of the lithosphere did not occur. Therefore, it is unlikely that the derivation of magma was from the continental lithosphere (e.g., Jung and Hoernes 2000).  Studies of extensional regimes (e.g., Jung and Masberg 1998; Jung and Hoernes 2000) use the B factor (the ratio of unstretched to stretched thickness of the lithosphere) to determine the location of melting (e.g., B > 3 needed for melting of the crust).  We calculate a B factor of ~1.4 for the MCE, and therefore it is unlikely that melting commenced in the crust.  We therefore suggest that a component of asthenospheric melting was responsible for genesis of the Marmot Formation.  This interpretation is consistent with the strong garnet signature envisioned for the source. Additionally, the suggestion that the source experienced some degrees of metasomatism and precipitation of phlogopite (as indicated primarily for the generation of primitive DVI deposits) requires constraining the depth of melting to 90 – 100 km (depth at which phlogopite is stable with peridotite; Green 1970).  Considering the estimated post rifting thickness of the lithosphere beneath the MCE to be ~90km (?) would limit the depth to the base of the subcontinental lithosphere.  Additionally, the subcontinental lithosphere is thought to be a viable source for small degrees of kimberlitic, ultrapotassic and undersaturated melts (e.g., Furman 1995, and references therein), which are genetically  142 similar to the DVI samples (e.g., Godwin and Price 1986; Appendix E).  Our proposed model for the location of the source of the DVI samples therefore is at the base of a previously metasomatized subcontinental lithosphere at depths between 90-100 km (fig. 3.18 upper panel).  A component of an OIB-like (HIMU?) asthenospheric source is also required for the genesis of the MFV and MFX samples, as these samples do not appear to require the significant pre-melting mantle metasomatism that is required in the origin of the DVI deposits.  Therefore, it is suggested that the source for these deposits (MFV and MFX) extends through the base of the subcontinental lithosphere into the uppermost portion of the asthenosphere.  Initiation of MCE rifting and the resultant lithospheric thinning promoted the rise of the lithospheric isotherms.  This was accompanied by asthenospheric upwelling, which in turn triggered decompression melting.  Stretching of the lithosphere and upwelling of hot asthenosphere initiated small degrees of decompression melting. Zhou (1996) suggests that decompression melting of the asthenosphere at depths of ~90 km can be initiated by a B factor of ~1.5 with a potential asthenospheric temperature of 1380oC.  If this heat was maintained in the upwelling mantle the presence of metasomatically altered subcontinental lithosphere with entrained H2O (phlogopite) would have been significantly reduced the temperature of melting (e.g., Green 1973; Hirose and Kawamoto 1995).  This reduction of temperature typically produces parental magmas by small degrees of partial melting.  Green (1973) show that pressures of 25-30 kbar coupled with temperatures of 1200-1300oC with 2-7% H2O (possibly hosted in phlogopite) are required for the genesis of highly alkaline magmas from typical garnet bearing mantle compositions. Batch melting constraints utilizing typical primitive mantle compositions  143 indicate that ~1% partial melting is required for the genesis of the Marmot Formation parental magmas.  Following initial partial melt extraction, crystal fractionation took place primarily involving CPX at depths similar to that envisaged for partial melting and then largely controlled the evolution of the magmas. The mechanism that promoted metasomatism and initial partial melting of the mantle is difficult to constrain (e.g., underlying mantle plume or mantle decompression). For example, if these processes were initiated due to the impingement of a mantle plume at the base of the lithosphere then the volume and composition of the melt would reflect mixing of lithospheric and asthenospheric sources.  Spath et al. (2001) suggest that an uprising mantle plume provided fluids necessary to trigger mantle metasomatism observed in lavas from the Kenya Rift.  These authors also speculate that the plume head may have provided enough heat to promote partial melting of the now metasomatized lithosphere and later provided an additional asthenosphere component.  The volume of Marmot Formation magma is much lower than is typical of mantle plume derived flood basalts especially large igneous provinces (LIPs).  For example, if even the entire MCE contained volcanic deposits prior to sub-Silurian erosion the entire collective surface area of these deposits would be approximately 15,000 km2 which is much less that the >50,000 km2 typical for LIPs (Sheth 2007).  Basalts that make up typical LIPs are typically tholeiitic in composition (Ernst and Buchan 2003), which is also dissimilar to the Marmot Formation alkaline volcanics.  Furthermore, stratigraphic evidence for domal uplift above a mantle plume would be expected (Rainbird and Ernst 2001; Ernst and Buchan 2003), together with regional radial drainages and development of erosional unconformities.  There is no stratigraphic evidence for a radial erosional unconformity  144 pre-Marmot Formation in the center of the MCE (e.g., Cecile 1982).  We, suggest, therefore, that the genesis of the Marmot Formation and rifting of the MCE was not a result of a mantle plume. The upwelling asthenosphere near the center of the basin may have reached to shallower levels, and therefore, melting occurred at similar depths to that of the DVI deposits (fig. 3.18 upper panel).  The effects of overlying decompression process (e.g., faulting) may have been less controlling near the center and as such partial melting was controlled by upwelling asthenosphere.  The presence of these faults and localized volcaniclastic melilitic diatremes on the basin margin reflects partial melting as a result of decompression induced by faulting and therefore small degrees of melting. Although the volume of Marmot Formation igneous rocks is relatively small, massive mafic volcanic lavas with thicknesses up to 300 meters are observed locally near the center of the basin (e.g., sample 1508).  Therefore, there were periods of eruption of significant volumes of lava associated with local subsidence.  Rifting and asthenospheric upwelling likely concluded at this point, as oceanic crust was never formed.  Therefore, the formation of the MCE is likely analogous to the ‘rift stage’ of the East African Rift and more similar to the evolution of the Rhine Graben.  The concentration of DVI deposits along the eastern margin of the rift basin suggests that in some cases melts may have utilized normal faults along this rift margin as conduits to rise to the surface.  The lavas and volcaniclastic deposits observed near the center of the basin were emplaced at the same time as the diatremes; however, they ascended through relatively intact down dropped lithosphere.   145  Figure 3.18. Schematic model cross-section of the western margin of Laurentia through the lithosphere in Early Cambrian (lower panel) and Middle Ordovician (upper panel) showing inferred model for the evolution of the MCE discussed in text. The question mark in the lower panel indicates the uncertainty as to when metasomatism actually took place. The upper crust in the vicinity of the Misty Creek Embayment is vertically exaggerated. SCLM; sub-continental lithospheric mantle.  In summary, the model envisaged for the petrogenesis of the Marmot Formation involves both continental extension and asthenospheric upwelling.  Homogeneous mantle metasomatism at the base of the subcontinental lithosphere may have occurred during initial Cambrian rifting beneath the MCE but was not associated with any volcanism (fig. 3.18 lower panel).  Upwelling asthenosphere and melting/homogenizing of the metasomatized lithosphere followed this initial rifting (fig. 3.18 upper panel).  The upwelling is more pronounced near the center of the basin than near the basin edges (fig. 3.18 upper panel).  This contrast between melting of metasomatized lithospheric mantle  146 versus melting of asthenospheric plus metasomatized lithospheric mantle likely explains the differences in the petrography, geochemistry, isotopic signature and the volume of DVI deposits and the MFV deposits.  Following initial partial melting the evolution of the Marmot Formation magmas was largely controlled by crystal fractionation processe.  3.11. CONCLUSIONS  Timing constraints obtained from Ar-Ar dating of phlogopite indicates that volcanism in the MCE spanned ~16 m.y.  It is possible that volcanism in the MCE may have persisted until Early Devonian time (Cecile 1982); however, we have not studied the petrogenesis of this possible younger period of volcanism.  There are no obvious spatial controls on age of volcanism, and the two main eruption stages identified in this study, at 444 and 460 Ma, both occurred throughout the MCE.  There is also no apparent geochemical variations with age, suggesting that there was not a time-integrated depletion of the source (e.g., all samples reflect similar incompatible element enrichment).  The occurrence of phlogopite throughout the Marmot Formation suggests a hydrous component was present within the source.  Phlogopite comprises a higher modal component in the DVI samples (e.g., Mountain Diatreme) and may suggest a greater role of metasomatism by fluids and volatiles in the genesis of the DVI deposits.  Furthermore, the initial parental magmas for the Marmot Formation was likely generated by low degrees of partial melting.  The genesis of the MFV and MFX suites of the Marmot Formation was likely at the base of the subcontinental lithosphere with a minor asthenospheric input.  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Results of our detrital zircon dating study (Chapter 3) are applicable to all detrital zircon studies aimed at understanding the provenance of individual sedimentary units that accumulated within a long-lived passive margin.  For example, because we analyzed detrital zircons from the bases and tops of several sedimentary units we can therefore speculate on the relative roles of sediment recycling and tectonic controls on local sedimentation.  These are processes that are likely to influence any passive margin setting. We have also shown that making inferences about the provenance of sedimentary sequences along the length of the miogeocline should be avoided owing to these local tectonic controls on sedimentation.  Furthermore, we have shown that the source of sediments for individual sedimentary units is not necessarily located peripheral to the depositional basin.  By comparing our extensive U-Pb detrital zircon age data set with data reported by other workers from sedimentary units from eastern North America we have concluded that it is likely that in Neoproterozoic to Early Cambrian time sediments  159 were dispersed from eastern North America and accumulated along the western margin of Laurentia. This hypothesis, although somewhat controversial, can be rigorously tested with additional future research.  Felsic igneous rocks of the same age (within error) as the “anomalous” zircon ages that we have obtained in our study are known and represent magmatism in a variety of paleotectonic settings in eastern and southern North America therefore attempting to identify unique sources for detrital zircons of a specific age range is difficult.  Therefore, it would be highly beneficial to determine not only the U-Pb age of detrital zircons but also the Hf isotopic composition.  Hafnium in zircon can be used as a tracer of crustal and mantle processes (e.g., Belousova et al. 2006) and therefore zircons in igneous rocks from different tectonic settings will retain a specific range of Hf isotopic compositions.  A combination of U-Pb ages for specific zircons together with their Hf isotopic compositions would therefore provide a more specific “fingerprint’ for determining the provenance of zircons.  It would be necessary to also obtain Hf isotopic compositions from igneous zircons from known igneous sources (such as those exposed in eastern and southern North America) in order to use zircon Hf isotopic compositions as an additional provenance tracer. Additional future research directions for this detrital zircon study that could be undertaken include; (1) detailed morphological classifications for all analyzed detrital zircons to use as proxy for residency times in river systems; (2) petrographic analyses of thee samples from which the detrital zircons were derived to confirm sediment maturity; (3) a detailed examination of the sedimentary units in situ to obtain robust paleocurrent  160 measurements; and (4) use ages of other detrital minerals as provenance indicators (e.g., mica, monazite, and titanite).  Results of our study of Early Paleozoic alkaline magmatism in the Misty Creek Embayment has important implications regarding the mantle composition beneath the western margin of Laurentia and the tectonic evolution of this margin in Early Paleozoic time. Petrographic studies together with whole rock, mineral chemistry and Nd isotopic studies indicate that the Ordovician alkaline magmatism was likely generated at the base of the subcontinental lithosphere with a minor asthenospheric input and some magmas were generated by smaller degrees of partial melting primarily in metasomatized subcontinental lithosphere.  If all the Ordovician alkaline deposits identified by Goodfellow et al. (1995) are genetically related to the Marmot Formation then it would appear that the mantle in the Middle Ordovician was regionally homogeneous.  This study, however, has shown that local mantle heterogeneities did exist along the margin of Laurentia in Early Paleozoic time.  This study was reconnaissance in nature and additional future research could be done to further characterize this phase of magmatism. The Marmot Formation and associated diatreme bodies are well exposed; and a detailed study of the Marmot Formation could be undertaken to shed new light on both the nature and environment of this particular magmatic event, and more generally a relationship between the diatreme bodies and the associated epiclastic deposits.  For example, examination of the physical properties of the volcanic xenoliths (e.g., rounding, size and density) and the host volcanic units themselves (e.g., density, porosity, and viscosity) can yield considerable information regarding the explosiveness of the eruptions and the nature and evolution of magma ascent, fragmentation and eruption.  