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Jura-triassic magmatism and porphyry Au-Cu mineralization at the Pine Deposit, Toodoggone District, North-central… Dickinson, Jenni Michelle 2006

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JURA-TRIASSIC MAGMATISM AND PORPHYRY AU-CU MINERALIZATION AT THE PINE DEPOSIT, TOODOGGONE DISTRICT, NORTH-CENTRAL BRITISH COLUMBIA By Jenni Michelle Dickinson B.Sc. (Hons), University of St. Andrews, 2004 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 1JNWERSITY OF BRITISH COLUMBIA December 2006 © Jenni Michelle Dickinson 2006 ABSTRACT The Pine, Fin, and Mex porphyry Au-Cu ± Mo systems are all located within a 10 km2 area in the Toodoggone district, along the eastern margin of the Stikine terrane in British Columbia. Multiple episodes of porphyry-style mineralization are associated with these three magmatic centres. The Fin monzogranite is the oldest dated pluton in the district, with a U-Pb zircon emplacement age of 217.8 ± 0.6 Ma. A cross-cutting main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein gives an older Re-Os molybdenite mineralization age of 221.0 ± 1.4 Ma. Hence the vein probably cross-cuts a slightly older, undated magmatic phase. Lead isotope values for sulphide minerals from main-stage veins indicate that magmatic-derived fluids interacted with country rocks and possibly other fluids. The Pine quartz monzonite (U-Pb zircon emplacement age of 197.6 ± 0.5 Ma) intrudes, alters, and locally mineralizes Toodoggone Formation Duncan Member andesite tuff (U-Pb zircon age of 200.9 ± 0.4 Ma). High-grade (0.57 g/t Au and 0.15% Cu) mineralization occurs in main-stage quartz-magnetite chalcopyrite-pyrite veins and disseminated throughout the potassic alteration zone in the Pine quartz monzonite stock. Late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite, quartz-pyrite ± chalcopyrite, and pyrite ± chalcopyrite veins and related phyllic alteration cross-cut earlier veins. Propylitic alteration occurs distal to the potassic core of Pine in the Fin monzogranite and Duncan Member andesite. Limited fluid inclusion data, in combination with S and Pb isotope values for veins and host rocks, suggest that the main-stage fluid was magmatic-derived and deposited metals at 430 to 550 °C and depths of about 5.5 km. Late-stage fluids were also probably derived from the Pine quartz monzonite but interacted with Takla Group country rock prior to metal deposition. Metals were deposited at temperatures of 330 to 430 °C and depths of about 5.0 to 5.5 km. The final mineralization phase of the Pine porphyry system is temporally constrained by the emplacement of weakiy mineralized syenite dykes (U-Pb zircon age of 193.8 ± 0.5 Ma). The final stage of magmatism in the Pine-Fin-Mex area is defined by the emplacement of rhyolite dykes (U-Pb zircon age of 193.6±0.4 Ma). 11 TABLE OF CONTENTS Abstract ii Table of Contents iii List of Tables vi List of Figures vii Acknowledgements viii Dedication x Co-Authorship Statement xi CHAPTER I 1 General Introduction 1 Existing studies on the geology and mineralization of the Toodoggone district 1 Current NSERC-Industry funded research on porphyry and epithermal deposits in the Toodoggone district 2 Methodology 3 Geological mapping and drill core logging 3 Lithogeochemical study 3 Petrographic and mineralogical studies 4 U-Pb geochronological study 4 Re-Os geochronological study 5 Al-in-homblende geobarometry study 6 Fluid inclusion study 6 Sulphur isotope study 7 Carbon and oxygen isotope study 7 Lead isotope study 7 Thesis structure 8 References 9 CHAPTER II 12 Introduction 12 Regional geological setting 17 Geology of the Pine-Fin-Mex area 19 Lithological units 19 Toodoggone Formation 19 Fin monzogranite stock 19 Black Lake intrusive suite monzonite-granodiorite 19 111 Diorite dykes .25 Monzonite-syenite-rhyolite dykes 25 Gabbro dykes 26 Structure 26 Faults 26 Joints 27 Hydrothermal breccias 27 Geochemistry of protoliths 28 li-Pb isotope geochronology 32 Data and interpretation 32 Al-in homblende geobarometry 40 Vein paragenesis and alteration styles in the Pine-Fin-Mex area 43 Vein paragenesis and hypogene alteration styles at Pine 43 Early-stage veins and alteration 43 Main-stage veins and alteration 43 Late-stage veins and alteration 53 Propylitic alteration 54 Post-mineralization veins 55 Vein paragenesis and hypogene alteration styles at Fm 55 Vein paragenesis and hypogene alteration styles at Mex 61 Supergene alteration 61 Fluid inclusion microthermometry 62 Sample selection 62 Fluid inclusion petrography 63 Microthermometric results 66 Freezing data 66 Heating data 68 Interpretation of microthermometric data 70 Isotopes 73 Sulphur isotope data 73 Sulphide-sulphate isotope geothermometry 73 Oxygen and carbon isotope data 76 Lead isotope data 76 Discussion 80 Constraints on magmatism in the Pine-Fin-Mex area 80 Geochemistry of intrusive and extrusive rocks 80 iv Oxidation state of themagm.80 Depth of emplacement 81 Physical and chemical properties of the Pine ore fluid 81 Temperature constraints 82 Pressure and depth estimates 83 Constraints on the source(s) of metals and sulphur in the hydrothermal ore fluid at Pine 84 Sulphur isotope source constraints 84 Initial sulphur isotope composition and 50421HSratio of the Pine late-stage hydrothermal fluid ...85 Pb isotope interpretation 88 Carbon and oxygen isotope source constraints 93 Summary of S, C, 0, and Pb isotope constraints for the Fin and Pine systems 93 Integrated temporal and isotopic model for the genesis of the Pine, Fin, and Mex systems 97 Triassic magmatism and porphyry-style Au-Cu-Mo mineralization at Fin 97 Late Triassic-Early-Jurassic volcanism and faulting 99 Late Triassic-Early-Jurassic plutonism and porphyry-style Au-Cu-Mo mineralization at Pine 99 Late-stage mineralization at Pine 101 The last stages of magmatism at Pine, Fin and Mex: evidence for mafic and felsic dykes 102 Exploration implications 103 Conclusions 105 References 107 CHAPTER III 114 Conclusions 114 Further work 115 Appendix A Geochemistry results 117 Appendix B Microprobe data 122 v LIST OF TABLES Table 2.1 Mineralogy and textures of magmatic rock types in the Pine-Fin-Mex area 20 Table 2.2 Representative major- and trace-element geochemistry for magmatic rock types in the Pine-Fin-Mex area 29 Table 2.3 U-Pb geochronology data 34 Table 2.4 Summary of ages calculated from U-Pb geochronology 35 Table 2.5 Summary of microprobe analyses of homblende from intrusive phases 42 Table 2.6 Summary of mineralogy, textures, and alteration of vein types at the Pine prospect 44 Table 2.7 Summary of mineralogy, textures, and alteration of vein types at the Fin prospect 56 Table 2.8 Re-Os geochronology data 57 Table 2.9 Summary of petrographic characteristics of fluid inclusions in main- and late-stage veins 64 Table 2.10 Fluid inclusion microthermometric results from main- and late-stage veins 67 Table 2.11 Hypogene 634S isotope values and geothermometry from the Pine and Fin prospects. 74 Table 2.12 C and 0 isotope values from post-mineralization veins at the Pine deposit 77 Table 2.13 Pb isotopic values for sulphides and host rock feldspars in the Pine-Fin-Mex area 78 Table A. 1 Whole-rock major- and trace-element geochemistry for igneous rocks 117 Table B. 1 Microprobe analyses of amphiboles in the igneous rocks of the Pine-Fin-Mex area.... 122 vi LIST OF FIGURES CHAPTER II Figure 2.1 Location, terrane, and geological maps of the Toodoggone district 13 Figure 2.2 Geological map of the Pine-Fin-Mex area 15 Figure 2.3 Deposit scale cross-sections through the Pine and Fin prospects 21 Figure 2.4 Photomicrographs of rock type thin sections 22 Figure 2.5 Geochemical discrimination diagrams for igneous rocks in the Pine-Fin-Mex area 30 Figure 2.6 Tectonic setting discrimination diagrams for igneous rocks in the Pine-Fin-Mex area 33 Figure 2.7 Photomicrographs of zircon grains collected dated by U-Pb ID-TIMS 36 Figure 2.8 U-Pb concordia plots for zircon analyses 37 Figure 2.9 Vein paragenesis for the Pine prospect 45 Figure 2.10 Photoplate of vein types from the Pine prospect 46 Figure 2.11 Photomicrographs of vein types from the Pine prospect 47 Figure 2.12 Diamond drill-hole data for selected holes located throughout the Pine prospect 48 Figure 2.13 Rose diagram of relative dip of veins from the Pine prospect 50 Figure 2.14 Interpreted hypogene alteration map for the Pine-Fin-Mex area 52 Figure 2.15 Vein paragenesis for the Fin prospect 58 Figure 2.16 Diamond drill hole data for selected holes located in the Fin prospect 59 Figure 2.17 Photoplate showing vein types in hand specimens from the Fin prospect 60 Figure 2.18 Photomicrographs of fluid inclusions 65 Figure 2.19 Total homogenization by vapour to liquid versus total homogenization by halite dissolution diagrams 69 Figure 2.20 Total homogenization versus salinity diagrams 70 Figure 2.21 Estimated pressure-temperature conditions based on concordant isochore for main- and late-stage veins at the Pine deposit 72 Figure 2.22 Sulphur isotope data for pyrite and anhydrite from the Pine prospect and suiphide Figure minerals from the Fin prospect 74 Figure 2.23 634S versus34SiJphate-sulphide diagram for coexisting sulphate-sulphide mineral pairs 87 Figure 2.24 Pb isotope compositions of sulphide samples and feldspars from the Pine-Fin-Mex area on the uranogenic and thorogenic diagrams 89 Figure 2.25 207Pb/6 versus 208Pb/°6diagram 91 Figure 2.26 Oxygen and carbon data for post-mineralization veins at the Pine prospect 94 Figure 2.27 Proposed schematic model for the formation of the Pine and Mex systems 96 Figure 2.28 Detailed genetic model for the Pine-Fin-Mex area 98 vii ACKNOWLEDGEMENTS First and foremost, Dr. Paul Duuring is thanked for his guidance in both laboratory work and thesis writing, extensively editing previous versions of this work, sharing his expertise on structural geology in the field, and never shutting his door on the many questions I had. Dr. Steve Rowins provided supervision and research direction, and is thanked for leading the wider Toodoggone project. His knowledge and experience in fluid inclusion studies, porphyry systematics and economic geology have been of great benefit. Dr. Larry Diakow of the BCGS shared his immense knowledge of Toodoggone geology and discussions with him are much appreciated. At Stealth Minerals Ltd., Bill McWilliams is thanked for the financial commitment, and Dave Kuran and Ken Dawson provided support in the field. Ariadne Hiller is thanked for all of her help during the 2005 field season, especially for discussions regarding Fin and Mex. Northgate Minerals financially supported this project and are thanked for their hospitality and tour of mine geology. Andrew Orr proved an excellent field assistant (and fisherman), and his help and company are much appreciated. Cohn Smith is thanked for his contribution to the laboratory studies and the drafting of the summary regional map. At UBC, Dr. Richard Friedman quickly processed and obtained the U-Pb dates in the fall of 2005, which have been central to this work, and Janet Gabites obtained much Pb isotope data for this project. Dr. Jim Mortensen demonstrated staining for K-feldspar. Dr. Mati Raudsepp and Sasha Wilson provided guidance and instruction with probe work. Dr. Rob Creaser at the University of Alberta did the Re-Os dating. During writing of this thesis, I am very grateful for Dr. Jim Mortensen’s help with the Pb isotope interpretation and for reviewing the discussion, and Dr. Lee Groat’s editorial contributions to the whole thesis. Dr. Larry Diakow is also thanked for reviewing this work and providing helpful suggestions. Dr. Greg Dipple helped with the interpretation of C and 0 isotope data. It has been great going through this degree with Brad McKinley, and I have benefited both viii from his support and discussions on our projects. My officemates, James Thom and Steve Moss have lightened up many days. To the friends I have made since coming to Vancouver, thank you for all our good times and putting up with me in stressful times. My parents have always inspired and supported me, and I am so grateful to them for everything, from buying my boots to providing wonderful relaxing times when I go home. The late Lillian and Stanley Dickinson provided financial support when I needed it the most and enabled me to come to Canada and pursue this degree. Finally, thank you to Scott, who is always there for me. I appreciate everything, including help with the interpretation and reviewing of this work, and all the happiness that you have brought me here. ix To my parents For keeping my options open... x CO-AUTHORSHIP STATEMENT Dr. Paul Duuring contributed significantly to the interpretation of data in this project and the resulting genetic model. Dr. Duuring also provided advice in the field and instruction on fluid inclusion microthermometry and isotope sample preparation. Dr. Steve Rowins also contributed to the interpretation and discussion. Bulk rock samples were given to Dr. Richard Friedman at UBC for U-Pb dating. Selected rock samples were submitted to Dr. Rob Creaser at the University of Alberta for Re-Os dating. Crushed mineral separates were submitted to the University of Ottawa G.G. Hatch Laboratory for S, C and 0 isotope analysis. Mineral separates were submitted to Janet Gabites at UBC for Pb isotope analysis. xi CHAPTER I GENERAL INTRODUCTION Existing studies on the geology and mineralization of the Toodoggone district Geological mapping at 1:250,000 scale was first undertaken in the Toodoggone district by Lord (1948) from the Geological Survey of Canada. Reconnaissance mapping was completed in 1977 by Gabrielse et al (1977). Since then, more detailed geological mapping has been performed (Diakow 1985, 1990; Diakow 2001, 2004, 2006; Diakow et al. 1991; Diakow et al. 1993; Diakow et al. 2005; Roger and Houle 1998). The British Columbia Geological Survey (BCGS) completed a regional geochemical survey of the Toodoggone and adjacent McConnell Creek map areas in 1996 (Jackaman 1997), which has helped with porphyry-epithermal target selection (Diakow and Shives 2004). The BCGS and Geological Survey of Canada (GSC) have also released an airborne geophysical survey of the district (Shives et al. 2004). Previous published work on the mineral deposits in the district has focused on epithermal deposits (Clark and William-Jones 1986, 1989, 1991; Marsden 1990; Peter 1983; Thiersch and Williams-Jones 1990; Vulimiri et al. 1986). Diakow et al. (1991) published a summary on the epithermal systems in the Toodoggone region including data from the previous deposit studies (e.g., temperature and salinity of the hydrothermal fluids, oxygen isotope results, and K-Ar dates for hydrothermal activity). Diakow et al. (1991) proposed a depth zoning model indicating that these prospects formed in grabens and haif-grabens during district-scale extensional faulting. Previous work on porphyry systems in the district include short summaries of the geology and mineralization at the Kemess South, Kemess North, and Pine systems (Rebagliati et al. 1995; Rebagliati et al. 1995; Roger and Houle 1998). Subsequent to these studies there has been significant exploration and development in the Toodoggone district, including the recognition of other potential porphyry systems (e.g., Sofia), epithermal systems (e.g., Griz-Sickle), and the continued development of the Kemess South mine. Despite these existing exploration efforts in the district, there are few detailed deposit studies that integrate geochronology, isotope, and fluid inclusion studies with geological mapping. Current NSERC-Industry funded research on porphyry and epithermal deposits in the Toodoggone district In order to address the lack of detailed deposit studies that integrate modem laboratory techniques with detailed mapping, a collaborative three-year project was formed that aims to investigate the genesis of and links between, porphyry and epithermal deposits in the Toodoggone district. This program involves a post-doctoral Fellow (Dr. Paul Duuring) and two graduate students (Bradley McKinley and the author), supervised by Dr. Stephen Rowins of the Department of Earth and Ocean Sciences at the University of British Columbia (now working at the Saskatchewan Research Council, Saskatoon). The program commenced in June 2004 and is now in its final stages with the completion of the two M.Sc. projects in December 2006 and will be concluded in June 2007. Funding has been provided by an NSERC grant to Dr. Rowins, as well as contributions from Northgate Exploration Limited and Stealth Minerals Limited. The ultimate goal of the project is to develop a predictive exploration model for exploring for porphyry and epithermal systems in the Toodoggone district. In order to achieve this, detailed deposit studies have been completed on representative porphyry and epithermal (low- and high-sulphidation) systems in the district. Porphyry deposits/prospects studies include Kemess South (Duuring), Kemess North (McKinley), Pine (this study), and Sofia (Duuring). Epithermal deposits/prospects studied by Dr. Duuring include Baker, Lawyers, Shasta, and Griz-Sickle. This mineral deposits research is being performed in conjunction with a two-year regional mapping program in the Toodoggone district, led by Dr. Larry Diakow from the BCGS. This mapping effort is supported by several mining and exploration companies, including Northgate Exploration Limited, Stealth Minerals Limited, Findlay Minerals, Bishop Resources, and Sable Resources. The mapping project has recently been completed. The specific aims of this MSc study on the Pine porphyry Au-Cu-Mo prospect include: — Documenting the distribution, petrology, geochemistry, and timing relationships between lithologies present in the Pine-Fin-Mex area, 2 — Studying the mineralization- and alteration-styles present in the Pine system, including the relative timing of magmatic-hydrothermal events, as well as the physical and chemical properties of the ore fluid and its probable source(s), — Discerning the possible genetic relationship between the Pine, Fin, and Mex systems, and — integrating these studies into a temporal model for the genesis of the Pine system. METHODOLOGY In order to achieve the objectives of the research on the Pine-Fin-Mex area, this study combines detailed geological mapping and core logging with modem analytical techniques, including: (1) petrography and mineralogy work using transmitted light microscopes and the scanning electron microscope (SEM); (2) lithogeochemical analysis; (3) U-Pb and Re-Os geochronology; (3) Al-in hornblende geobarometry using the microprobe; (4) fluid inclusion microthermometry; and (5) S, C, 0, and Pb isotope analysis. Geological mapping and drill core logging Geological interpretation of the Pine, and adjacent Fin and Mex properties is based on surface mapping at 1:5000 scale (less than 5% exposure level) combined with detailed logging of 14 diamond drill-holes (2800 m of core from throughout the Pine deposit). Logging of drill core was done during the summers of 2004 and 2005; geological mapping was conducted over the summer of 2005. Approximately 700 hand specimens were taken from drill core and outcrop for laboratory analysis. Lithogeochemical study Major-element oxide and trace-element analyses were performed on 51 samples comprising six intrusive rocks, two volcanic rocks, and nine dykes. The masses of samples submitted for geochemical analysis ranged from 200 to 500 g. Major-element oxides were analyzed by X-ray fluorescence (XRF) and trace-elements by inductively coupled plasma mass spectrometry (ICP-MS) techniques at ALS 3 Chemex (Vancouver). Precision was estimated at l0 % by comparing multiple analyses of the same sample. Petrographic and mineralogical studies A total of 106 samples of representative rock- and vein types from the Pine-Fin-Mex area were submitted for polished and doubly polished (for fluid inclusion studies) thin sections. Optical mineralogy, and to a lesser extent, scanning electron microscopy and the electron microprobe, were used to fully characterize: (1) the mineralogy and textures of intrusive and extrusive rocks; (2) alteration in these intrusive and extrusive rocks; and (3) the paragenesis, mineralogy, and textures of hydrothermal veins and their associated alteration selvages. The Pine, Fin, and Mex intrusive and extrusive rocks were classified using the nomenclature of the International Union of the Geological Sciences based on data obtained from thin section petrography, point-counting of hydrofluoric acid etched and sodium-cobalt nitrate stained slabs, and major- and trace-element geochemistry. U-Pb geochronological study Absolute ages were determined from five U-Pb analyses of zircon from stratigraphically key rock types. All zircon U-Pb ID-TIMS sample preparation, geochemical separations, and mass spectrometry were done at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) in the Department of Earth and Ocean Sciences, University of British Columbia. Magmatic zircon was separated from samples using conventional crushing, grinding, and Wilfley table techniques, followed by final concentration using heavy liquids and magnetic separations. Mineral fractions for analysis were selected on the basis of grain quality, size, magnetic susceptibility, and morphology. All zircon crystals were air abraded prior to dissolution to minimize the effects of post-crystallization Pb loss, using the technique of Krogh (1982). Single zircon grains were dissolved in sub-boiled 48% HF and 14 M HNO3 (ratio of —‘1 0:1, respectively) in the presence of a mixed 233235U-05Pb tracer for 40 hours at 240 °C in 300 iL PTFE or PFA microcapsules contained in high-pressure vessels (ParrTM acid digestion vessels with 125 mL PTFE liners). Sample solutions were then dried to salts at —1 30 °C. Zircon residues were redissolved in —‘100 E.LL 4 of sub-boiled 6.2 M HC1 for 12 hours at 210 °C in high-pressure vessels. These solutions were transferred to 7 mL PFA beakers, and were dried to a small droplet after addition of 2 EIL of 1 M H3P04.Samples were then loaded on single, degassed zone-refined Re filaments in 5 jiL of a silica gel phosphoric acid emitter (SiC14). Isotopic ratios were measured using a modified single collector VG-54R thermal ionization mass spectrometer equipped with an analogue Daly photomultiplier. Measurements were done in peak-switching mode on the Daly detector. Analytical blanks during the course of this study were 1 pg for U and for Pb in the range of 2-3 pg. Uranium fractionation was determined directly for individual experiments using the 233235U tracer, and Pb isotopic ratios were corrected for fractionation of 0.28%, based on replicate analyses of the NBS-982 Pb standard and the values recommended by Thiriwall (2000). Reported precisions for Pb/U and Pb/Pb dates were determined by numerically propagating all analytical uncertainties through the entire age calculation using the technique of Roddick (1987). Standard concordia diagrams were constructed and regression intercepts, concordia ages and weighted means were calculated with Isoplot 3.00 (Ludwig 2003). Re-Os geochronological study An absolute age was determined by Re-Os dating of molybdenite from a vein at the Fin prospect. Whole-rock samples of molybdenite in veins were sent to the University of Alberta for selection by the laboratory. Detailed methods used for molybdenite geochronology are described in detail by Selby and Creaser (2004). The samples were processed by metal-free milling and grinding, followed by density and magnetic separation to produce a molybdenite concentrate. The 187Re and ‘870s concentrations in molybdenite were determined by isotope dilution mass spectrometry using Carius-tube, solvent extraction, anion chromatography, and negative thermal ionization mass spectrometry techniques. A mixed double spike containing known amounts of isotopically enriched ‘85Re, t900s, and 1880s was used. Isotopic analysis was done using a Micromass Sector 54 mass spectrometer with a Faraday collector. Total procedural blanks for Re and Os were less than 3 picograms and 1 picogram (<20 fg 1s7Os), respectively. These procedural blanks are insignificant in comparison to the Re and Os concentrations in molybdenite. The Chinese molybdenite powder HLP-5 (Markey et al. 1998), which is used as an in-house 5 control sample by AIRIE, Colorado State University, is also routinely analyzed at the University of Alberta. For this control sample a Re-Os date of 220.0 ± 1.0 Ma was determined at the time of the Fin molybdenite analysis. This age is identical