Also, additional  161 mineral chemistry, whole-rock geochemistry and isotopic compositions could be used to further constrain initial interpretations of the petrogenesis of this rock suite.  Finally, the anomalous Early Devonian volcanism that may represent a later manifestation of Misty Creek Embayment magmatism should be investigated to constrain the relationship of this event with the bulk of the Ordovician Marmot Formation.                     162 REFERENCES CITED  Belousova, E.A., Griffin, W.L., and O’Reilly, S.Y., 2006, Zircon Crystal Morphology, Trace Element Signatures and Hf Isotope Composition as a Tool for Petrogenetic Modeling: Examples From Eastern Australian Granitoids: Journal of Petrology, v. 47, p. 329-353.  Goodfellow, W.D., Cecile, M.P., and Leybourne, M.I., 1995, Geochemistry, petrogenesis, and tectonic setting of lower Paleozoic alkalic and potassic volcanic rocks, Northern Canadian Cordilleran Miogeocline: Canadian Journal of Earth Sciences, v. 32, p. 1236-1254.                                     163     APPENDICES      APPENDIX A  Chapter 3 Analytical Methods                             164  Mineral Chemistry  Clinopyroxene and phlogopite crystals were analysed in situ from polished thin sections prepared by Vancouver Petrographics Ltd. heavy minerals were separated from a ~1 kg sample of volcaniclastic basalt from sample 1502c using conventional crushing and grinding, Wilfley table, and magnetic separation methods.  A random selection of spinel grains from this separate was mounted in an epoxy puck and was brought to a high polish for microprobe analysis.  Quantitative elemental analysis of clinopyroxene, spinel and phlogopite was performed using the fully automated CAMECA SX-50 electron microprobe (Electron Microbeam/X-ray Diffraction Facility, Earth and Ocean Sciences, University of British Columbia, Vancouver, Canada). Analytical conditions were 15.03 kV accelerating voltage, a 20.05 nA primary electron beam current, and peak and background counting times of 60 and 30 s respectively. We used a 5 µm beam diameter.  Lithogeochemistry  Two samples of Mineral Deposit Research Unit in-house standard BAS-1 (basalt from near Cheakmus, British Columbia) for replicate analyses were submitted along with our whole rock samples in order to examine the analytical precision of lithogeochemical data obtained from ALS Chemex.  Analytical results obtained from these samples were compared to a mean of 5 previous repeat analyses of this in-house standard and are presented in Table A.1.  Duplicate analyses of these samples are precise and are within two standard deviations of the standard values.  Analytical procedures at ALS Chemex are reported below.   165 Table A.1. Mean values and duplicate analyses of standard BAS-1   *BAS-1 (n=5)    BAS-1-A BAS-1-B   mean  std deviation  This Study  Major elements (wt. %) SiO2 53.56 0.36  53.00 52.84 TiO2 1.31 0.01  1.38 1.38 Al2O3 15.12 0.07  15.86 15.83 Fe2O3 11.16 0.05  10.66 10.68 FeO 8.86 0.12  8.53 8.54 MnO 0.14 0.00  0.14 0.15 MgO 7.35 0.05  7.08 7.10 CaO 8.28 0.05  7.97 8.00 Na2O 3.28 0.04  3.13 3.15 K2O 0.56 0.02  0.52 0.53 P2O5 0.22 0.00  0.21 0.21 TOTAL 100.96   99.93 99.86  Trace and rare earth elements (ppm) Ag 0.28 0.08  <1 <1 Ba 194.00 18.55  181.50 187.50 Ce 21.80 0.75  19.60 19.90 Co 42.20 0.40  50.40 51.80 Cs 0.10 0.01  0.21 0.15 Cu 59.00 0.80  64.00 63.00 Dy 3.30 0.15  3.14 3.24 Er 1.50 0.08  1.61 1.49 Eu 1.30 0.07  1.18 1.16 Ga 19.60 1.02  20.50 20.40 Gd 3.80 0.11  3.36 3.38 Hf 2.40 0.08  2.70 2.70 Ho 0.60 0.02  0.62 0.59 La 9.30 0.44  9.10 9.20 Lu 0.20 0.01  0.19 0.18 Nb 8.20 0.49  7.90 8.10 Nd 13.60 0.80  13.00 13.00 Ni 172.00 1.90  178.00 179.00 Pb 4.40 4.32  <5 <5 Pr 3.00 0.13  2.76 2.74 Rb 7.00 0.14  7.40 6.80 Sm 3.50 0.48  3.24 3.25 Sn 1.20 0.21  1.00 1.00 Sr 502.00 7.48  469.00 461.00 Ag 0.50 0.02  0.50 0.60 Tb 0.60 0.03  0.55 0.52 Th 0.80 0.03  1.04 0.81 Tl 0.10 0.00  <0.5 <0.5 Tm 0.20 0.01  0.22 0.20 U 0.30 0.01  0.32 0.29 V 152.00 1.10  171.00 177.00 Y 18.40 1.02  14.50 14.80 Yb 1.40 0.06  1.38 1.28 Zn 91.40 1.02  102.00 108.00 Zr 94.50 2.18   89.00 94.00 BAS-1 is a UBC standard * Mean values of BAS-1 are based on the average of 5 previous repeat analyses. A and B are duplicate analyses of the standards that were used to test for precision.       166      167      168      169      170      171      172 Neodymium Isotopes  Samples were analyzed at the Pacific Center for Isotope and Geochemical Research (PCIGR) facility at the University of British Columbia.  Analytical protocol for leaching techniques are described in Weis et al. (2005) and the analytical techniques and instrumentation are described in Weis et al. (2006).  40Ar/39Ar Geochronology Methodology Phlogopite crystals were manually extracted from hand samples from eight Marmot Formation occurrences across the Misty Creek Embayment.  Attempts were made to select the smallest grain size observed. Mineral separates were wrapped in aluminum foil and stacked in an irradiation capsule with similar-aged samples and neutron flux monitors (Fish Canyon Tuff sanidine (FCs), 28.02 Ma (Renne et al. 1998). The samples were irradiated on March 26, 2008 at the McMaster Nuclear Reactor in Hamilton, Ontario, for 45 MWH, with a neutron flux of approximately 6x1013 neutrons/cm2/s. Analyses (n=30) of 6 neutron flux monitor positions produced errors of <0.5% in the J value. The samples were analyzed on February 13 and April 8 through April 9, 2008 at the Noble Gas Laboratory, Pacific Centre for Isotopic and Geochemical Research, University of British Columbia, Vancouver, BC, Canada. The mineral separates were step-heated at incrementally higher powers in the defocused beam of a 10W CO2 laser (New Wave Research MIR10) until fused. The gas evolved from each step was analyzed by a VG5400 mass spectrometer equipped with an ion-counting electron multiplier. All  173 measurements were corrected for total system blank, mass spectrometer sensitivity, mass discrimination, radioactive decay during and subsequent to irradiation, as well as interfering Ar from atmospheric contamination and the irradiation of Ca, Cl and K (Isotope production ratios: (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)).  Results Details of the analyses, including plateau (spectrum) and inverse correlation plots, are presented in Excel spreadsheets. The plateau and correlation ages were calculated using Isoplot ver.3.09 (Ludwig 2003).  Errors are quoted at the 2-sigma (95% confidence) level and are propagated from all sources except mass spectrometer sensitivity and age of the flux monitor. The best statistically-justified plateau and plateau age were picked based on the following criteria:  1. Three or more contiguous steps comprising at least 50% of the 39Ar; 2. Probability of fit of the weighted mean age greater than 5%; 3. Slope of the error-weighted line through the plateau ages equals zero at 95% confidence level; 4. Ages of the two outermost steps on a plateau are not significantly different from the weighted-mean plateau age (at 1.8"# six or more steps only); 5. Outermost two steps on either side of a plateau must not have nonzero slopes with the same sign (at 1.8"# nine or more steps only)   174 REFERENCES CITED   Ludwig, K.R., 2003, Isoplot 3.09, A Geochronological Toolkit for Microsoft Excel: Berkeley Geochronology Center, Special Publication No. 4.  Renne, P.R., Swisher, C.C., III, Deino, A.L., Karner, D.B., Owens, T., and DePaolo, D.J., 1998, Intercalibration of standards, absolute ages and uncertainties in 40Ar/39Ar dating: Chemical Geolology, v. 145, p.117-152.  Weis, D., Keiffer, B., Maerchalk, C., Pretorius, C., and Barling, J., 2005, High-precision Pb-Sr-Nd-Hf isotopic characterization of USGS BHVO-1 and BHVO-2 reference materials: Geochemistry Geophysics Geosystems, v. 6, 10 p.  Weis, D., Keiffer, B., Maerchalk, C., Barling, J., de Jong, J., Williams, G.A., Hanano, D., Prestorius, W., Mattielli, N., Scoates, J.S., Goolaerts, A., Friedman, R.M., and Mahoney, J.B., 2006, High-precision isotopic characterization of USGS reference materials by TIMS and MC-ICP-MS: Geochemistry Geophysics Geosystems, v. 9, 30 p.                             175  APPENDIX B  Chapter 2 Analytical Data                                         176 Table B.1. U-Pb geochronological data Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. Sample 07CL-1442A, Katherine Group ( 7153132N, 518443E) 1c 0.3215 0.0016 4.5447 0.0744 0.1021 0.0009  1797 8 1739 14 1662 17 -9.3 1d 0.1866 0.0023 2.0056 0.0802 0.0775 0.0021  1103 13 1117 27 1133 52 2.9 1e 0.1985 0.0015 2.1495 0.0544 0.0750 0.0012  1167 8 1165 18 1068 33 -10.2 1f 0.2980 0.0037 4.1946 0.1769 0.1019 0.0022  1682 18 1673 35 1660 40 -1.5 1g 0.2358 0.0024 2.6659 0.0870 0.0888 0.0018  1365 13 1319 24 1400 38 2.8 1h 0.1968 0.0011 2.0621 0.0348 0.0784 0.0009  1158 6 1136 12 1158 23 0.0 1k 0.2498 0.0011 3.1098 0.0401 0.0922 0.0008  1437 6 1435 10 1470 16 2.5 1l 0.1830 0.0011 1.9254 0.0357 0.0767 0.0010  1083 6 1090 12 1112 25 2.8 1m 0.2571 0.0019 3.0767 0.0762 0.0895 0.0013  1475 10 1427 19 1414 28 -4.8 1p 0.2565 0.0010 3.6492 0.0382 0.1042 0.0007  1472 5 1560 8 1700 13 15.0 1q 0.2085 0.0009 2.2863 0.0274 0.0815 0.0007  1221 5 1208 8 1234 16 1.2 1r 0.5172 0.0024 12.9549 0.2086 0.1845 0.0014  2687 10 2676 15 2694 12 0.3 1s 0.2327 0.0010 2.6147 0.0321 0.0853 0.0007  1349 5 1305 9 1322 16 -2.3 1t 0.2997 0.0019 3.9355 0.0828 0.0995 0.0012  1690 10 1621 17 1614 22 -5.4 2a 0.2956 0.0014 3.8895 0.0570 0.1007 0.0009  1669 7 1612 12 1638 17 -2.2 2b 0.2911 0.0012 3.9698 0.0473 0.1014 0.0008  1647 6 1628 10 1649 14 0.1 2d 0.2523 0.0013 3.0011 0.0490 0.0880 0.0009  1450 7 1408 12 1382 20 -5.6 2g 0.1760 0.0027 1.8022 0.0883 0.0754 0.0025  1045 15 1046 32 1078 66 3.3 2h 0.2486 0.0012 3.2587 0.0456 0.0939 0.0009  1432 6 1471 11 1507 17 5.6 2i 0.1869 0.0008 2.0012 0.0230 0.0759 0.0007  1105 4 1116 8 1093 17 -1.2 2k 0.2099 0.0016 2.3777 0.0578 0.0844 0.0013  1228 9 1236 17 1302 30 6.2 2l 0.1756 0.0010 1.7548 0.0300 0.0727 0.0009  1043 5 1029 11 1005 25 -4.1 2m 0.3565 0.0018 5.9488 0.0940 0.1211 0.0011  1966 8 1968 14 1972 16 0.4 2n 0.2189 0.0016 2.7116 0.0620 0.0902 0.0013  1276 8 1332 17 1430 27 11.9 2o 0.2234 0.0010 2.8251 0.0343 0.0924 0.0008  1300 5 1362 9 1475 16 13.1 2p 0.3988 0.0033 8.7806 0.2476 0.1694 0.0022  2163 15 2316 26 2552 21 17.9 2q 0.1761 0.0010 1.7833 0.0318 0.0733 0.0009  1046 6 1039 12 1022 26 -2.6 2r 0.2156 0.0016 2.6307 0.0627 0.0839 0.0013  1259 8 1309 18 1291 29 2.7 2s 0.2359 0.0014 2.7600 0.0495 0.0867 0.0010  1365 7 1345 13 1353 23 -1.0 3a 0.2313 0.0014 2.7210 0.0527 0.0878 0.0011  1341 8 1334 14 1379 24 3.0 3b 0.1922 0.0011 2.2107 0.0372 0.0819 0.0010  1133 6 1184 12 1244 23 9.7 3c 0.1982 0.0013 2.0884 0.0437 0.0745 0.0011  1166 7 1145 14 1056 29 -11.4 3d 0.2330 0.0019 2.7545 0.0701 0.0878 0.0014  1350 10 1343 19 1377 30 2.2 3e 0.3031 0.0028 4.3426 0.1336 0.1063 0.0018  1707 14 1702 25 1737 30 2.0 3g 0.2158 0.0012 2.4806 0.0394 0.0855 0.0010  1260 6 1266 12 1327 21 5.6 3h 0.2899 0.0015 3.9590 0.0652 0.0995 0.0011  1641 8 1626 13 1615 19 -1.8 3j 0.2000 0.0023 2.1039 0.0793 0.0761 0.0019  1175 12 1150 26 1098 49 -7.7 3l 0.2680 0.0015 3.3798 0.0604 0.0908 0.0011  1530 8 1500 14 1442 22 -6.9 3n 0.1853 0.0019 1.8088 0.0569 0.0728 0.0016  1096 10 1049 21 1008 43 -9.4 3o 0.5160 0.0022 13.3070 0.1792 0.1868 0.0016  2682 10 2702 13 2714 14 1.4 3p 0.2566 0.0012 3.2223 0.0461 0.0933 0.0010  1473 6 1463 11 1495 19 1.7 3q 0.2532 0.0012 3.2790 0.0449 0.0955 0.0010  1455 6 1476 11 1538 19 6.1 3r 0.1909 0.0012 2.0619 0.0378 0.0788 0.0010  1126 6 1136 13 1168 26 3.9 4d 0.2515 0.0013 3.2247 0.0482 0.0899 0.0010  1446 6 1463 12 1423 20 -1.8 4e 0.2609 0.0018 3.1496 0.0680 0.0926 0.0013  1495 9 1445 17 1479 27 -1.2  Sample RAS06-188A, Little Dal Group (7099166N, 540923E) 1a 0.3202 0.0012 4.5318 0.0545 0.1045 0.0007  1791 6 1737 10 1705 12 -5.8 1b 0.2627 0.0020 3.2498 0.0812 0.0918 0.0013  1504 10 1469 19 1463 27 -3.1 1c 0.5459 0.0027 14.4918 0.2582 0.1889 0.0013  2808 11 2783 17 2733 11 -3.4 1d 0.3483 0.0034 5.6769 0.1879 0.1208 0.0019  1927 16 1928 29 1967 27 2.4 1e 0.2583 0.0010 3.2536 0.0383 0.0914 0.0007  1481 5 1470 9 1455 14 -2.0 1f 0.3490 0.0014 5.8528 0.0793 0.1177 0.0008  1930 7 1954 12 1921 12 -0.5 1g 0.3637 0.0011 5.7815 0.0552 0.1156 0.0006  1999 5 1944 8 1889 10 -6.8 1h 0.2927 0.0009 4.0735 0.0380 0.1030 0.0006  1655 5 1649 8 1679 11 1.6 1i 0.1941 0.0009 2.0951 0.0277 0.0801 0.0007  1144 5 1147 9 1199 17 5.1 1j 0.2333 0.0007 2.6578 0.0223 0.0840 0.0005  1352 4 1317 6 1293 11 -5.1 1k 0.1921 0.0009 2.0272 0.0284 0.0771 0.0007  1133 5 1125 10 1123 19 -1.0 1l 0.2856 0.0011 4.0261 0.0486 0.1015 0.0007  1620 6 1640 10 1651 13 2.2 1m 0.2048 0.0005 2.2468 0.0157 0.0795 0.0004  1201 3 1196 5 1185 10 -1.5 1n 0.2263 0.0009 2.6781 0.0343 0.0851 0.0007  1315 5 1322 9 1318 16 0.3 1o 0.1940 0.0015 1.9492 0.0480 0.0757 0.0013  1143 8 1098 17 1088 33 -5.5 1p 0.1997 0.0007 2.1493 0.0211 0.0785 0.0005  1174 4 1165 7 1159 14 -1.4 1q 0.3419 0.0014 5.3941 0.0704 0.1153 0.0008  1896 7 1884 11 1885 12 -0.7 1r 0.2982 0.0012 4.0945 0.0502 0.1000 0.0007  1683 6 1653 10 1624 13 -4.1  177 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 1s 0.3438 0.0010 5.4327 0.0495 0.1153 0.0006  1905 5 1890 8 1885 10 -1.2 1t 0.2086 0.0015 2.3208 0.0536 0.0820 0.0012  1221 8 1219 16 1245 29 2.1 2a 0.2468 0.0010 3.1595 0.0386 0.0937 0.0007  1422 5 1447 9 1503 14 6.0 2b 0.2541 0.0010 3.1621 0.0387 0.0921 0.0007  1460 5 1448 9 1468 14 0.7 2c 0.2504 0.0011 3.3101 0.0454 0.0913 0.0008  1440 6 1483 11 1453 16 1.0 2d 0.1841 0.0009 1.8967 0.0283 0.0765 0.0008  1089 5 1080 10 1109 20 1.9 2e 0.2007 0.0009 2.2090 0.0285 0.0794 0.0007  1179 5 1184 9 1183 17 0.3 2f 0.1959 0.0008 2.0955 0.0245 0.0786 0.0006  1153 4 1147 8 1162 16 0.8 2g 0.2837 0.0008 3.8700 0.0294 0.0994 0.0005  1610 4 1607 6 1613 10 0.2 2h 0.2296 0.0008 2.7966 0.0276 0.0878 0.0006  1332 4 1355 7 1378 13 3.6 2i 0.2824 0.0012 3.8382 0.0530 0.0996 0.0008  1603 6 1601 11 1617 15 1.0 2j 0.1986 0.0007 2.0741 0.0229 0.0769 0.0006  1168 4 1140 8 1118 15 -4.9 2k 0.1948 0.0012 2.0285 0.0392 0.0772 0.0010  1147 7 1125 13 1127 26 -2.0 2l 0.1821 0.0008 1.9653 0.0271 0.0781 0.0007  1079 5 1104 9 1150 19 6.8 2m 0.5279 0.0020 13.4990 0.1745 0.1865 0.0011  2733 8 2715 12 2711 10 -1.0 2n 0.3124 0.0012 4.7341 0.0543 0.1092 0.0007  1753 6 1773 10 1786 12 2.1 2o 0.2034 0.0023 2.1903 0.0781 0.0816 0.0019  1193 12 1178 25 1237 44 3.8 2p 0.1994 0.0011 2.0340 0.0343 0.0771 0.0009  1172 6 1127 11 1124 23 -4.7 2q 0.4062 0.0015 8.6665 0.1036 0.1485 0.0009  2198 7 2304 11 2329 11 6.7 2r 0.2303 0.0008 2.6828 0.0264 0.0863 0.0006  1336 4 1324 7 1344 13 0.6 2s 0.1827 0.0006 1.9263 0.0178 0.0765 0.0005  1082 3 1090 6 1109 14 2.6 2t 0.2441 0.0011 3.0114 0.0409 0.0906 0.0008  1408 6 1411 10 1437 16 2.3 3a 0.2516 0.0012 3.1631 0.0493 0.0906 0.0009  1447 6 1448 12 1439 18 -0.6 3b 0.1890 0.0011 2.0199 0.0345 0.0775 0.0009  1116 6 1122 12 1135 23 1.8 3c 0.3599 0.0012 5.9425 0.0595 0.1216 0.0007  1982 6 1968 9 1980 11 -0.1 3d 0.2024 0.0007 2.1320 0.0194 0.0775 0.0005  1188 3 1159 6 1133 13 -5.4 3e 0.2302 0.0010 2.8057 0.0353 0.0868 0.0007  1335 5 1357 9 1356 16 1.7 3f 0.2362 0.0007 2.9868 0.0260 0.0929 0.0006  1367 4 1404 7 1486 12 8.9 3g 0.2201 0.0008 2.6148 0.0268 0.0851 0.0006  1283 4 1305 8 1318 14 3.0 3h 0.1696 0.0009 1.9281 0.0283 0.0822 0.0008  1010 5 1091 10 1250 20 20.8 3i 0.2393 0.0065 2.7605 0.2438 0.0896 0.0048  1383 34 1345 66 1416 99 2.6 3j 0.2383 0.0032 2.9382 0.1286 0.0898 0.0023  1378 17 1392 33 1421 48 3.3 3k 0.2117 0.0011 2.3401 0.0389 0.0811 0.0009  1238 6 1225 12 1224 21 -1.2 3l 0.2146 0.0012 2.6653 0.0469 0.0865 0.0010  1253 6 1319 13 1350 21 7.9 3m 0.2476 0.0014 3.1520 0.0570 0.0908 0.0010  1426 7 1446 14 1443 21 1.3 3n 0.1705 0.0012 1.7900 0.0397 0.0732 0.0011  1015 7 1042 14 1019 31 0.4 3o 0.5083 0.0020 13.3685 0.1739 0.1920 0.0012  2650 8 2706 12 2759 11 4.8 3p 0.1871 0.0008 2.0643 0.0272 0.0805 0.0008  1106 5 1137 9 1209 18 9.3 3q 0.4777 0.0019 11.4967 0.1485 0.1729 0.0011  2517 8 2564 12 2586 11 3.2 3r 0.1924 0.0014 1.9705 0.0438 0.0768 0.0012  1134 7 1106 15 1116 30 -1.8 3s 0.1824 0.0013 1.8767 0.0405 0.0766 0.0011  1080 7 1073 14 1112 29 3.1 3t 0.1906 0.0008 1.9333 0.0234 0.0770 0.0007  1125 4 1093 8 1121 17 -0.3 4a 0.2306 0.0007 2.7372 0.0224 0.0858 0.0005  1338 4 1339 6 1334 12 -0.3 4b 0.1827 0.0009 1.9077 0.0271 0.0771 0.0008  1082 5 1084 9 1123 20 4.0 4c 0.5012 0.0027 12.1610 0.2317 0.1835 0.0015  2619 12 2617 18 2684 14 3.0 4d 0.2416 0.0010 3.0542 0.0362 0.0921 0.0007  1395 5 1421 9 1470 15 5.6 4e 0.2400 0.0011 3.0369 0.0414 0.0921 0.0008  1387 6 1417 10 1470 17 6.3  Sample CL06-48A, Keele Formation (7083127N, 522129E) 1a 0.4446 0.0040 9.0959 0.3100 0.1538 0.0022  2371 18 2348 31 2389 24 0.9 1b 0.3569 0.0090 5.4168 0.4919 0.1181 0.0050  1968 43 1888 78 1928 74 -2.4 1c 0.4136 0.0040 8.3117 0.3067 0.1438 0.0022  2231 18 2266 33 2274 27 2.2 1d 0.2151 0.0044 2.1847 0.1489 0.0790 0.0034  1256 23 1176 47 1171 83 -8.0 1e 0.1934 0.0012 2.2669 0.0438 0.0837 0.0011  1140 7 1202 14 1285 25 12.4 1f 0.4518 0.0038 8.9620 0.2725 0.1573 0.0021  2403 17 2334 28 2426 22 1.1 1j 0.1762 0.0032 1.8723 0.1085 0.0802 0.0031  1046 18 1071 38 1202 74 14.0 1k 0.2029 0.0045 1.9228 0.1458 0.0722 0.0036  1191 24 1089 51 993 98 -21.9 1l 0.2157 0.0038 1.8972 0.1069 0.0782 0.0029  1259 20 1080 37 1153 72 -10.1 1o 0.1600 0.0020 1.6986 0.0684 0.0773 0.0022  957 11 1008 26 1128 55 16.3 1r 0.2077 0.0026 2.2623 0.0914 0.0844 0.0022  1217 14 1201 28 1302 49 7.2 2a 0.2253 0.0024 2.5613 0.0887 0.0846 0.0018  1310 12 1290 25 1305 41 -0.4 2b 0.3351 0.0048 4.9067 0.2549 0.1107 0.0028  1863 23 1803 44 1810 46 -3.4 2c 0.3330 0.0040 5.0813 0.2228 0.1110 0.0024  1853 19 1833 37 1816 39 -2.3 2d 0.2031 0.0030 2.2560 0.1145 0.0767 0.0024  1192 16 1199 36 1113 62 -7.8 2h 0.3369 0.0042 4.7440 0.2094 0.1072 0.0024  1872 20 1775 37 1751 40 -7.9 2j 0.2125 0.0051 1.9245 0.1467 0.0804 0.0041  1242 27 1090 51 1208 97 -3.2 2k 0.3396 0.0027 5.1848 0.1425 0.1084 0.0016  1885 13 1850 23 1773 27 -7.3 2l 0.5344 0.0072 13.3576 0.7123 0.1870 0.0036  2760 30 2705 50 2716 31 -2.0 2m 0.2203 0.0023 2.7463 0.0953 0.0879 0.0019  1283 12 1341 26 1381 40 7.8  178 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 2r 0.6797 0.0062 25.9706 0.9353 0.2940 0.0037  3343 24 3345 35 3439 20 3.6 3a 0.1840 0.0033 2.0496 0.1134 0.0842 0.0031  1089 18 1132 38 1297 70 17.4 3b 0.3576 0.0038 5.5794 0.2199 0.1098 0.0021  1971 18 1913 34 1795 35 -11.4 3c 0.5674 0.0104 12.5261 0.9268 0.1734 0.0046  2897 43 2645 70 2591 43 -14.7 3d 0.2984 0.0043 4.1030 0.1996 0.1025 0.0027  1684 21 1655 40 1671 48 -0.9 3e 0.5168 0.0078 12.8232 0.7348 0.2056 0.0044  2686 33 2667 54 2871 34 7.9 3f 0.3317 0.0042 5.1609 0.2312 0.1111 0.0025  1846 20 1846 38 1818 40 -1.8 3g 0.3101 0.0016 4.3562 0.0601 0.1036 0.0011  1741 8 1704 11 1689 19 -3.5 3h 0.2042 0.0012 2.1027 0.0340 0.0780 0.0010  1198 6 1150 11 1147 24 -4.9 3i 0.4404 0.0043 8.6466 0.3099 0.1451 0.0024  2353 19 2302 33 2289 28 -3.3 3j 0.5145 0.0039 12.7967 0.3598 0.1801 0.0024  2676 17 2665 26 2653 22 -1.0 3k 0.3327 0.0030 4.8684 0.1486 0.1088 0.0018  1852 14 1797 26 1779 31 -4.7 3l 0.4514 0.0057 8.5760 0.3975 0.1457 0.0030  2402 25 2294 42 2296 35 -5.5 3m 0.6471 0.0067 27.4499 1.1821 0.2881 0.0043  3217 26 3400 42 3408 23 7.1 3r 0.4368 0.0083 8.0001 0.5586 0.1479 0.0044  2336 37 2231 63 2322 50 -0.7 4a 0.3050 0.0020 4.3308 0.0854 0.1047 0.0014  1716 10 1699 16 1710 25 -0.4 4c 0.3350 0.0023 4.8987 0.1028 0.1129 0.0016  1863 11 1802 18 1846 25 -1.0 4e 0.3742 0.0030 5.7582 0.1523 0.1136 0.0018  2049 14 1940 23 1858 28 -12.0 5a 0.3496 0.0014 5.4712 0.0715 0.1144 0.0009  1933 7 1896 11 1870 14 -3.9 5b 0.2012 0.0008 2.1164 0.0247 0.0770 0.0006  1182 4 1154 8 1120 17 -6.0 5c 0.3376 0.0013 5.2144 0.0614 0.1110 0.0008  1875 6 1855 10 1817 13 -3.7 5d 0.3443 0.0027 5.1494 0.1424 0.1153 0.0015  1907 13 1844 24 1885 24 -1.3 5e 0.2237 0.0023 2.8318 0.1040 0.0853 0.0018  1302 12 1364 28 1322 40 1.7 5f 0.3203 0.0024 4.8042 0.1293 0.1078 0.0014  1791 12 1786 23 1762 24 -1.9 5g 0.3201 0.0020 4.7112 0.1010 0.1099 0.0012  1790 10 1769 18 1798 20 0.5 5h 0.3977 0.0014 8.0660 0.0832 0.1478 0.0010  2159 6 2238 9 2321 12 8.2 5i 0.3557 0.0018 5.9231 0.1065 0.1217 0.0012  1962 9 1965 16 1982 17 1.2 5j 0.2249 0.0017 2.5527 0.0616 0.0868 0.0013  1308 9 1287 18 1357 28 4.0 5k 0.4325 0.0023 8.8172 0.1729 0.1473 0.0014  2317 10 2319 18 2315 16 -0.1 5l 0.2249 0.0009 2.7096 0.0326 0.0865 0.0007  1308 5 1331 9 1349 17 3.4 5m 0.3253 0.0016 4.7071 0.0778 0.1087 0.0011  1816 8 1769 14 1777 18 -2.5 5n 0.3435 0.0016 5.4300 0.0864 0.1146 0.0011  1904 8 1890 14 1874 17 -1.8 5o 0.3353 0.0018 5.1720 0.0955 0.1120 0.0012  1864 9 1848 16 1831 19 -2.1 5p 0.2006 0.0012 2.0962 0.0381 0.0781 0.0010  1179 6 1148 12 1150 24 -2.7 5q 0.1753 0.0008 1.7403 0.0240 0.0717 0.0007  1041 4 1024 9 977 21 -7.2 5r 0.1690 0.0007 1.7232 0.0216 0.0738 0.0007  1006 4 1017 8 1037 19 3.2 5s 0.3723 0.0034 6.3544 0.2353 0.1178 0.0019  2040 16 2026 32 1923 29 -7.1 5t 0.5174 0.0031 13.1807 0.2950 0.1820 0.0019  2688 13 2693 21 2671 17 -0.8 6a 0.3540 0.0019 5.4929 0.0935 0.1162 0.0013  1954 9 1900 15 1898 19 -3.4 6b 0.1820 0.0018 1.8417 0.0600 0.0740 0.0016  1078 10 1061 21 1041 44 -3.9 6c 0.3379 0.0026 4.9016 0.1306 0.1100 0.0016  1877 12 1803 22 1800 26 -4.9 6d 0.3546 0.0024 5.6280 0.1337 0.1175 0.0016  1957 12 1920 20 1919 24 -2.3 6e 0.3452 0.0059 5.4964 0.3457 0.1180 0.0034  1912 28 1900 54 1927 51 0.9 6f 0.2018 0.0016 2.2596 0.0581 0.0779 0.0014  1185 8 1200 18 1144 34 -3.9 6g 0.2131 0.0042 2.5935 0.1673 0.0949 0.0038  1245 22 1299 47 1526 73 20.2 6h 0.4042 0.0026 7.5098 0.1671 0.1395 0.0018  2188 12 2174 20 2220 22 1.7 6i 0.3437 0.0026 4.9971 0.1352 0.1108 0.0017  1905 13 1819 23 1813 27 -5.8 6j 0.2184 0.0016 2.5099 0.0606 0.0825 0.0013  1273 9 1275 18 1258 31 -1.3  Sample CL06-47A, lower member Backbone Ranges Formation (7085058N, 519346E) 1a 0.3291 0.0031 5.2811 0.1805 0.1157 0.0019  1834 15 1866 29 1891 29 3.4 1b 0.3231 0.0059 5.5319 0.3891 0.1133 0.0036  1805 29 1906 60 1852 57 2.9 1d 0.3440 0.0017 5.6633 0.0925 0.1184 0.0010  1906 8 1926 14 1933 15 1.6 1e 0.4183 0.0032 8.6795 0.2446 0.1479 0.0017  2253 15 2305 26 2321 20 3.5 1f 0.3225 0.0059 5.0013 0.3461 0.1072 0.0035  1802 29 1820 59 1753 59 -3.2 1h 0.3400 0.0047 4.7285 0.2337 0.1132 0.0027  1887 23 1772 41 1851 43 -2.2 1i 0.3528 0.0017 5.6724 0.0876 0.1173 0.0010  1948 8 1927 13 1915 14 -2.0 1j 0.3448 0.0020 5.4867 0.1053 0.1204 0.0012  1910 10 1899 16 1962 17 3.1 1k 0.3513 0.0020 5.5876 0.1039 0.1190 0.0011  1941 9 1914 16 1942 17 0.0 1l 0.1791 0.0010 1.7715 0.0306 0.0741 0.0009  1062 6 1035 11 1045 24 -1.8 1m 0.3390 0.0039 5.4443 0.2256 0.1176 0.0023  1882 19 1892 36 1920 34 2.3 1n 0.3654 0.0065 6.1896 0.4220 0.1204 0.0036  2008 31 2003 60 1962 52 -2.7 1o 0.5471 0.0056 12.3736 0.4943 0.1810 0.0026  2813 23 2633 38 2662 23 -7.0 1p 0.3454 0.0018 5.4240 0.0877 0.1197 0.0010  1913 8 1889 14 1951 15 2.3 1q 0.3921 0.0022 6.9978 0.1338 0.1290 0.0012  2132 10 2111 17 2085 16 -2.7 1r 0.3646 0.0025 5.9736 0.1368 0.1251 0.0014  2004 12 1972 20 2029 19 1.5 1s 0.6174 0.0045 20.0044 0.6014 0.2325 0.0023  3099 18 3092 29 3070 15 -1.2 1t 0.3593 0.0026 5.5797 0.1383 0.1153 0.0014  1979 12 1913 21 1885 22 -5.8 2a 0.3445 0.0026 5.2428 0.1361 0.1137 0.0015  1908 12 1860 22 1859 23 -3.0  179 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 2b 0.5673 0.0055 15.0189 0.5877 0.1900 0.0025  2897 23 2816 37 2743 21 -7.0 2d 0.5157 0.0030 13.1027 0.2806 0.1841 0.0016  2681 13 2687 20 2690 14 0.4 2e 0.5025 0.0028 11.4604 0.2292 0.1717 0.0015  2624 12 2561 19 2574 14 -2.4 2f 0.3226 0.0027 4.8696 0.1383 0.1155 0.0017  1802 13 1797 24 1888 26 5.2 2g 0.3511 0.0021 5.5968 0.1082 0.1172 0.0012  1940 10 1916 17 1914 18 -1.6 2j 0.5072 0.0053 12.9356 0.5243 0.1920 0.0028  2645 23 2675 38 2759 24 5.1 2k 0.3480 0.0022 5.4304 0.1137 0.1169 0.0013  1925 11 1890 18 1910 19 -1.0 2l 0.3317 0.0030 4.5493 0.1331 0.1152 0.0018  1847 15 1740 24 1883 27 2.2 2n 0.3326 0.0026 5.0765 0.1320 0.1166 0.0016  1851 13 1832 22 1904 24 3.2 2o 0.3481 0.0025 5.4366 0.1296 0.1148 0.0014  1926 12 1891 20 1877 22 -3.0 2q 0.3498 0.0032 5.4654 0.1688 0.1182 0.0018  1934 15 1895 27 1929 27 -0.3 2r 0.3422 0.0026 5.1775 0.1345 0.1130 0.0015  1897 12 1849 22 1849 24 -3.1 2t 0.3397 0.0035 5.0075 0.1742 0.1165 0.0021  1885 17 1821 29 1902 31 1.0 3a 0.3591 0.0047 5.7140 0.2721 0.1174 0.0026  1978 22 1934 41 1917 39 -3.7 3b 0.3355 0.0043 5.2799 0.2517 0.1136 0.0026  1865 21 1866 41 1858 41 -0.4 3c 0.2080 0.0011 2.2761 0.0337 0.0828 0.0008  1218 6 1205 10 1264 20 3.9 3d 0.4833 0.0028 11.7762 0.2383 0.1786 0.0016  2542 12 2587 19 2640 15 4.5 3e 0.5323 0.0040 13.3244 0.3485 0.1976 0.0021  2751 17 2703 25 2806 17 2.4 3f 0.3374 0.0033 5.0708 0.1685 0.1162 0.0019  1874 16 1831 28 1899 30 1.5 3g 0.3317 0.0020 4.8256 0.0919 0.1132 0.0012  1847 10 1789 16 1851 19 0.3 3h 0.3429 0.0031 4.8490 0.1519 0.1108 0.0018  1901 15 1793 26 1813 29 -5.6 3i 0.3385 0.0025 5.4556 0.1341 0.1199 0.0015  1880 12 1894 21 1954 22 4.4 3j 0.3311 0.0030 4.6690 0.1436 0.1084 0.0017  1844 14 1762 26 1773 29 -4.6 3k 0.3545 0.0018 6.0718 0.0977 0.1229 0.0011  1956 9 1986 14 1999 16 2.5 3l 0.3346 0.0031 4.6062 0.1401 0.1135 0.0018  1861 15 1750 25 1856 29 -0.3 3n 0.3126 0.0023 4.7387 0.1160 0.1150 0.0015  1754 11 1774 21 1880 24 7.7 3o 0.3533 0.0039 6.1049 0.2417 0.1292 0.0024  1951 19 1991 35 2086 32 7.5 3p 0.4781 0.0031 12.5671 0.2826 0.1904 0.0019  2519 13 2648 21 2745 16 9.9 3q 0.3497 0.0062 5.3590 0.3572 0.1068 0.0033  1933 29 1878 57 1746 56 -12.4 4a 0.3708 0.0057 5.9772 0.3433 0.1184 0.0031  2033 27 1973 50 1932 46 -6.1 4c 0.3474 0.0026 5.4526 0.1403 0.1195 0.0016  1922 13 1893 22 1949 24 1.6 4d 0.3353 0.0079 4.9412 0.3921 0.1118 0.0043  1864 38 1809 67 1829 68 -2.2 4e 0.5079 0.0029 12.5304 0.2519 0.1817 0.0018  2648 13 2645 19 2668 16 1.0  Sample 07CL-1553A, upper member Backbone Ranges Formation (7138019N, 481960E) 1a 0.3699 0.0012 6.1203 0.0627 0.1179 0.0007  2029 6 1993 9 1925 10 -6.3 1b 0.3784 0.0035 5.9423 0.1969 0.1132 0.0017  2069 16 1967 29 1851 27 -13.8 1c 0.3644 0.0015 6.1958 0.0878 0.1211 0.0009  2003 7 2004 12 1973 13 -1.8 1d 0.3130 0.0020 5.0616 0.1074 0.1191 0.0013  1755 10 1830 18 1943 19 11.0 1e 0.3619 0.0026 5.7764 0.1447 0.1131 0.0013  1991 12 1943 22 1850 21 -8.8 1f 0.5179 0.0014 13.3239 0.1045 0.1849 0.0008  2690 6 2703 7 2697 7 0.3 1g 0.4542 0.0013 9.4875 0.0788 0.1531 0.0007  2414 6 2386 8 2381 8 -1.7 1h 0.3359 0.0011 5.2002 0.0525 0.1118 0.0007  1867 5 1853 9 1829 11 -2.4 1i 0.3555 0.0013 5.6608 0.0677 0.1178 0.0008  1961 6 1925 10 1924 12 -2.2 1j 0.4908 0.0016 11.8265 0.1218 0.1737 0.0009  2574 7 2591 10 2593 9 0.9 1l 0.3402 0.0014 5.2657 0.0721 0.1105 0.0008  1888 7 1863 12 1807 13 -5.1 1m 0.3312 0.0010 5.0327 0.0469 0.1114 0.0006  1844 5 1825 8 1822 10 -1.4 1n 0.3401 0.0010 5.2617 0.0425 0.1126 0.0006  1887 5 1863 7 1842 9 -2.8 1o 0.3370 0.0017 5.5988 0.0964 0.1186 0.0010  1872 8 1916 15 1935 15 3.8 1p 0.5320 0.0016 13.2786 0.1282 0.1801 0.0009  2750 7 2700 9 2654 8 -4.4 1q 0.3438 0.0010 5.3689 0.0475 0.1117 0.0006  1905 5 1880 8 1828 10 -4.9 1r 0.3269 0.0010 4.9886 0.0424 0.1114 0.0006  1823 5 1817 7 1823 10 0.0 1s 0.6076 0.0020 18.7038 0.2100 0.2251 0.0012  3061 8 3027 11 3017 8 -1.8 1t 0.3727 0.0011 6.2679 0.0537 0.1218 0.0006  2042 5 2014 8 1983 9 -3.5 2a 0.3514 0.0010 5.6351 0.0432 0.1176 0.0006  1941 5 1922 7 1919 9 -1.3 2b 0.5720 0.0020 16.9018 0.1977 0.2185 0.0012  2916 8 2929 11 2969 9 2.2 2c 0.3265 0.0015 5.0043 0.0766 0.1114 0.0009  1821 7 1820 13 1822 15 0.0 2d 0.3280 0.0020 4.8649 0.0986 0.1130 0.0012  1829 10 1796 17 1849 19 1.2 2e 0.3329 0.0025 5.1960 0.1352 0.1104 0.0014  1852 12 1852 22 1806 23 -3.0 2f 0.3276 0.0020 4.9894 0.0990 0.1132 0.0012  1827 9 1818 17 1851 18 1.5 2g 0.5098 0.0022 13.0942 0.1981 0.1877 0.0012  2656 9 2687 14 2722 11 3.0 2h 0.3793 0.0015 6.7890 0.0848 0.1266 0.0009  2073 7 2084 11 2051 12 -1.3 2i 0.3399 0.0012 5.3853 0.0572 0.1143 0.0007  1886 6 1883 9 1869 11 -1.1 2j 0.3570 0.0012 5.8524 0.0589 0.1164 0.0007  1968 6 1954 9 1902 11 -4.1 2k 0.3714 0.0019 5.8781 0.0998 0.1184 0.0010  2036 9 1958 15 1932 15 -6.3 2l 0.3428 0.0013 5.2449 0.0621 0.1123 0.0008  1900 6 1860 10 1838 12 -3.9 2m 0.3791 0.0010 6.7392 0.0510 0.1292 0.0007  2072 5 2078 7 2087 9 0.8 2n 0.3419 0.0012 5.5648 0.0613 0.1171 0.0008  1896 6 1911 9 1913 12 1.0 2o 0.4219 0.0012 8.3754 0.0661 0.1448 0.0008  2269 5 2273 7 2285 9 0.8  180 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 2p 0.3748 0.0014 6.7206 0.0806 0.1276 0.0009  2052 7 2075 11 2065 12 0.7 2q 0.5955 0.0026 17.9381 0.2778 0.2206 0.0014  3012 10 2986 15 2985 10 -1.1 2r 0.3089 0.0014 4.5955 0.0683 0.1059 0.0009  1735 7 1749 12 1730 15 -0.4 2s 0.3218 0.0018 5.0046 0.0910 0.1137 0.0011  1798 9 1820 15 1859 17 3.8 2t 0.5469 0.0024 14.1006 0.2179 0.1864 0.0013  2812 10 2757 15 2710 11 -4.6 3a 0.3481 0.0018 5.8659 0.1024 0.1170 0.0011  1925 9 1956 15 1912 16 -0.8 3b 0.3456 0.0015 5.4585 0.0750 0.1152 0.0009  1913 7 1894 12 1884 14 -1.8 3c 0.3356 0.0014 5.2980 0.0681 0.1120 0.0008  1865 7 1869 11 1832 13 -2.1 3d 0.3532 0.0011 5.5728 0.0494 0.1157 0.0007  1950 5 1912 8 1890 11 -3.6 3e 0.6208 0.0025 19.6705 0.2694 0.2293 0.0015  3113 10 3075 13 3047 10 -2.7 3f 0.5555 0.0036 14.9104 0.3527 0.1980 0.0018  2848 15 2810 23 2810 15 -1.7 3g 0.5362 0.0019 13.2926 0.1555 0.1849 0.0012  2767 8 2701 11 2697 10 -3.2 3h 0.3530 0.0013 5.8028 0.0677 0.1194 0.0008  1949 6 1947 10 1948 13 -0.1 3i 0.3005 0.0012 4.0677 0.0494 0.0994 0.0008  1694 6 1648 10 1613 14 -5.7 3j 0.3219 0.0014 5.1404 0.0696 0.1138 0.0009  1799 7 1843 12 1862 14 3.9 3k 0.3356 0.0023 5.1657 0.1177 0.1137 0.0013  1866 11 1847 19 1859 