to that reported by Markey et al. (1998). Al-in-hornblende geobarometry study Five least-altered samples with the appropriate mineralogy were selected for amphibole microprobe studies from throughout the Pine-Fin-Mex area. Two of these samples were from plutons, the Giegerich granodiorite and Mex monzonite. The other three samples were from Type 1 syenite dykes from the Pine, Fin, and Mex areas. No other rock types provided suitable material for study, mainly because the hornblende was chloritized. A Philips XL3O electron microscope with Princeton Gamma- Tech energy-dispersion X-ray spectrometer was used for qualitative analysis prior to quantitative analysis on the electron microprobe to check for zoning and alteration and to locate suitable spots for analysis. Electron-probe micro-analyses of amphiboles were obtained from polished, carbon-coated slides with a fully-automated CAMECA SX-50 instrument at the University of British Columbia, operating in the wavelength-dispersion mode with the following operating conditions: excitation voltage, 15 kV; beam current, 20 nA; peak count time, 20 s (40 s for Al, F, Cl); background count-time, 10 s (20 s for F, Cl); spot diameter, 5 urn. Data reduction was done using the ‘PAP’ p(pZ) method (Pouchou and Pichoir 1985). For the elements considered, the following standards, X-ray lines and crystals were used: synthetic phlogopite, FKc PC 1 (W/Si); albite, NaKa TAP; kyanite, A1K TAP; diopside, MgKa, TAP; diopside, SiKa, TAP; scapolite, CIKa, PET; orthoclase, KKaç PET; diopside, CaKa, PET; rutile, TiKu PET; synthetic magnesiochrornite, CrKa’, LIF; synthetic rhodonite, MnKa’, LIF; synthetic fayalite, FeKc LIF. The detection limit of the instrument is approximately 0.2 wt. % as oxides. Fluid inclusion study Microthermometric measurements were made using a Liukam THMSG 600 stage at the University of British Columbia. Precision is ± 0.2 °C for temperatures below 30 °C and ± 3.0 °C for temperatures above 30 °C (Macdonald and Spooner 1981). NaCI equivalent salinities were calculated by 6 measuring the dissolution temperature of NaC1 in liquid-rich halite-saturated inclusions. In vapour-rich (halite-undersaturated) inclusions, salinities were calculated by measuring the final melting temperature of ice. Weight percent NaCl equivalent (wt. % NaC1 equiv.) salinities were calculated using the Bodnar and Vityk (1994) equation of state for halite-saturated inclusions and Bodnar et al (1985) for halite undersaturated inclusions using the MacFlinCor software (Brown and Hagemann 1994). Sulphur isotope study Pyrite, chalcopyrite, sphalerite, and anhydrite were extracted from vein samples and analyzed for ö34S at the G.G. Hatch Isotope Laboratories, University of Ottawa, Canada. Samples were weighed into tin capsules with tungsten oxide (WO3) and loaded into a VarioEL III elemental analyzer to be flash combusted at 1800 °C in the presence of oxygen. Released gases were carried by ultra-pure helium through the instrument to be cleaned and separated. The SO2 gas was then carried into the DeltaPlus isotope ratio mass spectrometer (ThermoFinnigan, Germany) for analysis. The routine precision (2a) is ± 0.2 per mil. Data are reported as per mil deviations relative to the Canyon Diablo Troilite (CDT) standard. Carbon and oxygen isotope study Calcite was extracted from vein samples and analyzed for ö’3C and I8O at the G.G. Hatch Isotope Laboratories, University of Ottawa, Canada, using a Delta XP and a Gas Bench II following the methods of Coplen et al. (1983) and Al-Asam et al. (1990). Analytical precision (2G) for sample analyses is ± 0.1 %o. Carbon and oxygen isotope data are reported as per mil deviations relative to the Vienna Pee Dee Belemnite (V-PDB) and the Vienna Standard Mean Ocean Water (V-SMOW) standards, respectively. Lead isotope study Pyrite and feldspar crystals were extracted from vein and rock samples and analyzed for Pb at the PCIGR at the University of British Columbia, Canada. For pyrite separates, 10-50 mg of each sample was leached in dilute hydrochloric acid to remove surface contamination before dissolution in nitric acid. For 7 feldspar separates, 10-50 mg of each sample was leached in dilute hydrochloric acid, followed by washing in dilute hydrofluoric/hydrobromic acids to remove surface contamination, before dissolution in hydrofluoric acid. Separation and purification of Pb employed ion-exchange column techniques. The samples were converted to bromide, the solution was passed through ion exchange columns in hydrobromic acid, and the Pb eluted in 6N hydrochloric acid. Approximately 10-25 ng of the Pb in chloride form was loaded on Rh filaments using a phosphoric acid-silica gel emitter, and isotopic compositions were determined in peak-switching mode using a modified VG54R thermal ionization mass spectrometer. The measured ratios were corrected for instrumental mass fractionation of 0.12% (Faraday collector) per mass unit based on repeated measurements of the N.B.S. SRM 981 Standard Isotopic Reference Material and the values recommended by Thirlwall (2000). Errors were numerically propagated, including all mass fractionation and analytical errors, using the technique of Roddick (1987). All errors are quoted at the 2a level. The total procedural blank on the trace lead chemistry was 110 pg. THESIS STRUCTURE Research results for the Pine-Fin-Mex porphyry Au-Cu ± Mo systems are presented here as a manuscript that is intended for submission to an international journal. The paper provides a temporally and isotopically-constrained genetic model for magmatism and mineralization in the Pine-Fin-Mex area. The integration of geological mapping and core logging, U-Pb and Re-Os geochronology, S, Pb, C, and 0 isotope studies, and fluid inclusion microthermometry provides constraints on the evolution of the Pine Fin-Mex porphyry systems, including the potential source areas for intrusions and ore fluids. This manuscript is prepared for submission to Mineralium Deposita; headings and referencing formats are compatible with this journal. 8 REFERENCES Al-Aasm IS, Taylor BE, South B (1990) Stable isotope analysis of multiple carbonate samples using selective acid extraction. Chemical Geology: Isotope Geoscience section 80:119-125 Blann DE, Kuran D (2004) Prospecting, Geological, Geophysical, Geochemical, Trenching and Diamond Drilling Report on the Pine Property, Toodoggone Project. Stealth Minerals Limited, Canada Bodnar RJ, Vityk MO (1994) Interpretation of microthermometric data for H20-NaC1 fluid inclusions. Virginia Tech, Blacksburg, VA, p Bodnar RJ, Burnham CW, Sterner SM (1985) Synthetic fluid inclusions in natural quartz; III, Determination of phase equilibrium properties in the systemH20-NaCl to 1000 degrees C and 1500 bars. Geochimica et Cosmochimica Acta 49:1861-1873 Brown PB, Hagemann SG (1994) MacFlinCor: A computer program for fluid inclusion data reduction and manipulation. Virginia Tech, Blacksburg, VA, p Clark JR. William-Jones AE (1986) Geology and Genesis of Epithermal Gold-Barite Mineralization, Verrenass Deposit, Toodoggone District, British Columbia. Geological Association of Canada, Program with Abstracts 11:57 Clark JR, William-Jones AE (1989) New K-Ar Isotopic Ages of Epithermal Alteration from the Toodoggone River Area, British Columbia (94E). Geological Fieldwork 1988, BC Ministry of Energy, Mines and Petroleum Resources Clark JR, William-Jones AE (1991) 4OAr/39Ar Ages of Bpithermal Alteration and Volcanic Rocks in the Toodoggone Au-Ag District, North-central British Columbia (94E). Geological Fieldwork 1990, BC Ministry of Energy, Mines and Petmleum Resources Paper 1991-1:207-216 Coplen TB, Kendall C, Hopple J (1983) Comparison of stable isotope reference samples. Nature 302:236- 238 Diakow U (1985) Potassium-argon age detreminations from biotite and hornblende in Toodoggone volcanic rocks. British Columbia Ministry Energy, Mines Petroleum Research Geological Fieldwork 1984:298-300 Diakow U (1990) Volcanism and evolution of the Early and Middle Jurassic Toodoggone Formation, Toodoggone Mining District, British Columbia. University of Western Ontario Diakow U (2001) Geology of the Southern Toodoggone River and Northern McComiell Creek Map Areas, North-Central British Columbia. B.C. Ministry of Energy and Mines, BCMBM Geoscience Map 2001-1, 1:50 000 scale Diakow U (2004) Geology of the Samuel Black Range Between the Finlay River and Toodoggone River, Toodoggone River Map Area, North-central British Columbia, Parts of NTS 94E/2,6 and 7. B.C. Ministry of Energy and Mines, Open File Map 2004- 4, 1:50 000 scale Diakow U (2006) Geology between the Finlay River and Chukachida Lake, Central Toodoggone River Map Area, North-central British Columbia (Parts of NTS 94E/2, 6, 7, 10 and 11). B.C. Ministry of Energy, Mines and Petroleum Resources, Open File Map 2006-4, 1:50 000 scale 9 Diakow U, Shives RBK (2004) Geoscience partnerships in the Toodoggone River and McConnell Creek map areas, north-central British Columbia. British Columbia Ministry of Energy, Mines and Petroleum Resources, Victoria, Canada Diakow U, Panteleyev A, Schroeter T (1991) Jurassic Epithermal Prospects in the Toodoggone River Area, Northern British Columbia: Examples of Well Preserved, Volcanic-Hosted, Precious Metal Mineralization. Economic Geology 86:529-554 Diakow U, Panteleyev A, Schroeter TG (1993) Geology of the Early Jurassic Toodoggone Formation and gold-silver deposits in the Toodoggone River map area, northern British Columbia. British Columbia Ministry of Energy, Mines and Petroleum Resources, 86, Victoria, Canada Diakow U, Nixon GT, Rhodes R, Lane B (2005) Geology Between the Finlay and Toodoggone Rivers, Toodoggone River Map Area, North-central British Columbia (Parts of NTS 94E/2, 6 and 7). B.C. Ministry of Energy and Mines, Open File Map 2005-3, 1:50 000 scale Gabrielse H, Dodds CJ, Mansy JL, Eisbacher GH (1977) Geology of the Toodoggone River (94E) and Ware West-Half (94F). Geological Survey of Canada, Open File 483 Hedenquist JW, Arribas A, Reynolds TJ (1998) Evolution of an intrusion-centered hydrothermal system; Far Southeast-Lepanto porphyry and epithermal Cu-Au deposits, Philippines. Economic Geology 93:373-404 Jackaman W (1997) British Columbia Regional Geochemical Survey, NTS 94D-McConnell Creek. B.C. Ministry of Employment and Investments, Victoia, Canada Krogh TE (1982) Improved accuracy of U-Pb zircon ages by the creation of more concordant systems using an air abrasion technique. Geochimica et Cosmochimica Acta 46:637-649 Lord C (1948) McConnell Creek Map-Area, Cassiar District, British Columbia. Geological Survey of Canada Memoir 25 1:1-12 Ludwig K (2003) Isoplot 3.00 — a geochronological toolkit for Microsoft Excel. Macdonald AJ, Spooner ETC (1981) Calibration of a Linkam TH 600 programmable heating-cooling stage for microthermometric examination of fluid inclusions. Economic Geology 76:1248-1258 Markey RJ, Stein HJ, Morgan JW (1998) Highly precise Re-Os dating for molybdenite using alkaline fusion and NTIMS. Talanta 45:935-946 Marsden H (1990) Stratigraphic, Structural and Tectonic Setting of the Shasta Au-Ag Deposit, North- central British Columbia. Unpublished M.Sc. thesis, Carleton University McKinley B (2006) Geological Characteristics and Genesis of the Kemess North Porphyry Au-Cu-Mo Deposit, Toodoggone District, north-central British Columbia, Canada. Unpublished M.Sc. thesis, University of British Columbia Peter JM (1983) Mineralogy, Wall Rock Alteration, and Geochemistry of the Baker Ag-Au Mine, Toodoggone River Area, North-central B.C. Unpublished B.Sc. thesis, University of British Columbia Pouchou J-U, Pichoir F (1985) Determination by X-Ray Microanalysis of Thickness and Composition of Thin Coatings. J Microsc Spectrosc Electron 10:279-290 10 Rebagliati CM, Bowen BK, Copeland DJ (1995) The Pine Property gold-copper and copper- molybdenum prophyry prospects, Kemess-Toodoggone district, northern British Columbia. In: Schroeter TG (ed) Porphyry Deposits of the Northwestern Cordillera of North America. Canadian Institute of Mining and Metallurgy, p.436-440 Rebagliati CM, Bowen BK, Copeland DJ, Niosi DWA (1995) Kemess South and Kemess North porphyry gold-copper deposits, northern British Columbia. In: Schroeter TG (ed) Porphyry Deposits of the Northwestern Cordillera of North America. Canadian Institute of Mining and Metallurgy, p.377- 396 Roddick JC (1987) Generalized numerical error analysis with applications to geochronology and thermodynamics. Geochimica et Cosmochimica Acta 51:2129-2135 Roger C, Houle J (1998) Geological Setting of the Kemess South Au-Cu- Porphyry Deposit and Local Geology between Kemess Creek and Bicknell Lake. Geological Fieldwork for 1998, B.C. Ministry of Energy Mines and Resources, Paper 1999-1, Victoria, Canada Shives RBK, Carson 3M, Ford KL, Holman PB, Diakow L (2004) Toodoggone Multisensor Geophysical Survey (Parts of NTS 094D/1 5, E/2,3 ,6,7, 10,11): Helicopter-borne gamma ray spectrometric and magnetic total field geophysical survey, Toodoggone Area, British Columbia. BCMEMPR Open File 2004-0 8 / GSC Open Files 4606-4613, Toodoggone Multisensor Geophysical Survey (Parts of NTS 094D/1 5, E/2,3,6,7, 10,11): Helicopter-borne gamma ray spectrometric and magnetic total field geophysical survey, Toodoggone Area, British Columbia, Thiersch PC, Williams-Jones AE (1990) Paragenesis and Ore Controls of the Shasta Ag-Au Deposit Toodoggone River Area, British Columbia (94E). Geological Fieldwork 1989, BC Ministry of Energy, Mines and Petroleum Resources Paper 1990-1:315-322 Thirlwall MF (2000) Inter-laboratory and other errors in Pb isotope analyses investigated using a 207Pb- 204Pb double spike. Chemical Geology 163:299-322 Vulimiri MR, Tegart P, Stammers MA (1986) Lawyer Gold-Silver Deposits, British Columbia. In: Morn JA (ed) Mineral Deposits of the Northern Cordillera, Canadian Institute of Mining and Metallurgy. p.191 -201 11 CHAPTER II JURA-TRIASSIC MAGMATISM AND PORPHYRY AU-CU MINERALIZATION AT THE PINE DEPOSIT, TOODOGGONE DISTRICT, NORTH-CENTRAL BRITISH COLUMBIA1 INTRODUCTION The Pine porphyry Au-Cu prospect is situated in the Toodoggone district on the eastern margin of the Stikine island arc terrane of B.C., approximately 450 km north of Prince George and 20 km northeast of Thutade Lake (Fig. 2.1). The Stikine terrane hosts significant caic-alkaline and alkaline porphyry-style Au-Cu ± Mo mineralization, with major deposits including Kemess South (68 Mt at 0.65 g/t Au and 0.2 1% Cu, O’Connor 2005) and Galore Creek (proven reserve of 239.5 Mt at 0.343 g/t Au, 6.01 g/t Ag and 0.625 % Cu, Rustad et al. 2006). The Toodoggone district is currently being explored for porphyry Au-Cu ± Mo and low- and high-suiphidation epithermal Au-Ag systems. Recognized deposits in the district include the Kemess South (porphyry Au-Cu-Mo), Kemess North (porphyry Au-Cu ± Mo), Baker (low-suiphidation epithermal Au), and Shasta (low-suiphidation epithermal Ag-Au) deposits. The porphyry Au-Cu ± Mo systems are commonly hosted by Late Triassic (ca. 202-200 Ma, Diakow 2001; Mortensen et al. 1995) intermediate to felsic plutonic rocks that intrude Late Triassic-Early Jurassic intermediate to felsic volcanic tuffs and flows (Diakow et al. 1991; Diakow et al. 1993; Rebagliati et al. 1995). In comparison, the epithermal high- and low-suiphidation Au-Ag systems are younger than the porphyry systems although they overlap the later phases of plutonism in the district (Diakow et al. 1991). 1A version of this chapter has been prepared for submission to Mineralium Deposita. Dickinson, Jenni M.,Duuring,Paul, Rowins, Stephen M., Friedman, Richard M., Creaser, Robert A., Diakow, Larry, J. 12 Mid Cretaceous Z•Cu >(u 5ø .= øCDCu cuzcowo O d !.9 • I “ Late Penn sylvanian to Early Permian — 57°1O’ N Giegerich Pkgn Sustut Group Hazelton Group Toodoggone Formation Saunders Member ---a, Metsantan Member • Duncan Member , Takia Group 5 Asitka Group Wrench fault Reverse fault • Porphyry deposit Black Lake Intrusive Suite • Epithermal deposit Fig. 2.1 A. Summary geological map of the Toodoggone district south of the Toodoggone River. Note the location of the main porphyry and epithermal occurrences in the Toodoggone district and the northwest trend of major faults (after Diakow 2001; 2004; 2006; Diakow et al. 2005). Legend shows stratigraphic relationships of the main lithological sub-divisions. B. Map of British Columbia showing the location of panel C. C. Terrane map showing the location of the lower Toodoggone district (black outline) and surrounding area. Solid line represents the Pinchi-Ingenika-Finlay fault system that forms the boundary between the Stikine and Quesnel terranes. Square shows the location of the Pine deposit. 13 Interestingly, coeval mineralized and unmineralized stocks occur in proximity to one another throughout the district (based on U-Pb dating by Diakow 2001, 2004, 2006; Diakow et al. 2005) although these stocks are petrologically and geochemically similar. For example at Kemess North, the mineralized Kemess North stock occurs proximal to the coeval, but barren, Sovereign pluton (Diakow 2001; Mortensen et al. 1995). The 5 m.y. older, barren Duncan pluton intrudes the Kemess North stratigraphy and displays comparable trace-element geochemistry to the older plutons (McKinley 2006). Historically, the Pine-Fin-Mex area has been subdivided by exploration geologists into three main zones of exploration: the Pine Cu-Au-Mo, Fin Cu-Mo-Au, and Mex Cu-Au prospects based on mapping, surface soil geochemical sampling, and geophysics. The Pine and Fin prospects are superficially divided by Fin Creek, whereas the Mex prospect is located on a ridge about 2 km to the southeast of Pine and Fin (Fig. 2.2). The Pine and Fin prospects were first defined in 1968; since then 33 diamond drill-holes, 30 short percussion drill holes, outcrop mapping, geochemical soil, silt and rock sampling programs, and geophysical magnetometry and induced polarization (IP) surveys have been completed by six exploration companies. The result of this work has been the estimation of a geological reserve for the Pine area of 40 Mt at 0.57 g/t Au and 0.15% Cu (Rebagliati et al. 1995). Stealth Minerals Ltd. has explored the area since 1997, and in 2005 a subsidiary company, Cascadero Copper Corp., completed diamond drilling at the Fin and Mex prospects. The only compiled geological map that exists for the Pine-Fin-Mex area was completed by Rebagliati et al (1995); this map summarizes outcrop and drill data collected by exploration companies. Importantly, no prior study on the Pine-Fin-Mex area describes the genesis of or relationships between the Pine, Fin, and Mex prospects. This study investigates the temporal and genetic relationships between igneous rocks and porphyry-style mineralization in the Pine-Fin-Mex area. The paper provides the results and interpretation of new geologic maps and sections, in combination with geochemical, U-Pb and Re-Os geochronology, S and Pb isotopes, fluid inclusion microthermometry, and Al-in-hornblende geobarometry studies. The resultant genetic model for the Pine-Fin-Mex area has implications for wider Au-Cu ± Mo porphyry exploration in the Toodoggone district. 14 LATE TRIASSIC Fin monzogranite Diorite dyke Type 1 syenite dyke Type 2 syenite dyke Rhyolite dyke Fig. 2.2 Geological map of the Pine-Fin-Mex area based on detailed mapping and drill core logging. A. Note the three main exploration areas (Pine, Fin, and Mex) and the distribution of drill holes logged in the Pine deposit. See Fig. 2.14 for distribution of all holes drilled. Faults and dykes commonly strike northwest or kilometers F SYMBOLSLimit of mappingGeological contact (approximate)Fault (approximate) (inferred) — — — Drillhole collar with number •981 EARLY JURASSIC HAZELTON GROUP TOODOGGONE FORMATION Duncan Member _______ andesite tuff U-Pb geochronology 200.9 ± 0.4 Ma sample location with age DYKES AND SILLS LITHOLOGIES EARLY JURASSIC BLACK LAKE INTRUSIVE SUITE BLd Giegerich granodiorite Pine granodiorite ______ Pine quartz monzonite ffI Mex monzonite 15 northeast. The stereonet inset shows poles to planes of all joints from both intrusive rocks and dykes in the map area and demonstrates a dominant north to northeast strike. The solid lines AB and CD represent section lines through the study area shown in Fig. 2.3. B. Enlargement of the main zone of mineralization at Pine, showing the distribution ofdrill holes logged. 16 REGIONAL GEOLOGICAL SETTING The Intermontane tectonic belt of British Columbia (B.C.) comprises several allocthonous Mesozoic arc terranes that host significant caic-alkaline and alkaline porphyry-style Au-Cu ± Mo mineralization. The Toodoggone district of north-central B.C. is located at the northeastern extent of the Jurassic-Cretaceous Dowser Basin and contains Permian-Jurassic island arc assemblages of the Stikine and Quesnel terranes. The Stikine and Quesnel terranes are both predominantly island arc terranes that formed outboard of ancestral North America and are presently juxtaposed across the Pinchi-Ingenika Finlay fault system (Fig. 2. 1B). The history and evolution of these terranes is still controversial (i.e., see Dostal et a!. 1999) and there is much debate about how the Cretaceous amalgamation of these two arcs occurred (for different models see Coney 1989; Gehrels and Kapp 1998; Mihalynuk et al. 1994; Mortensen 1992; Nelson and Mihalynuk 1993; Tempelman-Kluit 1979; Wemicke and Kiepacki 1988). The Toodoggone district is an elongate, northwest trending, intra-arc basin that is approximately 100 km in length and 30 kin wide. This regional volcano-tectonic depression formed between the Stikine and Skeena arches, initiated by steep, oblique, westward subduction (Diakow et al. 1993). Mineralization in the Toodoggone district includes caic-alkaline porphyry- and epithermal-style occurrences, with the Kemess South porphyry Au-Cu-Mo open-pit mine (Northgate Exploration Limited) representing the present largest Au producer in B.C., with 272 Mt mined, 102 Mt milled, and almost 1.7 million ounces of Au and 417 million lbs of Cu produced between 2000 and 2005 (O’Connor 2005). The mine has a current proven reserve of 68 Mt @ 0.65 g/t Au and 0.21% Cu (O’Connor 2005). The Kemess North deposit, which is 6 km north of the Kemess South deposit, has proven resources of 300 Mt of ore at 0.30 g/t Au and 0.16% Cu (Gray and Edmunds 2005). Regional mapping (Diakow 2001, 2004, 2006; Diakow et al. 1993; Diakow et al. 2005) has identified five major lithological subdivisions (Fig. 2.lB): (1) the Late Pennsylvanian-Early Permian Asitka Group (ca. 308 Ma, Diakow 2001) comprises rhyolitic and basaltic volcanic rocks and limestone; (2) the Late Triassic Takla Group contains dominantly basaltic flows with volcaniclastic rocks and is dated the presence of local macro-fossils as Late Carnian to Early Norian (ca. 216 ± 2 Ma) (Monger and Church 1977) and in the Toodoggone by a granitic clast in a rare exposure of 17 basal conglomerate at less than Ca. 236.5 Ma (Diakow 2001); (3) the Toodoggone Formation (ca. 200-186 Ma, Diakow 2001, 2006; Diakow et a!. 1991) of the Stikinia-wide Hazelton Group is dominated by dacitic to andesitic tuffs and flows; (4) the Late Triassic-Early Jurassic Black Lake intrusive suite (ca. 202-190 Ma, Diakow 2001, 2006) is comprised of felsic-intermediate plutons and stocks, including the unmineralized Duncan and Giegerich plutons and the mineralized stocks at Kemess South, Kemess North, and Pine (Fig. 2.1); (5) the older volcanic successions were blanketed by flat-lying fluviotile continental Sustut Group sediments (part of the Overlap Assemblage, Fig. 2. 1B) during the mid-Cretaceous. The Sustut Group comprised stratigraphic successions which presumably capped and contributed to the preservation of mineralized stratigraphy (Diakow et a!. 1993). Multiple steeply dipping, northeast trending normal faults cross-cut earlier, regionally predominant, north-west trending normal fault systems (Diakow et al. 1993). The northwest trending faults have controlled the emplacement of intrusive rocks and possibly the associated porphyry and epithermal styles of mineralization (Diakow et al. 1993). Rare outcrop-scale, recumbent and inclined plunging folds that are genetically related to contraction occur in Asitka and Takia Group rocks, whereas parts of the Sustut Group define large-scale, east-verging folds (Diakow et al. 1993). Regional zeolite-grade metamorphism affects all rocks in the district (Diakow et a!. 1993). 18 GEOLOGY OF THE PINE-FIN-MEX AREA Lithological units Toodoggone Formation The Toodoggone Formation is represented in the Pine-Fin-Mex area by the Duncan Member andesitic tuff which occurs over at least 4 km2 (Fig. 2.2). It is mapped intermittently over a vertical elevation of approximately 400 m (Fig. 2.3). The andesite is dark grey and porphyritic with approximately 60% randomly orientated phenocrysts consisting of subhedral plagioclase (40%), anhedral embayed quartz (10%), euhedral amphibole (10%), and trace subhedral K-feldspar in a fine-grained groundmass (40%) (Table 2.1 and Fig. 2.4A). Although no bedding indicators were observed, Diakow (2001) described andesite dipping shallowly (<300) to the northwest southwest of the Pine-Fin-Mex area. The inferred contact between the Fin monzogranite stock and the Duncan Member andesite tuff strikes northwest. A 60 to 110 m wide, northwest striking felsic dyke occurs 30 m northeast of the contact area. Fin monzogranite stock Fin monzogranite crops out over an area of at least 2 km2 in the northeast of the map area (Fig. 2.2) and has been intersected in drill core to depths of 300 m below the surface (Fig. 2.3B). The monzogranite is grey to pale-pink in hand specimen, coarse-grained, and equigranular. Primary minerals include plagioclase (45%), K-feldspar (30%), quartz (20%), and homblende (pseudomorphed by chlorite) (5%) (Table 2.1 and Fig. 2.4B). The Fin stock contains Cu-Mo-Au mineralization. Black Lake intrusive suite monzonite-granodiorite The Black Lake intrusive suite regionally consists of granodiorite to quartz monzonite plugs, stocks, and plutons. The suite is represented in the Pine-Fin-Mex area by porphyritic to equigranular intrusive rocks ranging in composition from granodiorite to monzonite. These rocks are petrologically very similar to Black Lake intrusive suite rocks that occur throughout the Toodoggone district described by Diakow (2001; 2004; 2006) and Diakow et al. (2005). 19 T ab le 2. 1 M od al m in er al og y a n d te xt ur es o fp lu to ni c a n d v o lc an ic ro c ks fr om po in t- co un tin g an d th in se c tio n pe tr og ra ph y Ro ck u n it Te xt ur e Ph en o’ Pr im ar y m in er al s’ (% )A cc es so ry O th er co m m en ts (% ) pl ag ks p qt z hb l m in er al s D un ca n an de sit e tu ff po rp hy rit ic 60 40 2 < 2 10 10 m ag gi ne -g ra in ed gr ou nd m as sc o m pr ise s4 0% ,q tz u su al ly em ba ye d Fi n m o n zo gr an ite eq ui gr an ul ar n /a 45 30 20 5 m ag ,t it, ap . . . 2 2 Pi ne qu ar tz m o n zo n ite po rp hy nt ic — 30 35 35 20 10 m ag ,a p ex tr em el y al te re d - pr im ar y te xt ur es de str oy ed M ex m o n zo n ite eq ui gr an ul ar n /a 40 40 10 10 bi o, m ag ,t it, ap so m e ks p as o ik oc ry sts Pi ne gr an od io rit e po rp hy rit ic — 30 45 15 30 10 m ag ks p o fte n as an he dr al o ik oc ry sts G ie ge ric h gr an od io Ti te se n at e — 30 . 50 15 20 15 bi o, tit , m ag ,a p ks p o fte n as an he dr al o ik oc ry st s en cl os in g eu he dr al ch ad ac ry sts o f p la g, hb l, tit ,q tz an d m ag D io rit e dy ke ap ha ni tic n /a 50 25 53 5 m ag (5% ) w ea k tr ac hy tic te xt ur e G ab br o dy ke po rp hy rit ic 5 70 5 — 15 m ag (10 %) w ea k tr ac hy tic te xt ur e M on zo ni te dy ke po rp hy rit ic 20 -4 0 30 25 25 20 em ba ye d qt z w ith cr ys ta lli ne gr ou nd m as s Ty pe I sy en ite dy ke po rp hy rit ic 30 -5 0 25 60 5 10 m ag ,t it m ic ro gr ap hi c te xt ur e (qt z i n ks p) Ty pe 2 sy en ite dy ke po rp hy rit ic — 30 30 60 — 10 m ag ,a p em ba ye d qt z, so m e sp he ru lit es (< 0. 05 m m di am et er ) Rh yo lit e dy ke po rp hy rit ic 5 to 10 5 85 5 5 m ag m o st ly co m pr ise d o fs ph er ul ite s (0. 2-0 .5 m m di am et er ) A bb re vi at io ns : a p = ap at ite ,b io = bi ot ite , h bl = ho rn bl en de , k sp = K -fe ld sp ar ,m ag = m a gn et ite ,p he no = ph en oc ry st, pl ag = pl ag io cl as e, qt z = qu ar tz, tit = tit an ite 2 to o se ric iti ze d to ge t a n ac u ra te es tim at e m ic ro gr ap hi c qt z- ks p; n .h . f or di or ite ks p pr es en t a s bo th in di vi du al m in er al s an d m ic ro gr ap hi c qt z- ks p t. J C 1800 1600 1400 1200 1000 Gabbro dyke 193.6 ± 0.4 Rhyolite dyke — Type 2 syenite dyke 193.8 ± 0.5 + Type 1 syenite dyke None BLd Giegerich quartz diorite Diorite dyke Weak Pine granodiorite Mex monzonite Strong — 197.5 ± 0.6 _________ Pine quartz monzonite Weak 2009 ± 04 TD. Duncan Member andesite tuff Strong 217.8 ± 0.6 Fin monzogranite — — uTTa Takla Group basalt 4 U-Pb (Zircon) date Fig. 2.3 Sections through the Pine, Fin, and Mex prospects showing rock types, alteration, faults, and locations and projections of drill holes. A. The east-west section shows the four main Black Lake suite intrusions in the area. Hydrothermal alteration is zoned, with potassic alteration proximal to the Pine quartz monzonite intrusion and grading outwards into phyllic and then propylitic alteration in distal areas. Note that the faults are cut by the late syenite dykes. B. The northeast-southwest section shows the spatial distribution of the Fin pluton, Pine quartz monzonite, and Pine granodiorite. Potassic alteration surrounds the Pine granodiorite stock, whereas the Fin monzogranite displays propylitic alteration. Legend shows relative relationships between rock-types and mineralization intensity. 1800 E 1600 1400 a) 1200 1000 1800 .. 1600 1400 1200 C Southwest Pine OCJ9 0 0 Northeast D Fin 1800 1600 1400 LITHOLOGIES AND MINERALIZATION Mineralization Age Intensity (Ma) ____ None ALTERATION P Potassic (K-feldspar-magnetite) M Phyllic (quartz-sericite-pyrite) E Propylitic (epidote-chlorite) SYMBOLS Geological contact.. ______ Fault Alteration extent — .. —. Drillhole 21 Fig. 2.4 Photomicrographs of thin sections taken in cross-polarized transmitted light. A. Duncan Member andesite tuff (sample JDO5-279), showing embayed quartz and sericite-altered plagioclase. B. Fin pluton monzogranite (sample JDO5-269) showing epidote completely replacing homblende. C. Pine quartz monzonite (sample 92-3 8 162.5) showing extremely altered groundmass and phenocrysts. D. Mex monzonite (sample JDO5- 134) showing unaltered hornblende and sericite-altered plagioclase. E. Pine granodiorite (sample JDO5-288) showing plagioclase phenocrysts in groundmass ofplagioclase, quartz, and K-feldspar. F. Giegerich granodiorite (sample JDO5- 11 7a) showing K-feldspar oikocrysts and unaltered hornblende. 22 Fig. 2.4 cont. Diorite dyke (sample 92-40 1 67B) showing plagioclase laths and blebs ofmicrographic quartz- K-feldspar. H. Gabbro dyke (sample 98-04 48.3) showing aligned plagioclase laths. I. Monzonite dyke (sample F05-05 166.9) showing altered magnetite, micrographic quartz-K-feldspar, and epidote alteration of homblende. J. Type 1 syenite dyke (sample JDO5-282) showing sericite-altered plagioclase phenocrysts. K. Type 2 syenite dyke (sample JDO5-286). L. Rhyolite dyke showing spherulites. Abbreviations: chi chlorite, epi epidote, plag = plagioclase, mag = magnetite, qtz = quartz. 23 The dominant intrusive phase at Pine, and the main host to Au-Cu mineralization, is a porphyritic quartz monzonite (with local quartz-poor variants). The Pine quartz monzonite outcrops intermittently over 1 km2 and is present in drill core to depths of 290 m below the surface (Figs. 2.3A and B). It is intrusive into the basal member of the Toodoggone Formation, the Duncan Member andesite tuff. The Pine quartz monzonite is pink to orange and porphyritic with approximately 25% phenocrysts of primary plagioclase (15%) and quartz (10%). The groundmass comprises plagioclase (20%), K-feldspar (35%), quartz (10%), and homblende (pseudomorphed by chlorite) (10%) (Table 2.1 and Fig. 2.4C). Locally, a second phase of quartz monzonite, which is compositionally identical to the first, intrudes the mineralized quartz monzonite and brecciates earlier quartz veins. Porphyritic granodiorite intrudes the Pine quartz monzonite and occurs as a stock and dykes in the southern and western parts of the Pine deposit area (Figs. 2.2 and 2.3). Euhedral plagioclase phenocrysts (1 to 3.5 mm in length) comprise about 30% of the rock, with the fine- to medium-grained groundmass comprised of quartz (30%), plagioclase (15%), K-feldspar (15%), and homblende (10%) (Table 2.1 and Fig. 2.4E). Rounded, 1 to 2 cm diameter, mafic inclusions are commonly interspersed in the Pine granodiorite and appear to be partially assimilated by the granodiorite. The inclusions are dominantly comprised of plagioclase laths and hornblende, and locally display a weak trachytic texture. The Pine granodiorite does not contain significant Au-Cu-Mo mineralization. Monzonite outcrops in the eastern part of the map area on Mex ridge and extends for at least I km2 in plan (Fig. 2.2) and 300 m below the surface (Fig. 2.3A). The Mex monzonite is coarse-grained and equigranular, with primary plagioclase (40%), K-feldspar (40%), hornblende (10%), quartz (7%), magnetite (2%), and biotite (1%) (Table 2.1 and Fig. 2.4D). The monzonite hosts Au-Cu mineralization. The contact between the monzonite and andesite to the northwest is obscured by scree. The Giegerich granodiorite outcrops to the east of the Mex fault (Figs. 2.2 and 2.3B) and comprises a small portion of a larger unmineralized pluton to the east (Fig. 2.1). A phase of this pluton is dated to 197.5 ± 2.0 Ma (Diakow 2001) at Giegerich Peak, 6 km southeast of the Pine-Fin-Mex area. The Giegerich granodiorite exposed in the map area displays senate texture and primary minerals include plagioclase (50%), quartz (20%), K-feldspar (15%), and hornblende (15%) (Table 2.1 and Fig. 2.4F). 24 Diorite dykes Melanocratic, black, aphanitic diorite dykes cross-cut the Pine quartz monzonite in drill core and locally on surface. The dykes strike to northwest or northeast and are commonly 30 cm to 7 m wide. The dykes are cross-cut by the intermediate-felsic dykes described below. The diorite dykes consist of euhedral tabular plagioclase (50%) and magnetite (5%), with interstitial chlorite (15%) and phenocrysts of K-feldspar (20%) and micrographic quartz-K-feldspar (10%) (Table 2.1 and Fig. 2.4G). Chlorite may be a replacement mineral after primary hornblende. Monzonite-syenite-rhyolite dykes A series of four late monzomte to rhyolite dykes cross-cut the Pine-Fin-Mex area. The dykes mainly strike northeast, with northwest-striking dykes occurring locally. The earliest dykes are 5-10 m thick and of monzonitic composition; they are mainly exposed in the Fin area, where they cross-cut Fin monzogranite. The monzonite dykes comprise plagioclase phenocrysts (30%), micrographic K-feldspar- quartz (50%), and hornblende (20%). The three younger dykes are porphyritic, with the percentage and size of phenocrysts decreasing over time. The oldest syenite dykes (Type 1) are present throughout the Pine-Fin-Mex area and cross-cut all stocks and the diorite dykes. Type I syenite dykes are maroon-pink and comprise approximately 40% phenocrysts of euhedral tabular plagioclase (25%) (up to 10 mm in length), with lesser euhedral homblende (10%) and quartz (up to 5%) surrounded by a hematized K-feldspar-dominated (60%) aphanitic groundmass with magnetite (2%) (Table 2.1 and Fig. 2.4J). Rounded, 1-3 cm diameter fme grained mafic (plagioclase-amphibole) inclusions constitute up to 5% of the dyke and are similar in composition and texture to the inclusions present in the Pine granodiorite. The Type I syenite dykes locally contain minor (< 5%) disseminated pyrite ± chalcopyrite. Type 2 syenite dykes have fine-grained chilled margins in locations where they intrude the Type 1 syenite dykes, quartz monzonite, and granodiorite. Phenocrysts of bladed plagioclase (30%), homblende (pseudomorphed by chlorite) (10%), and rare quartz (1%) are surrounded by a groundmass of K-feldspar (60%), some of which occurs as 0.5 mm diameter spherulites (5%) (Table 2.1 and Fig. 2.4K). Phenocryst 25 size increases from 1 mm diameters at the margins, to 5 mm diameters about 40 cm away from the margin; the groundmass changes from maroon-pink to brown over this range. Rhyolite dykes cross-cut quartz monzonite, granodiorite, and Type 1 and 2 syenite dykes. The rhyolite dykes have 10% phenocrysts comprised of bladed plagioclase (5%) and chlorite (5%) that replaces elongate, euhedral hornblende crystals (up to 9 mm in length). The groundmass is comprised of 0.3 to 0.5 mm diameter spherulites of hematized K-feldspar (85%) and anhedral quartz (5%) (Table 2.1 and Fig. 2.4L). Gabbro dykes Aphanitic gabbro dykes cross-cut the Pine quartz monzonite and granodiorite stocks and syenite dykes and are grey-green and 50 cm to 2 m wide. Primary minerals include bladed plagioclase (70%) and K-feldspar (5%), with interstitial chlorite (15%) and magnetite (10%). The gabbro locally displays a trachytic texture (Table 2.1 and Fig. 2.411). Chlorite is probably a replacement mineral after primary homblende. Structure Faults Three major northwest and northeast trending faults can be mapped; they are labelled the Fin Lake, Mex, and Mex Creek faults (Figs. 2.2 and 2.3). The Fin Lake fault is observed in outcrop along the southeast shore of Fin Lake. The fault trends northeast (045°) and is marked by a sharp change in elevation along the southeast side of Fin Lake. Andesite tuff that is exposed along the northeast side of Fin Lake displays intense fractures that are northeast-striking and steeply dipping. The fault zone is 10 m wide and coincides with oxidized quartz-pyrite-sericite alteration. Late Type 1 syenite dykes outcrop along the southeast shore of Fin Lake and are unaffected by faulting. The Mex Creek fault occurs in the central map area and strikes northwest (315°). The fault is exposed as a shallowly dipping (30°) fabric in a stream bed that cross-cuts quartz-sericite-pyrite-altered Duncan Member andesite tuff. Several other creeks in the Pine-Fin-Mex area are also northwest-striking 26 and rock exposures within these creeks locally display a northwest-trending fabric. A lack of outcrop to the north of Fin Lake, due to the fluvial and glacial deposits that overlie bedrock, prevents the determination of the relative timing between the Fin Lake and Mex Creek faults. The regional-scale northwest striking Black fault occurs to the north of the Pine-Fin-Mex area and displays the same orientation as the Mex Creek fault (Diakow 2004). Elsewhere in the Toodoggone district, northwest- striking faults are cross-cut by northeast-striking faults (Diakow et al. 1993), however, these timing relationships have not been established in the Pine-Fin-Mex area. Available geophysical data are useful for the identification of large-scale faults in the Toodoggone district, however the resolution is poor at the scale of the Pine deposit and structures can not be confidently defined. The Mex fault trends north-northeast (026°), is steeply dipping (90°), and defines the contact between the unmineralized Giegerich granodiorite and the Au-Cu-Mo bearing Mex monzonite and Duncan member volcanic rocks on Mex ridge. Host rocks are intensely fractured and display a network of 1 mm-wide quartz veinlets within 10 m of the Mex fault. Chloritized mafic dykes that occur within the fault are fractured, whereas Type 1 syenite dykes cross-cut the fault zone and are undeformed. Joints Joints are developed in all observed rock types and occur throughout the Pine-Fin-Mex area. The joints display a wide range of orientations, however they mainly strike north to northeast and are moderate to steeply dipping (30 to 90°) (Fig. 2.2). A less common joint set strikes northwest and dips steeply (60 to 90°) to the northeast. Joints have a more randomly oriented strike but are steeply dipping in the intermediate-felsic dykes. Hydrothermal breccias Hydrothermal breccias occur extensively on Mex ridge and locally in the Pine-Fin area. Near surface hydrothermal breccias at Mex are matrix-supported, with the matrix comprised of yellow goethite-jarosite and about 5% pyrite (presently expressed as a weathered “boxwork” texture). Clasts within the breccia zones are 4 to 15 cm in diameter, sub-angular, and include fragments of the surrounding quartz-sericite-pyrite altered host rock. Locally on Mex ridge, the clasts are 40 to 50 cm in 27 diameter, sub-angular, and contain numerous microfractures. Hydrothermal breccias intersected in drill holes at Mex indicate that the breccia zone is funnel shaped (A. Hiller, pers. comm). Local hydrothennal brecciation of potassic-altered Pine quartz monzonite and phyllic-altered andesite tuff occurs locally seen in contact zones between Pine quartz monzonite and andesite tuff in drill-holes. Geochemistry of protoliths Major- and trace-element data for selected intrusive and extrusive rock types from the Pine-Fin Mex area are presented in Table 2.2, and all data are presented in Appendix A. 1. Most rocks in the area are moderately to intensely hydrothermally altered. Hence, only least-altered samples were selected for geochemical analysis. Sample analyses that contain loss on ignition (LOl) components more than 5% are not included in the protolith discrimination diagrams as they are considered to be significantly altered. The exception is the Pine quartz monzonite, which is universally intensely altered by hypogene fluids. For all data, the LOl component was removed and the data normalized to 100 percent before plotting. Discrimination diagrams that use mobile major elements (i.e., Si, Ca, Na, and K), such as the total alkali- silica (TAS) diagram of Le Bas et al. (1986) or the R1-R2 diagram of De la Roche et al. (1980), are unreliable for protolith classification. Therefore, discrimination diagrams that use elements that are relatively immobile during hydrothermal alteration (e.g., Zr, Ti02 Nb, and Y) are preferred. Representative rock types were plotted on the Zr-Ti02 versus NbIY diagram of Winchester and Floyd (1977) (Figs. 2.5A-C). Duncan and Saunders Members (the Saunders Member is a sample from outside the Pine-Fin-Mex area) volcanic rocks plot in the andesite field (Fig. 2.5A), whereas all the Black Lake suite intrusions all plot in the equivalent diorite intrusive field (Fig. 2.5B). Significantly, Fin pluton samples plot as a separate group in the monzodiorite field (Fig. 2.5B) and are thus markedly different from any other analyzed granitoids in the Pine-Fin-Mex area. Least-altered samples of all rock types for the Pine-Fin-Mex area and two regional samples (Duncan pluton quartz monzonite and Saunders Member dacite tuft) were plotted on the chondrite-normalized spider diagram of McDonough and Sun (1995) (Fig. 2.5D). 28 T ab le 2. 2 R en re se nt at iv e w ho le -r oc k m ajo r- an d tr ac e- el em en t g eo ch em is try o fi gn eo us ro ck s fr om th e Pi ne -F m -M ex ar ea V ol ca ni c u n its In tru siv e ro ck s dy ke s Ro ck an de ai te tu ff da ci te tu & m o n zo gr an ite qt z m o n z o n ite m o n z o n ite gr an od io rit e gr an od io rit e m o n zo n ite di or ite ga bb ro sy en ite I sy en ite 2 rh yo lit e Fi el d N o. JD O 5- 27 9 JD O 4- 10 2 JD O5 -2 69 JD OS -2 80 JD O 5- l3 4 JD OS -2 88 JD O 5- I I Th F0 5- 06 15 5. 2 SM RO 5- I 1 0 98 -0 4 48 .3 JD O5 -2 82 JD O 5- 28 6 JD OS -2 83 La b N o. N 13 96 59 N 13 96 53 N 13 96 64 N 13 96 97 N 13 96 71 N 13 96 74 N 13 96 55 N 13 96 91 N 13 96 73 N 13 96 84 N 13 96 81 N I3 96 78 N 13 96 87 Pr op er ty Pi ne /M ex Ti ge r N ot ch Fi n Pi ne M ex Pi ne G ie ge ric h Fi n Pi ne Pi ne Pi ne Pi ne Pi ne Lo ng . ( UT M ) 64 03 16 64 04 93 63 79 85 64 07 49 63 85 70 64 14 46 64 02 95 63 96 44 63 87 27 63 82 94 63 82 94 63 86 95 La t. (U TM ) 63 43 22 0 63 45 26 3 63 43 35 5 63 42 70 6 63 43 46 8 63 41 84 2 63 44 70 5 63 43 21 7 63 43 48 4 63 43 33 7 63 43 33 7 63 43 37 5 A lte ra tio n ep i+ py (se r) he m se r+ ch l+ he m se e (+ ch l) se r+ ch l+ py (se r+ ch l) ep i+ se r+ ch l ep i-i -p y ch l+ se r he m he m +c hl he m +c hl m in er al s’ (+ ch l+s er) + se r + se r Si O , 59 .7 6 62 .2 9 68 .5 0 65 .9 2 63 .9 0 64 .6 0 63 .9 7 63 .4 2 58 .3 3 51 .6 5 65 .7 8 64 .8 6 73 .9 5 T i0 2 0. 53 0. 52 0.3 1 0. 35 0. 59 0. 45 0. 59 0. 49 0. 72 0.8 1 0. 35 0. 50 0. 17 A l, 0 3 16 .7 8 15 .8 6 15 .27 14 .7 8 16 .0 3 14 .9 9 15 .8 8 15 .6 0 15 .4 0 17 .54 14 .53 15 .37 12 .9 7 F e 203 5. 06 5. 28 2. 96 5.5 9 4. 58 4. 91 4. 80 4. 80 8. 05 8. 62 3. 56 4. 08 1. 34 M oO 0. 26 0. 12 0. 13 0. 09 0. 16 0. 16 0. 13 0. 23 0. 40 0. 25 0. 09 0. 11 0. 03 M gO 1. 89 1. 96 0. 81 0. 57 1. 67 1.7 3 1. 39 1. 54 3. 30 2. 96 1. 29 1. 42 0. 24 Ca O 4. 39 3. 95 1. 49 0. 51 2. 41 3. 29 4. 54 4. 92 5. 32 5. 28 2. 78 2. 07 0. 58 N a 50 2. 71 3. 72 4. 23 3. 55 4. 13 3. 08 3. 55 2. 66 1. 08 5. 17 3. 54 4. 82 2. 73 l( 0 4. 24 3. 12 3. 39 6. 26 3. 68 2. 63 2. 90 3. 27 2. 52 2. 49 3.5 1 3. 30 5. 46 P ,0 5 0. 24 0. 21 0. 13 0. 16 0. 20 0. 14 0. 20 0. 19 0. 12 0. 32 0. 10 0. 13 0. 02 C r 203 < 0. 01 < 0. 01 0. 01 < 0. 01 0.0 1 0. 02 0. 01 < 0. 01 < 0. 01 < 0. 01 0. 0! 0. 01 0.0 1 St O 0. 07 0. 07 0. 06 0. 04 0. 05 0. 06 0. 07 0. 08 0. 06 0. 02 0. 04 0. 03 0. 02 Ba O 0. 28 0. 15 0. 15 0. 29 0. 18 0.2 1 0. 19 0. 19 0. 15 0. 12 0. 20 0. 17 0. 27 LO t 2. 00 1. 68 1. 06 0. 72 1. 24 2. 35 0. 66 1. 36 2. 73 3. 04 2. 59 1.8 5 0. 99 To ta l 98 .1 9 98 .9 3 98 .4 8 98 .8 2 98 .8 4 98 .6 2 98 .8 6 98 .7 5 98 .1 9 98 .2 5 98 .3 7 98 .7 1 98 .7 9 V 15 2 18 8 61 7! 10 1 11 6 12 4 97 18 7 21 4 84 99 16 Cr 20 20 40 20 20 80 20 30 40 < 10 10 50 10 Ni II 12 9 13 11 6 II 6 14 < 5 5 7 5 Cu 15 10 32 7 10 2 36 64 <5 5 35 5 16 17 24 6 Zn 15 6 79 10 4 11 7 23 2 12 3 63 56 1 28 3 14 3 49 61 27 G a 19 18 21 16 19 16 20 16 20 19 15 17 12 Rb 95 96 11 0 14 9 90 56 8! 56 11 2 53 96 88 14 1 Sr 60 0 66 3 52 5 36 1 43 7 62 9 66 7 68 8 50 5 22 2 45 6 33 9 25 3 Y 20 22 10 12 17 14 25 18 24 25 12 19 14 Zr 11 4 12 2 12 2 11 9 14 5 10 7 12 7 98 98 96 93 14 2 13 4 N b 6 7 4 7 6 5 7 6 4 5 5 6 8 Ba 26 60 15 00 15 65 29 40 17 25 20 30 18 85 16 95 15 20 11 30 19 30 16 30 26 80 La 18 23 12 14 15 17 23 19 II 19 17 22 26 Ce 35 43 23 26 30 30 43 35 24 40 30 41 42 N d 17 20 II 12 16 12 20 15 14 22 12 18 14 Y b 2 3 1 1 2 2 3 2 2 3 I 2 2 Pb 24 7 12 20 8 20 6 16 73 5 7 < 5 <5 Th 5 8 2 5 4 7 7 7 2 3 7 10 14 U 2 3 1 2 I 3 3 3 1 1 3 5 5 M ajo r e le m en ts ar e in w t. % ;t ra ce el em en ts ar e in pp m ‘ ch l = ch lo rit e, ep i= ep id ot e, he m = he m at ite , p y = py rit e. se c = a e ric ite . Pa re nt he si s de no te m in or pr es en ce . sa m pl ed fro m th e Pi ne -F in -M ex ar ea ‘ . 