21 -0.4 3l 0.3449 0.0015 5.3487 0.0751 0.1143 0.0009  1910 7 1877 12 1869 15 -2.6 3m 0.5301 0.0024 13.3578 0.2089 0.1833 0.0014  2742 10 2705 15 2683 12 -2.7 3n 0.3314 0.0012 5.2491 0.0565 0.1146 0.0008  1845 6 1861 9 1873 13 1.7 3o 0.5530 0.0020 14.0945 0.1638 0.1872 0.0012  2838 8 2756 11 2718 11 -5.5 3p 0.4694 0.0020 11.3266 0.1558 0.1805 0.0013  2481 9 2550 13 2658 12 8.0 3q 0.3209 0.0025 4.9313 0.1293 0.1083 0.0015  1794 12 1808 22 1770 25 -1.5 3r 0.3423 0.0018 5.1615 0.0901 0.1116 0.0011  1898 9 1846 15 1825 17 -4.6 3s 0.5385 0.0025 13.9638 0.2262 0.1898 0.0015  2777 11 2747 15 2740 13 -1.6 3t 0.4524 0.0029 9.3800 0.2119 0.1545 0.0016  2406 13 2376 21 2396 17 -0.5 4a 0.3357 0.0013 5.1157 0.0584 0.1118 0.0008  1866 6 1839 10 1828 13 -2.4 4b 0.3351 0.0013 5.3464 0.0651 0.1154 0.0009  1863 6 1876 10 1885 14 1.4 4c 0.3460 0.0018 5.1522 0.0886 0.1111 0.0011  1915 9 1845 15 1817 17 -6.2 4d 0.3662 0.0014 5.8928 0.0695 0.1171 0.0009  2012 7 1960 10 1912 13 -6.0 4e 0.5029 0.0017 11.6805 0.1140 0.1711 0.0011  2626 7 2579 9 2569 11 -2.7  Sample CL06-44A, Sekwi Formation (7060521N, 501706E) 1a 0.3410 0.0019 5.2840 0.0983 0.1142 0.0011  1892 9 1866 16 1868 17 -1.5 1aa 0.3577 0.0025 5.5249 0.1317 0.1185 0.0014  1971 12 1905 21 1933 21 -2.3 1b 0.5248 0.0028 11.8667 0.2273 0.1727 0.0014  2720 12 2594 18 2584 13 -6.5 1bb 0.5210 0.0027 12.5815 0.2234 0.1815 0.0015  2703 11 2649 17 2667 14 -1.7 1cc 0.2431 0.0015 2.9008 0.0577 0.0888 0.0011  1403 8 1382 15 1400 23 -0.2 1d 0.1798 0.0013 1.7289 0.0370 0.0742 0.0011  1066 7 1019 14 1047 29 -1.9 1dd 0.5034 0.0021 12.5623 0.1594 0.1816 0.0013  2628 9 2647 12 2668 12 1.8 1e 0.1883 0.0015 1.9341 0.0480 0.0775 0.0013  1112 8 1093 17 1134 33 2.1 1f 0.2088 0.0013 2.3968 0.0440 0.0836 0.0010  1222 7 1242 13 1282 23 5.1 1g 0.4703 0.0020 10.3397 0.1381 0.1626 0.0011  2485 9 2466 12 2483 11 -0.1 1h 0.2499 0.0014 3.0144 0.0538 0.0883 0.0010  1438 7 1411 14 1388 20 -4.0 1i 0.2596 0.0011 3.2442 0.0386 0.0913 0.0007  1488 6 1468 9 1454 14 -2.6 1j 0.1904 0.0012 2.0000 0.0382 0.0775 0.0010  1124 6 1116 13 1135 25 1.1 1k 0.2352 0.0021 2.9903 0.0869 0.0915 0.0016  1362 11 1405 22 1458 32 7.3 1l 0.2904 0.0014 3.9617 0.0585 0.1005 0.0009  1643 7 1626 12 1634 16 -0.7 1m 0.2499 0.0016 3.2155 0.0662 0.0963 0.0012  1438 8 1461 16 1554 23 8.3 1n 0.3369 0.0032 4.9705 0.1641 0.1149 0.0019  1872 15 1814 28 1878 29 0.4 1p 0.4081 0.0018 9.5268 0.1295 0.1715 0.0012  2206 8 2390 13 2573 12 16.8 1q 0.2040 0.0013 2.1036 0.0402 0.0787 0.0010  1197 7 1150 13 1164 25 -3.1 1r 0.3007 0.0021 3.8870 0.0888 0.1027 0.0013  1695 11 1611 18 1673 23 -1.5 1s 0.2477 0.0023 2.7207 0.0792 0.0918 0.0016  1427 12 1334 22 1462 33 2.7 1t 0.1940 0.0012 2.0867 0.0376 0.0826 0.0010  1143 6 1144 12 1259 23 10.1 1u 0.1762 0.0015 1.9178 0.0515 0.0806 0.0015  1046 8 1087 18 1213 35 14.9 1v 0.2987 0.0020 4.2908 0.0965 0.1032 0.0013  1685 10 1692 19 1682 22 -0.2 1w 0.2707 0.0023 3.3283 0.0957 0.0929 0.0015  1544 12 1488 22 1485 31 -4.5 1x 0.1772 0.0013 1.8201 0.0401 0.0758 0.0012  1052 7 1053 14 1089 30 3.7 1y 0.2510 0.0023 3.1407 0.0958 0.0943 0.0017  1444 12 1443 24 1514 33 5.2 1z 0.2678 0.0025 3.1241 0.0962 0.0919 0.0016  1529 13 1439 24 1464 34 -5.0 2a 0.1599 0.0010 1.4684 0.0276 0.0725 0.0010  956 6 918 11 999 28 4.6 2b 0.3251 0.0024 4.7980 0.1217 0.1061 0.0014  1815 11 1785 21 1733 24 -5.4 2c 0.2957 0.0031 4.0127 0.1474 0.1002 0.0020  1670 15 1637 30 1628 36 -3.0 2d 0.1807 0.0008 1.7789 0.0212 0.0729 0.0007  1071 4 1038 8 1012 18 -6.3 2f 0.1966 0.0012 2.1184 0.0384 0.0792 0.0010  1157 6 1155 13 1177 25 1.9 2g 0.2516 0.0014 3.1739 0.0546 0.0932 0.0010  1447 7 1451 13 1492 21 3.4 2i 0.4273 0.0033 8.5525 0.2390 0.1536 0.0019  2293 15 2292 25 2387 21 4.6 2k 0.2405 0.0020 2.5381 0.0659 0.0866 0.0015  1389 10 1283 19 1351 32 -3.1 2l 0.2257 0.0011 2.6230 0.0364 0.0856 0.0009  1312 6 1307 10 1330 19 1.5  181 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 2m 0.2566 0.0015 3.1134 0.0541 0.0904 0.0010  1473 7 1436 13 1434 21 -3.0 2n 0.2587 0.0022 3.1600 0.0884 0.0936 0.0016  1483 11 1447 22 1500 31 1.2 2o 0.1933 0.0012 1.8838 0.0345 0.0733 0.0010  1139 6 1075 12 1023 26 -12.4 2p 0.2553 0.0018 3.0179 0.0695 0.0877 0.0013  1466 9 1412 18 1375 27 -7.4 3a 0.1880 0.0016 1.9093 0.0541 0.0754 0.0014  1110 9 1084 19 1080 36 -3.0 3b 0.2588 0.0012 3.3128 0.0515 0.0927 0.0008  1484 6 1484 12 1482 17 -0.2 3c 0.1813 0.0011 1.9163 0.0372 0.0763 0.0010  1074 6 1087 13 1103 25 2.9 3d 0.2028 0.0034 2.5164 0.1428 0.0886 0.0031  1190 18 1277 41 1396 64 16.1 3e 0.2101 0.0010 2.3535 0.0351 0.0837 0.0008  1230 5 1229 11 1285 18 4.7 3f 0.2316 0.0009 2.8636 0.0328 0.0896 0.0007  1343 4 1372 9 1417 14 5.8 3g 0.2814 0.0010 3.9947 0.0472 0.1016 0.0007  1599 5 1633 10 1653 13 3.7 3i 0.2588 0.0008 3.1924 0.0290 0.0904 0.0006  1484 4 1455 7 1434 12 -3.9 3j 0.3078 0.0009 4.4540 0.0402 0.1050 0.0006  1730 5 1722 7 1714 11 -1.1 3k 0.3036 0.0010 4.6076 0.0474 0.1077 0.0007  1709 5 1751 9 1762 12 3.4 3l 0.2069 0.0010 2.3065 0.0351 0.0813 0.0008  1212 5 1214 11 1228 19 1.4 3m 0.1781 0.0019 1.7459 0.0610 0.0714 0.0016  1057 10 1026 23 968 46 -10.0 3n 0.1814 0.0014 2.0447 0.0531 0.0800 0.0013  1074 8 1131 18 1198 32 11.2 3o 0.2476 0.0010 3.0232 0.0363 0.0905 0.0007  1426 5 1414 9 1436 15 0.8 3p 0.2527 0.0011 3.2604 0.0470 0.0901 0.0008  1452 6 1472 11 1428 16 -1.9 3q 0.2950 0.0015 4.2196 0.0741 0.1051 0.0010  1666 8 1678 14 1716 18 3.3 3r 0.2837 0.0018 3.9390 0.0880 0.1014 0.0012  1610 9 1622 18 1650 22 2.7 3s 0.2430 0.0014 2.9452 0.0539 0.0921 0.0010  1402 7 1394 14 1469 21 5.0 3t 0.2256 0.0013 2.7014 0.0511 0.0834 0.0010  1312 7 1329 14 1278 22 -3.0 4a 0.2239 0.0020 2.6182 0.0794 0.0870 0.0016  1303 11 1306 22 1359 34 4.6 4b 0.2540 0.0015 3.2629 0.0645 0.0907 0.0011  1459 8 1472 15 1440 22 -1.5 4c 0.1728 0.0011 1.8859 0.0369 0.0768 0.0010  1027 6 1076 13 1115 26 8.5 4d 0.2576 0.0018 3.7080 0.0914 0.0985 0.0013  1477 9 1573 20 1595 25 8.3 4e 0.1827 0.0012 1.8545 0.0382 0.0743 0.0010  1082 6 1065 14 1051 28 -3.2 4f 0.1922 0.0008 2.1120 0.0261 0.0781 0.0007  1133 4 1153 9 1149 18 1.5 4g 0.2464 0.0021 3.2594 0.0984 0.0920 0.0016  1420 11 1471 23 1468 32 3.7 4h 0.1778 0.0016 1.8480 0.0546 0.0758 0.0015  1055 9 1063 19 1091 40 3.6 4i 0.2095 0.0009 2.2866 0.0313 0.0798 0.0008  1226 5 1208 10 1192 19 -3.1 4j 0.1955 0.0009 2.0804 0.0290 0.0766 0.0008  1151 5 1142 10 1110 20 -4.0  Sample EM06-113B, Franklin Mountain Formation (7135557N, 516721E) 1a 0.1949 0.0019 2.0241 0.0649 0.0763 0.0016  1148 10 1124 22 1104 42 -4.4 1b 0.2170 0.0010 2.4989 0.0322 0.0829 0.0007  1266 5 1272 9 1266 17 0.0 1c 0.2069 0.0008 2.1585 0.0209 0.0772 0.0006  1212 4 1168 7 1126 15 -8.4 1d 0.6677 0.0031 21.3193 0.3316 0.2442 0.0018  3297 12 3153 15 3148 11 -6.1 1e 0.2037 0.0040 2.0583 0.1340 0.0781 0.0033  1195 21 1135 45 1149 83 -4.4 1f 0.2234 0.0026 2.3531 0.0918 0.0799 0.0019  1300 14 1229 28 1194 47 -9.8 1g 0.2404 0.0012 2.8774 0.0431 0.0894 0.0009  1389 6 1376 11 1412 19 1.8 1h 0.2048 0.0014 2.1267 0.0457 0.0785 0.0011  1201 8 1158 15 1160 28 -3.8 1i 0.2200 0.0009 2.4012 0.0270 0.0815 0.0007  1282 5 1243 8 1234 16 -4.2 1j 0.2121 0.0008 2.3836 0.0240 0.0819 0.0006  1240 4 1238 7 1244 15 0.3 1k 0.2408 0.0024 3.0182 0.1010 0.0916 0.0018  1391 12 1412 26 1458 37 5.1 1l 0.1716 0.0008 1.7281 0.0236 0.0738 0.0007  1021 5 1019 9 1037 20 1.6 1m 0.1996 0.0009 2.0761 0.0259 0.0777 0.0007  1173 5 1141 9 1138 18 -3.4 1n 0.2074 0.0021 2.1343 0.0678 0.0812 0.0017  1215 11 1160 22 1227 40 1.1 1p 0.2035 0.0023 2.1381 0.0771 0.0813 0.0019  1194 12 1161 25 1228 46 3.1 1q 0.2292 0.0015 2.5011 0.0500 0.0829 0.0011  1330 8 1272 14 1267 25 -5.5 1r 0.2658 0.0018 3.1306 0.0690 0.0902 0.0012  1520 9 1440 17 1430 25 -7.0 1t 0.5154 0.0030 13.3798 0.2772 0.1855 0.0017  2680 13 2707 20 2703 15 1.1 2a 0.2480 0.0015 3.1579 0.0608 0.0918 0.0011  1428 8 1447 15 1462 23 2.6 2b 0.5341 0.0036 13.0996 0.3166 0.1861 0.0019  2759 15 2687 23 2708 17 -2.3 2c 0.5759 0.0044 13.7971 0.3914 0.1894 0.0021  2932 18 2736 27 2737 18 -8.9 2d 0.2232 0.0025 2.5458 0.0961 0.0838 0.0020  1299 13 1285 28 1288 45 -0.9 2e 0.3651 0.0028 5.8517 0.1534 0.1209 0.0016  2006 13 1954 23 1969 23 -2.2 2f 0.2124 0.0018 2.3539 0.0621 0.0821 0.0014  1242 9 1229 19 1247 33 0.5 2g 0.1910 0.0012 1.9823 0.0374 0.0766 0.0010  1127 6 1110 13 1112 26 -1.5 2h 0.3082 0.0018 4.0565 0.0748 0.0992 0.0011  1732 9 1646 15 1609 21 -8.7 2i 0.2554 0.0017 3.1719 0.0687 0.0936 0.0013  1466 9 1450 17 1500 25 2.5 2j 0.1961 0.0012 2.0608 0.0386 0.0783 0.0010  1155 7 1136 13 1154 26 -0.1 2k 0.1880 0.0011 1.8911 0.0340 0.0764 0.0010  1111 6 1078 12 1105 26 -0.5 2l 0.2538 0.0014 3.1040 0.0527 0.0909 0.0010  1458 7 1434 13 1445 21 -1.0 2m 0.2080 0.0017 2.5740 0.0690 0.0868 0.0015  1218 9 1293 20 1357 32 11.2 2n 0.2238 0.0012 2.7100 0.0421 0.0893 0.0010  1302 6 1331 12 1411 21 8.5 2p 0.2268 0.0023 2.6241 0.0865 0.0850 0.0017  1318 12 1307 24 1316 39 -0.1 2q 0.2495 0.0018 2.8882 0.0673 0.0892 0.0013  1436 9 1379 18 1408 28 -2.2  182 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 2r 0.2498 0.0014 3.0176 0.0494 0.0892 0.0010  1437 7 1412 12 1409 21 -2.3 2s 0.5596 0.0052 13.2538 0.4627 0.1926 0.0026  2865 21 2698 33 2765 22 -4.5 2t 0.2285 0.0036 2.4421 0.1290 0.0803 0.0026  1327 19 1255 38 1205 63 -11.2 3a 0.1903 0.0022 1.6846 0.0611 0.0715 0.0018  1123 12 1003 23 972 51 -16.8 3b 0.2029 0.0025 2.0181 0.0789 0.0799 0.0021  1191 13 1122 27 1195 50 0.4 3f 0.3654 0.0022 5.7053 0.1086 0.1185 0.0013  2008 10 1932 16 1934 20 -4.4 3g 0.2454 0.0015 2.9457 0.0550 0.0878 0.0011  1415 8 1394 14 1379 24 -2.9 3h 0.1940 0.0014 2.0353 0.0451 0.0776 0.0012  1143 8 1127 15 1136 31 -0.6 3i 0.1932 0.0011 2.0586 0.0328 0.0781 0.0009  1139 6 1135 11 1150 24 1.1 3j 0.1846 0.0024 1.7651 0.0746 0.0716 0.0021  1092 13 1033 27 973 58 -13.3 3m 0.3576 0.0018 5.6139 0.0811 0.1172 0.0012  1971 8 1918 12 1914 18 -3.5 3o 0.2167 0.0020 2.3954 0.0703 0.0863 0.0017  1265 11 1241 21 1345 37 6.6 3p 0.2648 0.0016 3.3569 0.0604 0.0911 0.0011  1514 8 1494 14 1448 23 -5.1 3r 0.4913 0.0053 9.5343 0.3827 0.1580 0.0028  2576 23 2391 37 2434 29 -7.1 3s 0.2576 0.0017 3.1629 0.0650 0.0911 0.0012  1477 9 1448 16 1448 26 -2.2 3t 0.2581 0.0022 3.2792 0.0943 0.0916 0.0016  1480 11 1476 22 1460 33 -1.6 4a 0.4948 0.0038 10.9021 0.3009 0.1655 0.0022  2592 16 2515 26 2513 22 -3.8 4b 0.3487 0.0029 5.4154 0.1548 0.1126 0.0018  1929 14 1887 25 1841 28 -5.5 4c 0.1858 0.0012 1.8547 0.0344 0.0750 0.0011  1099 6 1065 12 1069 28 -3.1 4e 0.2035 0.0026 2.1981 0.0894 0.0812 0.0022  1194 14 1180 28 1226 52 2.9  Sample EM06-70A, Tsichu Group (7062212N, 520096E) 1a 0.2079 0.0005 2.2066 0.0129 0.0777 0.0003  1218 3 1183 4 1140 9 -7.5 1b 0.1873 0.0004 1.8990 0.0103 0.0739 0.0003  1107 2 1081 4 1038 8 -7.2 1c 0.1817 0.0004 1.8421 0.0084 0.0737 0.0003  1076 2 1061 3 1034 8 -4.5 1d 0.2781 0.0009 3.6225 0.0336 0.0953 0.0005  1582 5 1555 7 1534 11 -3.5 1e 0.2456 0.0006 2.9879 0.0157 0.0887 0.0003  1416 3 1405 4 1398 7 -1.4 1f 0.2656 0.0007 3.3821 0.0211 0.0929 0.0004  1518 3 1500 5 1486 8 -2.4 1g 0.5335 0.0019 13.6519 0.1655 0.1866 0.0009  2756 8 2726 11 2712 8 -2.0 1h 0.3273 0.0008 4.7809 0.0259 0.1060 0.0004  1826 4 1782 5 1732 7 -6.2 1i 0.1874 0.0005 1.9406 0.0128 0.0756 0.0004  1107 3 1095 4 1084 10 -2.3 1j 0.5427 0.0012 13.9188 0.0693 0.1902 0.0006  2795 5 2744 5 2744 6 -2.3 1k 0.3350 0.0008 5.0336 0.0328 0.1098 0.0005  1863 4 1825 6 1797 8 -4.2 1l 0.3154 0.0009 4.6064 0.0376 0.1050 0.0005  1767 4 1750 7 1714 9 -3.6 1m 0.2337 0.0005 2.7265 0.0121 0.0847 0.0003  1354 3 1336 3 1309 7 -3.8 1n 0.1911 0.0005 1.9507 0.0105 0.0750 0.0003  1127 2 1099 4 1067 8 -6.1 1o 0.1961 0.0005 2.0260 0.0130 0.0763 0.0004  1154 3 1124 4 1103 9 -5.1 1p 0.2591 0.0007 3.2831 0.0227 0.0926 0.0004  1485 4 1477 5 1480 9 -0.4 1q 0.2995 0.0007 4.1486 0.0230 0.1012 0.0004  1689 3 1664 5 1647 7 -2.9 1r 0.2752 0.0009 3.4377 0.0332 0.0934 0.0006  1567 5 1513 8 1497 11 -5.3 1s 0.1785 0.0006 1.8676 0.0155 0.0752 0.0005  1059 3 1070 6 1073 12 1.4 1t 0.1954 0.0007 2.1446 0.0222 0.0785 0.0006  1150 4 1163 7 1160 14 0.9 2a 0.2902 0.0007 3.9500 0.0253 0.0998 0.0004  1643 4 1624 5 1620 8 -1.6 2b 0.5883 0.0016 16.4354 0.1367 0.2057 0.0008  2983 7 2903 8 2872 7 -4.8 2c 0.1935 0.0006 2.1129 0.0157 0.0788 0.0004  1141 3 1153 5 1168 11 2.5 2d 0.2834 0.0007 3.9142 0.0249 0.1004 0.0004  1609 4 1617 5 1632 8 1.6 2e 0.2306 0.0006 2.7025 0.0151 0.0853 0.0004  1338 3 1329 4 1323 8 -1.2 2f 0.2900 0.0008 3.9237 0.0275 0.0997 0.0005  1642 4 1619 6 1618 9 -1.6 2g 0.2017 0.0010 2.4488 0.0372 0.0888 0.0009  1185 6 1257 11 1400 19 16.8 2h 0.1742 0.0008 1.8004 0.0259 0.0740 0.0008  1035 5 1046 9 1041 20 0.6 2i 0.2925 0.0008 4.0954 0.0320 0.1012 0.0005  1654 4 1653 6 1646 9 -0.6 2j 0.2186 0.0007 2.6015 0.0226 0.0852 0.0005  1275 4 1301 6 1320 11 3.8 2k 0.3249 0.0007 5.1621 0.0254 0.1173 0.0004  1814 4 1846 4 1915 7 6.1 2l 0.3570 0.0008 5.9070 0.0315 0.1205 0.0005  1968 4 1962 5 1963 7 -0.3 2m 0.2824 0.0009 3.8533 0.0321 0.0989 0.0005  1603 4 1604 7 1604 10 0.0 2n 0.1815 0.0008 1.9384 0.0253 0.0759 0.0007  1075 4 1094 9 1092 18 1.7 2o 0.3317 0.0009 5.1122 0.0381 0.1125 0.0005  1847 4 1838 6 1841 8 -0.4 2p 0.1914 0.0007 1.9700 0.0211 0.0749 0.0006  1129 4 1105 7 1065 15 -6.5 2q 0.1589 0.0006 1.5384 0.0176 0.0697 0.0006  951 4 946 7 920 17 2r 0.0684 0.0003 0.5304 0.0055 0.0565 0.0005  427 2 432 4 469 21 2s 0.0701 0.0002 0.5325 0.0033 0.0553 0.0003  437 1 434 2 426 12 2t 0.1782 0.0007 1.7976 0.0188 0.0741 0.0006  1057 4 1045 7 1044 15 -1.4 3a 0.3102 0.0009 4.5431 0.0337 0.1056 0.0005  1742 4 1739 6 1725 9 -1.1 3b 0.2798 0.0011 4.0474 0.0466 0.1048 0.0007  1590 5 1644 9 1710 12 7.9 3c 0.2056 0.0006 2.2995 0.0191 0.0819 0.0005  1205 3 1212 6 1244 11 3.4 3d 0.0694 0.0002 0.5266 0.0036 0.0549 0.0003  433 1 430 2 408 14 3e 0.2606 0.0007 3.3771 0.0223 0.0938 0.0004  1493 4 1499 5 1505 9 0.9 3f 0.2122 0.0007 2.3472 0.0197 0.0811 0.0005  1241 4 1227 6 1223 12 -1.6 3g 0.1876 0.0021 2.0298 0.0696 0.0794 0.0019  1108 11 1126 23 1181 46 6.7  183 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 3h 0.3008 0.0009 4.3279 0.0324 0.1043 0.0005  1695 4 1699 6 1703 9 0.5 3i 0.1728 0.0006 1.7694 0.0157 0.0736 0.0005  1028 3 1034 6 1031 13 0.3 3j 0.4289 0.0029 8.7571 0.2011 0.1518 0.0015  2301 13 2313 21 2367 17 3.3 3k 0.1727 0.0006 1.8041 0.0167 0.0754 0.0005  1027 3 1047 6 1078 14 5.2 3l 0.5551 0.0017 14.6313 0.1378 0.1929 0.0009  2847 7 2792 9 2767 8 -3.5 3m 0.1752 0.0005 1.8144 0.0136 0.0745 0.0004  1041 3 1051 5 1055 12 1.5 3n 0.1928 0.0006 2.0148 0.0153 0.0760 0.0004  1137 3 1121 5 1094 11 -4.3 3o 0.2857 0.0010 3.8446 0.0362 0.0982 0.0006  1620 5 1602 8 1591 11 -2.1 3p 0.1910 0.0005 2.0516 0.0134 0.0773 0.0004  1127 3 1133 4 1129 10 0.2 3q 0.3656 0.0010 6.1054 0.0395 0.1209 0.0006  2009 4 1991 6 1970 8 -2.3 3r 0.1940 0.0007 2.0464 0.0219 0.0763 0.0006  1143 4 1131 7 1103 15 -4.0 3s 0.3298 0.0009 5.1304 0.0374 0.1118 0.0006  1837 4 1841 6 