0 0.1 0.1 o N N 0.01 0.01 0.001 0.01 1000 100 C 0 10 14 12 10 6 210 200 Emplacement Age (Ma) • Duncan andesite tuff Pine granodiorite A Diorite dyke • Mex Type I syenite dykeQ Saunders dacite tuff ‘* Mex monzonite V Gabbro dyke • Pine Type I syenite dykelJa Fin monzogranite X Duncan pluton quartz monzonite EJ Monzonite dyke • Type 2 syenite dyke Pine quartz monzonite Giegerich pluton granodiorite Fin Type 1 syenite dyke Rhyolite dyke Fig. 2.5 Geochemical discrimination diagrams for igneous rocks in the Pine-Fin-Mex area. A. Toodoggone volcanic rocks define a tight cluster in the andesite field on the Winchester and Floyd (1977) plot. B. Black Lake suite data cluster in the diorite field on the Winchester and Floyd (1977) plot, whereas Fin monzogranite data plot in a separate cluster in the monzodiorite field. C. Mafic and felsic dykes show a range of compositions on the Winchester and Floyd (1977) plot. D. Chondrite-normalized multi-element plot normalized to the values ofMcDonough and Sun (1995). All samples show enriched concentrations ofLILE and depleted concentrations of Nb, P and Ti. B. REE multi-element plot normalized to the values of McDonough and Sun (1995). Note close correspondence of all samples. The Fin monzogranite has a slightly lower concentration of REE. Light REE are enriched compared to the heavy REE. F. REE fractionation over time. The Fin monzogranite is highly fractionated. G Enlargement of the early Jurassic part of the previous plot. Black Lake suite intrusive rocks show increased fractionation with time, in contrast to the dykes which do not show a consistent relationship with time. Rhyolite\ TrachyteRhyodacite - /Dacite iteTraYan \nde_/Basal/’ IAIkaline Subalkaline basalt basalt 0.1 0 N 0.01 0.1 NbIY 0.001 1 10 0.01 Gra\ e - Granodiorite—-... ./onaIite Modiôrt Diorite I Monzonhi’’ Alkali Gabi2/ gabbroj 0.1 Nb!Y 10 0,1 Nb/Y 10 Area of panel C dyke Fin monzogranite + Black Lake intrusive sui c Duncan sndesitetI) Lki 4— 220 190 Emplacement Age (Ma) 30 All rocks display similar profiles, with strong enrichment of LILE elements Rb, Ba, K, and Sr, which is typical for volcanic-arc rocks. Elsewhere, this signature has been explained, particularly for andesitic magmas, by the dehydration of a subducting slab and the subsequent metasomatism of the overlying mantle wedge (Best 1975; Dostal et al. 1977; Lopez-Escobar et al. 1977; Thorpe and Francis 1979; Thorpe et a!. 1976). All samples have small negative Nb and Ti anomalies relative to REE, which is typical for calc-alkaline arc magmas (cf. Saunders et al. 1980). Zircon, which is another high field strength element (HFSE), is not depleted. Thirwall et al. (1994) showed that depletion of the HFSE Nb, Zr, and Ti exists in the mantle wedge before partial melting due a flux produced from the subducted slab. Relative to the REE, Ti depletion is enhanced during melting and fractionation of magnetite and amphibole whereas Zr depletion is reduced (Pearce and Norry 1979). Saunders Member dacite, Duncan pluton, Pine granodiorite, Giegerich pluton, monzonite and syenite dykes (i.e., the non- to weakly-mineralized volcanic and intrusive rocks) have higher concentrations of U and Th than the Duncan member andesite, Fin monzogranite, Pine quartz monzonite, and Mex monzonite (i.e., the mineralized rocks). The strong negative P anomaly is typical of granitic rocks because P partitions into apatite early during crystallization. All volcanic rocks, Black Lake suite plutons, and dykes display the same pattern of light rare earth element (LREE) enrichment (Fig. 2.5E), which can be attributed to the presence of amphibole in felsic liquids (Rollinson 1993) and is typical for calc-alkaline arc-related rocks generated from a mantle source containing amphibole (e.g., McKenzie and O’Nions 1991). The Fin monzogranite displays a similar pattern but slightly lower concentrations of all REE. The mafic dykes have flatter REE profiles with only a minor enrichment in the LREE and higher concentrations of heavy rare earth elements (HREE) than the felsic rocks. This is common for less evolved rocks due to the lack of incompatible elements (Rollinson 1993). LREE/HREE ratios were plotted versus time for the felsic stocks, dykes, and volcanic rocks to test for magmatic fractionation trends with time (Figs. 2.5F and G). Ideally, a more-fractionated rock would display a higher (LaN/YbN)/CeN ratio than a less-fractionated rock (cf. Rollinson 1993). Figure 2.5F shows that the Fin monzogranite is commonly more fractionated than the younger Black Lake intrusive 31 suite (Fig. 2.5F), whereas fractionation ratios are comparable between intrusive rocks of the Black Lake suite and later cross-cutting dykes. Within the Black Lake suite, there appears to be an increasing fractionation trend with time (Fig. 2.5G), whereas, the dykes do not show a clear fractionation trend, probably due to their rapid emplacement. Felsic plutons and dykes all cluster in the volcanic-arc granite field of the Y+Nb vs Rb tectonic setting discrimination diagram of Pearce et al. (1984) (Fig. 2.6A). In the arc rock type discrimination diagram of Peccerillo and Taylor (1976) (Fig. 2.6B), least-altered intrusive and extrusive rocks plot in the high-K calc-alkaline field, plotting mainly in the dacite field, whereas the Fin monzogranite plots in the calc-alkaline to high-K cale-alkaline field. The Duncan and Giegerich plutons are metaluminous (Fig. 2.6C) [(K + Na) <Al < (K + Na + 2Ca)], which is typical for biotite granites and granodiorites (cf. Hall 1996). Least-altered Fin monzogranite, Pine granodiorite, and Mex monzonite are weakly peraluminous [A1/(K + Na + 2Ca) > 1], which is also typical for biotite-hornblende granites and could indicate the involvement of some crustal material (Hall 1996). All samples of Pine quartz monzonite are too hydrothermally altered to be classified using these mobile element discrimination diagrams. U-Pb isotope geochronology Zircon U-Pb isotope dilution thermal-ionization mass spectrometry (ID-TIMS) dating was completed on samples of Toodoggone Formation andesite, Pine quartz monzonite, Fin monzogranite, and Type 1 syenite dyke, and rhyolite dyke to constrain the timing of magmatic and Au-Cu-Mo mineralization events (Tables 2.3 and 2.4). Data and interpretation Zircon grains extracted from all five rock types mentioned above are colourless to pink, 0.2 to 0.8 mm long, and display euhedral prismatic morphologies (Fig. 2.7). No inherited zircon cores were detected using transmitted light microscopy (Fig. 2.7). All U-Pb data are concordant (with the exception of one analysis from the rhyolite dyke), with only minor dispersion along the concordia line most likely caused by Pb loss (Fig. 2.8). All U-Pb ID-TIMS data are presented in Table 4. The samples and their age results are described below in chronological order. 32 7Fig. 2.6 A. Tectonic setting discrimination diagram for host intrusive rocks after Pearce et al. (1984). All rocks plot as volcanic-arc granites. B. Arc rock types, after Peccerillo and Taylor (1976) showing Black Lake intrusive suite and Toodoggone volcanic rocks are high-K calc-alkaline series. The two Pine quartz monzonite samples that plot in the shoshonitic field are highly-altered. C. Shand Index classification for the intrusive rocks. Least-altered samples are metaluminous to weakly peraluminous. .0 Syn-collisional / granites /‘Iithin-plate granites Z-ridge Volcanic-arc granites granites 11111111 I II 11111 I 11111111 0 Basa’t BasAnd’ Mdeste Da Rhyolite 5 *i Shoshonjtic Serjes -4r-j • ‘— :_4’+ I )1th-K Calc’ vi Calc-Alkaline Series’ I I Ar Tholeiiteries 2 V 1000 100 10 3 2 + z 0 0.5 10 100 Y+Nb 1000 0 45 50 55 60 Sb2 65 70 75 K> Duncan andesite tuff Saunders dacite tuff + Fin monzogranite Pine quartz monzonite Pine granodiorite * Mex monzonite X Duncan pluton quartz monzonbte Giegerich pluton granodiorite A Diorite dyke V Gabbrodyke LI Monzonite dyke Fin Type I syenite dyke Mex Type 1 syenite dyke • Pine Type 1 syenite dyke • Type 2 syenite dyke o Rhyolite dyke1.0 1.5 2.0Al/(Ca+Na+K) 33 T ab le 2. 3 U -P b TI M S an al yt ic al da ta fo rm ag m at ic z irc on s fr om th e Pi ne de po si ta n d Fi n pr os pe ct Sa m pl e n o . U TM A lte ra tio n Fr ac tio n’ W t U2 P b 3 2 0 6 P b 4 P b 5 T hJ U 6 Is ot op ic ra tio s (1 a,% ) A pp ar en ta ge s (2 a,M a) & ro ck -ty pe Co or di na te s (m g) (pp m) (pp m) 2°4P b (pg ) 2 0 6 P b / 3 8 U 2 0 7 P b / 3 5 U 2 0 7 P b f °6 b 2 0 6 P b / 3 8 U 2 0 7 P b / 3 5 U 2 0 7 P b /6b JD O5 -2 81 64 02 13 E ep i a lt o f A , 10 49 1.8 33 4 3. 3 0. 48 0. 03 44 0 (0. 43 ) 0. 23 75 (2. 7) 0. 05 00 8 (2. 5) 21 8. 0 (1. 8) 21 6 (10 ) 19 9 (11 2/1 20 ) Fi n 63 45 18 3N hb l, so m e B, 1 11 46 1.6 52 8 2.1 0. 37 0. 03 43 1 (0. 29 ) 0. 23 91 (2. 7) 0. 05 05 4 (2. 5) 21 7. 5 (1. 3) 21 8 (10 ) 22 0( 11 2/1 21 ) m o n zo - se ro fp la g C, 1 12 71 2. 5 63 1 3. 0 0. 44 0. 03 43 8 (0. 17 ) 0. 23 93 (1. 2) 0. 05 04 9 (1. 1) 21 7. 9 (0. 7) 21 7. 9 (4. 7) 21 8 (52 /54 ) gr an ite ph en os D ,2 13 95 3. 3 12 21 2. 2 0. 39 0. 03 43 1 (0. 36 ) 0. 23 68 (0. 88 ) 0. 05 00 5 (0. 76 ) 21 7. 5 (1. 5) 21 5. 8 (3. 4) 19 8 (35 /36 ) JD O 5- 27 9 64 03 16 E ep ia lt o f A , 1 47 22 4 6. 9 29 34 7.1 0. 28 0. 03 16 1 (0. 12 ) 0. 21 82 (0 .22 ) 0. 05 00 6( 0.1 6) 20 0. 6( 0.5 ) 20 0. 4( 0.8 ) 19 7. 8( 7.2 /7. 3) D un ca n 63 43 22 0N hb l B, 1 25 26 3 8. 2 24 73 5. 3 0. 28 0. 03 16 9 (0. 14 ) 0. 21 98 (0. 44 ) 0. 05 03 1 (0. 39 ) 20 1. 1 (0. 6) 20 1. 7 (1. 6) 20 9 (18 ) an de sit e C, 1 27 31 4 9. 8 25 26 6. 6 0. 27 0. 03 20 2 (0. 12 ) 0. 22 22 (0. 22 ) 0. 05 03 2 (0. 16 ) 20 3. 2 (0. 5) 20 3. 7 (0. 8) 20 9. 9 (7. 6) tu ff D, 2 17 26 3 8. 2 15 78 5. 6 0. 33 0. 03 16 7 (0. 10 ) 0. 21 93 (0. 29 ) 0. 05 02 1 (0. 25 ) 20 1. 0 (0. 4) 20 1. 3 (1. 1) 20 5 (11 /12 ) JD O 5- 28 0 63 79 85 E ks p al to f A ,3 14 52 9 16 .3 72 26 2. 0 0. 32 0. 03 11 6 (0. 17 ) 0. 21 54 (0. 22 ) 0. 05 01 3 (0. 16 ) 19 7. 8 (0. 7) 19 8.1 (0. 8) 20 1. 3 (7. 4/7 .5) Pi ne 63 43 35 5N gm ,s er al t B, 3 11 49 1 15 .1 21 31 4. 9 0. 32 0. 03 10 1 (0. 31 ) 0. 21 30 (0. 88 ) 0. 04 98 1 (0. 78 ) 19 6. 9 (1. 2) 19 6.1 (3. 2) 18 6 (36 /37 ) qu ar tz o fp la g C, 3 5 49 5 15 .3 14 97 3. 2 0. 33 0. 03 10 9 (0. 26 ) 0. 21 52 (0. 58 ) 0. 05 02 1 (0. 53 ) 19 7. 4 (1. 0) 19 7. 9 (2. 1) 20 5 (25 ) m o n zo n ite ph en os D ,9 16 41 0 12 .7 24 89 5.1 0. 32 0. 03 11 4( 0.2 2) 0. 21 52 (0 .45 ) 0. 05 01 2( 0.3 9) 19 7. 7( 0.8 ) 19 7. 9( 1.6 ) 20 1 (18 ) JD O 5- 28 2 63 82 94 E se r al t A , 1 23 29 5 9. 2 23 68 5. 7 0. 26 0. 03 19 7 (0. 15 ) 0. 22 19 (0. 25 ) 0. 05 03 3 (0. 18 ) 20 2. 9 (0. 6) 20 3. 5 (0. 9) 21 0. 3 (8. 1/8 .2) Ty pe 1 63 43 33 7N o fp la g B, 1 20 33 4 10 .3 19 89 6. 3 0. 40 0. 03 05 0 (0. 19 ) 0. 21 01 (0. 57 ) 0. 04 99 6 (0. 52 ) 19 3. 7 (0. 7) 19 3. 7 (2. 0) 19 3 (24 ) sy en ite dy ke ph en os C, 1 18 41 5 12 .9 23 95 5. 9 0. 42 0. 03 05 4 (0. 17 ) 0. 21 05 (0. 29 ) 0. 04 99 8 (0. 21 ) 19 4. 0 (0. 7) 19 3. 7 (1. 0) 19 4. 0 (9. 8) D, 1 18 45 1.6 32 8 5. 6 0.4 1 0. 03 44 4 (0. 51 ) 0. 22 00 (3. 9) 0. 05 26 5 (3. 7) 21 8. 3 (2. 2) 22 7 (16 ) 31 4 (15 9/1 76 ) JD OS -2 83 63 86 95 E se ra lt A , 1 19 33 5 10 .9 23 38 5. 2 0.6 1 0. 03 04 3 (0. 21 ) 0. 21 07 (0. 50 ) 0. 05 02 1 (0. 45 ) 19 3. 3 (0. 8) 19 4.1 (1. 8) 20 5 (21 ) Ty pe 3 63 43 37 5N o fp la g B, 1 32 28 1 8. 7 23 11 7. 3 0. 43 0. 03 05 0 (0. 12 ) 0. 21 00 (0. 34 ) 0. 04 99 3 (0. 30 ) 19 3. 7 (0. 5) 19 3. 5 (1. 2) 19 2 (14 ) rh yo lit e dy ke ph en os C, 1 28 29 3 9. 3 21 22 7. 3 0. 50 0. 03 05 6 (0. 26 ) 0. 21 03 (0. 44 ) 0. 04 99 1 (0. 32 ) 19 4.1 (1. 0) 19 3. 8 (1. 6) 19 1 (15 ) A bb re vi at io ns :a lt = al te ra tio n, ep i= ep id ot e, gm = gr ou nd m as s, hb l= ho rn bl en de ,p la g = pl ag io cl as e, ph en os = ph en oc ry st s, se r = se ric te . ‘ A ll z irc on gr ai ns se le ct ed fo ra n al ys is w er e ai ra br ad ed pr io rt o di ss ol ut io n. Fr ac tio n ID (ca pit al le tte r) fo llo w ed by th e n u m be ro f g ra in s. 2 U bl an k co rr ec tio n o f l pg ± 20 % ;U fra ct io na tio n co rr ec tio ns w er e m ea su re d fo re ac h ru n w ith a do ub le 2 3 3 U -5 sp ik e. 3R ad io ge ni c Pb . 4M ea su re d ra tio co rr ec te d fo rs pi ke an d Pb fra ct io na tio n o f0 .0 02 8/ am u ± 20 % (D aly co lle ct or ), w hi ch w as de te rm in ed by re pe at ed an al ys is o f N B S Pb 98 2 st an da rd th ro ug ho ut th e co u rs e o f t hi s st ud y. co m m o n Pb in an al ys is ba se d o n bl an k iso to pi c co m po sit io n. Th /U de riv ed fro m ra di og en ic 2 08 P b an d th e 2 0 7 P b / °6 b o f f ra ct io n. 7B la nk an d co m m o n Pb co rr ec te d; Pb pr oc ed ur al bl an ks w er e — 2 pg an d U <1 pg . Co m m on Pb co n ce n tr at io ns ar e ba se d o n St ac ey -K ra m er s (19 75 )m o de lP b at th e in te rp re te d ag e o ft he ro ck o r th e 2 0 7 P b / 5 6b ag e o f t he ro ck . U -P b ID -T IM S an al yt ic al te ch ni qu es ar e lis te d in Fr ie dm an et al. (20 01 ex ce pt th at cu rr en t da ta w er e pr oc es se d w ith ou tc o lu m n ch em ist ry . Table 2.4 Interpreted ages from U-Pb ID-TIMS dating of magmatic zircon from the Pine deposit and Fin prospect Sample Unit Age & 2a (Ma) Comments JDO5-279 Toodoggone Fm. andesite 200.9 ± 0.4 3-point concordia age JDOS-280 Pine quartz monzonite 197.6 ± 0.5 4-point concordia age JDO5-281 Fin monzogranite 217.8 ± 0.6 4-point concordia age JDO5-282 Type 1 syenite dyke 193.8 ± 0.5 2-point weighted average 206Pb/38Uage JDO5-283 Rhyolite dyke 193.6 Jr 0.4 3-point concordia age 35 Fig. 2.7 Photomicrographs taken in transmitted plane polarized light of zircon grains collected from the five rock types at the Pine and Fin prospects and analyzed by U-Pb ID-TIMS. Samples analyzed were a subset of those shown. A. Duncan Member andesite tuff, sample JDO5-279. B. Fin monzogranite, sample JDO5-28 1. C. Pine quartz monzonite, sample JDO5-280. D. Type 1 syenite dyke, sample JDO5-282. E. Rhyolite dyke, sample JDO5-283. Photographs taken by Dr. Friedman. 36 Toodoggone Fm. andesite JDO5-279 0.21 0.23 0.25 0.27 0.29 207Pb/35U Fig. 2.8 U-Pb concordia plots for zircon analyses (data presented in Table 2.3). Plots, ages, and errors were calculated using the ISOPLOT program ofLudwig (2003). Data point ellipses are 2ó. Grey ellipses represent the statistical ‘concordia age’ as calculated by ISOPLOT. The black line with white ellipses represents the concordia curve with time (Ma) markers. In Fig. 2.7D, the zircon fractions are labeled (Ato D) to indicate the data used to calculate the weighted average age for the Type I syenite dyke. 207Pb/35U 207Pb/35U Type 1 syenite dyke JDO5-282 0.035 0.033 D 0.031 0.029 — 0.19 Weighted average age: 193.8±0.5 0(2 points) 220 210 ,‘ apparent discordia line A 200 C 190 B 207Pb/35U 0.211 257Pb/35U 37 Fin monzogranite A 20 kg bulk sample of least-altered monzogranite was collected from outcrop in the Fin area (sample location shown in Fig. 2.2). The concordia plot for the Fin monzogranite displays four zircon analyses from sample JDO5-281 (Fig. 2.8A). The zircon analyses show good overlap, giving a crystallization age of 217.8 ± 0.6 Ma for the Fin monzogranite. This represents the oldest magmatic event in the Pine-Fin-Mex area and also in the Toodoggone district. Toodoggone Formation andesite tuff About 25 kg of least-altered andesite tuff was sampled from outcrop located 1.5 km to the east- southeast of the Pine potassic zone (sample location shown in Fig. 2.2). The concordia plot for the andesite tuff sample (JDO5-279) shows that three of the four zircon analyses overlap (Fig. 2.8B) and gives an age of 200.9 ± 0.4 Ma. The fourth analysis gives an age of ca. 203 Ma (not shown in Fig. 2.8B) and represents the sampling of an older zircon population in the volcaniclastic rock. Hence, the younger three zircons give a maximum depositional age for the Toodoggone Formation andesite tuff of 200.9 ± 0.4 Ma. This date for the andesite tuff at Pine is comparable to the depositional age for the Duncan Member dated elsewhere in the Toodoggone district, i.e., 200.4 ± 0.3 Ma for dacite tuff northeast of Kemess North (Diakow 2001). Pine quartz monzonite A 25 kg sample of quartz monzonite (sample JDO5-280) was collected from the centre of the Au- Cu-Mo mineralized, potassic-altered Pine porphyry system (sample location shown in Fig. 2.2). Four overlapping zircon analyses in the concordia plot (Fig. 2.8C) define a crystallization age of 197.6 ± 0.5 Ma for the Pine quartz monzonite stock. This crystallization age for the granitoid host to Au-Cu-Mo mineralization gives a maximum age for mineralization at Pine. Type 1 syenite dyke A 25 kg sample of the earliest-forming syenite dyke (sample JDO5-282) was collected from drill hole 98-2, which is located in the central area of the Pine deposit (Fig. 2.2) where the dyke cross-cuts Au 38 Cu-Mo mineralized, potassic-altered (K-feldspar-magnetite) quartz monzonite. Two of the four zircon analyses from the sample overlap (analyses B and C in Fig. 2.8D) and define a 2-point weighted average age of 193.8 ± 0.5 Ma for the crystallization of Type 1 syenite dykes. The other two analyses (analyses A and D in Fig. 2.8D) are around 203 and 218 Ma and suggest an inherited component. As Type 1 syenite dykes locally contain minor disseminated pyrite and chalcopyrite, the age of the Type 1 syenite dykes provides a minimum age for porphyry-related Au-Cu-Mo mineralization at Pine of 193.8 ± 0.5 Ma. Hence, the timing of mineralization at Pine is bracketed by the emplacement of the Pine quartz monzonite host (197.6 ± 0.5 Ma) and the intrusion of late, cross-cutting, Type 1 syenite dykes (193.8 ± 0.5 Ma). Rhyolite dyke A 25 kg sample was obtained of the rhyolite dyke from outcrop in the Pine area (sample location shown in Fig. 2.2). The sampled rhyolite dyke cross-cuts phyllic-altered (quartz-sericite-pyrite) quartz monzonite, as well as Type 1 and Type 2 syenite dykes. Three zircon analyses from the sample (Fig. 2.7E) overlap and define a crystallization age of 193.6 ± 0.4 Ma for rhyolite dykes (Fig. 2.8E). This age represents the cessation of magmatism in the Pine-Fin-Mex area. 39 Al-in hornblende geobarometry The Giegerich granodiorite, Mex monzonite, and Type 1 syenite dykes from the Pine, Fin, and Mex areas contain the required mineral assemblage for Al-in-homblende geobarometry and contain unaltered amphibole. The compositions of 25 amphibole crystals from these rock types were determined (representative analyses are shown in Table 2.5 and full analyses are given in Appendix B. 1). All of the amphiboles have Ca + Na contents greater than 2.00, indicating that they are calcic amphiboles according to the classification of Leake (1978). Most amphiboles plot in the homblende field of the (Na + K)A versus Si diagram of Leake (1978), with several of the Giegerich granodiorite amphiboles plotting in the edenite field. The Fe3 and Fe2 contents were calculated using the method of Droop (1987) for calcic amphiboles. All homblende oxygen fhgacities are greater than 0.25 (Table 2.5), which is the range suggested by Anderson and Smith (1995) as suitable for the Al-in-hornblende barometer. Likewise, hornblende Fe0/(Fe+ Mg) ratios range from 0.46 to 0.62 (Table 2.5) and fall within the range suggested by Anderson and Smith (1995). The amphibole analyses have oxide weight percents that sum to between 95.25 and 98.75 percent (the remaining percent is probably the H20 content). Mineral formulas were calculated on an anhydrous basis of 23 (0) equivalents [i.e., 22(0) + 2(OH)]. Cation totals are between 15.03 and 15.55 per 23-oxygen anhydrous formula unit after the Fe correction has been applied. Theoretically, the Al-content in amphiboles increases with pressure due to tschermaks substitution; the Na- and Al-content increases with temperature due to edenite substitution. Using these substitutions, Hammerstrom and Zen (1986) generated an equation that relates Al-content in homblende to pressures of magmatic crystallization: P = 5.03 Alt0t— 3.92 kbar (Hammarstrom and Zen 1986) Following their study, several other calibrations were made, both empirically- (Hollister et al. 1987) and experimentally-derived (Johnson and Rutherford 1989; Ruther et al. 1989; Schmidt 1992; Thomas and Ernst 1990), some of which are listed below: P = 5.64 Alt0t— 4.76 kbar (Hollister et al. 1987) P = 4.28 Alt0t— 3.54 kbar (Johnson and Rutherford 1989) P = 4.76 Al’°t— 3.01 kbar (Schmidt 1992) 40 The Johnson and Rutherford (1989) equation is used in this study to calculate the crystallization pressures of analyzed hornblendes in magmatic rocks from the Pine-Fin-Mex area (Table 2.4). The experimentally-derived equation of Schmidt (1992) was not used as it was calibrated at higher pressures (2.5 to 13 kbar). The experimentally-derived equation Johnson and Rutherford (1989) is preferred because they showed that the Al content of hornblendes in equilibrium with the required phase assemblage is greater for a given total pressure than results from empirically-derived equations; they used both volcanic and plutonic rocks for their experiments; and they reduced the uncertainty associated with the geobarometer. Microprobe results that were unreliable or anomalous (i.e., those that were significantly different from the majority of analyses in a grain) were removed from the data set. Pressure estimates have a 0.5 kbar calculated uncertainty based on the uncertainty associated with each AlT (Johnson and Rutherford 1989). Application of the geobarometer to the Giegerich granodiorite (mean AlT = 1.386 ± 0.096, n = 14) gives a pressure estimate of 2.4 kbar. Analyses from the Mex monzonite (mean AlT = 1.182 ± 0.150, n = 9) gives a pressure of 1.5 kbar. Type 1 syenite dykes from the Pine area (mean AlT 1.205 ± 0.069, n = 15) produces a pressure of 1.6 kbar. Type I syenite dykes from the Fin area (mean AlT 1.402 ± 0.078, n = 23) gives a pressure of 2.3 kbar. Type 1 syenite dykes from the Mex area (mean AlT = 1.4 11 ± 0.069, n = 22) gives a pressure of 2.6 kbar. 41 TabJe 2.5 Representative microprobe analyses of amphiboles in the igneous rocks of the Pine-Fin-Mex area Giegerich Mex Mex Fin Pine granodiorite monzonite syenite dyke syenite dyke syenite dyke Sample JDO5-l 17a (8) JDO5-134 (20) JD12OB (17) F05-05 166 (23) 92-38 195 (24) Si02 Ti02 A1203 Cr203 Fe203 FeO MnO MgO CaO Na20 K20 F Total 43.67 1.44 8.21 0.00 6.66 9.56 0.73 12.05 11.37 1.31 0.97 0.25 96.24 46.16 1.78 6.73 0.00 7.63 5.98 0.50 14.83 11.18 1.64 0.51 0.25 97.19 Number of ions on the basis of 23 oxygen equivalents Si Ti Al Cr Fe3+ Fe2+ Mn Mg Ca Na K Total Fe3/(Fe-f-F2) Fe0/(Fe10+ Mg) 6.600 0.164 1.461 0.000 0.757 1.208 0.094 2.715 1.84 1 0.384 0.187 15.413 0.39 0.57 6.765 0.197 1.162 0.000 0.841 0.733 0.062 3.240 1.755 0.467 0.096 15.318 0.53 0.48 44.62 1.75 7.94 0.00 6.92 7.38 0.50 13.47 11.17 1.40 0.81 0.06 96.03 6.654 0.196 1.395 0.000 0.777 0.921 0.064 2.994 1.785 0.405 0.153 15.343 0.46 0.52 44.59 1.82 7.79 0.00 6.64 7.23 0.51 13.71 11.13 1.49 0.85 0.16 95.92 6.660 0.204 1.371 0.000 0.746 0.903 0.065 3.052 1.781 0.431 0.162 15.374 0.45 0.50 45.36 2.17 6.77 0.00 6.39 7.16 0.65 14.35 11.41 1.49 0.64 0.40 96.78 6.724 0.242 1.182 0.000 0.713 0.888 0.082 3.170 1.812 0.429 0.121 15.362 0.45 0.49 42 VEIN PARAGENESIS AND ALTERATION STYLES IN THE PINE-FIN-MEX AREA Vein paragenesis and hypogene alteration styles at Pine Ten major vein types that cross-cut quartz monzonite and granodiorite are identified at Pine (Fig. 2.9). These veins are divided into four stages of formation with respect to Au-Cu-Mo mineralization (i.e., early-, main-, late-, and post-mineralization), based on their vein mineralogy, texture, and relative cross cutting relationships (Table 2.6). Early-stage veins and alteration Early-stage veins include (1) massive quartz-rich veins and (2) magnetite veinlets (Fig. 2.9). Quartz veins are less than 10 mm wide and contain trace microscopic magnetite, chalcopyrite, and pyrite (Fig. 2.1OA). Magnetite veins are less than 3 mm wide and are abundant in the potassic core of the prospect (Fig. 2.1 OB). Early-stage veins do not display alteration selvages and are not associated with significant Au or Cu concentrations (Fig. 2.12). Main-stage veins and alteration The main phase of Au-Cu mineralization is associated with 3 to 15 mm wide quartz-magnetite chalcopyrite-pyrite veins. Primary quartz in these veins is commonly recrystallized and fine-grained (particularly at vein margins); only rarely is quartz undeformed, coarse-grained and comb-textured. Magnetite mainly occurs along vein margins and locally between comb-textured quartz crystals (Figs. 2. 1OC and 1 IB). Subhedral chalcopyrite and magnetite are the earliest-formed metallic minerals (Figs. 2.1 1A and B) and are commonly cross-cut by subhedral to euhedral pyrite grains. These veins are associated with the highest grades of Au-Cu mineralization (Fig. 2.12). Where exposed on surface, these veins commonly strike northeast and dip shallowly to the northwest (Fig. 2.1 3B). Vein angles measured from unoriented core dip between 0 and 30° relative to the horizontal (Fig. 2.1 3A). 43 Ta bl e 2. 6 D es cr ip tio ns o fv ei n ty pe s at th e Pi ne pr os pe ct , b as ed o n ha nd sp ec im en s an d th in se ct io n pe tro gr ap hy V ei n D es cr ip tio n A lte ra tio n se lv ag e St ag e’ M in er al s 2 Te xt ur e W id th (m m) D ip ’ M in er al s 2 W id th (m m) O th er co m m en ts E qt z- ric h m as siv e 2- 10 40 -6 0 E m ag m as siv e 1- 3 30 -6 0 M qt z- m ag -c py -p y So m e co m b, ty pi ca lly re cr ys ta lli ze d, m ag at m ar gi l 3- 15 30 -5 0 m ag -k sp pe rv as siv e cp y as so ci at ed w ith m ag L an hy -p y ± sp ec ± cp y (a) m as siv e, v u gg y, py ±c py at m ar gi n & th ro ug ho u 25 0- 3 5 0 20 -5 0 qt z- se r-p y < 40 an hy do m in an t o v er su lp hi de s (b) ba nd ed 2- 5 20 -5 0 qt z- se r-p y < 5 an hy -p y pr op or tio ns sim ila i L qt z- py ± cp y ce n tr al se am o fs u lp hi de s 3- 20 25 -4 5 qt z- se r-p y± ch i < 20 L py ± cp y m as siv e 2- 4 30 -5 0 qt z- se r-p y < 20 P ep ± ch l± qt z ba nd ed 2 10 -4 0 n o n e an at st om isi ng , o fte n ex pl oi t e ar lie rv ei ns P ze o m as siv e 1- 3 40 -5 5 n o n e P an hy -g yp m as siv e 2- 10 0 al l n o n e st ro ng es t i n pe rip he ra l a re as P ca ic -d ol o m as siv e 1- 3 sh al lo w n o n e E = ea rly -, M = m ai n, I= La te , p = po st- m in er al iz at io n- sta ge 2 an hy = an hy dr ite , c al c = ca lc ite , c hl = ch lo rit e, cp y = ch al co py rit e, do lo = do lo m ite , e pi = ep id ot e, ks p = K -fe ld sp ar , m ag = m ag ne tit e, py = py rit e, qt z = qu ar tz ,z eo = ze o lit e D ip an gl e in de gr ee s an d re la tiv e to th e ho riz on ta l VEIN TYPE EARLY MAIN LATE POST-MIN Quartz-rich — Magnetite Quartz-magnetite-chalcopyrite-pyrite Anhydrite-pyrite ± specular hematite ± chalcopyrite Quartz-pyrite ± chalcopyrite Pyrite ± chalcopyrite Epidote ± chlorite — Zeolite — Anhydrite-gypsum — Calcite-dolomite — Fig. 2.9 Vein paragenesis for the Pine prospect. The veins are divided into early-, main-, late-, and post mineralization-stage veins based on their vein mineralization, textures, and relative cross-cutting relationships. 