1828 9 -0.6 3t 0.2862 0.0009 3.9177 0.0352 0.1001 0.0006  1623 5 1617 7 1625 11 0.2 4a 0.2578 0.0007 3.1636 0.0205 0.0899 0.0004  1478 4 1448 5 1423 9 -4.3 4b 0.3198 0.0008 4.6207 0.0283 0.1058 0.0005  1789 4 1753 5 1729 8 -4.0 4c 0.1755 0.0007 1.7873 0.0202 0.0746 0.0006  1042 4 1041 7 1056 17 1.4 4d 0.3232 0.0012 4.7219 0.0544 0.1084 0.0007  1805 6 1771 10 1773 12 -2.1 4e 0.1762 0.0005 1.7501 0.0109 0.0725 0.0004  1046 3 1027 4 1000 10 -5.0  Sample DL07-01, Cretaceous clastic foredeep (7054655N, 516721E) 1b 0.0768 0.0005 0.5572 0.0105 0.0525 0.0009  477 3 450 7 305 38 1d 0.2728 0.0017 3.3654 0.0668 0.0922 0.0011  1555 9 1496 16 1472 22 -6.3 1f 0.1874 0.0008 1.9288 0.0246 0.0732 0.0007  1107 5 1091 9 1021 18 -9.2 1g 0.2000 0.0011 2.1787 0.0367 0.0776 0.0009  1175 6 1174 12 1137 22 -3.7 1h 0.3007 0.0014 3.7777 0.0530 0.0934 0.0008  1695 7 1588 11 1495 16 -15.2 1j 0.1776 0.0020 1.5925 0.0582 0.0662 0.0017  1054 11 967 23 812 52 -32.4 1k 0.2080 0.0012 1.9983 0.0332 0.0741 0.0009  1218 6 1115 11 1043 23 -18.4 1m 0.0696 0.0006 0.4850 0.0119 0.0529 0.0012  434 4 402 8 325 50 1n 0.1824 0.0009 1.8139 0.0255 0.0729 0.0007  1080 5 1051 9 1010 20 -7.5 1q 0.3069 0.0034 3.7761 0.1382 0.0994 0.0020  1725 17 1588 29 1613 37 -8.0 1t 0.1803 0.0013 1.6960 0.0383 0.0704 0.0011  1068 7 1007 14 940 32 -14.9 2a 0.0719 0.0012 0.5207 0.0254 0.0548 0.0024  448 7 426 17 403 95 2d 0.5332 0.0057 14.4572 0.6054 0.1879 0.0027  2755 24 2780 40 2724 24 -1.4 2e 0.3395 0.0036 5.3395 0.1965 0.1154 0.0021  1884 17 1875 31 1886 32 0.1 2f 0.2125 0.0009 2.2017 0.0273 0.0760 0.0007  1242 5 1182 9 1095 18 -14.7 2g 0.1881 0.0023 1.8421 0.0743 0.0701 0.0019  1111 13 1061 27 932 55 -21.0 2h 0.1726 0.0017 1.7801 0.0535 0.0762 0.0016  1027 9 1038 20 1100 41 7.2 2k 0.2525 0.0034 3.8295 0.1737 0.1022 0.0025  1451 17 1599 37 1665 44 14.3 2m 0.3099 0.0028 4.1613 0.1291 0.0970 0.0016  1740 14 1666 25 1568 31 -12.5 2n 0.2312 0.0018 2.6266 0.0673 0.0834 0.0013  1341 10 1308 19 1278 31 -5.4 2p 0.0766 0.0007 0.5614 0.0145 0.0568 0.0013  476 4 452 9 483 50 2r 0.2774 0.0011 3.4570 0.0405 0.0932 0.0008  1578 6 1518 9 1491 15 -6.6 2s 0.0681 0.0007 0.5285 0.0164 0.0564 0.0016  425 4 431 11 465 61 2t 0.5497 0.0020 14.6399 0.1638 0.1954 0.0014  2824 8 2792 11 2788 11 -1.6 3a 0.0659 0.0006 0.4662 0.0117 0.0527 0.0012  411 3 389 8 314 51 3b 0.0683 0.0008 0.5257 0.0188 0.0528 0.0017  426 5 429 13 320 71 3c 0.3889 0.0024 6.7082 0.1407 0.1337 0.0015  2118 11 2074 19 2147 19 1.6 3f 0.2679 0.0016 3.0546 0.0582 0.0863 0.0011  1530 8 1421 15 1346 23 -15.4 3g 0.3241 0.0038 4.8402 0.2016 0.1092 0.0025  1810 19 1792 35 1786 41 -1.5 3h 0.2099 0.0010 2.2471 0.0300 0.0779 0.0008  1228 5 1196 9 1144 19 -8.1 3i 0.3517 0.0017 5.7083 0.0894 0.1224 0.0012  1943 8 1933 14 1992 17 2.9 3j 0.2216 0.0018 2.5190 0.0655 0.0825 0.0014  1290 9 1278 19 1257 32 -2.9 3l 0.1947 0.0008 2.1272 0.0260 0.0775 0.0007  1147 5 1158 8 1133 19 -1.3 3n 0.2234 0.0029 2.3683 0.0971 0.0842 0.0022  1300 15 1233 29 1298 50 -0.2 3p 0.2842 0.0014 3.9034 0.0592 0.1014 0.0010  1612 7 1614 12 1650 19 2.6 3q 0.0659 0.0004 0.5020 0.0077 0.0538 0.0008  412 2 413 5 362 31 3r 0.2708 0.0010 4.6215 0.0435 0.1231 0.0010  1545 5 1753 8 2002 14 25.6 3s 0.0679 0.0005 0.5009 0.0119 0.0537 0.0012  424 3 412 8 360 48 3t 0.2502 0.0013 2.9952 0.0463 0.0874 0.0010  1440 7 1406 12 1369 21 -5.7 4a 0.0626 0.0004 0.4848 0.0084 0.0562 0.0009  392 2 401 6 461 35 4b 0.3065 0.0018 5.1653 0.0972 0.1225 0.0014  1723 9 1847 16 1993 20 15.4 4c 0.2071 0.0009 2.1731 0.0277 0.0772 0.0008  1213 5 1173 9 1126 20 -8.5 4d 0.0729 0.0005 0.5852 0.0121 0.0577 0.0011  454 3 468 8 519 40 4e 0.2130 0.0012 2.1525 0.0379 0.0761 0.0010  1245 7 1166 12 1099 25 -14.6 5a 0.2212 0.0006 2.4368 0.0159 0.0792 0.0004  1288 3 1254 5 1178 11 -10.3 5b 0.4963 0.0020 11.2607 0.1535 0.1638 0.0011  2598 9 2545 13 2495 11 -5.0 5c 0.5529 0.0016 14.4891 0.1225 0.1895 0.0010  2837 7 2782 8 2738 9 -4.5 5d 0.0711 0.0003 0.5471 0.0051 0.0555 0.0005  443 2 443 3 433 18 5e 0.5392 0.0027 14.3078 0.2515 0.1918 0.0014  2780 11 2770 17 2758 12 -1.0  184 Grain   Isotopic ratios (±1")      Apparent ages (±1",Ma)  % No. 206Pb/238U 207Pb/235U 207Pb/206Pb   206Pb/238U 207Pb/235U 207Pb/206Pb Disc. 5f 0.3020 0.0009 4.1918 0.0329 0.1012 0.0006  1701 4 1672 6 1646 11 -3.9 5g 0.3044 0.0018 4.1792 0.0793 0.1014 0.0011  1713 9 1670 16 1650 20 -4.4 5h 0.3135 0.0008 4.5473 0.0293 0.1062 0.0006  1758 4 1740 5 1735 10 -1.5 5i 0.3262 0.0009 4.8948 0.0344 0.1097 0.0006  1820 4 1801 6 1794 10 -1.7 5j 0.2004 0.0006 2.1339 0.0152 0.0777 0.0005  1177 3 1160 5 1139 12 -3.6 5k 0.2213 0.0007 2.4826 0.0217 0.0824 0.0005  1289 4 1267 6 1254 13 -3.1 5l 0.1984 0.0011 2.0918 0.0331 0.0783 0.0009  1167 6 1146 11 1154 22 -1.2 5m 0.2309 0.0008 2.5976 0.0262 0.0859 0.0006  1339 4 1300 7 1336 14 -0.3 5n 0.5302 0.0015 13.6881 0.1029 0.1891 0.0010  2742 6 2728 7 2734 9 -0.4 5o 0.2648 0.0008 3.3698 0.0262 0.0931 0.0006  1514 4 1497 6 1489 11 -1.9 5p 0.2944 0.0009 4.0325 0.0351 0.1004 0.0006  1664 5 1641 7 1632 12 -2.2 5q 0.2976 0.0010 4.3472 0.0379 0.1084 0.0007  1679 5 1702 7 1773 12 6.0 5r 0.3348 0.0012 5.2406 0.0544 0.1131 0.0008  1862 6 1859 9 1850 12 -0.7 5s 0.4802 0.0015 10.7348 0.0881 0.1643 0.0010  2528 6 2500 8 2501 10 -1.3 5t 0.1774 0.0006 1.8144 0.0160 0.0751 0.0005  1053 3 1051 6 1071 14 1.8 6a 0.3623 0.0017 6.1801 0.0937 0.1216 0.0011  1993 8 2002 13 1980 15 -0.8 6b 0.3252 0.0010 4.6786 0.0336 0.1067 0.0007  1815 5 1763 6 1744 11 -4.7 6c 0.3610 0.0012 5.7609 0.0499 0.1167 0.0008  1987 6 1941 8 1906 12 -4.9 6d 0.5178 0.0016 12.6658 0.1041 0.1811 0.0012  2690 7 2655 8 2663 11 -1.2 6e 0.2111 0.0015 2.4662 0.0520 0.0853 0.0012  1235 8 1262 15 1321 27 7.2 6f 0.5672 0.0021 16.7947 0.1817 0.2142 0.0015  2896 8 2923 10 2938 11 1.7 6g 0.1825 0.0007 1.8863 0.0185 0.0755 0.0006  1081 4 1076 7 1081 16 0.0 6h 0.2487 0.0011 3.0857 0.0387 0.0899 0.0008  1432 6 1429 10 1422 17 -0.7 6i 0.5307 0.0022 13.7069 0.1782 0.1861 0.0014  2745 9 2730 12 2708 13 -1.6 6j 0.1836 0.0008 1.9017 0.0215 0.0757 0.0007  1086 4 1082 8 1087 18 0.0                               185     APPENDIX C  Chapter 3 Analytical data                                        186  Table C.1. Microprobe analysis of clinopyroxene from sample 1508  Sample 1508       1508 Position core    rim  Average Min Max  Average Min Max n 26       29 Major elements (wt.%) Na2O 0.31 0.25 0.37  0.41 0.31 0.50 MgO 15.21 14.28 16.33  13.09 12.28 13.91 Al2O3 3.66 2.42 5.03  6.34 4.83 7.59 SIO2 49.01 46.18 52.17  45.39 42.69 47.04 CaO 23.23 22.34 23.65  22.69 21.78 23.27 TiO2 1.37 0.83 1.94  2.72 2.14 3.32 Cr2O3 0.28 0.00 0.92  0.05 0.00 0.17 MnO 0.09 0.06 0.16  0.15 0.08 0.26 FeO 1.48 0.00 3.63  2.24 0.46 4.05 NiO 0.02 0.00 0.08  0.02 0.00 0.07 Total 98.99 96.52 100.72  98.74 97.72 100.24  Structural Formula (6 oxygen basis) Na 0.02 0.02 0.03  0.03 0.02 0.04 Mg 0.85 0.79 0.90  0.74 0.69 0.79 Al 0.16 0.10 0.22  0.28 0.22 0.34 Si 1.83 1.76 1.91  1.72 1.64 1.77 Ca 0.93 0.91 0.96  0.92 0.89 0.94 Ti 0.04 0.02 0.05  0.08 0.06 0.09 Cr 0.01 0.00 0.03  0.00 0.00 0.01 Mn 0.00 0.00 0.01  0.00 0.00 0.01 Fe2+ 0.05 0.00 0.11  0.07 0.02 0.13 Fe3+* 0.12 0.05 0.19  0.16 0.10 0.23 Ni 0.00 0.00 0.00  0.00 0.00 0.00 Mg # 83.46 78.58 86.78  76.11 72.27 78.97  Wo 47.84 46.62 49.07  48.68 46.38 50.16 En 43.51 39.96 45.98  39.06 37.09 40.99 Fs 8.65 7.00 11.97  12.26 10.83 14.23  * calculated from stoichiometry Mg # = Mg/(Mg+Fe) Analytical methods given in Appendix A         187     Table C.2. Microprobe analysis of spinel from sample 1502C Sample 1052 Position macrocryst  average min max n 11 Major elements (wt.%) SiO2 0.19 0.08 0.90 TiO2 0.91 0.56 1.18 Al2O3 28.15 18.65 38.32 Cr2O3 36.74 26.87 47.36 FeO 12.26 10.77 15.11 Fe2O3 5.13 4.62 5.68 MnO 0.13 0.10 0.22 MgO 16.39 15.13 18.22 Total 99.91 99.30 100.30  Structural Formula (4 oxygen basis) Si 0.01 0.00 0.03 Ti 0.02 0.01 0.03 Al 0.97 0.68 1.27 Cr 0.86 0.60 1.15 Fe2+ 0.30 0.25 0.37 Fe3+* 0.11 0.10 0.13 Mn 0.00 0.00 0.01 Mg 0.72 0.67 0.76  Mg # 70.40 64.10 75.10 Cr # 46.94 31.99 63.01 Fe2+/Fe3+ 2.67 2.26 3.37  * calculated from stoichiometry Mg # = Mg/(Mg+Fe) Cr # = Cr/(Cr+Al)             188 Table C.3. Micoprobe analysis of phlogopite from selected DVI and MFV samples  Mountain Diatreme Sample 1494       1494       1494      1504       1504 Position macrocryst   rim    groundmass   macrocryst   groundmass  average min max  average min max  average min max  average min max  average min max n 31       6       4       12       6 Major elements (wt.%) SiO2 37.25 36.61 37.78  36.92 36.27 39.06  33.77 33.24 34.87  39.98 36.66 47.77  34.69 29.70 48.68 TiO2 4.43 4.06 4.71  3.28 1.37 4.31  2.90 2.43 3.18  4.90 3.94 6.03  0.77 0.02 2.64 Al2O3 16.29 15.93 16.60  14.79 12.48 16.21  14.62 14.38 14.76  14.06 11.90 15.90  13.97 11.77 16.37 FeO 7.04 6.44 9.65  11.64 10.93 12.32  16.47 15.86 16.99  9.62 8.60 10.99  18.98 9.68 21.66 MnO 0.04 0.00 0.11  0.03 0.00 0.08  0.04 0.03 0.06  0.09 0.01 0.20  0.23 0.20 0.25 MgO 20.12 18.03 20.60  17.95 16.24 20.47  13.77 13.22 14.99  14.71 6.16 18.18  17.69 7.88 20.26 CaO 0.04 0.00 0.10  0.43 0.02 2.06  0.12 0.08 0.21  7.24 0.00 13.21  1.98 0.10 9.75 Na2O 0.38 0.27 0.50  0.16 0.05 0.27  0.06 0.04 0.08  1.47 0.48 2.13  0.74 0.03 4.17 K2O 9.55 9.32 9.78  9.11 8.14 9.83  7.76 7.44 8.49  4.93 1.69 9.53  0.33 0.04 0.91 Cr2O3 0.05 0.00 0.19  0.07 0.00 0.12  0.02 0.00 0.05  0.07 0.00 0.26  0.17 0.00 0.40 F 0.27 0.08 0.51  0.27 0.20 0.39  0.42 0.33 0.58  0.33 0.00 0.92  0.11 0.02 0.35 Cl 0.02 0.00 0.04  0.02 0.00 0.03  0.01 0.00 0.02  0.02 0.00 0.05  0.01 0.00 0.02 Total 95.49 94.41 96.41  94.67 93.11 96.53  89.96 89.11 91.93  97.40 95.40 99.41  89.67 88.01 96.11  Structural Formula (11 oxygen basis) Si 2.67 2.64 2.70  2.73 2.67 2.87  2.68 2.67 2.70  2.82 2.66 3.30  2.69 2.39 3.38 Al (IV) 1.33 1.30 1.36  1.27 1.13 1.33  1.32 1.30 1.33  1.18 0.70 1.34  1.31 0.62 1.61 Al (VI) 0.05 0.02 0.08  0.02 -0.05 0.07  0.05 0.01 0.07  0.00 -0.24 0.53  -0.02 -0.15 0.34 Ti 0.24 0.22 0.26  0.18 0.08 0.23  0.17 0.14 0.19  0.26 0.20 0.32  0.04 0.00 0.14 Fe2+ 0.42 0.39 0.59  0.72 0.66 0.76  1.09 1.03 1.13  0.57 0.50 0.65  1.25 0.56 1.46 Mn 0.00 0.00 0.01  0.00 0.00 0.01  0.00 0.00 0.00  0.01 0.00 0.01  0.02 0.01 0.02 Mg 2.15 1.96 2.20  1.98 1.79 2.24  1.63 1.57 1.73  1.55 0.63 1.96  2.09 0.82 2.41 Ca 0.00 0.00 0.01  0.03 0.00 0.17  0.01 0.01 0.02  0.54 0.00 0.98  0.15 0.01 0.73 K 0.87 0.85 0.90  0.86 0.76 0.91  0.79 0.76 0.84  0.45 0.15 0.87  0.03 0.00 0.08 Na 0.05 0.04 0.07  0.02 0.01 0.04  0.01 0.01 0.01  0.20 0.07 0.29  0.10 0.01 0.56 F 0.06 0.02 0.11  0.06 0.05 0.09  0.10 0.08 0.14  0.07 0.00 0.20  0.03 0.01 0.08 Cl 0.00 0.00 0.01  0.00 0.00 0.00  0.00 0.00 0.00  0.00 0.00 0.01  0.00 0.00 0.00 OH- 1.94 1.88 1.98  1.93 1.91 1.95  1.89 1.86 1.92  1.92 1.80 2.00  1.97 1.92 1.99 Mg # 83.58 76.90 84.62  73.28 70.15 74.81  59.81 58.10 62.76  72.23 55.89 78.96  62.13 59.20 64.16  Mg # = Mg/(Mg+Fe) Analytical methods given in Appendix A    189 Table C.4. Major and trace element data from the Marmot Formation and related occureces SAMPLE 1500B 1501B 1501D 1502A 1502C 1503A 1504A UTME 432424 430041 430041 433522 433522 433751 437077 UTMN 7110721 7111609 7111609 7138569 7138569 7139931 7115929 Suite MFV MFV MFV MFV MFV MFV MFV  Major elements (wt.%) SiO2 40.38 47.88 40.31 40.98 39 37.84 35.88 Al2O3 16.67 11.43 12.66 12.07 10.1 11.29 12.13 FeO 8.904 4.192 2.08 8.424 7.184 8.12 9.128 Fe2O3 2.226 1.048 0.52 2.106 1.796 2.03 2.282 CaO 6.93 11.58 17.79 6.44 14.88 10.71 11.64 MgO 7.59 3.33 0.86 13.9 12.88 11.16 9.24 Na2O 4.21 4.58 4.16 0.16 1.65 0.07 1.08 K2O 0.37 0.12 2.92 1.56 0.03 0.94 2.33 Cr2O3 0.13 0.11 0.02 0.11 0.1 0.1 0.06 TiO2 2.55 3.32 2.83 2.43 2.05 2.3 4.32 MnO 0.07 0.06 0.05 0.15 0.23 0.14 0.13 P2O5 0.487 0.089 0.669 0.652 0.516 0.673 0.446 SrO 0.01 0.07 0.05 0.08 0.1 0.07 0.02 BaO 0.02 0.04 0.18 0.18 0.02 0.21 0.94 LOI 9.04 10.55 13.2 10.05 8.91 13.55 8.67 Total 99.59 98.4 98.29 99.3 99.44 99.2 98.29 Mg # cation 60.31 58.61 42.43 74.63 76.16 71.01 64.34  Trace and rare earth elements (ppm) La 28.1 25.4 36.2 81 68.5 70.2 47.6 Ce 52.1 55.6 66.5 146 123.5 125.5 93.4 Pr 6.62 7.16 8.42 16.2 14 14.2 11.55 Nd 27.5 28.1 33.8 59.6 51.2 52.9 47.4 Sm 5.64 5.39 6.66 9.5 8.12 8.29 8.48 Eu 1.81 1.54 1.66 2.81 2.38 2.48 2.63 Gd 6.22 4.45 6.47 9.1 7.38 8.45 8.35 Tb 0.82 0.55 0.94 1.09 0.94 1 1.08 Dy 4.38 2.57 4.91 5.08 4.32 4.88 5.25 Ho 0.82 0.47 0.99 0.95 0.8 0.91 0.98 Er 2.01 1.23 2.28 2.46 2.13 2.4 2.33 Tm 0.26 0.14 0.28 0.3 0.27 0.31 0.28 Yb 1.52 0.9 1.83 1.98 1.56 1.74 1.7 Lu 0.22 0.11 0.24 0.26 0.21 0.24 0.22 Th 5.29 4.55 5.42 10.25 8.76 8.75 5.8 Nb 71.4 64 81.7 124.5 111.5 113 80.3 Zr 207 281 342 221 219 210 257 Y 17.5 9.9 23.3 22.1 18.8 20.1 21.5 Cr 1000 880 130 860 810 810 490 Ni 237 214 26 307 260 291 189 Co 39.9 60.7 24.5 63 59.3 60 58.9 V 275 67 192 280 229 267 460 Cu 224 46 29 68 35 61 112 Mo <2 <2 5 2 4 3 5 Cs 0.69 2.74 1.58 10.95 7.16 7.52 1.65 Pb <5 <5 <5 <5 <5 <5 <5 Zn 88 89 40 97 73 97 124 Rb 8.9 2.4 39.4 67.2 4.3 25.8 39.4 Ba 475 363 1620 1625 153 1870 8500 Sr 129.5 711 500 876 1065 778 166 Ga 20.4 13.6 12.9 17.4 13.9 16.5 19.4 Ta 4.1 3.7 4.9 7.1 5.8 6.3 4.8 Hf 5.8 7.4 8.5 5.9 5.7 5.7 7.3 U 1.46 0.22 0.96 2.19 1.96 2 1.53     190 Table C.4. Cont. SAMPLE 1505I 1506F 1507A 1508B 1508D 1500A4 1500A1 UTME 436097 420455 403384 399309 399309 432424 432424 UTMN 7114816 7137286 7148035 7141753 7141753 7110721 7110721 Suite MFV MFV MFV MFV MFV MFX MFX  Major elements (wt.%) SiO2 45.14 42.35 42.4 45.66 44.62 45.89 46.12 Al2O3 16.44 13.88 13.04 19.18 13.94 10.7 11.05 FeO 8.272 8.784 9.632 7.064 10.952 5.936 6.648 Fe2O3 2.068 2.196 2.408 1.766 2.738 1.484 1.662 CaO 3.52 7.62 10.32 4.11 9.09 14.56 13.17 MgO 7.86 6.69 6.47 2.9 7.34 9.08 9.1 Na2O 1.8 1.49 2.09 1.12 2.92 0.98 1.85 K2O 4.9 2.38 2.24 7.75 1.14 2.78 1.72 Cr2O3 0.01 0.01 0.02 0 0.03 0.04 0.02 TiO2 3.36 3.22 3.39 3.02 2.94 2.12 3.43 MnO 0.16 0.16 0.17 0.11 0.2 0.15 0.18 P2O5 0.839 0.715 0.62 0.671 0.501 0.423 0.714 SrO 0.05 0.03 0.08 0.02 0.05 0.08 0.07 BaO 0.73 0.53 0.26 0.42 0.06 1.34 2.46 LOI 4.72 8.95 6.69 5.81 3.35 2.76 1.46 Total 99.86 99 99.84 99.6 99.87 98.32 99.65 Mg # cation 62.87 57.58 54.49 42.25 54.43 73.16 70.93  Trace and rare earth elements (ppm) La 60.5 73.3 63.3 43 94.8 52.5 68.2 Ce 120.5 128.5 121 94.7 140.5 99.7 134 Pr 15.8 15.25 14.75 12.25 15.35 11.5 15.9 Nd 62.8 59.2 59.8 49.4 55.7 43.6 64.6 Sm 12.05 10.25 11.15 10.6 9.68 7.34 11.4 Eu 3.56 2.82 3.4 3.05 2.98 3.19 3.14 Gd 10.95 10.05 10.5 10.2 9.52 7.32 11.5 Tb 1.49 1.32 1.35 1.46 1.2 0.95 1.42 Dy 7.63 6.52 6.96 7.96 6.02 4.7 7.13 Ho 1.39 1.22 1.24 1.61 1.13 0.91 1.38 Er 3.6 3.24 3.19 4.26 2.88 2.38 3.45 Tm 0.47 0.38 0.41 0.62 0.35 0.32 0.45 Yb 2.75 2.47 2.51 3.66 2.18 2.03 2.68 Lu 0.36 0.34 0.32 0.5 0.28 0.27 0.36 Th 7.84 8.62 8.3 16 6.22 8.28 8.51 Nb 100 110.5 100 197.5 81.1 83.1 111.5 Zr 382 355 323 602 289 245 366 Y 32.5 28.3 28.3 38.4 25.9 21 30.1 Cr 60 50 180 30 240 290 170 Ni 17 37 57 12 74 61 61 Co 31.8 37.9 47.8 18 49.2 38.2 39 V 285 330 372 212 318 230 330 Cu 16 93 125 5 54 598 309 Mo <2 3 2 2 <2 <2 30 Cs 1.11 1.62 10.1 1.21 0.48 1.42 1.16 Pb <5 <5 5 <5 <5 <5 <5 Zn 146 133 167 66 94 74 126 Rb 90.4 35.3 106 129 20.2 32.7 41.3 Ba 6530 4820 2320 3840 586 10000 10000 Sr 483 319 799 150.5 481 735 739 Ga 26.6 23.2 23.1 32.6 25.2 11.5 14.9 Ta 5.9 6.8 6 12 4.5 4.5 6.5 Hf 9.7 8.5 9.3 13.5 7.6 6.1 9.8 U 1.99 2 1.61 3.38 1.18 2.06 1.84     191 Table C.4. Cont. SAMPLE 1500A4 1500A1 1500E 1505E 1505G 1475A 1494F UTME 432424 432424 432424 436097 436097 472802 476723 UTMN 7110721 7110721 7110721 7114816 7114816 7129586 7123556 Suite MFX MFX MFX MFX MFX DVI DVI  Major elements (wt.