45 Fig. 2.10 Photoplate of vein types from the Pine prospect. A. Early-stage quartz-rich veinlets show no alteration selvages. B. Early-stage magnetite veinlets are cut by main-stage veins. C. Main-stage quartz magnetite-chalcopyrite-pyrite vein displays magnetite along the vein margins and chalcopyrite and pyrite distributed throughout. The quartz monzonite is pervasively altered by K-feldspar and magnetite. The main- stage vein is cut by a late-stage pyrite ± chalcopyrite vein. D. Late-stage, comb-textured, quartz-pyrite ± chalcopyrite vein in chlorite and hematite-altered granodiorite. E. Late-stage massive quartz-pyrite ± chalcopyrite vein with a quartz-sericite-pyrite alteration selvage (hematized) is cut by post-mineralization anhydrite veinlets. F. Late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite vein in quartz-sericite pyrite altered quartz monzonite, with banded specular hematite and pyrite distributed throughout the vein. G Late-stage pyrite ± chalcopyrite stringer veins with quartz-sericite-pyrite alteration selvages. H. Post mineralization epidote ± chlorite ± quartz veinlets cut main-stage quartz-magnetite-chalcopyrite-pyrite veins. I. Post-mineralization anhydrite-gypsum veinlets cut a late-stage quartz-pyrite ± chalcopyrite vein and associated (hematized) quartz-sericite-pyrite alteration selvage. 46 Fig. 2.11 Photomicrographs of vein types from the Pine prospect as seen in reflected plane polarized light except where otherwise specified. A. Main-stage quartz-magnetite-chalcopyrite-pyrite vein showing a close correlation between chalcopyrite and magnetite. B. Main-stage quartz-magnetite-chalcopyrite-pyrite vein showing chalcopyrite, magnetite, and pyrite distributed between quartz crystals in the centre of the vein. C. Main-stage quartz-magnetite-chalcopyrite-pyrite vein (highlighted) with abundant magnetite and minor chalcopyrite disseminated in the groundmass. D. Aggregate of magnetite grains in potassic-altered quartz monzonite. B. K-feldspar alteration in the potassic-altered quartz monzonite (as seen in cross-polarized transmitted light). F. Late-stage quartz ± pyrite ± chalcopyrite vein that is cut by a post-mineralization anhydrite vein. G Deformed pyrite ± chalcopyrite vein that shows cataclasis of pyrite with chalcopyrite filling fractures. H. Less-deformed pyrite and chalcopyrite in a late-stage pyrite-chalcopyrite stringer vein showing anhedral chalcopyrite grains surrounded by pyrite. I. Sericite alteration that surrounds a late-stage pyrite-chalcopyrite stringer vein (as seen in cross-polarized transmitted light). 47 VEINS Qtz-rich — Mag — Qtz-mag-cpy-py —Anhy-py±spec±cpy — Qtz-py ± cpy Py ± cpy —Epi±chl±qtz — Zeo — Anhy-gyp — Caic-dolo LITHOLOGIES Overburden [jPine granodiorite l] Pine quartz monzonite Type 1 syenitedyke Type 2 syenite dyke Rhyolite dyke Gabbro dyke Fault zone Fig. 2.12 Diamond drill hole data for selected holes located throughout the Pine deposit (Fig. 2.2) show the distribution of lithologies, veins, and metals (Cu, Au ± Ag, Mo, Pb, Zn) and other trace-element concentrations. A. Hole 92-40 (collar located at 638 116E, 6343565N, drilled at -65° towards 270°) is 200 m 50 m Fe P (glt) (ppm) (ppm) (ppm) (ppm) 48 long and demonstrates the close association between Cu and Au values with main-stage quartz-magnetite chalcopyrite-pyrite veins (nibbled zone) and the decrease in Cu and Au values in the syenite dykes. Cu andAu values are highest at the top of the hole. B. Hole 98-2 (collar located at 638294E, 6343337N, drilled at 600 towards 2700) is 246 m long and demonstrates the close association between Cu and Au values with main stage quartz-magnetite-chalcopyrite-pyrite, late-stage quartz-pyrite ± chalcopyrite, and late-stage pyrite ± chalcopyrite veins. Copper and Au values decrease in the granodiorite at the base of the hole. C and D. Hole 97-04 (collar located at 6384 lOB, 6343528N, drilled at 7O0 towards 270°) is 192 m long. C demonstrates the close association of Cu and Au values with late-stage quartz-pyrite ± chalcopyrite and pyrite ± chalcopyrite veins, and the association between high Cu, Au, Ag, and Mo values, with Pb and Zn showing separate peaks. Fig. 2.1 2D demonstrates a correlation between high Cu andAu values and increased Fe, Mn, P andAl. 49 Fig 2.13 A. Rose diagrams showing angles ofveins from the Pine prospect, measured from core relative to the horizontal. Early-stage veins are oriented 0 to 500 with most oriented at 0 and 30° to 45°. Main-stage veins are oriented 0 to 70° with most oriented between 0 and 25°. Late-stage veins (including anhydrite-pyrite ± specular hematite ± chalcopyrite veins) are oriented 0 to 75° with most oriented between 25 and 35°. Post- mineralization veins are oriented 0 to 90° with most oriented between 0 and 10°. B. Equal area (Schmidt) stereonet showing poles to planes of main-stage veins from the Pine and Fin prospects. Veins from the Pine prospect mostly dip shallowly to the northwest. 0 z + Pine qtz-mag-cpy-py vein + Fin qtz-py-cpy ± moly ± sph vein n = 17 50 Quartz-magnetite-chalcopyrite-pyrite veins do not display discrete alteration mineral selvages, but they are spatially associated with broad K-feldspar-magnetite-chalcopyrite-pyrite alteration in quartz monzonite (Figs. 2.1 OA to C, 2.11 B to B). Fine-grained (< 0.25 mm) hydrothermal K-feldspar and magnetite are disseminated in the groundmass of the quartz monzomte. The magnetite occurs as irregularly shaped, 2 mm diameter aggregates comprised of smaller (-0. 1 mm) subhedral magnetite grains (Fig. 2.1 1D). Magnetite preferentially replaces homblende crystals. Potassic alteration only occurs in the quartz monzonite (Fig. 2.14). Based on outcrop and drill hole intersections beneath glacial till, the near-surface potassic alteration zone extends for a horizontal area of approximately 300 m2 and continues to at least 200 m below surface (i.e., to the base of drilling) in the centre of the Pine area (Figs. 2.3 and 2.12). Interestingly, K-feldspar-magnetite alteration does not occur in the Pine granodiorite that occurs to the southwest and beneath the intensely potassic-altered quartz monzonite (Fig. 2.3). Rather, the granodiorite that intrudes the quartz monzonite is relatively fresh in the centre of the stock, with minor alteration defined by the chioritization of igneous hornblende and biotite and sercitization of K-feldspar and plagioclase increasing towards the margins. Dykes of granodiorite are often completely hematized, with more intense sericite and chlorite alteration. High-grade Au and Cu mineralization is centred on the potassic (K-feldspar-magnetite) altered quartz monzonite. For example, diamond drill hole 92-40 intersects more than 140 m of quartz monzonite with an average grade of 0.95 ppm Au and 0.18% Cu (Figs. 2.3 and 2.12); the highest grades occur near surface (Fig. 2.1 2A). Gold concentrations show a positive relationship to Cu grades (Fig. 2.12). The granodiorite stock to the southwest contains low concentrations of Au and Cu, with diamond drill hole 97-12 intersecting 120 m of granodiorite with an average grade of 0.02 ppm Au and 0.02% Cu. Diorite and Type 1 syenite dykes locally host disseminated and fracture-filling chalcopyrite and pyrite. Diamond drill hole 92-40 intersects 2 m of diorite dyke that has an average grade of 0.09 ppm Au and 0.03% Cu. The hole also intersects 5 m of Type 1 syenite dyke that has an average grade of 0.03 ppm Au and 0.004% Cu. It is uncertain if the pyrite and chalcopyrite precipitated during the crystallization of the Type 1 syenite dykes or if the sulphide minerals were remobilized from the proximal Pine porphyry system. 51 Fig. 2.14 Interpreted hypogene alteration map for the Pine, Fin, and Mex areas based on outcrop and diamond drill hole data. Potassic alteration in the Pine area corresponds to the highest density of main-stage quartz magnetite-chalcopyrite-pyrite veins and Au-Cu mineralization, and minor potassic alteration occurs at Mex. Potassic zones are surrounded by a zone ofphyllic alteration, which is associated with late-stage quartz-pyrite ± chalcopyrite and pyrite ± chalcopyrite veins and quartz-sericite-pyrite alteration and lower grades ofAu-Cu mineralization. Phyllic alteration also occurs at the Mex area in a region that is intensely brecciated. Propylitic alteration, defined by epidote ± chlorite and low-grade mineralization, occurs as a peripheral zone around phyllic alteration at Pine and Mex, and throughout the Fin area. Fresh rock flanks propylitic alteration at Pine and Mex and contains negligible veins and Au-Cu mineralization. Note the distribution of drill holes throughout the three exploration areas, including those drilled in 2005 by Cascadero Copper (numbered). ALTERATION SYMBOLS Potassic (K-feldspar-magnetite) Limit of mapping _____ Approximate property limits — —Phyllic (quartz-sericite-pynte) Alteration extent (inferred) — — — - Propylitic (epidote-chiorite) Diamond drillhole collar (logged) Fresh to hematized Diamond drillhole collar (drilled by Cascadero Copper 2005) . Diamond drillhole collar Percussion drillhole collar 52 Late-stage veins and alteration Late-stage veins include (1) anhydrite-pyrite ± specular hematite ± chalcopyrite, (2) quartz-pyrite ± chalcopyrite and (3) pyrite-chalcopyrite veins. All late-stage veins are associated with quartz-sericite pyrite (i.e., phyllic) alteration and only minor Au-Cu mineralization (Fig. 2.10). Anhydrite-pyrite ± specular hematite ± chalcopyrite veins Anhydrite-pyrite ± specular hematite ± chalcopyrite veins and their associated 5 to 40 mm wide quartz-sericite-pyrite alteration selvages replace earlier potassic (K-feldspar) alteration (Figs. 2.1 OF and 2.1 iF). The veins are most common in areas peripheral to the potassic-altered core of the Pine quartz monzonite but they are also locally developed within the central potassic zone. Vein angles measured from unoriented core dip between 10 and 45° relative to the horizontal (Fig. 2.1 3A). These veins are divided into two sub-types, including (1) massive, vuggy, anhydrite-dominated veins up to 350 mm wide containing pyrite and chalcopyrite throughout the vein; and (2) banded, 2 to 5 mm wide veins that contain equal proportions of anhydrite and pyrite ± chalcopyrite. Quartz-pyrite ± chalcopyrite veins Quartz-pyrite ± chalcopyrite veins contain fine-grained quartz and are often extremely recrystallized (Fig. 2.1 OE). Where primary textures are locally preserved, quartz is anhedral or comb- textured (Fig. 2.1 OD), with anhedral to subhedral pyrite and chalcopyrite located between euhedral quartz crystals in the centre of the vein. Pyrite commonly rims or fills fractures in chalcopyrite. Quartz-pyrite ± chalcopyrite veins are surrounded by 5 to 20 mm wide quartz-sericite-pyrite (i.e., phyllic) alteration selvages. Vein angles measured from unoriented core dip between 0 and 40° relative to the horizontal (Fig. 2.13A). Pyrite ± chalcopyrite veins Pyrite ± chalcopyrite stringer veins are 2 to 4 mm wide and are surrounded by quartz-sericite ± hematite alteration selvages that are up to 20 mm wide (Figs. 2.IOG and 2.11G to I). The pyrite ± chalcopyrite stringer veins occur as discrete veins or as microveins that cross-cut earlier main- and late 53 stage veins perpendicular to their vein margins. Vein angles measured from unoriented core dip between o and 400 relative to the horizontal (Fig. 2.1 3A). Late-stage pyrite-chalcopyrite veinlets also locally occur along contacts between diorite dykes and quartz monzonite. Phyllic alteration All three types of late-stage veins have associated phyllic alteration in the quartz monzonite and andesite tuff. Phyllic-altered rock crops out to the northeast, east, and southeast of the potassic alteration core (Fig. 2.10) and is defined by pervasive, fme-grained quartz-sericite-pyrite alteration. Phyllic alteration replaces earlier potassic alteration and is most pronounced where late-stage veins are the dominant vein type. In these intensely-altered areas, primary igneous textures are destroyed and the igneous protolith is difficult to identify in hand specimen. Altered volcaniclastic rocks can be distinguished from intrusive rocks in the area in thin-section by the presence of embayments in primary quartz phenocrysts. Phyllic-altered quartz monzonite at Pine generally hosts lower grades of Au-Cu mineralization relative to potassic-altered zones. For example, diamond drill hole 98-4 intersects 110 m of phyllic-altered quartz monzonite that has an average grade of 0.17 ppm Au and 0.026% Cu. Over a similar sample distance (i.e., 141 m) in the same hole, potassic-altered quartz monzonite has a higher average grade of 0.31 ppm Au and 0.14% Cu. Propylitic alteration Phyllic-related quartz-sericite-pyrite alteration is replaced by epidote-chiorite (propylitic) alteration to the east and northeast of the phyllic zone (approximately 1 km from the potassic core) (Fig. 2.14). Propylitic alteration is defined by epidote ± chlorite alteration of primary hornblende in the quartz monzonite, the Toodoggone Formation andesite tuff, and the Fin monzogranite. No macro-scale veins were identified that are responsible for the propylitic alteration. Rather, it is interpreted that microfractures acted as conduits to fluids and caused the broadly distributed propylitic alteration in the rocks. Propylitic alteration is not associated with mineralization, with metal concentrations in these rocks mostly below about 0.01 ppm Au and 0.02% Cu. 54 Post-mineralization veins Post-mineralization veins cross-cut all rock and vein types. The earliest epidote ± chlorite ± quartz veins are cross-cut by zeolite veins, which are in turn cross-cut by anhydrite-gypsum, and late calcite-dolomite veins. Vein angles measured from unoriented core dip between 0 and 200 relative to the horizontal (Fig. 2.l3A). Vein paragenesis and hypogene alteration styles at Fin Eight generations of veins cut the Fin monzogranite (Table 2.7). Early-stage, massive quartz veins that are 0.5 to 2 mm wide cross-cut the Fin monzogranite (Fig. 2.15). No significant alteration or Cu-Mo-Au mineralization is associated with these veins (Fig. 2.16). Main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite veins are 1 to 30 mm wide and commonly display crustiform textures, with pyrite and chalcopyrite occurring in the centre of the vein while molybdenite and sphalerite define 1 mm wide bands along vein margins (Figs. 2.1 7A and B). The few veins that occur on surface strike to the north or east. These veins are surrounded by pervasive epidote and chlorite alteration and are responsible for most of the Cu, Au, Mo and Zn mineralization at Fin (Fig. 2.16). Molybdenite from a main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein that cross-cuts the Fin monzogranite gives a Re-Os model age of 221 ± 1.4 Ma (Table 2.8). This age is slightly older but comparable to the 217.8 ± 0.6 Ma (U-Pb on zircon) age for the emplacement of the Fin monzogranite, and indicates that the vein probably cross-cuts a slightly older, undated magmatic phase. Main-stage epidote-pyrite-chalcopyrite ± molybdenite ± sphalerite veins are 1 to 20 mm wide and cross-cut the main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite veins (Fig. 2. l7C). The epidote-pyrite-chalcopyrite ± molybdenite ± sphalerite veins are also associated with epidote and chlorite alteration in the Fin monzogranite as well as Cu, Au, Mo, and Zn mineralization. Late-stage pyrite ± chalcopyrite ± epidote stringer veins and molybdenite ± sphalerite ± pyrite microveins cross-cut the main-stage veins and are responsible for minor Cu-Mo mineralization (Figs. 2.1 7D and E). Post-mineralization epidote-chiorite, zeolite, and anhydrite-gypsum veins cross-cut all other vein types (Fig. 2.17F). 55 Ta bl e 2. 7 D es cr ip tio ns o fv ei n ty pe sa tt he Fi n pr os pe ct ,b as ed o n ha nd sp ec im en sa n d th in se ct io n pe tro gr ap hy V ei n D es cr ip tio n A lte ra tio n se lv ag e St ag e’ M in er al s 2 Te xt ur e W id th (m m) D ip ’ M in er al s2 W id th (m m) O th er co m m en ts E qt z m as siv e 0. 5- 2 0- 30 n o n e n /a M qt z- py -c py ± m o ly ± sp h co m b- te xt ur ed ,b an de d 1- 40 10 -4 0 ep i-c hl w ho le ro ck ? su ip hi de so fte n in ce n tr e M ep i-p y- cp y * m o ly ± sp li ba nd ed to m as siv e 2- 30 10 -4 0 ep i-c hl w ho le ro ck ? L m o ly ± sp h ± py fra ct ur e co at in g 1- 3 al l n o n e n /a L py ±c py m as siv e 1- 10 20 -5 0 qt z- se r 2- 20 P ep i± ch l ba nd ed 1- 5 al l n o n e n /a P ze o m as siv e 0. 5- I al l n o n e n /a P an hy -g yp m as siv e 1- 10 0- 45 n o n e n /a E = ea rly -, M = m ai n, I= La te , p = po st -m in er al iz at io n- st ag e 2 an hy = an hy dr ite ,c hl = ch lo rit e, cp y = ch al co py rit e, ep i= ep id ot e, ks p = K -fe ld sp ar ,m ag = m ag ne tit e, m o ly = m o ly bd en ite , p y = py rit e, qt z = qu ar tz , sp h = sp ha le rit e, ze o = ze o lit e D ip an gl e in de gr ee sa n d re la tiv e to th e ho riz on ta l U I Table 2.8 Re-Os data for Fin Main-stage vein Sample no. Vein-type Re ± 2o ‘870s ± 2 Total common Model age ± 2G ppm abs ppb abs Os pg Ma absl F05-06-103 qtz-py-cpy±moly±sph 211 1.1 489.3 1.1 1.6 221 1.4 uncertainty included 57 VEIN TYPE EARLY MAIN LATE POST-MIN Qtz Qtz-py-cpy ± moly ± sph Epi-py-cpy ± moly ± sph Py ± cpy MoIy±sph±py Epidote ± chlorite Zeolite Anhydrite-gypsum Fig. 2.15 Vein paragenesis for the Fin prospect. The veins are divided into early-, main-, late- and post mineralization-stage veins based on their vein mineralizations, textures, and relative cross-cutting relationships. 58 Fi g. 2. 16 D ia m on d dr ill ho le da ta fo rs el ec te d ho le sl oc at ed in th e Fi n pr os pe ct sh ow th e di str ib ut io n o f l ith ol og ie s, v ei ns ,a n d m et al s ( Cu ,A u, A g, M o, Pb , Z n). D ia m on d dr ill ho le F0 5- 05 (64 02 90 E, 63 44 70 5N ,d ril le d a t - 55 °t ow ar ds 30 00 ) i s3 25 m lo ng ,F 05 -0 6 (64 03 00 E, 63 44 70 5N ,d ril le d at - 50 °t ow ar ds 12 00 ) i s2 09 m lo ng an d F0 5- 02 (64 06 00 E, 63 44 51 9N ,d ril le d at 75 ° to w ar ds 27 0° )i s 2 90 m lo ng .A ll dr ill ho le s sh ow a co rr el at io n be tw ee n m ai n- sta ge qu ar tz -p yr ite ch al co py rit e ± m o ly bd en ite ± sp ha le rit e an d ep id ot e- py rit e- ch al co py rit e± m o ly bd en ite ± sp ha le rit e v ei ns an d hi gh Cu ,A u, an d M o v al ue s. M on zo ni te dy ke s co n ta in lo w v al ue s o fC u, A u, an dM o. D ril lh ol eF 05 -0 6 sh ow sa cl os ec o rr el at io nb et w ee nM o an dC u v al ue s. F0 5- 05 F0 5- 06 F0 5- 02 Zn (pp m) M (pp m) (% ) (pp m) (pp m) Zn (pp m) LI TH O LO G IE S Fi n m o n z o gr an ite M on zo ni te dy ke Fa ul tz o n e V EI N S — — Qt z Qt z-p y- cp y± m o ly ± sp h Ep i-p y- cp y ± m o ly ± sp h M oI y± sp h± py P y± cp y Ep i-c hl — Z eo 50 m A nh y U i Fig. 2.17 Photoplate showing vein types in hand specimens from the Fin prospect. All veins are cutting monzogranite. A. Main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein in epidote-chlorite altered monzogranite. Pyrite and chalcopyrite occur in the centre of the vein, molybdenite and sphalerite occur as bands along vein margins. B. Main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein that contains pyrite and chalcopyrite in the centre of the vein. C. Epidote-pyrite-chalcopyrite ± molybdenite ± sphalerite vein with weak epidote-chlorite alteration. D. Late-stage pyrite ± chalcopyrite vein with quartz sericite alteration selvage in epidote-chlorite altered monzogranite. B. Late-stage molybdenite ± sphalerite ± pyrite vein with no significant alteration. F. Main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein that is cut by post-mineralization epidote ± chlorite veinlets. 60 Vein paragenesis and hypogene alteration styles at Mex Quartz-magnetite-pyrite ± chalcopyrite veins that are 3 to 10 mm wide cross-cut pervasively K feldspar-magnetite altered monzonite at Mex. These veins do not display discrete alteration selvages and it is unclear if these veins are responsible for the 50 m2 potassic alteration that crops out on the northeast slope of the Mex ridge (Fig. 2.14). Pervasive quartz-pyrite-sericite alteration outcrops over 500 m2 along Mex ridge (Fig. 2.14). The intense phyllic alteration replaces primary igneous textures in the monzonite. Disseminated pyrite is euhedral and extensive zones of brecciation containing clasts of phyllic-altered monzonite occur in the phyllic alteration zone. Chiorite-epidote altered monzonite crops out to the northwest of Mex ridge and proximal to the Mex fault (Fig. 2.14). Propylitic alteration is defined by epidote ± chlorite alteration of primary hornblende. Supergene alteration Supergene alteration affects all rock types in all areas within about 150 m of the present surface. Anhydrite associated with the post-mineralization veins is hydrated by supergene fluids to form gypsum. The volume increase caused by this supergene alteration results in a highly broken and permeable zone. Hematite occurs in fractures and pervasively alters rocks and existing hypogene alteration mineral assemblages. Locally, malachite occurs along joint planes in the potassic-altered Pine quartz monzonite and along joint planes cross-cutting Type 1 syenite dykes. Neotocite occurs along joint planes that cross cut Type 1 syenite dykes own to vertical depths of about 200 m. 61 FLUID INCLUSION MICROTHERMOMETRY Microthermometric studies were undertaken on main- and late-stage veins at Pine and Fin in an attempt to obtain and compare their respective pressure-temperature-compositional (P-T-X) conditions of ore deposition. Main-stage quartz-magnetite-chalcopyrite-pyrite and late-stage quartz-pyrite ± chalcopyrite veins were chosen because of their temporal and spatial association with Au and Cu (Fig. 2.12). The remaining vein types were unsuitable for fluid inclusion analysis because they either did not contain minerals suitable for fluid inclusion studies or were recrystallized. Sample selection Thirty-four doubly polished sections of veins were examined from the Pine and Fin prospects. Of these sections, five main-stage quartz-magnetite-chalcopyrite-pyrite and five late-stage quartz-pyrite ± chalcopyrite veins from the Pine area were used for microthermometry. These samples were collected from seven diamond drill holes located throughout the prospect, from depths of 28 to 251 m below surface. Only one main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein from Fin proved suitable for microthermometry, the other veins were intensely recrystallized. All fluid inclusions measured from Fin veins decrepitated at low temperatures (<250 °C); these veins are not further discussed. 62 Fluid inclusion petrography Quartz grains in the main-stage and late-stage veins from Pine are mostly fine-grained (< 0.5 mm) and display undulose extinction, indicating that they are deformed (Fig. 2.1 8A). Only rarely are quartz grains with straight extinction preserved. In these exceptional cases, the quartz grains are partially or completely surrounded by chalcopyrite or pyrite and are interpreted to host fluid inclusions that are representative of the ore fluid that was responsible for Au and Cu deposition. Within these less-deformed quartz grains, fluid inclusions that are 2 to 54 jim long and randomly distributed in groups or trails that terminate before grain boundaries are interpreted to be primary or pseudosecondary inclusions (Figs. 2.1 8B-D). The absence of mineral growth zonation in quartz prevents positive identification of primary inclusions. Fluid inclusions that are smaller (< 1 jim long) and cross-cut grain boundaries are interpreted to be secondary inclusions (Fig. 2.1 8B). The primary or pseudosecondary fluid inclusions are divided into three groups based on the phases present at 25 °C (Table 2.9). Type 1 inclusions are liquid-rich and have a liquid to vapour volume ratio of more than 9:1 (Fig. 2.1 8E), Type 2 inclusions are ‘mixed’ in that they contain between 10 and 90 vol. % liquid (Fig. 2.1 8F), whereas Type 3 inclusions are vapour-rich and have a liquid to vapour ratio of less than 1:9 (Fig. 2.1 8F). The mixed inclusions give meaninglessly high temperatures, therefore only liquid- and vapour-rich inclusions were used in this study (cf. Bodnar and Vityk 1994). Liquid-rich inclusions may contain two to seven phases. The inclusions are divided into two halite-undersaturated and three halite-saturated subgroups based on the presence of various solid phases, such as halite crystals and opaque minerals (pyrite or chalcopyrite) (Table 2.9). For main- and late-stage veins at Pine, liquid-rich halite-undersaturated inclusions are 2.5 to 17 jim long (mean = 6.9 ± 4.3 jim, n = 22) and 2 to 10 jim wide (mean 3.5 ± 1.9 jim, n = 22) (subtypes A and B, Table 2.9). Liquid-rich halite-saturated inclusions are 4 to 54 jim long (mean = 9.5 ± 6.5 jim, n = 111) and 2 to 40 jim wide (mean = 5.8 ± 4.3 jim, n = 111) (subtypes c to e, Table 1). Vapour-rich inclusions are slightly larger, and are 5 to 25 jim long (mean = 10.3 ± 4.0 jim, n = 40) and 2 to 15 jim wide (mean 6.4 ± 2.4 jim, n = 40). Where present, halite crystals constitute between 20 and 25 vol. % of the inclusion, whereas chalcopyrite or pyrite represent only about 6 vol. % of the inclusion (Table 2.8). 