%) SiO2 45.89 46.12 40.79 31.74 40.09 23.75 26.85 Al2O3 10.7 11.05 10.28 12.53 16.21 7.18 7.59 FeO 5.936 6.648 8.136 5.584 6.336 4.464 6.192 Fe2O3 1.484 1.662 2.034 1.396 1.584 1.116 1.548 CaO 14.56 13.17 15.66 17.47 9.3 17.56 19 MgO 9.08 9.1 13.16 6.64 6.74 12.23 12.9 Na2O 0.98 1.85 0.8 1.89 1.85 0.09 0.07 K2O 2.78 1.72 0.72 1.83 3.65 2.91 3.65 Cr2O3 0.04 0.02 0.05 0.2 0.08 0.01 0.02 TiO2 2.12 3.43 2.23 1.53 2.02 2.33 2.48 MnO 0.15 0.18 0.21 0.26 0.14 0.12 0.16 P2O5 0.423 0.714 0.67 0.223 0.26 1.124 0.93 SrO 0.08 0.07 0.06 0.04 0.04 0.02 0.08 BaO 1.34 2.46 0.2 0.08 0.15 0.05 0.12 LOI 2.76 1.46 4.31 17.6 11.25 25.3 16.7 Total 98.32 99.65 99.31 99.01 99.69 98.26 98.29 Mg # cation 73.16 70.93 74.25 67.94 65.47 83.00 78.78  Trace and rare earth elements (ppm) La 52.5 68.2 58.3 46.5 46.8 132.5 121 Ce 99.7 134 107 84.6 79.9 242 223 Pr 11.5 15.9 12.55 8.98 8.77 27.8 25.3 Nd 43.6 64.6 47.6 34.2 31.9 99.7 93.2 Sm 7.34 11.4 7.97 6.17 5.9 15.15 13.9 Eu 3.19 3.14 2.31 1.51 1.84 3.73 3.35 Gd 7.32 11.5 8.17 5.94 6.28 12 11.25 Tb 0.95 1.42 1.01 0.77 0.83 1.28 1.16 Dy 4.7 7.13 4.9 3.76 4.4 5.41 4.99 Ho 0.91 1.38 0.85 0.69 0.84 0.94 0.83 Er 2.38 3.45 2.18 1.77 2.14 2.35 2.15 Tm 0.32 0.45 0.27 0.22 0.25 0.28 0.25 Yb 2.03 2.68 1.59 1.36 1.57 1.55 1.4 Lu 0.27 0.36 0.23 0.18 0.22 0.22 0.2 Th 8.28 8.51 5.77 3.65 7.2 17.3 13.05 Nb 83.1 111.5 91.8 42.7 74.9 158.5 137 Zr 245 366 184 122 202 316 259 Y 21 30.1 19.5 16.6 19.7 25 22 Cr 290 170 380 1610 680 60 100 Ni 61 61 215 238 225 41 77 Co 38.2 39 62.3 37.4 42 24.6 44.4 V 230 330 241 220 195 236 209 Cu 598 309 132 39 <5 54 178 Mo <2 30 8 <2 4 8 35 Cs 1.42 1.16 1.35 0.54 0.83 3.45 3.98 Pb <5 <5 18 <5 <5 82 8 Zn 74 126 105 91 105 221 82 Rb 32.7 41.3 20.5 33.9 83.3 57.8 137 Ba 10000 10000 1815 739 1330 505 1080 Sr 735 739 652 387 379 199 908 Ga 11.5 14.9 16.7 15.2 19.8 12.5 10.4 Ta 4.5 6.5 4.7 2.4 4 8 7.5 Hf 6.1 9.8 4.8 3.4 5 6.6 5.8 U 2.06 1.84 1.36 0.93 1.89 4.31 2.4        192 Table C.4. Cont. SAMPLE 1496A 1497A 1498B 1499B 1542B UTME 479548 480062 478719 480574 466227 UTMN 7122103 7124416 7124665 7102139 7136196 Suite DVI DVI DVI DVI DVI  Major elements (wt.%) SiO2 15.31 26.11 36.59 31.63 8.21 Al2O3 3.78 8.1 9.32 9.96 3.66 FeO 1.576 7.768 7.064 8.408 2.232 Fe2O3 0.394 1.942 1.766 2.102 0.558 CaO 22.37 11.73 10.63 5.45 41.71 MgO 16.5 18.66 15.61 23.54 4.95 Na2O 0.09 0.02 0.39 0.02 0.07 K2O 2.05 0.24 3.1 0.05 0.68 Cr2O3 0.02 0.06 0.11 0.03 0.03 TiO2 0.73 2.77 2.68 2.95 0.79 MnO 0.02 0.17 0.13 0.24 0.11 P2O5 0.421 1.159 0.576 1.417 0.528 SrO 0.01 0.03 0.03 0.04 0.06 BaO 0.01 0.05 0.17 0.16 0.08 LOI 33.9 20 11.05 12.55 34.8 Total 97.19 98.8 99.22 98.55 98.46 Mg # cation 94.91 81.07 79.75 83.31 79.81  Trace and rare earth elements (ppm) La 45.6 147.5 66.6 176 167 Ce 83.8 271 129 317 257 Pr 9.38 30.9 15.05 36.1 25.2 Nd 33.6 110 56.9 133.5 83.2 Sm 4.96 15.5 9.3 19 10.85 Eu 1.32 4.75 2.72 4.97 2.96 Gd 5.34 15.6 8.49 16.85 11.2 Tb 0.58 1.59 0.96 1.78 1.13 Dy 2.81 6.41 4.39 7.42 4.45 Ho 0.49 1.06 0.71 1.23 0.79 Er 1.4 2.77 1.76 3.2 2.07 Tm 0.17 0.3 0.19 0.35 0.23 Yb 1.1 1.76 1.2 2.1 1.4 Lu 0.15 0.23 0.16 0.26 0.18 Th 7.07 19 8.09 20.7 9.77 Nb 48.8 174 97.9 213 122.5 Zr 152 327 217 310 94 Y 11.9 23.4 16.1 28.4 18.1 Cr 150 450 910 240 260 Ni 42 149 276 114 141 Co 17.7 54.2 59 45.6 21.3 V 94 283 272 383 107 Cu 20 78 33 70 20 Mo 2 27 2 6 4 Cs 1.03 2.41 38.8 3.5 0.52 Pb 13 <5 6 5 <5 Zn 39 121 90 117 52 Rb 29.7 12.6 203 5.4 28.4 Ba 85.2 390 1490 1405 666 Sr 128.5 267 258 391 673 Ga 7.7 16.1 15.1 16.8 7.1 Ta 2.4 8.8 5.1 9.3 2.5 Hf 3.7 7.5 6.1 7.3 2.3 U 2.52 3.92 1.92 4.35 1.7     193 Table C.5. Neodymium data from selected DVI, MFV and MFX occurences Sample 1494F 1498B 1499B 1500A1 1500E 1502C 1504A 1505G 1505I 1506F 1507A 1508B 1508D UTME 476723 478719 480574 432424 432424 433522 437077 436097 436097 420455 403384 399309 399309 UTMN 7123556 7124665 7102139 7110721 7110721 7138569 7115929 7114816 7114816 7137286 7148035 7141753 7141753 Suite DVI DVI DVI MFX MFX MFV MFV MFX MFV MFV MFV MFV MFV  Nd (ppm) 93.20 56.90 133.50 64.60 47.60 51.20 47.40 31.90 62.80 59.20 59.80 49.40 55.70 Sm (ppm) 13.90 9.30 19.00 11.40 7.97 8.12 8.48 5.90 12.05 10.25 11.15 10.60 9.68 147Sm/144Nd 0.09016 0.09881 0.08604 0.10669 0.10123 0.09588 0.10816 0.11182 0.11600 0.10468 0.11272 0.12973 0.10506 143Nd/144Nd (today) 0.512493 0.512494 0.512458 0.512707 0.512682 0.512546 0.512680 0.512690 0.512717 0.512707 0.512696 0.512900 0.512609 2" 0.000008 0.000005 0.000008 0.000006 0.000007 0.000006 0.000005 0.000007 0.000006 0.000006 0.000006 0.000006 0.000009 143Nd/144Nd (initial) 0.512225 0.512200 0.512201 0.512389 0.512380 0.512261 0.512357 0.512357 0.512371 0.512395 0.512360 0.512513 0.512296 !Nd (initial) $ 3.2 2.7 2.7 6.3 6.2 3.8 5.7 5.7 6.0 6.5 5.8 8.7 4.5 tDM (Ma)  811 870 827 632 635 781 680 689 677 620 685 455 759  Epsilon values and model ages were calculated using 143Nd/144NdDM = 0.5135 and 147Nd/144NdDM = 0.2137 (Salters and Stracke 2004); and 143Nd/144NdCHUR = 0.512638 (DePaolo and Wasserburg 1976).               194 Table C.6. 40Ar/39Ar age data for phlogopite from selected DVI and MFV samples  Step Cum. 39Ar 40Ar/39Ar 2" 38Ar/39Ar 2" 37Ar/39Ar 2" 36Ar/39Ar 2" 40Ar*/39ArK 2" Ca/K Cl/K Age (Ma) error (2")   1475A1 (DVI) 1 0.0 43.453 0.033 0.002 21.182 -0.056 0.593 0.097 0.216 26.908 10.585 0.555 -0.018 395 139 2 0.2 22.585 0.031 0.026 0.892 0.120 0.332 0.027 0.608 16.781 5.464 1.674 0.001 256 78 3 0.5 24.513 0.013 0.022 0.160 0.006 0.759 0.012 0.211 22.182 0.812 0.172 0.001 331 11 4 1.6 31.379 0.006 0.023 0.096 -0.001 5.452 0.008 0.130 29.432 0.367 0.027 0.002 428 5 5 4.1 31.659 0.006 0.023 0.030 -0.001 1.366 0.002 0.631 31.215 0.429 0.004 0.002 450 5 6 9.7 31.572 0.005 0.022 0.048 0.000 1.040 0.001 0.156 31.379 0.176 0.004 0.002 453 2 7 21.3 32.010 0.006 0.021 0.029 0.000 1.071 0.001 0.233 31.849 0.182 0.005 0.002 459 2 8 27.9 31.883 0.005 0.021 0.038 0.000 8.414 0.001 0.099 31.713 0.160 0.006 0.002 457 2 9 36.9 32.080 0.005 0.021 0.056 0.000 1.613 0.000 0.237 31.944 0.175 0.005 0.002 460 2 10 45.7 32.039 0.004 0.022 0.055 0.000 9.834 0.000 0.236 31.906 0.143 0.005 0.002 459 2 11 52.5 32.038 0.005 0.022 0.069 0.000 5.302 0.001 0.201 31.847 0.166 0.005 0.002 459 2 12 58.4 32.041 0.006 0.023 0.041 0.000 1.987 0.001 0.184 31.875 0.204 0.005 0.002 459 3 13 68.3 32.097 0.005 0.023 0.052 0.000 0.848 0.000 0.190 31.974 0.148 0.006 0.002 460 2 14 77.0 32.142 0.005 0.023 0.072 0.000 3.072 0.001 0.327 32.002 0.167 0.006 0.002 460 2 15 87.8 32.133 0.004 0.022 0.025 0.000 1.575 0.000 0.108 32.008 0.144 0.002 0.002 461 2 16 93.0 32.085 0.007 0.022 0.045 0.000 1.607 0.001 0.187 31.911 0.236 0.005 0.002 459 3 17 97.0 32.158 0.006 0.022 0.067 0.000 2.207 0.001 0.193 31.934 0.209 0.007 0.002 460 3 18 100.0 32.136 0.007 0.021 0.112 0.000 1.290 0.001 0.233 31.888 0.233 0.009 0.002 459 3 Integrated age = 457.6±0.9 Ma J= 0.009086±0.000014   Plateau age = 459.6±2.4 Ma (steps 9-18)  1494B1 (DVI) 1 0.0 426.733 0.231 0.368 0.920 0.005 33.286 1.434 0.240 -15.060 36.035 0.191 0.191 -142 355 2 0.1 261.959 0.132 0.079 2.545   0.818 0.201 3.842 38.689 0.175 0.175 35 345 3 0.3 61.273 0.051 0.035 0.633 0.783 0.099 0.104 0.196 28.158 6.214 1.911 1.911 239 49 4 1.1 71.894 0.027 0.037 0.350 0.077 0.168 0.089 0.046 44.643 1.673 0.194 0.194 366 12 5 5.1 61.449 0.011 0.026 0.093 0.010 0.213 0.014 0.125 57.162 0.813 0.026 0.026 457 6 6 18.5 57.207 0.008 0.022 0.056 0.003 0.202 0.003 0.074 56.163 0.432 0.008 0.008 450 3 7 54.4 56.702 0.006 0.021 0.089 0.002 0.378 0.001 0.204 56.214 0.362 0.005 0.005 450 3 8 75.7 56.923 0.005 0.021 0.073 0.001 0.356 0.001 0.270 56.569 0.301 0.003 0.003 453 2 9 92.8 57.149 0.006 0.022 0.068 0.001 1.027 0.001 0.247 56.683 0.376 0.002 0.002 453 3 10 96.5 58.633 0.009 0.020 0.168 0.002 1.550 0.004 0.242 57.117 0.616 0.007 0.007 456 4 11 99.1 58.805 0.023 0.027 0.268 0.000 7.844 0.006 0.199 56.594 1.366 0.003 0.003 453 10 12 100.0 62.355 0.023 0.029 0.225 0.010 0.754 0.024 0.247 54.553 2.188 0.031 0.031 438 16 Integrated age = 450.2±1.6 Ma J= 0.005041±0.000012   Plateau age = 451.9±2.6 Ma (steps 5-9) Analytical methods are reported in Appendix A   195  Table C.6. Cont.  Step Cum. 39Ar 40Ar/39Ar 2" 38Ar/39Ar 2" 37Ar/39Ar 2" 36Ar/39Ar 2" 40Ar*/39ArK 2" Ca/K Cl/K Age (Ma) error (2")   1496B1 (DVI) 1 0.1 154.880 0.029 0.141 0.166 0.233 0.173 0.517 0.041 3.113 5.464 3.348 0.007 50 87 2 0.3 35.869 0.021 0.055 0.275 1.575 0.034 0.095 0.036 9.614 0.972 18.245 0.005 151 15 3 0.9 37.134 0.015 0.041 0.093 0.879 0.031 0.051 0.053 23.257 0.895 9.569 0.004 346 12 4 1.5 41.123 0.010 0.030 0.119 0.020 0.203 0.039 0.063 30.665 0.824 0.300 0.002 444 11 5 3.4 34.886 0.009 0.023 0.146 0.008 0.166 0.015 0.051 30.712 0.363 0.107 0.001 444 5 6 11.4 32.202 0.011 0.021 0.036 0.003 0.154 0.004 0.031 31.048 0.334 0.042 0.001 449 4 7 29.1 31.876 0.006 0.020 0.040 0.001 0.290 0.001 0.105 31.519 0.195 0.016 0.001 455 2 8 47.1 31.832 0.005 0.020 0.035 0.001 0.323 0.001 0.086 31.598 0.157 0.008 0.001 456 2 9 58.4 31.770 0.006 0.020 0.053 0.001 0.464 0.001 0.128 31.564 0.197 0.009 0.001 455 3 10 77.0 31.724 0.004 0.019 0.050 0.001 0.291 0.001 0.099 31.496 0.139 0.011 0.001 454 2 11 80.7 31.410 0.007 0.019 0.074 -0.001 1.355 0.001 0.192 31.117 0.230 0.008 0.001 450 3 12 84.6 31.243 0.007 0.021 0.098   0.001 0.163 30.928 0.217 0.012 0.002 447 3 13 93.1 31.298 0.006 0.020 0.074 0.001 0.468 0.001 0.122 31.030 0.183 0.013 0.001 448 2 14 97.7 31.649 0.007 0.020 0.093 -0.001 0.487 0.001 0.171 31.331 0.246 0.001 0.001 452 3 15 100.0 31.752 0.007 0.020 0.081 -0.001 1.853 0.002 0.235 31.401 0.260 0.014 0.001 453 3 Integrated age = 451.4±1.1 Ma J= 0.009094±0.000012   Plateau age = 455.0±2.5 Ma (steps 7-10)  1497A1 (DVI) 1 1.4 42.097 0.026 0.041 0.236 0.524 0.054 0.079 0.088 19.212 2.046 5.654 0.003 291 29 2 22.5 32.269 0.006 0.022 0.033 0.004 0.118 0.005 0.057 30.613 0.213 0.032 0.002 443 3 3 53.1 30.938 0.005 0.020 0.062 0.002 0.110 0.001 0.081 30.636 0.167 0.012 0.001 444 2 4 81.9 31.021 0.005 0.020 0.071 0.002 0.209 0.001 0.138 30.760 0.146 0.007 0.001 445 2 5 97.4 31.241 0.005 0.021 0.101 0.002 0.126 0.001 0.159 30.816 0.179 0.008 0.002 446 2 6 99.5 32.263 0.010 0.024 0.186 0.010 0.587 0.007 0.172 30.472 0.448 0.000 0.002 441 6 7 100.0 37.982 0.023 0.026 0.317 0.052 0.162 0.029 0.256 30.309 2.310 0.021 0.003 439 30 Integrated age = 442.2±1.4 Ma J= 0.009097±0.000014   Plateau age = 444.4±2.5 Ma (steps 2-7)           196   Table C.6. Cont.  Step Cum. 39Ar 40Ar/39Ar 2" 38Ar/39Ar 2" 37Ar/39Ar 2" 36Ar/39Ar 2" 40Ar*/39ArK 2" Ca/K Cl/K Age (Ma) error (2")   1499B (DVI) 1 0.1 62.615 0.018 0.112 0.392 0.419 0.050 0.197 0.103 6.723 5.868 3.298 0.017 107 91 2 3.8 23.368 0.007 0.034 0.094 0.667 0.026 0.037 0.025 12.873 0.285 7.346 0.003 200 4 3 7.1 28.150 0.025 0.032 0.103 0.025 0.148 0.017 0.075 23.035 0.735 0.235 0.003 343 10 4 17.3 29.849 0.010 0.024 0.037 0.007 0.071 0.003 0.063 29.048 0.296 0.062 0.002 423 4 5 36.1 32.098 0.005 0.022 0.037 0.003 0.116 0.001 0.099 31.798 0.162 0.023 0.002 458 2 6 62.9 32.218 0.004 0.022 0.057 0.002 0.069 0.001 0.170 32.003 0.145 0.015 0.002 461 2 7 67.6 32.295 0.008 0.022 0.134 0.003 0.213 0.002 0.260 31.780 0.301 0.000 0.002 458 4 8 87.3 32.178 0.006 0.021 0.047 0.001 0.148 0.001 0.168 31.978 0.197 0.008 0.002 461 3 9 93.3 32.117 0.009 0.022 0.057 0.002 0.239 0.001 0.209 31.701 0.309 0.000 0.002 457 4 10 99.1 32.648 0.007 0.021 0.085 0.003 0.273 0.002 0.310 32.150 0.268 0.011 0.002 463 3 11 100.0 33.788 0.016 0.020 0.317 0.015 0.300 0.009 0.226 31.276 0.801 0.000 0.001 452 10 Integrated age = 443.2±1.2 Ma J= 0.009096±0.000014   Plateau age = 459.9±2.5 Ma (steps 5-10)  1504B1 (MFV) 1 0.2 49.120 0.024 0.069 0.262 0.079 0.372 0.140 0.060 8.571 2.324 0.387 0.008 135 35 2 1.5 19.517 0.019 0.028 0.194 0.024 0.100 0.027 0.058 11.764 0.499 0.188 0.002 183 7 3 3.4 29.689 0.009 0.025 0.111 0.008 0.191 0.010 0.081 26.653 0.344 0.042 0.002 391 5 4 7.0 32.849 0.006 0.025 0.152 0.005 0.266 0.004 0.071 31.694 0.214 0.026 0.002 457 3 5 15.2 32.935 0.006 0.023 0.089 0.003 0.199 0.002 0.107 32.395 0.203 0.017 0.002 466 3 6 28.0 32.686 0.005 0.022 0.036 0.004 0.175 0.001 0.116 32.339 0.179 0.031 0.002 465 2 7 41.8 32.520 0.012 0.022 0.037 0.003 0.127 0.001 0.109 32.232 0.403 0.029 0.002 464 5 8 46.4 32.507 0.007 0.024 0.105 0.003 0.077 0.002 0.085 32.025 0.236 0.014 0.002 461 3 9 54.7 32.334 0.009 0.023 0.021 0.003 0.110 0.001 0.161 31.992 0.284 0.026 0.002 461 4 10 73.7 32.300 0.004 0.023 0.021 0.004 0.071 0.001 0.136 32.040 0.144 0.038 0.002 461 2 11 88.2 32.142 0.005 0.022 0.028 0.002 0.213 0.001 0.177 31.890 0.164 0.013 0.002 459 2 12 94.1 32.210 0.007 0.024 0.070 0.002 0.278 0.002 0.068 31.626 0.216 0.008 0.002 456 3 13 98.2 32.250 0.006 0.022 0.068 0.003 0.208 0.002 0.193 31.647 0.226 0.005 0.002 456 3 14 100.0 32.419 0.014 0.024 0.151 0.005 0.304 0.003 0.177 31.529 0.464 0.003 0.002 455 6 Integrated age = 456.0±1.2 Ma J= 0.009090±0.000014   Plateau age = 460.6±2.6 Ma (steps 7-11)        197  Table C.6. Cont.  Step Cum. 39Ar 40Ar/39Ar 2" 38Ar/39Ar 2" 37Ar/39Ar 2" 36Ar/39Ar 2" 40Ar*/39ArK 2" Ca/K Cl/K Age (Ma) error (2")   1507D5 (MFV) 1 0.1 305.946 0.042 0.281 0.295 0.624 0.270 1.018 0.082 0.429 21.285 1.509 0.025 4 193 2 0.3 94.379 0.050 0.107 0.340 0.217 0.179 0.261 0.066 15.142 3.656 0.531 0.014 133 31 3 0.5 88.709 0.026 0.044 0.466 0.127 0.168 0.155 0.088 41.512 4.145 0.312 0.003 343 31 4 2.2 63.393 0.018 0.027 0.062 0.023 0.133 0.026 0.043 55.303 1.093 0.055 0.002 444 8 5 10.8 58.027 0.008 0.023 0.049 0.007 0.153 0.004 0.059 56.764 0.462 0.017 0.002 454 3 6 18.9 57.728 0.006 0.022 0.069 0.006 0.222 0.002 0.196 57.127 0.365 0.014 0.002 457 3 7 35.6 55.563 0.010 0.022 0.067 0.004 0.056 0.001 0.145 55.127 0.568 0.010 0.002 442 4 8 51.8 55.384 0.016 0.022 0.036 0.002 0.185 0.001 0.100 55.085 0.886 0.006 0.002 442 6 9 64.0 56.129 0.007 0.020 0.043 0.001 0.388 0.001 0.182 55.772 0.404 0.003 0.001 447 3 10 85.5 55.496 0.006 0.021 0.042 0.001 0.266 0.001 0.258 55.248 0.312 0.003 0.002 443 2 11 94.2 56.026 0.006 0.023 0.105 0.000 1.777 0.001 0.157 55.549 0.366 0.001 0.002 445 3 12 97.8 57.282 0.010 0.022 0.109 0.001 1.383 0.002 0.250 56.446 0.574 0.003 0.002 452 4 13 100.0 56.829 0.018 0.022 0.102 0.003 0.754 0.005 0.118 55.256 1.031 0.009 0.002 443 7 Integrated age = 444.8±1.8 Ma J= 0.005041±0.000012   Plateau age = 444.5±2.6 Ma (steps 7-11)  1542C1 (DVI) 1 0.6 24.697 0.023 0.146 0.094 0.131 0.159 0.066 0.054 5.542 1.006 1.567 0.028 89 16 2 1.4 31.735 0.017 0.055 0.107 0.485 0.053 0.035 0.036 22.228 0.540 5.863 0.008 332 7 3 3.4 34.085 0.009 0.030 0.051 0.086 0.039 0.011 0.080 31.017 0.389 1.030 0.003 448 5 4 10.0 33.302 0.008 0.022 0.051 0.003 0.202 0.005 0.079 31.972 0.278 0.029 0.002 460 4 5 35.7 32.417 0.007 0.019 0.052 0.001 0.189 0.001 0.072 32.152 0.228 0.012 0.001 462 3 6 51.8 32.272 0.005 0.020 0.038 0.001 0.158 0.001 0.152 32.043 0.162 0.010 0.001 461 2 7 58.7 32.173 0.007 0.019 0.067 0.001 0.395 0.001 0.164 31.885 0.238 0.011 0.001 459 3 8 70.1 32.185 0.006 0.020 0.043 0.001 0.200 0.001 0.162 31.926 0.181 0.012 0.001 459 2 9 79.6 32.235 0.005 0.022 0.030 0.001 0.519 0.001 0.201 31.964 0.167 0.009 0.002 460 2 10 88.7 32.121 0.006 0.021 0.037 0.001 0.370 0.001 0.173 31.845 0.201 0.010 0.002 458 3 11 95.1 32.173 0.005 0.022 0.074 0.001 0.670 0.001 0.150 31.813 0.184 0.006 0.002 458 2 12 100.0 32.274 0.005 0.023 0.051 0.001 1.360 0.002 0.216 31.786 0.209 0.002 0.002 457 3 Integrated age = 456.8±1.2 Ma J= 0.009077±0.000012  Plateau age = 459.7±2.5 Ma (steps 4-10)    198 REFERENCES CITED  DePaolo, D.J., and Wasserburg, G.J., 1976, Nd isotopic variations and petrogenetic models: Geophysical Research Letters, v. 3, p. 249-252.  Salters, V., and Stracke, A., 2004, Composition of the depleted mantle: Geochemistry, Geophysics, Geosystems, v. 5, 27   199 APPENDIX D  Marmot Formation photo documentation                        200    201 Figure D.1. (previous page) Field photographs of the Marmot Formation. (a) View of sample location 1501; red dashed lines outline the boundaries of what is interpreted to be a volcanic conduit. (b) Massive Marmot Formation mafic volcanic rocks underlying basinal shales of the Ordovician Duo Lakes Formation. (c) Sorted limestone and volcanic cobble clast volcanogenic conglomerate from sample location 1500. (d) Range of compositions and shapes of xenoliths in a volcaniclastic deposit of sample location 1505. (e) Cross-bedded epiclastic facies deposit or a base surge deposit near sample location 1505 (f) Carbonate and chlorite amygdaloidal basalt from sample location 1501. (g) Hyaloclastite from sample location 1501. (h) Pillow lava from sample location 1503.     202         203 Figure D.2. (previoius page) Photomicrographs and one field photograph showing the petrographic range of compositions of the MFV and MFX suites.  (a) Clinopyroxene porphyritic mafic volcanic rock of MFV sample 1507E. Identifying the groundmass composition in this and other samples is difficult owing its fine nature therefore cannot be distinguished between tephrite or nephelinite. (b) Carbonate, serpentine and quartz altered tephrite/nephelinite of MFV sample 1504B. (c) Brecciated mafic volcanic rock from MFV sample 1508E. (d) Intergranular mafic volcancic rock with carbonate and sericite pseudomorphed clinopyroxene phenocrysts. (e) Volcanic xenolith from MFX sample location 1500 showing a clinopyroxene porphyritic texture (note the similarity with ‘a’). (f) Volcanic xenolith from MFX sample location 1500 showing a relatively unaltered version of ‘d’.  All photomicrographs are in crossed polarized light.               204         205 Figure D.3. (previous page) Photomicrographs and one SEM image of important minerals used in microprobe and dating studies. (a) Pervasive sieve texture [S] in a clinopyroxene macrocryst from MFV sample 1508E. (b) Zoned clinopyroxene macrocryst showing a rim [R] and core [C] from MFV sample 1508E. (c) Complex clinopyroxene macrocryst with a resorbed core [C] and a euhedral rim [R] from MFV sample 1508E. (d) SEM image of a phlogopite macrocryst with an alteration rim [R] and an unaltered core [C] from DVI sample 1494. (e) Groundmass phlogopite [P] from DVI sample 1498. (f) Phlogopite macrocryst [P] and a carbonate after olivine macrocryst with spinel [S] inclusions from DVI sample 1475. Photomicrographs (a), (b), and (c) are in cross polarized light whereas (e) and (f) are in plane polarized light.            206  APPENDIX E  Volcaniclastic Diatremes3                                     3. A version of this appendix will be published as Chapter 4.2 Volcaniclastic Diatremes in Martel et al. (2009 in prep); Geology of central Mackenzie Mountains; Sekwi Mountain (105P), Mount Eduni (106A), and northwest part of Wrigley Lake (95M) map areas, Northwest Territories: NTGO Report xxxx.   207 VOLCANICLASTIC DIATREMES    To the south and west of the Mackenzie Platform, the platformal strata thin and transition to thick siliciclastic basinal facies deposits comprising the Selwyn Basin (Gordey and Anderson, 1993). The margin between the Mackenzie Platform and the Selwyn Basin contains an irregular, 100 x 150 kilometer basin that records rifting in Early to Middle Cambrian and late Early Ordovician to Middle Ordovician time, and is termed the Misty Creek Embayment (Cecile 1978, 1982; Cecile et al. 1997). This basin contains widespread occurrences of Ordovician mafic volcaniclastic diatremes on it’s eastern side emplaced over a short period of time during the second rifting event. Discrete diatreme occurrences occur throughout the southwest corner of NTS map sheet 106A (Fig. E.1). These volcanic rocks were emplaced into a shallow water carbonate platform environment along the eastern margin of the Misty Creek Embayment and are thought to be related to crustal extension and crustal scale normal faulting. Enriched mantle-derived melts (e.g., Cecile 1982, Goodfellow et al. 1995) may have utilized these faults as conduits to rapidly ascend through the lithosphere and erupt along the western margin of Ancestral North America.  Diatremes of the Mackenzie Mountains have received limited petrographic work, largely because of the inaccessibility of the central Mackenzie Mountains. The discovery of the Mountain Diatreme (Fig. E.1 and E.2) in 1973 by Dynasty Explorations Ltd., however, spurred an interest for diamond exploration in the area. Regional exploration in the vicinity of the Mountain Diatreme by Archer Cathro Ltd. in the early 1990’s led to the discovery of many additional diatremes and associated volcanic and volcaniclastic   208 occurrences (Fig. E.1., Copland 1994). Detailed volcano-stratigraphic work and reconnaissance level petrographic and geochemical results for some volcanic rocks in the Misty Creek Embayment are presented by Goodfellow et al. (1980) and Cecile (1982). Results of additional petrographic and geochemical studies specifically focusing on the Mountain Diatreme are given by Oldershaw (1977), Goodfellow et al. (1995) and Godwin and Price (1986) report. Initial age constraints for the Mountain diatreme and associated volcanic rocks are reported in McArthur et al. (1980) and Godwin and Price (1986). A regional synthesis of the nature and geochemistry of Paleozoic volcanic rocks of the miogeocline in northwestern Canada is presented by Goodfellow et al. (1995). Here we present an overview of previously published data incorporated with more recent work (e.g., Leslie et al., 2008).  Diatreme Geology   The term diatreme is used herein as defined by Cas and Wright (1987), to describe a “pipe like volcanic conduit filled with volcaniclastic debris.” Diatreme locations in the Misty Creek Embayment are presented in Fig. E.1 and are summarized in Table E.1. Names assigned to individual diatremes are from historic assessment reports (e.g., Copland 1994). Most diatremes are observed to crosscut Late Cambrian to Early Ordovician carbonate rocks of the transitional Franklin Mountain Formation. Owing to differential weathering, diatremes are also locally observed to crosscut interbedded siltstones and carbonates of the Early Cambrian Sekwi Formation.   209  Diatreme fill in the study area (e.g., Mountain, Bear, Tim, Monica) is dominated with poorly sorted, matrix supported, accidental (non-volcanic) lithic clast rich volcaniclastic breccias (e.g., figure E.5). The exposed diameter of these diatreme bodies ranges from 50-600m. These occurrences typically contain angular to subrounded, pebble to boulder sized country rock xenoliths derived mainly from host carbonate rocks of the Franklin Mountain Formation. Lapilli sized juvenile (magma derived) clasts are also observed in most bodies (e.g., Mountain, Bear, Monica, and Tim diatremes and the Bits plug) and are typically cored by subhedral phlogopite macrocrysts and olivine phenocrysts (figure E.6).  These olivine phenocrysts are commonly pseudomorphed by carbonate +/- serpentine, set in a spinel bearing carbonate and clay altered aphanitic groundmass. The Mountain (figure E.2), Monica (figure E.4), and Tim (figure E.7) diatremes also contain abundant irregular shaped lapilli sized juvenile clasts composed of olivine and clinopyroxene phenocrysts pseudomorphed by carbonate and a low birefringence lath like mineral, all of which are set in a serpentine and sericite altered matrix (e.g., figure E.7). The lath like mineral may have originally been melilite or plagioclase that has subsequently been replaced by clay minerals and carbonate or mottled colorless serpentine with low birefringence. Amoeboid juvenile clast boundaries (figure E.7) suggest that these are of probable pyroclastic origin. Additionally, pelletal lapilli are observed and are cored by a carbonate replaced olivine phenocrysts with a thin altered magma jacket (figure E.6). The interclast matrix is composed of country rock xenoliths, broken olivine macrocrysts commonly pseudomorphed by carbonate, clinopyroxene microphenocrysts and opaque minerals set in a serpentine and sericite   210 altered groundmass. The original nature of the groundmass in these deposits is difficult to discern due to the pervasive alteration and replacement of primary minerals. Epiclastic deposits occur proximal to diatreme bodies and are highly variable. Juvenile clast bearing, crossbedded conglomerates, poorly sorted, graded bedded volcanic cobble clast conglomerates, and poorly sorted, clast supported, volcaniclastic reworked breccias are typically associated with diatreme bodies where post emplacement erosion was minimal (e.g., Mountain Diatreme). For instance, Godwin and Price (1986) report the presence of a volcaniclastic reworked tuff unit that occurs on an unconformity that marks the base of the Mount Kindle formation.  Phlogopite porphyritic tuff beds (figure E.9) contain minor carbonate after olivine cored pelletal lapilli and intensely carbonate altered groundmass assemblages. Phlogopite megacrysts are generally subhedral to euhedral and are remarkably fresh and unaltered.  These diatremes are difficult to classify using genetic criteria owing to the pervasive alteration of most primary mineral phases.  However, original crystal shapes are typically preserved and inferences on original mineralogy can be therefore be used to classify these bodies.  The Tim, Monica, and Mountain diatremes are interpreted to be of the melilitite clan as evidenced from mineralogy comprising the juvenile clasts (e.g., pseudomorphed melilite, figure E.7). The Tim and Mountain diatremes are classified as heterolithic volcaniclastic olivine melilitite breccias, whereas the Monica contains less olivine and is therefore a heterolithic volcaniclastic melilitite breccia. The Bits plug (figure E.3) contains a somewhat less fragmental texture and contains minor melilite and abundant groundmass and macrocrystal phlogopite (figure E.8). Compositionally this plug is classified as a fragmental alnöite. However, if the lath like minerals present in the   211 juvenile clasts were originally plagioclase then these deposits would be of basaltic affinity. The Bear and Jane diatremes are pervasively altered to carbonate and original crystal shapes are not readily identified therefore they cannot be genetically classified. A textural descriptive name such as lapilli-tuff volcaniclastic breccia is therefore adequate for the Bear and Jane diatremes. The remaining diatremes and associated rocks identified by Copland (1994) and Carne (1994) were not genetically classified by these workers. Furthermore, there have been no detailed petrographic investigations of these bodies as they were not sampled during regional mapping.  Age and Correlation   The diatreme bodies crosscut Franklin Mountain Formation, thus indicating a maximum age of Llandeilian (Middle Ordovician; Cecile 1982). The sub Mount Kindle Formation volcaniclastic reworked tuff of the Mountain Diatreme reported by Godwin and Price (1986) suggest the youngest age of diatreme emplacement is Late Ordovician (Ashgillian interval; Norford and Macqueen 1975; Godwin and Price 1986). Additionally, conodonts were recovered from a large limestone accidental clast from within the Mountain Diatreme (McArthur et al. 1980). Ages interpreted from these conodonts suggest that the diatreme must have intruded during or after the Llanvirian interval (Middle Ordovician). Initial isotopic dating (K-Ar and Rb-Sr) was performed on phlogopite macrocrysts and whole rock samples from two occurrences in the area comprising the Mountain Diatreme and the Bear Diatreme (Godwin and Price, 1986). Ages from these studies suggest that the Mountain Diatreme and related volcanic rocks   212 range in age from Late Ordovician (Ashgillian) to late Early Silurian as evidenced by a K-Ar ages of 445±17 Ma (Mountain Diatreme) and 424±15 Ma (Bear Diatreme). The post Franklin Mountain Formation to basal Mount Kindle Formation interval is Caradocian and/or Ashgillian in age, and therefore, these K-Ar ages may be too young for the emplacement of the Mountain Diatreme and associated volcanic rocks. Advanced 40Ar/39Ar phlogopite methods are now commonly used as these new robust data are geologically meaningful and can be used to date diatreme emplacement (e.g., Kelley and Wartho 2000). New 40Ar/39Ar age determinations (Leslie et al. 2008) indicate that these diatremes are older than initially suggested. An age of 452±2.6 Ma is reported for the age of emplacement of the Mountain Diatreme while an age of 460±2.5 is reported for the Bear Diatreme (Leslie et al. 2008).  Both ages are in agreement with the aforementioned stratigraphic and conodont age constraints.  Additionally geochemical studies indicate that all diatreme deposits represent low degree partial melts of a metasomatized garnet bearing mantle source (Leslie et al. 2008).  Therefore, diatreme emplacement may have occurred in response to tapping of a relatively long-lived enriched mantle source.  