63 Ta bl e 2. 9 Fl ui d in cl us io n gr ou ps , s iz es an d ph as e pr op or tio ns fro m M ai n- an d La te -s ta ge v ei ns at th e Pi ne pr os pe ct G ro up Ty pe 1h al ite -u nd er sa tu ra te d Ty pe 1h al ite -s at ur at ed Ty pe 3 Su bg ro up A B al l C D E al l F Ph as es pr es en t L+ V L+ V +o ’ — L+ V +h ’ L+ V +h +o 1 L+ V +h +h e* o’ — V +L 2 Le ng th Mm 2. 5 2. 5 2. 5 4 4 5 4 5 (jim ) M ax 17 .5 16 17 .5 22 54 19 54 25 M ea n 7 7 7 9 10 9 10 10 ± 4 5 4 4 8 4 6 4 W id th Mm 2 2 2 2 2. 5 3 2 2 (jim ) M ax 10 6.5 10 15 40 9 40 15 M ea n 3 4 4 5 6 5 6 6 & 2 2 2 3 6 2 4 2 Li qu id M m 90 70 70 40 35 35 35 5 (vo l. % ) M ax 95 93 95 90 90 90 90 10 M ea n 91 87 90 73 70 68 71 9 ± 2 9 5 11 12 13 12 2 V ap ou r Mm 5 5 5 2 2 5 2 85 (v ol. % ) M ax 10 10 10 20 20 25 25 95 M ea n 9 7 8 7 6 7 7 90 ± 2 7 2 11 3 5 12 2 So lid Mm 0 2 0 2 5 5 2 0 (vo l. % ) M ax 0 20 20 50 60 60 60 5 M ea n 0 7 2 21 23 25 23 0 ± 0 7 4 11 12 12 12 1 L = liq ui d, V = v ap ou r, h = ha lit e, o = o pa qu e m in er al , h e = he m at ite , S = so lid ‘ ± ca lc ite ± an hy dr ite 2 ± ca lc ite ± o pa qu e m in er al a’ Fig. 2.18: Photomicrographs of fluid inclusions as seen in plane polarized transmitted light, except for photomicrograph A, which is taken in cross-polarized transmitted light. A. Quartz grain labeled A.3 has straight extinction and is partially surrounded by pyrite. B. Close up of a quartz grain (surrounded by pyrite and magnetite) with large, isolated primary or pseudosecondary inclusions and secondary trails of smaller inclusions. C. Co-existing liquid- and vapour-rich two-phase inclusions in quartz. D. Primary liquid-rich inclusions with various solids that include chalcopyrite. B. Common four-phase liquid-rich inclusions (liquid-vapour-halite-opaque mineral). F. Co-genetic vapour-rich and mixed inclusions. traiI: I I -Vapour-rich hali% satu ih * : : bVapoiiS I 65 Microthermometric results Freezing data Main-stage and late-stage primary and pseudosecondary inclusions were cooled to -110 °C and then heated slowly to check for the presence of CO2 and to observe the eutectic temperature (Te) and the temperature of final ice melting Tm). Freezing data were collected from 17 main-stage and 26 late-stage pseudosecondary fluid inclusions (Table 2.10). Main-stage veins Measurements of Te in liquid-rich fluid inclusions range from -49.8 to -29.0 °C (n = 10) in main- stage veins. These Te measurements are below the eutectic temperature for theH20-NaC1 system (-20.8 °C), indicating the presence of cations other than Na, such as Ca, Mg, and Fe (Shepherd et al. 1985). Measurements of Te for vapour-rich fluid inclusions range from -55.0 to -23.0 °C (n = 7). The Te range for vapour-rich inclusions suggests the additional presence of Ca, Mg, and Fe, in addition to Na. Measurements of Tm(jce) in liquid-rich, halite-undersaturated inclusions range from -16.0 to -3.1 °C (mean = -9.1 ± 5.0 °C, n = 5). These measurements correspond to salinities between 5.0 and 19.4 wt. % NaCl equiv. (mean = 12.3 ± 5.6 wt. % NaCl equiv., n 5) for main-stage liquid-rich inclusions. Measurements of Tm(,ce) for vapour-rich inclusions range from -12.5 to -3.9 °C (mean -6.7 ± 3.4 °C, n = 5). These measurements correspond to salinities between 6.2 and 16.4 wt. % NaC1 equiv. (mean = 9.8 ± 4.0 wt. % NaC1 equiv., n = 5). Late-stage veins Measurements of Te in liquid-rich fluid inclusions range from -51.0 to -24.0 °C (n = 9) in late stage veins. These Te measurements are below the eutectic temperature for theH20-NaC1 system (-20.8 °C), indicating the presence of cations other than Na, such as Ca, Mg, and Fe (Shepherd et al. 1985). Measurements of Te for vapour-rich fluid inclusions range from -29.0 to -20.3 °C (n = 17). The lower Te range for vapour-rich fluid inclusions in late-stage veins suggests the presence of Na and K in the inclusions. 66 Table 2.10 Fluid inclusion microthennometry results from main- and late-stage veins at the Pine deposit Main-stage Late-stage Type 1 h-u Type 1 h-s Type 3 Type 1 h-u Type 1 h-s Type 3 Te Mm -49.8 -47.0 -55.0 -25.0 -51.0 -29.0 Max -39.0 -29.0 -23.0 -24.0 -25.0 -20.3 TmO) Mm -16.0 -8.8 -12.5 -4.4 -14.0 -13.1 Max -6.0 -3.1 -3.9 -1.6 0.0 -0.6 Mean -11.2 -6.0 -6.7 -2.9 -7.0 -5.7 ± 5.0 4.0 3.4 1.5 9.9 4.0 Wt. % NaCI equiv. Mm 9.2 6.2 2.6 1.0 Max 19.4 16.4 7.0 17.0 Mean 14.7 9.8 4.7 8.3 & 5.2 4.0 2.4 5.1 n 3 7 7 5 4 17 Abbreviations: h-u = halite-undersaturated; h-s = halite-saturated; Te = eutectic temperature; Tm(I) temperature of final ice melt; Wt. % NaC1 equiv. = weight % NaC1 equivalent, n = number ± calculated as the standard deviation from the mean 67 Measurements Of Tnice) in liquid-rich, halite-undersaturated inclusions range from -14.0 to 0.0 °C (mean = -4.3 ± 5.0 °C, n = 6) for late-stage veins. These measurements correspond to salinities between 0.0 and 17.8 wt. % NaC1 equiv. (mean = 6.1 ± 6.3 wt. % NaC1 equiv., n = 6) for late-stage liquid-rich inclusions. Measurements of Tm(ice) for vapour-rich inclusions range from -13.1 to -0.6 °C (mean -5.7 ± 4.0 °C, n = 15) for late-stage veins. These measurements correspond to salinities between 1.0 and 17.0 wt. % NaC1 equiv. (mean = 8.3 ± 5.1 wt. % NaCl equiv., n = 15) for late-stage vapour-rich inclusions. Heating data Main-stage and late-stage primary and pseudosecondary inclusions were heated to a maximum of 550 °C or until all fluid inclusions had decrepitated. Most heating data display a non-normal (i.e., skewed) distribution, therefore medians and quartile deviations are used, unless otherwise stated. Main-stage veins In halite-undersaturated liquid-rich inclusions, vapour to liquid homogenization temperatures range from 95 to 246 °C (163 ± 58 °C, n = 6). In these inclusions vapour to liquid homogenization ranges from 48 to 354 °C (median = 114 ± 51.8 °C, n 39) and halite dissolution occurs from 284 to 519 °C (median = 432 ± 25 °C, n = 39) (Fig. 2.19A). These temperatures correspond to salinities of 37.0 to 62.4 wt. % NaCl equiv. (median = 51.1 ± 3.0 wt. % NaC1 equiv., n = 39) (Fig. 2.20A). Homogenization in vapour-rich fluid inclusions occurs from 350 to 528 °C (median = 398 ± 22 °C, n = 9). Late-stage veins In halite-undersaturated liquid-rich inclusions, vapour to liquid homogenization temperatures range from 98.8 to 327 °C (mean = 206.2 ± 83.6 °C, n = 11). In these inclusions temperatures of vapour to liquid homogenization range from 52 to 308.6°C (median = 194.6 ± 49.4 °C, n = 22) halite dissolution occurs from 230 to 550 °C (median = 420 ± 84 °C, n = 22) (Fig. 2.1 9B). These temperatures correspond to salinities of 33.5 to 63.8 wt. % NaCl equiv. (median = 50 ± 9.1 wt. % NaC1 equiv., n = 22) for late stage veins (Fig. 2.20B). Homogenization in vapour-rich fluid inclusions occurs from 321 to 475 °C (mean=366±33°C, n 22). 68 Th(L)(°C) Th(L)(°C) Fig. 2.19 Total homogenization by vapour to liquid versus total homogenization by halite dissolution diagrams. Different shaded squares correspond to different quartz grain hosts to inclusions. A. Main-stage veins: all primary or pseudosecondary liquid-rich halite-saturated fluid inclusions homogenize by halite dissolution between 400 and 500 °C. The inclusions have corresponding vapour to liquid homogenization temperatures of 50 to 300 °C. B. Late-stage veins: all samples, with the exception of one, homogenize by halite dissolution between 220 to 530 °C. Corresponding vapour to liquid homogenization temperatures range from --5O to 220 °C. 0 300 -c I— 0 100 200 300 400 500 600 0 I I I I 100 200 300 400 500 600 69 L 60 .11 60 ç50 5O — 0 a 4O — Z o40 I z — .30 • 4. a20 • 20 10 • 0 0 • 10 0 •e • 0 100 200 300 400 500 600 0 I Total homogenIzation 0 100 200 300 400 500 600 (by ThILV) (vapour-rich) or Th11 (liquid-rich)) Total homogenization (by Th(L-V) (vapour-rich) orTh(hat) (Gquid-dch)) Fig. 2.20: Total homogenization versus salinity diagrams for primary or pseudosecondary fluid inclusions for main- and late-stage veins at Pine. A. Main-stage veins: liquid-rich halite-saturated inclusions dominantly have total homogenization temperatures between 400 and 520 °C, corresponding to salinities of’—’45-60 wt % NaC1 equivalent. Co-existing vapour-rich inclusions have similar total homogenization temperatures. B. Late-stage veins: liquid-rich halite-saturated inclusions have a wide range of total homogenization temperatures, between 280 and 550 °C, corresponding to salinities of —35-65 wt % NaC1 equivalent. Co existing vapour-rich inclusions have total homogenization temperatures ofbetween 300 and 400 °C. 70 Interpretation of microthermometric data Calculated isochores are shown for primary and pseudosecondary liquid-rich halite-saturated fluid inclusions in main- and late-stage veins in Figure 2.21. Only one isochore of 39 calculated for main- stage veins meets the requirement for realistic pressure estimation, which requires that the salinity calculated from the halite dissolution temperature must correspond to the appropriate halite liquidi (Bodnar 2003). This sole isochore gives a concordant pressure estimate of 1.5 kbar. By similar reasoning, only three of 22 calculated isochores are feasible for late-stage veins; these pressure estimates range from 1.5 to 1.7 kb. The concordant main-stage inclusion has a Th of 415 °C and a salinity of 49 wt. % NaC1 equiv. The three concordant late-stage inclusions have a Th between 230 and 340 °C and salinities between 33.5 and 41.5 wt. % NaC1 equiv. However, one of the inclusions decrepitated shortly after Th at 230 °C, therefore this estimate may be unreliable. Although the primary and pseudosecondary fluid inclusions measured in main- and late-stage veins do not show visible signs of leakage (e.g., halos of very fine-grained inclusions, Bodnar 2003), the calculated isochores and the wide range in liquid to vapour homogenization temperatures for liquid-rich fluid inclusions suggest that these fluid inclusions were affected by events after trapping. In addition, the high volume percent of daughter crystals (i.e., halite) and their high corresponding salinities, suggest that the inclusions might have become more saline due to the loss of water during fluid inclusion leakage. The recrystallized texture of most early- and late-stage quartz veins at Pine (as well as at Fin) and the high concentration of post-mineralization intermediate-felsic dykes, suggests that they were modified after their formation. The absence of a foliation in the rocks or other large-scale structures suggest that stress was not as important as thermal effects resulting from wide-spread contact metamorphism during the emplacement of the Pine granodiorite and numerous mafic to felsic dykes in the Pine-Fin-Mex area. Hence, most of the primary and pseudosecondary fluid inclusion data presented are deemed to be unreliable due to thermal modification of the inclusions after trapping. Estimated salinities of liquid-rich halite-saturated inclusions are probably over-estimates. 71 2. 1.5 (‘5 1 0 D U) U) ci) I 0 0.5 0. Fig. 2.21: Concordant isochores calculated using MacFlinCor software (Brown and Hagemann 1994). Sub- vertical lines represent the halite liquidi for different wt. % NaC1 (Bodnar, 1994). Names plotted next to isochores refer to the inclusion number. C.3 .8 is from a main-stage vein and the other three are from late-stage veins. The four isochores plotted are the only ones from the 61 successful analyses of liquid-rich halite- saturated inclusions from main- and late-stage veins that give corresonding temperature and pressure estimates between the salinity calculated from halite dissolution temperature and calculated halite liquidi. These four isochores give a minimum pressure estimate of between 1.5 and 1.7 kbar and a minimum temperature estimate ofbetween 200 and 400 °C. Temperature (°C) 72 ISOTOPES Sulphur isotope data Sulphur isotope values were obtained for 20 samples of pyrite and 13 samples of anhydrite from late-stage and post-mineralization veins from the Pine deposit. Four pyrite, two chalcopyrite and one sphalerite separates were analyzed from main-stage veins from the Fin prospect. Due to the low abundance and fine-grained nature of sulphide minerals in main-stage veins, and the tendency in zones of main-stage veining for multiple vein events, sulphides from main-stage quartz-magnetite-chalcopyrite pyrite veins could not be separated. The ö34S values of Pine late-stage pyrite separates are between -1.86 and 1.67% (average = 0.08 ± 1.10%, n = 17), late-stage anhydrite samples are between 12.91 and 1 6.09% (average = 14.08 ± 1.l2%o, n = 7) and post-mineralization anhydrite samples are between 13.25 to 19.89% (average = 14.74 ± 2.45%, n = 6) (Table 2.11 and Fig. 2.22). The 834S values of pyrite from Fin are between 0.37 and 5.84%o (average = 2.82 ± 2.64%o, n = 5), chalcopyrite samples are between -0.43 and 1.1 %o (average = 0.34 ± 1 .O8%o, n = 2) and the sphalerite sample is -O.94%o (Table 2.11 and Fig. 2.22). Suiphide-suiphate isotope geothermometry Anhydrite-pyrite (Ohmoto and Lasaga 1982), pyrite-chalcopyrite (Ohmoto and Rye 1979), and pyrite-sphalerite equilibrium (Ohmoto and Rye 1979) mineral pairs are used to estimate fluid temperatures during the deposition of these minerals. The uncertainty in the calculated temperature is about ± 25 at 450 °C and ± 10 at 300 °C for the anhydrite-pyrite geothermometer, about ± 35 at 300 °C for the pyrite-chalcopyrite geothermometer, and about ± 40 at 300 °C for the pyrite-sphalerite geothermometer (Ohmoto and Rye 1979). Anhydrite-pyrite pair temperatures calculated for late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite veins from the potassic-altered centre of the Pine deposit range from 410 to 460 °C (average = 453 ± 11 °C, n = 2). Where this vein type occurs in peripheral phyllic-altered areas of Pine, they define a lower temperature range of 360 to 405 °C (average = 378 ± 17 °C, n 5) (Table 2.11). 73 Table 2.11634ScDT values and geothermometric data for pyrite, chalcopyrite, sphalerite, and anhydrite, from the Pine and Fin prospects 634S (%) (V-CDT) Anhy-pyrite therm Py-cpy therm Py-sph therm . . 34 0 34 0 34Sample no. Location Vein-type py cpy sph anhy SY T ( C) A T ( C) A T C) 92-38 185.5 Pine CZ A-P-S-C 0.88 13.22 12.35 445 92-38 185.6 PineCZ A-P-S-C 1.59 13.49 11.90 460 97-02 107 PinePZ A-P-S-C -1.86 12.91 14.77 380 97-02 154 PinePZ A-P-S-C 0.22 13.93 13.71 405 97-09 107 PinePZ A-P-S-C 0.99 16.09 15.10 375 97-10 61 PinePZ A-P-S-C -0.52 15.08 15.60 360 97-1064 PinePZ A-P-S-C -1.35 13.84 15.19 370 98-04 151 Pine CZ A-P-S-C -0.34 92-38 147 Pine CZ Q-P-C 1.67 97-0470 Pine CZ Q-P-C -1.07 98-04 148 Pine CZ Q-P-C -0.64 98-04 181 Pine CZ Q-P-C -0.74 98-04 275 Pine CZ Q-P-C -0.98 92-38 147 Pine CZ P-C 0.86 92-40 165b Pine CZ P-C 0.42 97-08 170 Pine CZ P-C 1.25 97-09 111 Pine PZ P-C 1.02 92-38 132 Pine CZ dissem -1.46 92-39 150 Pine CZ breccia 0.26 97-08 212a Pine CZ breccia 1.23 92-38 140 Pine CZ A 13.5 92-40 165a PineCZ A 13.3 97-02 108 Pine PZ A 19.665 97-08212b PineCZ A 13.71 97-08 212c Pine CZ A 13.685 97-1069 Pine PZ A 14.58 F05-05 284 Fin Q-P-C-M-S 2.04 1.1 -0.94 0.94 420 2.98 45 F05-06 71.5 Fin Q-P-C-M-S 0.37 F05-06 71.5* Fin Q-P-C-M-S 5.84 F05-06 80.1 Fin Q-P-C-M-S 5.38 F05-06 119.8 Fin Q-P-C-M-S 0.47 -0.43 0.90 435 Note sample numbers coffespond to drill-hole numbers and the sample depth was taken from (i.e., 92-38 132 is from diamond drill hole 92-3 8 and from a sample depth of 132 m). * Quality control duplicate with significant difference. Vein abbreviations: A = anhydrite vein, A-P-S-C = anhyrite-pyrite ± specular hematite ± chalcopyrite vein, P-C = stringer vein, Q-P-C = quartz-py ± chalcopyrite vein, breccia = as fill between clasts in hydrothermal breccia, dissem = sulphide disseminated throughout groundmass. Mineral abbreviations: anhy = anhydrite, cpy = chalcopyrite, py = pyrite, sph = sphalerite. 74 12 Typical porphyry suiphides pyrite (A-P-S-C veins) anhyhrite (A-P-S-C veins) pyrite (Q-P-C veins) . 1111111 pyrite (P-C veins) pyrite (disseminated) : pyrite (breccia fill) anhydrite (A veins) pyrite (Q-P-C-M-S veins) chalcopyrite (Q-P-C-M-S veins) E. sphalerite(Q-P-C-M-S veins) 0 5 6S (%0CDT) Fig. 2.22 Sulphur isotope data for pyrite and anhydrite from the Pine prospect and sulphide minerals from the Fin prospect. Typical ranges are also shown for porphyry Au-Cu suiphides and suiphates (Rollinson 1993) and Jurassic seawater sulphur (Claypool et al. 1980). Vein abbreviations: A anhydrite vein, A-P-S-C = anhyrite-pyrite ± specular hematite ± chalcopyrite vein, P-C = pyrite ± chalcopyrite vein, Q-P-C = quartz pyrite ± chalcopyrite vein, Q-P-C-M-S quartz-pyrite-chalcopyrite ± molybdendite ± sphalerite vein, breccia = matrix from hydrothermal breccia, dissem = disseminated sulphide. a) E6 z 75 Pyrite-chalcopyrite suiphide pairs for main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite veins from the centre of the Fin prospect give temperature estimates of 420 and 435 °C (average = 428 ± 11 °C, n = 2) (Table 2.11). The pyrite-sphalerite sulphur isotope pair geothermometer for a main- stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite vein from the Fin prospect gives a temperature estimate of only 45 °C. This temperature is very low and it is interpreted that pyrite and sphalerite were not in isotopic equilibrium. Oxygen and carbon isotope data Calcite from post-mineralization veins that cross-cut syenite dykes located in the potassic-altered centre of the Pine deposit have 6’O values that range from 6.8%o to 9.O%o (average = -4.01 ± 0.93%, n = 4), while 613C values range from -2.7%o to -4.9% (average = 7.78 ± 0.98%o, n 4) (Table 2.12). Lead isotope data Lead isotopic compositions detennined for pyrite from late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite, quartz-pyrite ± chalcopyrite, and pyrite ± chalcopyrite veins from the Pine deposit show a narrow range. For all late-stage vein samples, 206Pb/4 ratios range from 18.73 to 18.86 (mean 18.80 ± 0.04, n = 14). Ratios of207Pb/°4range from 15.54 to 15.65 (mean = 15.59 ± 0.04, n 14) and 208Pb/4 ratios range from 38.21 to 38.55 (mean = 38.36 ± 0.12, n 14) (Table 2.13). Lead isotopic compositions determined for pyrite and chalcopyrite from main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite veins from the Fin prospect also show a narrow range. Ratios of 206Pb/4 range from 18.79 to 18.82 (mean = 18.81 ± 0.01, n 5). Ratios of207Pb/4 range from 15.58 to 15.62 (mean = 15.60 ± 0.02, n = 5) and 208Pb/4 ratios range from 38.31 to 38.52 (mean = 38.39 ± 0.08, n = 5) (Table 2.13). Lead isotopic compositions of feldspars in Triassic to Jurassic intrusions and volcanic rocks in the Pine-Fin-Mex area also show little variation; 206Pb/4 ratios range from 18.83 to 19.08 (average = 18.97 ± 0.08, n = 13), 207Pb/4 ratios range from 15.56 to 15.70 (average = 15.61 ± 0.04, ii = 3), while 208pb/4 ratios range from 38.33 to 38.59 (average = 38.43 ± 0.09, n 13). The dykes that cross-cut the 76 Table 2.12 Carbon and oxygen isotope data for post-mineralization calcite veins at the Pine deposit o’3c%o Sample no. Vein-type Host rock (V-PDB) (V-SMOW) JD98-3 43.5 calcite Rhyolite dyke -2,68 7.98 JD98-4 255.6 calcite Rhyolite dyke -4.86 9.06 JD98-4 271.7 calcite Rhyolite dyke -4.35 7.24 JD98-4 270.4 calcite Type 2 syenite dyke -4.13 6.83 77 Table 2.13 Pb isotope compositions of sulphide and feldspar samples from the Pine-Fin-Mex aret Results have been normalized using a fractionation factor of 0.12% based on multiple analyses of NBS98 1 standard lead, and the values in Thirlwall., cal = calcite, ksp K feldspar, py pyrite, plag plagioclase 2 = anhyrite-pyrite ± specular hematite ± chalcopyrite vein, breccia = as fill between clasts in hydrothermal breccia, dissem sulphide disseminated throughout groundmass, P-C = pyrite & chalcopyrite stringer vein, Q-P-C = quartz-pyrite ± chalcopyrite vein Asitka Asitka Group limestone, Dun and = Duncan Member andesite, Dun gd = Duncan pluton granodiorite, Dun qm Duncan quartz monzonite, Fin mzg = Fin pluton monzogrsnite, Gieg gd Giegerich pluton granodiorite, Mex monz = Mex stock monzonite, Pine gd = Pine stock granodiorite, Pine qm Pine stock quartz monzonite, Rhyolite = rhyolite dyke, Saunders Saunders Member dacite, TI syenite = Type 1 syenite dyke, Takla * Data from McKinley (2006) and Duuring writ.comm. (2006) Sample Mm1 Sample 206Pb/4 2o 207Pb/’°4b 2a 208Pb/’54b 2a ‘°7Pb/206 2 ‘°5Pb/20’Pb 2s type2 (%) (%) (%) (%) (%) Sulphides samples from veim 92-38 185.5 py A-P-S-C 18.802 0.11 15.589 0.15 38.310 0.21 0.829 0.05 2.038 0.11 92-38 185.6 py A-P-S-C 18.830 0.12 15.612 0.16 38.503 0.22 0.829 0.06 2.045 0.11 97-02 107 py A-P-S-C 18.856 0.10 15.653 0.15 38.536 0.21 0.830 0.05 2.044 0.10 97-02 154 py A-P-S-C 18.828 0.12 15.617 0.16 38.420 0.21 0.829 0.05 2.041 0.10 97-09 107 py A-P-S-C 18.785 0.10 15.581 0.15 38.322 0.21 0.829 0.05 2.040 0.10 97-10 64 py A-P-S-C 18.863 0.31 15.639 0.33 38.552 0.36 0.829 0.06 2.044 0.10 98-04 151 py A-P-S-C 18.760 0.108 15.593 0.157 38.351 0.208 0.831 0.053 2.044 0.103 93-38 147 py Q-P-C 18.750 0.12 15.557 0.16 38.223 0.21 0.830 0.05 2.039 0.10 98-04 148 py Q-P-C 18.802 0.11 15.597 0.16 38.342 0.21 0.830 0.05 2.039 0.10 98-04 181 py Q-P-C 18.787 0.11 15.565 0.16 38.300 0.21 0.828 0.05 2.039 0.10 98-04 275 py Q-P-C 18.728 0.12 15.537 0.16 38.247 0.21 0.830 0.05 2.042 0.10 92-40 165 py P-C 18.813 0.18 15.547 0.19 38.269 0.26 0.826 0.10 2.034 0.13 97-09 111 py P-C 18.760 0.10 15.546 0.15 38.209 0.20 0.829 0.05 2.037 0.10 97-08 212 py breccia 18.808 0.11 15.635 0.16 38.481 0.21 0.831 0.05 2.046 0.10 F05-05-284c cpy Q-P-C-M-S 18.792 0.010 0.055 15.581 0.012 0.078 38.309 0.040 0.105 0.829 F05-06-7l.5 py Q-P-C-M-S 18.818 0.010 0.056 15.617 0.012 0.078 38.404 0.041 0.106 0.830 F05-06-l19.8p py Q-P-C-M-S 18.813 0.020 0.105 15.583 0.014 0.092 38.378 0.055 0.143 0.828 F05-06-80.1 py Q-P-C-M-S 18.803 0.011 0.056 15.599 0.012 0.078 38.352 0.041 0.106 0.830 F05-05-284p py Q-P-C-M-S 18.801 0.014 0.074 15.623 0.015 0.093 38.524 0.046 0.119 0.831 Feldspar and calcite samples from rocki JDO5-2l7 plag Dun and 18.837 0.10 15.582 0.14 38.307 0.19 0.827 0.06 2.034 0.10 JDO5-279 plag Dunand 18.959 0.09 15.587 0.14 38.367 0.19 0.822 0.05 2.024 0.10 JD05-l65 plag Dungd 18.914 0.09 15.571 0.10 38.263 0.12 0.823 0.03 2.023 0.05 F05-06 189 plag Finmzg 19.018 0.05 15.563 0.08 38.388 0.10 0.818 0.03 2.019 0.05 F05-06 195.5 plag Fin mzg 18.993 0.09 15.610 0.14 38.404 0.18 0.822 0.05 2.022 0.09 JDO5-281 plag Fin mzg 18.930 0.10 15.593 0.14 38.395 0.19 0.824 0.05 2.028 0.09 JDO5-ll7a plag Gieggd 18.990 0.09 15.620 0.11 38.419 0.13 0.823 0.03 2.023 0.05 JDO5-ll7b plag Gieggd 18.969 0.09 15.597 0.14 38.397 0.18 0.822 0.05 2.024 0.09 JD05-l34 ksp Mexmonz 19.067 0.11 15.617 0.15 38.500 0.19 0.819 0.05 2.019 0.09 JDO5-134 plag Mex monz 19.079 0.09 15.598 0.14 38.477 0.18 0.818 0.05 2.017 0.09 97-10 plag Pinegd 18.827 0.06 15.576 0.08 38.327 0.11 0.827 0.03 2.036 0.05 97-12 309.6 plag Pine gd 18.908 0.10 15.612 0.14 38.493 0.18 0.826 0.05 2.036 0.09 92-38 162.5 ksp Pineqm 19.006 0.19 15.702 0.20 38.590 0.21 0.826 0.03 2.030 0.06 JDO5-280 plag Pineqm 19.014 0.11 15.681 0.10 38.573 0.15 0.825 0.07 2.029 0.08 98-02 42.4 plag TI syenite 19.419 0.09 15.624 0.14 38.756 0.18 0.805 0.05 1.996 0.09 JDO5-282 plag Ti syenite 19.041 0.34 15.529 0.33 38.548 0.40 0.816 0.15 2.025 0.15 98-03 44.8 ksp Rhyolite 19.923 0.09 15.652 0.14 39.037 0.18 0.786 0.05 1.959 0.09 JDO5-283 ksp Rhyolite 26.494 0.147 16.029 0.168 43.271 0.195 0.605 0.040 1.633 0.073 BL-1-K ksp Dunqm 18.882 0.018 0.09 15.598 0.021 0.14 38.389 0.070 0.18 0.826 BL-1-P5 plag Dunqm 18.877 0.017 0.09 15.602 0.021 0.14 38.415 0.0697 0.18 0.827 BL2_K* ksp Dunqm 18.856 0.017 0.09 15.604 0.021 0.14 38.382 0.070 0.18 0.828 BL_2_P* plag Dunqm 18.842 0.017 0.09 15.601 0.021 0.14 38.374 0.070 0.18 0.828 BL_2P* plag Dunqm 18.842 0.017 0.09 15.601 0.021 0.14 38.374 0.070 0.18 0.828 SF_4P* plag Takla 18.891 0.018 0.09 15.579 0.021 0.14 38.377 0.070 0.18 0.825 BFP-l p1* plag TalcIa 18.793 0.011 0.06 15.575 0.013 0.08 38.249 0.040 0.11 0.829 BFP2* ksp Takla 18.743 0.017 0.09 15.579 0.021 0.14 38.242 0.069 0.18 0.831 TN01P* plag Saunders 18.835 0.019 0.10 15.602 0.022 0.14 38.365 0.071 0.19 0.828 SF1P* plag Metsantan 18.908 0.017 0.09 15.579 0.021 0.14 38.391 0.070 0.18 0.824 5F-4P plag Metsantan 18.891 0.018 0.09 15.579 0.021 0.14 38.377 0.070 0.18 0.825 KE04.0238l.5c* carb Asitka 19.947 0.011 0.05 15.662 0.012 0.08 38.419 0.040 0.l1 0.785 KE0402382.0c* carb Asitka 20.221 0.015 0.07 15.652 0.013 0.08 38.425 0.046 0.12 0.774 78 Pine-Fin-Mex area show more variation in lead isotope ratios and they are mostly more radiogenic other intrusive rocks. For all dykes (with the exception of the Rhyolite dykes), 206Pb/4 ratios range from 19.04 to 26.5, 207Pb/4 ratios range from 15.53 to 16.30, and 208Pb/4 ratios range from 38.55 to 43.27. Rhyolite dykes have lead isotope ratios that are significantly more radiogenic than all other data ranges. 