Lower Paleozoic alkaline volcanic rocks are widespread through much of the northern Canadian Cordillera miogeocline and some are tentatively correlated with the Mountain Diatreme (Goodfellow et al. 1995). These include the Niddery, Itsi Lakes, and MacMillan volcanic rocks of eastern Yukon and the Vulcan volcanics located south of the Misty Creek Embayment of western North West Territories (Goodfellow et al. 1995). Additionally, Ordovician to Silurian Menzie Creek volcanic rocks of the Selwyn basin (e.g., Pigage 2004) may be also correlated; however, limited age constraints on the Menzie Creek Formation makes this correlation uncertain.  Diatremes have also been   213 identified in British Columbia, where they intrude Cambrian to Devonian miogeocline strata (Pell 1986). These diatremes include the Mark, HP, Joff and Summer pipes, and the Blackfoot diatreme of south eastern British Columbia and the Ospika pipe of north central British Columbia. On the basis of lithogeochemistry, the Mountain Diatreme and thick clinopyroxene phyric basalts and associated volcaniclastic rocks of the Marmot Formation (central Misty Creek Embayment) are suggested to be genetically related (Goodfellow et al. 1995; Leslie et al. 2008). The Marmot Formation comprises a spatially wide and diverse suite of submarine mafic volcanic occurrences with its type section (see section 17 of Cecile 1982) located in NTS map sheet 106B. Cecile (1982) identified a two-holed crinoid ossicle fossil within interbedded limestones with volcanogenic siltstones at the top of the Marmot Formation type section. The two-holed crinoid suggests a late Early to early Middle Devonian age constraint for the cessation of Marmot Formation volcanism. Therefore the diatremes of map sheet 106A may be correlative with the lower half of the Marmot Formation type section.  Economic Considerations   Godwin and Price’s (1986) report on the presence of micro diamonds in a bulk sample from the Mountain Diatreme sparked economic (diamond) interest in the Mackenzie Mountains. Unpublished investigations by Carne (1994) provided insight into the diamond potential of the Mackenzie Mountain diatremes; however, these ideas were never followed up. For example, Carne (1994) report garnet macrocrysts observed in the Bits plug, however, detailed mineral chemistry on these grains were never conducted.   214 Diamond indicator minerals were identified from stream and bulk samples in the region (Copland 1994) and were suggested to be of mantle origin, thus affirming the diamond potential (Carne 1994). However, new petrographic investigations (this study) did not identify significant olivine populations, any garnet (e.g., mantle derived xenocrysts), or mantle xenoliths (e.g., peridotite; diamond carriers), and thus the potential for diamonds is considered to be limited. The Mountain Diatreme was originally classified as a kimberlite on the basis of preliminary geochemical studies, indicator minerals recovered from stream samples (e.g., picroilmenite, pyrope garnet, and chrome diopside) and if alteration and mineral replacement is unaccounted (Godwin and Price, 1986). However, subsequent petrographic work by Scott Smith (1988) suggested that the Mountain Diatreme should not be classified as a kimberlite and has more affinity to an olivine melilitite. Classifying Misty Creek Embayment diatremes as kimberlites is negated by the lack of a significant olivine macrocryst and phenocryst population (35-55 modal % olivine is needed for most kimberlites). Mineral chemistry presented in Goodfellow et al. (1995) is also inconsistent with kimberlite affinity.  For example, phlogopite macrocrysts analyzed by these authors are titanium rich (4.49-8.26 wt.% TiO2), and therefore are unlike worldwide kimberlitic phlogopite (0-3 wt.% TiO2; Mitchell 1995). Additionally, pyroxenes analyzed by Godwin and Price (1986) from the Mountain Diatreme are aluminum rich and sodium poor distinctive of crustal equilibrated pyroxenes of non kimberlitic affinity (e.g., Morris et al. 2002).      215 Summary   Volcaniclastic diatremes were originally discovered in the Mackenzie Mountains in the early 1970s (e.g., Mountain Diatreme) and subsequent exploration in the early 1990s led to the discovery of many more diatreme and related volcanic occurrences (e.g., Copland 1994).  Diatremes exposed in the study area are unique to this area of Mackenzie Mountains where they are localized in a Lower Paleozoic rift basin termed the Misty Creek Embayment.  The eastern margin of this basin has been interpreted to be defined by crustal scale normal faults that would have acted as conduits for ascending melt. The abundance of country rock carbonate xenoliths with affinity to the Franklin Mountain Formation and the observed crosscutting relationships with the Franklin Mountain Formation suggest a maximum age of Middle Ordovician for this volcanic diatreme activity.  Volcaniclastic debris similar to the Mountain Diatreme at the base of the Mount Kindle Formation suggests the youngest age is Late Ordovician Ashgillian. New 40Ar/39Ar age determinations (Leslie et al. 2008) confirm these stratigraphic age constraints and suggest a Caradocian to Ashgillian age.  Additionally, irregular shapes of juvenile clasts suggest explosive fragmentation of the melt in the conduit followed by eruptive pyroclastic processes.  These juvenile clasts are composed of a lath like mineral that may represent pseudomorphed melilite; thus some of these deposits may belong to the melilitite clan. Alternatively if this mineral were pseudomorphed plagioclase it would suggest a basaltic affinity.  Pervasive carbonate alteration in some occurrences makes all groundmass mineral phases difficult to identify, and therefore, non-genetic textural descriptive names are most appropriate.   216 REFERENCES Carne, R.C., 1994,  Summary report on diamond potential of the Source Peaks, NWT property of NDU Resources Ltd.  Unpublished summary report; for NDU Resources Ltd.  Cas, R.A.F. and Wright, J.V., 1987,  Volcanic successions: modern and ancient. Allen and Unwin, London. 528 pp.  Cecile, M.P., 1978, Report on Road River stratigraphy and Misty Creek Embayment, Bonnet Plume (106B) and surrounding map areas, Northwest Territories; in Current Research, Part A; Geological Survey of Canada, paper 78-1A, p. 371-377.  Cecile, M.P., 1982,  The lower Paleozoic Misty Creek Embayment, Selwyn Basin, Yukon and Northwest Territories.  Geological Survey of Canada, Bull. 335, p.78  Copland, H., 1994,  Geology and geochemistry on prospecting permits 1604 to 1650. Unpublished assessment report 83438; for NDU Resources Ltd.  Godwin, C.I. and Price, B.J., 1986, Geology of the Mountain Diatreme, north-central Mackenzie Mountains, district of Mackenzie, NWT. CIM Special Volume 37, p. 298-310.  Goodfellow, W.D., Jonasson, I.R., and Cecile, M.P., 1980, Geochemistry and mineralogy of shales, cherts, carbonates, and volcanic rocks from the Road River Formation, Misty Creek embayment, Northwest Territories, Part 1; in Current research, Part B; Geological Survey of Canada, Paper 80-1B, P. 163-171.  Goodfellow, W.D., Cecile, M.P. and Leybourne, M.I., 1995, Geochemistry, petrogenesis, and tectonic setting of lower Paleozoic alkalic and potassic volcanic rocks, Northern Canadian Cordilleran Miogeocline. Canadian Journal of Earth Sciences, v. 32, no. 8, p. 1236-1254  Gordey, S.P. and Anderson, R.G., 1993,  Evolution of the Northern Cordilleran Miogeocline, Nahanni map area (105I), Yukon and Northwest Territories. Geological Survey of Canada, Memoir 428.  Kelley, S.P. and Wartho, J., 2000, Rapid kimberlite ascent and the significance of Ar-Ar ages in xenolith phlogopites. Science, 289, p. 609-611  Leslie, C.D., Sandeman, H.A. and Mortensen., J.K., 2008,  Lower Paleozoic rift related alkaline volcanic rocks, Mackenzie Mountains, NT. in 36th Annual Yellowknife Geoscience Forum Abstracts; compiled by Jackson, V and Irwin, D.; Northwest Territories Geoscience Office, Yellowknife, NT. YKGSF Abstracts Volume 2008, p. 40.  McArthur, M.L., Tipnis, R.S. and Godwin, C.I., 1980, Early and Middle Ordovician conodont fauna from the Mountain Diatreme, northern Mackenzie Mountains, N.W.T.; in Current Research, Part A; Geological Survey of Canada, Paper 80-1A, p. 363-368.   217  Mitchell, R.H., 1995, Kimberlites, Organeites and Related Rocks. Plenum, New York.  Morris, T.,F., Sage, R.P., Ayer, J.A. and Crabtree, D.C., 2002,  A study in clinopyroxene composition; implications for kimberlite exploration. Geochemistry, Exploration, Environment, Analysis, Vol. 2 2002, pp. 321-331.  Norford, B.S. and Macqueen, R.W., 1975,  Lower Paleozoic Franklin Mountain and Mount Kindle formations, District of Mackenzie: their type sections and regional development; Geological Survey of Canada, Paper 74-34  Oldershaw, A.E., 1981, A preliminary analysis of the Mountain and Keele diatremes, N.W.T.; in Mineral Industry Report 1977, Northwest Territories; Indian and Northern Affairs Canada, Report EGS 1981-11, pp. 148-154  Pell, J., 1986, Diatreme breccias in British Columbia (82,J N; 83C; 94B). British Columbia Ministry of Energy, Mines and Petroleum Resources, Geological Fieldwork, Paper 1986-1 pp. 243-253  Pigage, L.C., 2004, Bedrock geology compilation of the Anvil District (parts of NTS 105K/2, 3, 5, 6, 7 and 11), central Yukon.  Yukon Geological Survey, Bulletin 15, 103 p.  Scott Smith, B.H., 1988, Petrography of some samples submitted by C.E. Fipke. Scott- Smith Petrology Report No. SSP-88-20/2 14 p.                         218    Figure E.1. Distribution of Mackenzie Mountain diatremes and associated volcanic rocks.  Locations of unknown bodies (small stars) are from Copland (1994) and Carne (1994).        219  Figure E.2. Aerial view looking northeast of the extent of exposure of the Mountain Diatreme (dashed black and red line), dashed red line marks the location of the sub Mount Kindle Formation (MK) unconformity.  FM, Franklin Mountain Formation (UTM 476723 east, 7123556 north).      Figure E.3. Exposure of the Bits plug, looking southwest. Sk, Sekwi Formation (UTM 478719 east, 7124665 north).    220  Figure E.4. Exposure of the Monica Diatreme looking northeast. FM, Franklin Mountain Formation (UTM 472802 east, 7129586 north).    Figure E.5. Outcrop image of the Mountain Diatreme central volcaniclastic breccia. X, carbonate xenolith, MC, juvenile clast, and P, phlogopite macrocrysts (UTM 476723 east, 7123556 north).     221  Figure E.6. Photomicrograph of a phlogopite (P) cored juvenile clast with olivine (O) in the intraclast matrix and a small olivine cored pelletal lapilli (PL) from the Mountain Diatreme. (FOV, field of view; plane polarized light)    Figure E.7. Photomicrograph of amoeboid shaped juvenile clast (MC) from the Tim diatreme. The lath like mineral (M) of the juvenile clast is potentially melilite pseudomorphed by serpentine and clay minerals.  The non-lath like minerals are likely olivine (O) ± clinopyroxene pseudomorphed by carbonate. (FOV, field of view; plane polarized light).      222  Figure E.8. Photomicrograph of the Bits plug depicting an olivine phenocryst (O) replaced primarily by carbonate, clinopyroxene phenocryst (CPX) replaced by serpentine and carbonate, and groundmass phlogopite (P) crystals. (FOV, field of view; plane polarized light).     Figure E.9. Photomicrograph of a phlogopite (P) porphyritic tuff with olivine cored pelletal lapillus (PL) from the Reid occurrence. (FOV, field of view; plane polarized light).        223  Table E.1. Locations and descriptions of the Mackenzie Mountain diatremes and associated volcanic rocks Name Easting Northing Classification Accessibility Description  Genetic classification where applicable (this study) Monica 472802 7129586 diatreme yes, below steep ridge ~50 meters in diameter, poorly sorted, country rock xenolith rich, melilite bearing juvenile clast rich, phlogopite macrocrystic, lapilli-tuff volcaniclastic breccia with marginal rusty weathering reworked breccia heterolithic volcaniclastic melilitite breccia Mountain  476723 7123556 diatreme yes, valley floor and wall ~ 700 meters in diameter, texturally complex, country rock xenolith rich, olivine and phlogopite porphyritic juvenile clast rich, lapilli-tuff volcaniclastic breccia, associated marginal reworked breccia deposit, associated epiclastic tuff deposit located at sub Mount Kindle Fm. unconformity, and a central 10m wide clinopyroxene and phlogopite porphyritic central dyke  heterolithic volcaniclastic olivine melilitite breccia Tim 479031 7122502 diatreme no, cliff wall, observed talus ~30 meters in diameter, poorly sorted, exotic xenolith rich, melilite and olivine bearing juvenile clast rich, phlogopite macrocrystic, lapilli-tuff volcaniclastic breccia heterolithic volcaniclastic olivine melilitite breccia Jane 479548 7122103 diatreme no, cliff wall, observed talus ~80 meters in diameter, intensly carbonate altered, poorly sorted, exotic xenolith rich,  juvenile clast rich, phlogopite macrocrystic, lapilli-tuff volcaniclastic breccia Blurt 486490 7100862 diatreme *yes, base of cirque wall *green autolithic breccia Stib 476946 7125813 diatreme **no, extensive talus cover **diatreme of unknown size Bits 478719 7124665 plug/diatreme no, cliff wall, observed talus ~50 meters in diameter, country rock xenolith bearing, pseudomorphed olivine and clinopyroxene porphyritic, fragmental ultramafic plug, with columnar jointing at base  fragmental alnoite plug Bear 480574 7102139 diatreme yes, valley floor ~50 meters in diameter, pervasively carbonate altered, rounded country rock xenolith rich, phlogopite macrocrystic lapilli-tuff volcaniclastic breccia Tom 484884 7120818 dyke *yes, valley floor *phlogopite and plagioclase porphyritic dyke      224 Name Easting Northing Classification Accessibility Description  Genetic classification where applicable (this study) Tallis 480062 7124416 dyke yes, ridge saddle intensly carbonate and chlorite altered phlogopite porphyritic mafic dyke Golden 476589 7135089 dyke *yes, valley floor, boulder train *phlogopite and plagioclase porphyritic dyke and volcanic breccia/conglomerate Reid 467670 7142626 tuff yes, on ridge bedding concordant phlogopite porphyritic lapilli tuff with minor olivine cored pelletal lapilli and abundant country rock fragments at margins Dino 467860 7141980 dyke *yes, ridge saddle *rusty weathering phlogopite porphyritic dyke Clark 466778 7143964 dyke/sill *no, high on cliff *rusty weathering oriented phlogopite porphyritic sill Clergy 460178 7140753 dyke/sill *yes, valley floor *rusty weathering oriented phlogopite porphyritic sill * denotes descriptions from Copland (1994), ** denotes descriptions from Carne (1994))

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