79 DISCUSSION Constraints on magmatism in the Pine-Fin-Mex area This section summarizes the nature and conditions of magmatism in the Pine-Fin-Mex area. The interpretations are based on a combination of data from whole-rock geochemistry, petrology, and Al-in homblende geobarometry. Geochemistry ofintrusive and extrusive rocks All Late Triassic-Early Jurassic rocks are high-K caic-alkaline series and display a strong enrichment of LILE elements (Rb, Ba, K, and Sr), a depletion of the HFSE (Nb, Zr, and Ti), and LREE enrichment. The Duncan and Giegerich plutons are metaluminous but Pine granodiorite, and Mex monzonite are weakly peraluminous. These geochemical features are consistent with partial melting of the mantle wedge containing amphibole which was metasomatized by the dehydration of a subducting slab (e.g., McKenzie and O’Nions 1991). The weakly peraluminous nature of the Mex and Pine granodiorite stocks indicates some crustal contamination. The (LaN/YbN)/CeN ratios of Black Lake suite intrusions indicates increasing fractionation with time. The Fin monzogranite is caic-alkaline, peraluminous and more fractionated than the younger magmatic rocks but otherwise displays similar trace element patterns. Oxidation state ofthe magmas The presence of the equilibrium mineral assemblage quartz-magnetite-titanite in least altered samples of the Fin monzogranite, Giegerich granodiorite and Mex monzonite places the magmas above the quartz-magnetite-fayalite (QMF) buffer, indicating that they are oxidized. The Pine quartz monzonite is too altered to determine the primary mineral assemblage. The oxidation state is important in the genesis of porphyry Au-Cu deposits because at oxidation states of less than QFM +2 the crystallization of pyrrhotite and the partioning of chalcophile elements into the crystallized residue results in less Cu and Au being available for the magmatic-hydrothermal fluid (Candela and Piccoli 2005). 80 Depth ofemplacement The application of the Johnson and Rutherford (1989) Al-in-hornblende geobarometer to suitable igneous rocks from the Pine-Fin-Mex area indicates pressures of about 2.4 kbar for the Giegerich granodiorite pluton, 1.5 kbar for the Mex monzonite stock, 1.6 kbar for Type 1 syenite dykes emplaced in the Pine area, 2.3 kbar for Type 1 syenite dykes emplaced in the Fin area, and 2.6 kbar for Type 1 syenite dykes emplaced in the Mex area (Table 4). Assuming a lithostatic load and a pressure gradient of 3.3 kmJl kbar (cf. Hagemann and Brown 1996), these pressures correspond to crystallization depths of about 8 km for the Giegerich granodiorite and 5 km for the Mex monzonite. For the porphyritic dykes, these pressures correspond to magma reservoir depths of between 5 and 8 km. The absolute error associated with the Al-in-homblende geobarometer is around 0.5 kbar (Johnson and Rutherford 1989) which corresponds to around 1.7 km. Therefore the calculated depths are subject to significant errors but do show relative relationships, particularly when the relative, 2a is considered. The Mex stock and Giegerich pluton are currently juxtaposed at surface across the Mex fault. The relative depths of emplacement calculated by Al-in-homblende geobarometry for the Giegerich granodiorite and Mex monzonite correspond with normal movement along the Mex fault prior to the emplacement of the syenite dykes. The Al-in-homblende study indicates that the depth of the magma reservoir for the porphyritic Type 1 syenite dykes is of comparable depth to the crystallization depths of the Mex stock and Giegerich pluton. The porphyritic textures of these dykes indicate that they formed by rapid, shallow-level emplacement away from this magma chamber where homblende crystallization occurred. Therefore it can be inferred that erosion took place between the emplacement of the Giegerich pluton at Ca. 198 Ma and the emplacement of the syenite dykes at Ca. 194 Ma. Physical and chemical properties of the Pine ore fluid This section describes the physical and chemical conditions of the hydrothermal fluid responsible for Au-Cu-Mo mineralization at Pine deposit. The interpretations are based on a combination of data from alteration equilibrium mineral assemblages, fluid inclusion microthermometry data, and S and Pb isotope data for main- and late-stage veins at Pine. 81 Temperature constraints The temperature of main-stage vein formation and associated potassic alteration at Pine is constrained by the alteration mineral assemblage related to the potassic alteration event. Experimental mineral stability studies (cf. Frank et al. 1998; Hemley 1959; Hemley et al. 1980; Sverjensky et al. 1991) demonstrate that, at pressures of? 1 kbar, the K-feldspar-biotite ± magnetite assemblage forms within a temperature range of about 430 to 625 °C (Seedorff et al. 2005). At Pine, chalcopyrite is present in main- stage and, to a lesser extent, in late-stage veins. Copper-Fe sulphides commonly precipitate between about 300 and 550 °C (Field et al. 2005; Redmond et al. 2004), although at the Butte Cu porphyry, Montana, chalcopyrite-pyrite assemblages were deposited around 600 °C (Brimhall 1977; Field et al. 2005; Roberts 1975). Although vein textures and fluid inclusion observations for main- and late-stage veins at Pine demonstrate that most veins are deformed, rare inclusions that have not leaked give temperature estimates that are comparable with temperatures estimated from alteration mineral assemblages. For example, the single liquid-rich halite-saturated pseudosecondary fluid inclusion from a main-stage vein gives a minimum trapping temperature of 415 °C. The temperature of late-stage vein formation and associated phyllic alteration at Pine is constrained by the mineral stability of sericite in porphyry systems. Sericite is stable at temperatures between —‘300 and 450 °C (Seedorff et al. 2005). Anhydrite-pyrite equilibrium mineral pair temperatures for late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite veins from the potassic-altered centre of Pine range from 410 to 460 °C (n = 2). While veins in more peripheral phyllic-altered areas of Pine define a lower temperature range of 360 to 405 °C (mean = 378 ± 17 °C, n = 5) (Table 2.11). Three liquid-rich, halite-saturated pseudosecondary fluid inclusions from subsequent late-stage quartz-pyrite-chalcopyrite veins define a minimum trapping temperature of 230 to 340 °C. Pyrite-chalcopyrite suiphide isotope pairs for main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite veins from the centre of the Fin prospect give comparable temperature estimates of 420 and 435 °C (average = 428 ± 11 °C, n = 2) (Table 2.11). In summary, Au-Cu-Mo mineralization related to main-stage veins and related potassic-alteration in Pine quartz monzonite occurred at a temperature between 430 to 550 °C (i.e., 430 °C is given by the 82 lower temperature estimate for the K-feldspar-magnetite alteration; 550 °C is defined by the upper temperature for chalcopyrite deposition). Chalcopyrite ± Au continued to be deposited during the late- stage vein and phyllic alteration event. The first-forming, anhydrite-pyrite ± specular hematite ± chalcopyrite veins indicate that the hydrothermal fluid temperatures were still relatively hot (i.e., 410 to 460 °C) in the centre of the Pine porphyry system and probably decreased outwards (i.e., 360 to 405 °C) over approximately 200 m in horizontal distance. The later-forming quartz-pyrite-chalcopyrite veins suggest that the hydrothermal fluid cooled to about 230 to 340 °C. Hence, the fluid temperature data for the evolving Pine porphyry system is consistent with a retrograde temperature-time pathway, which is commonly described for other porphyry systems, world-wide (i.e., Hemley et al. 1992; Redmond et al. 2004). The only temperature estimate available for Fin is based on sulphide mineral pairs from main-stage veins, which indicate a temperature of about 430 °C. Pressure and depth estimates Accurate pressure constraints for main-stage and late-stage Au-Cu-Mo mineralization are scarce at Pine and Fin, principally because of the deformed nature of most primary and pseudosecondary fluid inclusions in the veins. No concordant isochores were generated from data collected from Fin main-stage veins. However, the single liquid-rich halite-saturated pseudosecondary fluid inclusion from the Pine main-stage vein gives a minimum pressure estimate of ‘—4.7 kbar (Fig. 2.22). Assuming lithostatic fluid pressures and a pressure gradient of 3.3 kml 1 kbar (Hagemann and Brown 1996), the corresponding minimum depth estimate for main-stage mineralization is 5.5 km. Three liquid-rich, halite-saturated pseudosecondary fluid inclusions from late-stage quartz-pyrite-chalcopyrite veins defme a minimum trapping pressure of ‘-—1.5 to 1.75 kbar (Fig. 2.22). These pressures correspond to a range of minimum depths of 5.0 to 5.5 kin for late-stage mineralization. Although the pressure estimates for these events are based on limited data, the depths are realistic considering that veins at Pine have a consistent northeast orientation rather than defining concentric or radial sets, which are characteristic of shallow-forming (1-3 km deep) porphyry systems (cf. Tosdal and Richards 2001). 83 The overlapping pressure and depth estimates for main- and late-stage mineralization events at Pine suggest that cooling between the main- and late-stage mineralization events in the Pine porphyry system did not correspond with a significant pressure decrease. The estimated depth range for mineralization at Pine is comparable with the deeper range of porphyry systems (i.e., 1 to 6 km, Seedorff et al. 2005). Constraints on the source(s) of metals and sulphur in the hydrothermal ore fluid at Pine Possible sources for metals in fluids responsible for Au-Cu-Mo mineralization at Pine are estimated from stable isotope (e.g., S, C, and 0) values for hydrothermal sulphide, sulphate and carbonate minerals, as well as radiogenic Pb isotope values for hydrothermal suiphide minerals and igneous feldspar crystals. Sulphur isotope source constraints Late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite veins at Pine have ö34S values for pyrite of -1.86 to 1.67% and 34S values for anhydrite of 12.91 to 16.09%o. These values are comparable to 34S data reported elsewhere for porphyry deposits (Fig. 2.21). The narrow 634S range for pyrite values (around 0%o) indicates that the oxidation state of the late-stage fluid was below the S02/H boundary or remained constant with respect to it (cf. Ohmoto and Rye 1979). Using the isotope fractionation equation of Ohmoto and Rye (1979) for pyrite and a temperature estimate of 380 °C for the late-stage anhydrite pyrite ± specular hematite ± chalcopyrite veins (see fluid temperature constraints above), the634SH25value of the Pine late-stage fluid is calculated to be between -2.8 and 0.65% and using a temperature estimate of 330 °C for the late-stage quartz-pyrite ± chalcopyrite veins (see fluid temperature constraints above), the 634SH25 value of the Pine late-stage fluid is calculated to be between -2.17 and O.57%o. Using the isotope fractionation equations of Ohmoto and Rye (1979) for pyrite, chalcopyrite and sphalerite and a temperature estimate of 430 °C for the main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite 84 veins (see fluid temperature constraints above), theö34SH2svalue of the Fin main-stage fluid is calculated to be between -0.45 and 5.02%. Main-stage quartz-pyrite-chalcopyrite ± molybdenite ± sphalerite veins at Fin have ö34S values for suiphides of 0.37 and 5.84%o. These values are isotopically heavier than typical ö34S compositions reported elsewhere for porphyry deposits (Fig. 2.21) and may suggest a heavier source of S has been contributed to the fluid. Initial sulphur isotope composition and SQl /H2Sratio ofthe Pine late-stage hydrothermalfluid The ö34S versus 34SsuIphatesu1phide diagram(34S-z) for coexisting sulphate-sulphide mineral pairs has been used in other studies to demonstrate both equilibrium fractionation (e.g., Field and Gustafson 1976) and disequilibrium fractionation of sulphur isotopes in an evolving hydrothermal fluid (e.g. Shelton and Rye 1982; Zheng 1991). Requirements for the calculation include: suiphide and sulphate minerals must be contemporaneous and deposited over a range of temperatures; there must be a relatively constant S042JH ratio; and isotopic equilibrium must be achieved between the mineral pairs (Field and Gustafson 1976; Ohmoto 1986; Shelton and Rye 1982). Sulphur isotope equilibrium is commonly attained between aqueous sulphur species (S042 and H2S) during the cooling of a single hydrothermal fluid at temperatures between 200 and 400 °C and low-pH (<3) (Rye 2005). Assuming equilibrium has been attained between aqueous sulphur species (S042 and H2S), the bulk sulphur isotopic composition (6S) of the system can be determined by plotting regression lines through the data sets. The convergence of these lines at A34S = 0 should defme the approximate value of 634S5 due to the temperature dependency of sulphur isotope fractionation (Field and Gustafson 1976; Kusakabe et al. 1984). Similarly, the slopes of the regression lines should define the mole fractions of oxidized (Xso42) and reduced (X25)of the system (Field and Gustafson 1976; Kusakabe et al. 1984). At Pine, coexisting anhydrite-pyrite mineral pairs have been established for late-stage anhydrite pyrite ± specular hematite ± chalcopyrite veins based on textural relationships as well as sulphate sulphide geothermometry data, which are compatible with fluid inclusion temperature estimates. Assuming that equilibrium relationships exist, anhydrite-pyrite ö3”S compositions from late-stage veins at 85 the Pine deposit are plotted on the634S-A diagram (Fig 2.23), and give a ö34S5 value of 9.1 %o with calculated values for Xso42- and X25 of 0.65 and 0.35, respectively. With increasing dominance of Xso42- or XH2s, and over a given temperature range, the ö34S values of that species will become less variable and closer to the 34Ss value (Ohmoto and Lasaga 1982). Field et al. (2005) note that simple estimates from the intercept and slopes should be viewed with caution due to the covariance of Xso42- with ö34S5 and suggest that the best estimates of 34Ss and Xso42. should be made by using the statistical data from the regression analysis to obtain the most conservative estimates of 34S5and X5042- using the following equations: X5042 =b34SH2s/(A34S5042 +/34SH2s) (1) where z is the isotopic spread between the highest and lowest634S%0values for each mineral (or the standard deviation). Field et al. (2005) show that by comparing the spread of the two data sets the X5042- value can be approximated. 34 — IA o Ss — o ,so42 — vxso42-—iisji.I - X so42 = 34SH2s- (/.so42-_H2 )(XHs— 1) (2) where634S5042- and634S25 are the mean values for sulphate and suiphide minerals, respectively, and their difference is A8042-.H5. Application of Eqs. (1) and (2) to the Pine anhydrite-pyrite data gives a preferred834SEs value of 8.6% and a X5042- value of 0.61. Uncontaminated granitic magmas have 34Smeit values of between -3 and +3%o and the fluids derived from such melts have ö34Sfl11d values of -3 to +7%o, with the higher values characterized by initially high oxidation states (Hedenquist and Richards 1998; Obmoto and Rye 1979). The Pine late stage fluid 34S5 value of at least 8.6%o indicates either contamination of the magma if the fluid is magmatic only, or mixing of an uncontaminated magmatic fluid with another source of heavier634S. 86 15 li_a-1 10 - n=5 X0 = 0.61-0.65 6S8.6-9.1%o Cl) C., • anhydrite • pyrite 0 5 10 15 20 suiphate-suiphide Fig. 2.23 Theö34S versus ASSuIPhat,uIPhjd. diagram(634S-AS) for coexisting suiphate-suiphide mineral pairs from late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite veins at Pine. The mineral pairs give a 34S, value of 9.1 %o with calculated values for X50z- and XHs of 0.65 and 0.35, respectively. 87 Pb isotope interpretation Lead isotopes can be used to determine the source(s) of Pb within ore minerals, assuming that Pb is derived from a common source area as ore metals (Au and Cu) and that they were transported and deposited from the same hydrothermal fluid (Tosdal et al. 1999). The similar behaviour of Pb and Cu in magmatic hydrothermal systems (Henley et al. 1984) substantiates this assumption. However, the Pb data presented for Pine is for late-stage veins which are associated with less Au and Cu than main-stage veins (Fig. 2.12). Additional regional data from McKinley (2006) and Duuring (2006, writ. comm.) have been used for the comparison of Pb data from Pine with Pb data from more regional rocks in the Toodoggone district. All data used for the interpretation of the Pb data from Pine are presented in Table 2.13. Lead isotope values for rocks All rocks in the Toodoggone district are Late Paleozoic to Mesozoic in age (ranging from Permian Asitka Group limestone to Jurassic plutons and dykes); hence, no age correction has to be applied to their corresponding Pb isotope values. On thorogenic(208Pb/4 versus 207Pb/4) and uranogenic(207Pb/4 versus 207Pb/4) diagrams Pb values for rock samples plot close to the curve representing the average Pb isotope values for the upper crust (cf. Zartman and Doe 1981) but far from the curve representing the average Pb isotope values for the mantle (cf. Zartman and Doe 1981) (Fig. 2.24). These relationships indicate that all rocks represented by the data set are relatively enriched in U and Th compared to the mantle. Volcanic and plutonic rocks that form in island arc settings are commonly depleted in U and Th compared to the composition of the average upper crust (e.g., Doe 1970; Tatsumoto 1978). Therefore, the Pb isotope values for rocks indicate enrichment by the addition of older radiogenic Pb (e.g., from subducted sediments/altered oceanic crust or from wallrock assimilation during ascent and ponding of the magma in the crust). Even a relatively small addition of crustal Pb is enough to significantly modify the composition of the magma, resulting in most granitic rocks having a similar Pb signature to the crust that they intruded (Davidson 1996; Farmer and DePaulo 1983). Diakow et al. (1993) observe that the style of volcanism and major-element compositions of the Toodoggone Formation suggest a continent-margin arc setting. 88 15.70 * Representative error Sulphide analyses Feldspar analyses 18.8 19.0 19.2 19.4 19.6 19.8 20.0 20.2 PbPPb uooPbIaoaPb Fig. 2.24 Pb isotope compositions of sulphide samples from mineralized veins and disseminations from the Pine and Fin prospects as well as feldspar samples from the Pine-Fin-Mex area. Regional samples are also shown for comparison; these values are from McKinley (2006) and P. Duuring (2006, writ. comm.). A. Uranogenic diagram showing crustal growth curves of Zartman and Doe (1981). All samples have values closer to that of the upper crust than the mantle. B. Uranogenic diagram showing distribution of samples. C. Thorogenic diagram. Values for Black Lake suite intrusive rocks and Toodoggone volcanic are radiogenic and plot in a similar area. The Pine quartz monzonite and Mex monzonite are even more enriched. Pyrite samples also plot in a similar field, showing a trend between the Pine quartz monzonite and the Takla Group volcanic rocks. Rhyolite dikes plot far away from any sample in both diagrams due to the likely presence of microinclusions. 38.2 38.1 * 15.65 15.60 15.55 * * A + 0 Orogerte Pbfl°4b a a- 40.0 38.9 38.8 38.7 j1 38.6 j 38.5 38.4 38.3 15.50 — 18.6 Lltholngy (calcite analyses): A Stlka limestone Utttotogles (feldspar analyses): <3c Takla • Duncan andesilu tuft Metsanlan andusite tuft C) Saunders docile tuft Fin munzvgranite * Pine quartz msnznnlte Pine grunsdlorite * Men msnzonite X Duncan pluton quartz munzunlte X Duncan piutun granodisrite Giegedch pluton granodiotite • Pine Type I syanite dptre O PIne rhyulitedyku Veins front Fin: (pyrue urchalcopynte analyses) + 0uaar-pyrnt5alcynte ± tnolybdenite ± sphalerite 4 Epldote.pyrtte-dtatcupynlte ± rttotybdenlte ± spltaterfte Veins front Pete )pprfte nttatyses): • Brenda . Mhydrea-pyrtte±speselar hetttatite ± chalcupyrite • Qsa,tz-pynte ± chalcopyree 0 Pyrite ± chalcopytite18.6 18.8 19.0 19.2 19.4 19.6 19.8 20.0 20.2 89 All syenite dyke samples yield distinctly more radiogenic compositions than the main field for rocks and suiphide minerals at Pine-Fin-Mex on the thorogenic and uranogenic plots (Figs. 2.24B and C). This is probably caused by the presence of micro-inclusions of zircon or titanite within the feldspar crystals (J. Mortensen, pers. comm), hence these samples are discounted. The least-radiogenic sample of Type 1 syenite dyke (Figs. 2.24B and C) is the best estimate for the Pb isotopic composition of the late syenite dykes. Plutons of the Black Lake intrusive suite that range in age from ca. 200 to 197 Ma together with Toodoggone volcanic rocks (ca. 201 to 194 Ma) show data fields that largely overlap within error. Interestingly, despite the Fin monzogranite being about 20 million years older than the Black Lake suite plutonic rocks, the Fin pluton also plots within this group. The Pine and Mex stocks have more radiogenic Pb values than most plutons and Toodoggone Formation rocks, with the Pine stock plotting above the upper crustal curve in the uranogenic diagram (Fig. 2.24B). This suggests that the Pine stock has been contaminated by more radiogenic Pb than other plutonic rocks in the Pine-Fin-Mex area. The Pine quartz monzonite stock sample from which the feldspars were separated had been affected by potassic alteration. Regional samples of Takla Group andesite (from Kemess North and Sofia) have the least radiogenic Pb values, whereas Asitka Group limestone samples (from Kemess South) have the most radiogenic values. Limestone is normally less radiogenic relative to the average crust (Jahn and Cuvellier 1999); however, the Asitka could be enriched in 206Pb due to the Pb composition of the Permian seawater they precipitated from, or because of the presence of traces of fine-grained clastic sediments. The 207Pb/6 versus 208Pb/6 diagram (Fig. 2.25), with its smaller errors associated with each Pb isotope analysis, confirms the interpretations made from the thorogenic and uranogenic diagrams. 90 207Pb/6 0.84 0.83 0.82 0.81 0.80 0.79 0.78 0.77 I I 1.88 1.90 1.92 1.94 Lithology (calcite analyses): • 1 .96 A Astika limestone -U Lithologies (feldspar analyses): • 1 .98 Takla * Duncan andesite tuff Metsantan andesita area of panel B Saunders dacite tuft 2.00 I 4). Fin monzogranite I Pine quartz monzonite 2.02 * Mex monzonite x Duncan pluton quartz monzonite I Pine granodiorite x Duncan pluton granodiorite ‘2.04 Giegerich pluton granodioritePine Type 1 syenite dyke I Representative error: 0 Rhyolite dyke I + Sulphide analyses + Feldspar and calcite analyses ‘2.06 Veins from Fin:(pyrite or chalcopyrite analyses) + Quartz-pyrite-chalcopyrite 2 00 ± molybdenite ± sphalerite— 4. Epidote-pyrite-chalcopyrite± molybdenite ± sphalerite Veins from Pine (pyrite analyses): * • Breccia + — 2.02 Anhydrite-pyrite ± specularhematite ± chalcopyrite x • Quartz-pyrite ± chalcopyrite • 0 Pyrite ± chalcopyrite *4 * -.2.03 -o Representative error: — 2.04 • + Sulphide analyses • Feldspar analyses • _______________________________________________ — 2.05 0.830 0.825 0.820 0.815 207Pbfl°5b Fig. 2.25 207Pb/6 versus 258Pb/06 diagram showing Pb isotope values of sulphide samples from mineralized veins and disseminations from the Pine and Fin prospects and of feldspars from the Pine-Fin Mex area rock types. Suiphide samples from Late-stage veins from Pine plot away from the host quartz monzonite pluton. Suiphide samples from Main-stage veins from Fin also plot away from the host monzogranite pluton. 91 Pb isotope values for suiphide minerals Suiphide minerals from main-stage veins related to potassic alteration at Pine are either too fine grained to separate or are replaced by later-stage sulphide minerals; hence, they were not used for Pb isotope analysis. However, the feldspars analyzed from Pine show that main-stage potassic alteration caused a radiogenic signature in the feldspars. Pyrite from late-stage veins associated with phyllic alteration at Pine has Pb isotope values that define a linear array on all Pb isotope diagrams (Figs. 2.24 and 2.25). The linear trend demonstrates that the Pb in pyrite within the veins becomes less radiogenic with time. Although the error for the y-axis is relatively large there is a distinct difference between late- stage anhydrite-pyrite ± specular hematite ± chalcopyrite veins and subsequent late-stage quartz-pyrite ± chalcopyrite and pyrite ± chalcopyrite veins, which plot closer to the Takia Group. The relatively heterogeneous distribution of Pb values suggests that fluid mixing occurred close to the site of ore deposition (cf. Tosdal et al. 1999). Surprisingly, the Pb isotope values for the late-stage veins are more similar to the Pb isotope values for the weakly mineralized Pine granodiorite, rather than the mineralized Pine quartz monzonite stock. The least radiogenic sulphide samples are similar to Pb values for Takia Group basalt from Kemess North and Sofia (data from McKinley 2006 and P. Duuring pers. comm.). These relationships suggest that Pb (and other metals) in late-stage veins, particularly the quartz-pyrite ± chalcopyrite and pyrite ± chalcopyrite veins, are not entirely derived from the Pine quartz monzonite. Rather, it is possible that the late-stage fluids circulated through Takia Group basalt countryrock at depth (or even other unidentified, deeper basement rocks in the district), causing the Pb values of the suiphide minerals to approximate the Pb values of the basalt and other country rocks. Pyrite and chalcopyrite from main-stage veins at the Fin prospect do not plot with the Fin monzogranite as expected. In all diagrams (Figs. 2.24 and 2.25) Pb isotope compositions of sulphide minerals from Fin plot between the Fin monzogranite and the Takla Group data, suggesting that the metals were not s derived from the Fin magma. It is possible that the main-stage fluids circulated through Takla Group countryrock or other unidentified, deeper basement rocks in the district, causing the Pb values of the sulphide minerals to approximate that of the country rocks. It is also possible that there are other smaller porphyritic stocks underlying the Fin monzogranite that contributed fluids and metals. 92 Porphyry evolution models commonly attribute the metal source to the magma from which the mineralizing fluids were derived (Burnham 1979; Candela and Holland 1986; Cline 1995). This model cannot be tested at Pine because the only Pb isotope data are from late-stage veins, with main-stage veins proving too difficult to sample. Porphyry models typically describe phyllic alteration as being caused by the mixing between meteoric fluids and the magmatic fluids that are responsible for potassic alteration (Gustafson and Hunt 1975). Alternatively, Shinohara and Hedenquist (1997) show that the fluid responsible for late-stage low-temperature phyllic alteration at Far Southeast was a low-salinity magmatic-only fluid that did not undergo phase separation. As a further alternative, Norten (1988) suggests that a non-magmatic or externally derived fluid can contribute most or all of metals to a porphyry system. Pb isotope data for the Pine late-stage veins and associated phyllic alteration indicates that the fluid did not have a magmatic-only source. Rather, it is likely that they formed by the interaction between magmatic-derived and external (meteoric) fluids, both of which probably scavenged metals from the Pine granodiorite and surrounding Takla Group volcanic rocks. A contribution to the Pb isotopic compositions of sulphides associated with phyllic, and even potassic is noted in other porphyry deposits (Bagdad porphyry Cu, Arizona, Bouse et al. 1999; e.g., Porgera porphyry Au, Papua New Guinea, Richards et al. 1991). Carbon and oxygen isotope source constraints The Pine 6180 and ö’3C data from post-mineralization carbonate veins (n = 4, Table 2.10) occupies a field that partially overlaps the field for primitive igneous rocks (i.e., the magmatic field of Sharp 2006) on a versus 613C diagram (Fig. 2.26). This trend suggests that the post-mineralization veins formed by a magmatic-derived fluid at a late stage in the magmatic-hydrothermal evolution of the Pine porphyry Au-Cu-Mo system. 93 86 Jurassic __________ seawater 2 unaltered marine - 0 limestone -2 Pine P L magmatic -12 ‘., 5 10 15 20 25 30 35 40 18 0 I/QO) v-sMow Fig. 2.26 Oxygen and carbon data for post-mineralization veins at the Pine prospect. Fields from Sharp, 2006, and Jurassic seawater after Veizer et al. (1999). Samples show that the fluids were probably derived a magmatic source, with a possible meteoric contribution. 94 Summary of5, C, 0, and Pb isotope constraintsfor the Fin and Pine systems Lead isotope data for all rocks shows that magmas in the district are contaminated by radiogenic lead, with the Takia Group volcanic rocks displaying the least radiogenic Pb isotope signatures. The 634ss value of the initial late-stage fluid also suggests radiogenic lead contamination of the source magma. Sulphur isotopes for main-stage veins at Fin have isotopically heavier ö34S values than typical porphyry suiphide compositions. Lead isotope data for the main-stage ore fluid at Fin suggests that the magmatic-derived fluids interacted with country rock and other fluids. No isotope constraints are available for late-stage fluids at Fin. At Pine, the main-stage veins associated with Au-Cu mineralization and potassic alteration of the quartz monzonite stock suggest that the ore fluids were mainly derived from a magmatic source. Sulphur and Pb isotope data for late-stage veins at Pine indicate that, although the fluids were likely derived from the Pine quartz monzonite host, there is a significant Pb contribution from other fluids that probably circulated through country rocks, such as Takla Group volcanic rocks. Normally, a cooling hydrothermal fluid precipitating suiphide minerals, decreasing the suiphidation state (e.g., Seedorff et a!. 2005) due to the reaction of SO2 with H20 to form H2S (Ohmoto and Rye 1979). However, at Pine, the presence of anhydrite and hematite in late-stage veins suggests an initial relative increase in sulphidation state relative to the main-stage fluid, indicating addition of S to the system (e.g., Seedorff et a!. 2005). The calculated S042-H ratio of 3:2 for the first-forming late-stage fluid indicates that the fluid was oxidized. In contrast, subsequently forming late-stage veins (e.g., pyrite ± chalcopyrite veins) indicate a relative decrease in the sulphidation state of the fluid during cooling. An increase in S in the initial late-stage fluid may have been caused by: (1) mafic magma injection and SO2 discharge into the Pine quartz monzonite magma chamber, (2) higher S content of a magma pulse that contributed the main source of fluid, andJor (3) scavenging of S from country rocks by the late-stage fluid (Fig. 2.27). The presence of contemporaneous mafic dykes and mafic inclusions within the Pine granodionte and Type 1 syenite dykes suggests the presence of a mafic magma at depth as well as the possible assimilation of mafic country rock (i.e., Takla Group rocks). Lastly, C and 0 isotope data for post-mineralization carbonate veins indicate that the post mineralization fluid was magmatically derived. 95 ca.198Ma surface L <<< K K < < < •4MetsantanMember.< K K., K << < < < ( < << <can orphyiy minerahzation M’ex, _______ 5 km b ‘.‘ ‘ ‘“ ‘ ‘ ‘— •‘‘ ‘—‘ I I “ dèiioIatiIIzatibn’ “ “ “ r ‘ “ ‘ i I “. “ \, ,. ‘, I ‘., \/ \/ \f , felsic magma, Takla’Group “ ‘‘ “ “ “ ‘‘ ‘, “ “ ‘‘ 10 ee ye ‘ rrfic magriii — Ifi I I I Fig. 2.27 Proposed model for the formation of the Pine and Mex systems. The growth of the parent magma chamber follows the model proposed by Candela and Piccoli (2005) of an outwardly propagating magma chamber where the crystal mush is progressively spread laterally to make way for subsequent batches of magma which cause multiple stocks. The remnants of the parent magma chamber could be the Giegerich granodiorite. A. At ca. 198 Ma devolatilization of the magma results in the emplacement of the Pine and Mex porphyry stocks. Recharge events may cause subsequent smaller stocks (cf. Candela and Piccoli 2005). The magma is assumed to have assimilated continental crust (i.e. basement ± Astika) on its ascent. B. By ca. 194, the Metsantan member has undergone erosion and the Saunders Member is also being deposited. Mafic magma present at depth may have interacted with the felsic magma causing the addition of SO2 (cf. Hattori and Keith 2001). In addition, Takia Group volcanic rocks may have been assimilated into the parent magma for the Pine granodiorite during a longer ponding time in the crust. surface 194 Ma S.,, ff , ‘devoIatiIiatioi V km ‘.- ‘., V L’ ieIsic’riaga I I I IL-j I I I I I 96 Integrated temporal and isotopic model for the genesis of the Pine, Fin, and Mex systems Triassic magmatism andporphyry-style Au-Cu-Mo mineralization at Fin The Fin monzogranite, with its associated porphyry-style Cu-Mo-Au mineralization, is clearly older and genetically unrelated to the Pine Au-Cu porphyry system. Relative to the Pine system, the Fin monzogramte is significantly older (i.e., emplacement and mineralization ages are Ca. 218-221 Ma) and the host monzogranite is more fractionated [i.e., the monzogranite has higher Si02, (La/Ce)/Y, Zr/Ti02 and Rb/Sr ratios]. Pyrite, chalcopyrite, molybdenite, sphalerite, and galena are the main metallic minerals at Fin, whereas pyrite, magnetite and chalcopyrite are dominant at Pine. Vein characteristics differ in that the Fin system contains wider early- and main-stage veins that more commonly display large crustiform quartz crystals. Concentric hypogene alteration zonation is absent at Fin, with only pervasive propylitic alteration occurring at surface. It is uncertain which country rocks the Fin monzonite intruded because only the Fin and later- forming Pine plutons and Toodoggone Formation rocks are presently exposed in the Pine-Fin-Mex area (Fig. 2.2). A quartz monzonite pluton is exposed < 1.5 km to the north of Fin monzogranite where it intrudes Asitka and Takia Group country rock and is unconformably overlain by Duncan and Metsantan Member volcanic rocks (Diakow et al. 2005). The Fin monzogranite may be part of this much larger pluton. Thus it is likely that the Fin monzogranite in the Pine-Fin-Mex area also intruded Asitka and Takla Group rocks (Fig. 2.28A). Lead isotope data show that Fin main-stage vein sulphides did not derive all of their Pb from the Fin monzogranite pluton since the Pb isotope values for the main-stage veins are more comparable to values for Takia Group rocks than for the Fin monzogranite (Figs. 2.24 and 2.25). Rather, it is likely that magmatic-hydrothermal fluids that exsolved from the Fin monzogranite stock scavenged metals from nearby Takla Group volcanic rocks. Alternatively, some metals could also have been sourced from other deeper plutons that are not presently exposed in the area. It is unclear how the Early Jurassic alteration event at Pine might affect the Pb isotope values for the earlier Fin monzogranite and main-stage veins. 97 ., \ % \‘ \F ‘/ \/ J ,/ \f . \f \., \/ \/ V I I 1kn A Fig 2.28 Detailed genetic model for the Pine-Fin-Mex area. Horizontal scale is exaggerated. A. Emplacement of the Fin monzogranite into Takia Group and Asitka Group country rocks and associated Cu-Mo mineralization. B. Uplift and erosion, exposing the Fin monzogranite. Unconformable deposition of the Duncan Member andesite tuff, followed by deposition of the Metsantan Member. C. Normal faulting along the Black fault to displace volcanic units. D. Emplacement ofthe Pine porphyry stock and potassic alteration (pink overlay) and Au-Cu mineralization into Duncan Member and Takia Group. E. Erosion followed by the emplacement of the Pine granodiorite and the waning of the hydrothermal system. Phyllic alteration (grey overlay) is focused along pre-existing faults. E Deposition ofthe Saunders Member and the emplacement ofa series ofintermediate-felsic dikes, the first occurring at the end ofthe mineralization event. —218 Ma: Emplacement of Fin monzogranite pluton — 198 Ma: Emplacement of Pine quartz monzonite and mineralization (potassic alteration) SWSW NE \•• S., S.’ ‘V “V ‘V “V ‘V ‘V Cu, Io.Au Zn. Pt, “V ‘‘ ‘V V ‘V ‘V ‘V - Sttm” ‘VTakGrp “‘ S., ..A - V V V “V ‘V fr V V S.’ V ‘V .S.’j SV \f ‘./ \V ‘V ‘.V j ‘V ‘VS.’ ‘V ‘ ,_. 5f. V.••%‘ * V ““ 1’ ‘‘ I V .\‘ V ‘V SW NE — 201 Ma: Unconformable deposition of Duncan Member <198 and >194 Ma: Emplacement of Pine granodiorite and continuation of mineralization (phyllic alteration) voIcicia’ — V: “: - -- -.,- --. - -‘,- - -.- -‘. - >198 Ma: Black fault normal faulting —194 Ma: Emplacement of dykes and end of mineralization 98 Late Triassic-Early Jurassic volcanism andfaulting The 217.8 ± 0.6 Ma Fin monzogranite was unroofed and exposed due to uplift and erosion by about 200.9 ± 0.4 Ma (constrained by the maximum depositional age for the Duncan Member volcaniclastic rocks; Fig. 2.28B). No volcanic eruptive centre is identified at Pine-Fin-Mex, however, regional-scale northeast and northwest trending faults elsewhere in the Toodoggone district are considered to have focussed ascending magmas (Diakow et al. 1991; Diakow et al. 1993). These faults also may define graben structures that are filled with volcaniclastic material during volcanism (Diakow et al. 1991; Diakow et al. 1993). The Mex Creek fault cross-cuts Toodoggone Formation rocks located between Pine and Fin and probably represents the most southerly extension of the regional-scale Black fault (based on the continuation of the Black fault lineament in aeromagnetic data). To the north of the Pine-Fin-Mex area, the Black fault displaces Metsantan Member rocks with normal movement but not the Black Lake suite plutonic rocks. Hence, it is interpreted that the Mex Creek fault displaced the Fin monzogranite and Duncan Member rocks in the Pine-Fin-Mex area prior to the emplacement of the Pine pluton (Fig. 2.28C). Late Triassic-Early Jurassic plutonism and porphyry-style Au-Cu-Mo mineralization at Pine The Pine quartz monzonite stock likely intruded Takla Group basalt and the base of the Duncan Member andesite at 197.6 ± 0.5 Ma. The Duncan Member presently exposed in the Pine-Fin-Mex area is cross-cut by mineralized veins associated with the Pine porphyry. The Mex monzonite stock is interpreted to have been emplaced contemporaneously with the Pine quartz monzonite based on their similar geochemistry and because both plutons intrude Duncan Member rocks and are cross-cut by Type 1 syenite dykes. The estimated depth of emplacement for the Pine pluton is assumed to be similar to the calculated depth for Mex (determined by Al-in-homblende geobarometry) of about —5.5 ± 1.7 km below surface (Fig. 2.28D). Assuming that this emplacement depth for the Pine pluton is realistic, the Toodoggone Formation rocks that overlie the Pine pluton must have been up to 5.0 km thick. Limited microthennometry data for Pine main-stage pseudosecondary fluid inclusions, as well as the potassic 99 alteration mineral assemblage, suggest that ore minerals were deposited from a relatively high- temperature (430 to 550 °C) magmatic fluid, at a depth of at least 5.5 km. Lead isotope data show that the Pine pluton assimilated significant crustal material, giving it the highest radiogenic Pb isotope ratios of all the analyzed rocks in the Pine-Fin-Mex area (Figs. 2.24 and 2.25). The magmatic-hydrothermal fluid exsolved from the magma resulted in early- and main-stage veins and corresponding potassic (K-feldspar + magnetite) alteration in the Pine quartz monzonite. Coeval high-salinity, liquid-rich fluids and low-salinity, vapour-rich fluid inclusions were trapped by quartz in these veins. The deformed nature of these veins prevents the interpretation of whether these fluids represent the trapping of a boiling fluid or the heterogeneous mixing of two or more fluids. The trigger for early Au and Cu deposition is poorly constrained at Pine. However, the restriction of most mineralization to the relatively hot main-stage fluids suggests that fluid cooling was important in precipitating the metals. In most porphyry systems where Au and Cu are transported in the magmatic hydrothermal fluid as bisulphide or chloride complexes (Hemley et al. 1992; Holland 1972), cooling and fluid boiling are considered to be important in destabilizing the metal complexes and causing gold and chalcopyrite precipitation (c.f., Burnham 1979). The minor occurrence of sulphides in main-stage veins is commonly expressed in high-temperature, potassic assemblages, whereas the relative abundance of magnetite is typical of deep, oxidized porphyry systems (Seedorff et al. 2005). Propylitic alteration occurs throughout Duncan Member andesite tuff to the east of the Pine central zone and is considered to be the coeval, distal expression of the potassic alteration event (cf. Lowell and Guilbert 1970). Fin monzogranite farther to the northeast of the Pine system also displays propylitic alteration. Given that the Fin monzogranite displays early- and main-stage veins that are overprinted by pervasive epidote-chiorite alteration, it is likely that the propylitic alteration was caused by the distal circulation of fluids related to the Pine system. The Fin Lake fault cross-cuts Toodoggone volcanic rocks close to the eastern margin of the Pine quartz monzonite stock and may have slightly displaced the ore body. 100 Late-stage mineralization at Pine The Pine granodiorite stock probably intruded the Pine quartz monzonite during the final stages of Au-Cu mineralization at Pine (Fig. 2.28E). Evidence includes: (1) the granodiorite cross-cuts the quartz monzonite, (2) the granodiorite does not contain early- or main-stage veins or significant Au, Cu, or Mo (which are present in the nearby quartz monzonite), and (3) the granodiorite does contain late-stage veins and associated weak chiorite-sericite (i.e., phyllic) alteration along its contact with the quartz monzonite. Phyllic alteration associated with late-stage veins is most intense along pre-existing faults, such as the Mex Creek fault, and suggests that fluids were channelled along these faults, probably away from the centre of the Pine porphyry system. The limited fluid inclusion data for the late-stage veins suggest that late-stage mineralization occurred at a similar depth to the main-stage event (i.e., at about 5.5 km). Temperatures of formation for the first-forming, late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite veins range from 410 to 460 °C in the centre of the Pine porphyry system to 360 to 405 °C in the peripheral areas of the system (based on sulphide-sulphate S isotope geothermometry). Later-forming late-stage quartz-pyrite-chalcopyrite veins suggest that the hydrothermal fluid cooled to about 230 to 340 °C (based on fluid inclusion total homogenization temperatures). The S042--H ratio for the initial late- stage fluid was likely 3:2 (based on sulphur isotope data from anhydrite-pyrite ± specular hematite ± chalcopyrite veins), suggesting that the fluid was initially relatively oxidized. Sulphur isotope data for late-stage veins suggest that the fluid was derived from a magmatic source. Lead isotope data for these veins trends away from values for the Pine quartz monzonite, towards a less-radiogenic value, such as that for the Takia Group rocks. The trend may also be caused by the fluid interacting with a fluid derived from the Pine granodiorite stock or through mixing with other circulating fluids. Such a mixing process is commonly described for porphyry systems whereby magmatic fluids mix with meteoric waters during the collapse of the hydrothermal system (e.g., Lowell and Guilbert 1970). The 34Ss for the late-stage fluid calculated from coeval pyrite and anhydrite minerals in anhydrite-pyrite ± specular hematite ± chalcopyrite veins is about 8.596g. Most felsic magmas associated with porphyry deposits have a 34S5 value of between 3 and 7%o (Ohmoto and Rye 1979). Hence, it is likely that the late-stage fluid at Pine exsolved from a porphyry that was significantly contaminated by 101 another magma or country rock that contained higher 634S values. Lead isotope data show that the Pine quartz monzonite has the most radiogenic Pb isotope values. The Pb isotope values for late-stage veins confirm the involvement of rock or fluids external to the Pine pluton. The last stages ofmagmatism at Pine, Fin and Mex. evidencefor mafic andfelsic dykes Diorite dykes cross-cut the Pine quartz monzonite and are weakly mineralized. Late-stage and post-mineralization veins cross-cut the diorite dykes and occur parallel to dyke margins. Locally, the Pine granodiorite and the Type I syenite dykes contain mafic inclusions, indicating the presence of mafic rocks at depth, probably sourced from Takia Group country rock. The occurrence of these inclusions in the Pine granodiorite and the Type I syenite dykes indicates that the granodiorite magma and dyke magma chamber ponded in the Takla Group rocks at depth and assimilated some of the country rock (Fig. 2.27). This may have altered the composition and isotopic nature of the magmas, and possibly even lowered the oxidation state below the QFM +2 value which inhibits Au and Cu being included in the magmatic hydrothermal fluid (Candela and Piccoli 2005). The significance of the mineralized diorite dykes is unclear at Pine. Elsewhere, mafic magmatism that is coeval with porphyry systems is considered to be an important source of heat and material (e.g.,S02-C0H0-Cl and metals that have an affinity with Cl and H20) (Candela and Piccoli 2005; Hattori and Keith 2001). In addition, the injection of a hot mafic melt into a felsic magma is interpreted to be the cause of volatile release from the mafic melt, causing overpressuring of the felsic magma and subsequent stockwork mineralization (Hattori and Keith 2001; Sparks et al. 1977). In deposits for which a mafic magma contribution to mineralization is invoked, a mantle-like sulphur isotope signature and evidence for anhydrite and immiscible suiphide blebs in the groundmass of felsic volcanic rocks are observed (e.g., Hattori and Keith 2001). At Pine, like most porphyry deposits, sulphide S isotope values are around 0%, however the total ö34S of the system and radiogenic values for Pb isotopes do not suggest a soley mantle contribution to the Pb isotope composition of sulphides. Type 1 syenite dykes that intruded at 193.8 ± 0.5 Ma contain minor disseminated pyrite and chalcopyrite and low concentrations of Au and Cu. Hence, the onset of syenite dykes in the Pine-Fin-Mex 102 is interpreted to overlap with the final stages of the Pine porphyry system (Fig. 2.28F). Type 2 and rhyolite dykes (the latter was emplaced at 193.6 ± 0.4 Ma) are not mineralized with respect to Au and Cu and cross-cut all other rock types and faults. Their consistent NE and NW trend suggests that they exploited existing structures in the rocks. The weakly to non-mineralized syenite dykes that cross-cut the Pine and Fin plutons cause the wide-spread thermal recrystallization of pre-existing veins. This causes difficulties in conducting fluid inclusion studies on mineralized veins related to Pine and Fin porphyry systems, with representative fluid inclusions being only locally preserved. The intrusion of the Pine pluton near the Fin pluton may have also contributed to the recrystallization of Fin porphyry veins. The difference between the emplacement age of the Pine quartz monzonite and the Type 1 syenite dykes suggest that magmatism and porphyry-style mineralization related to the Pine system continued for about 4 million years. The length of time that a porphyry hydrothermal system exists ranges from less than one million years for a single intrusion (Skinner 1979) to several to tens of millions of years for multiphase porphyry systems (e.g., Chesley 1999). Exploration implications To date, the Fin porphyry system represents the oldest known plutonic rock and porphyry system in the Toodoggone district, with most plutonism occurring between 202 and 190 Ma and porphyry-style mineralization occurring between ca. 202 and 198 Ma (Duuring et al. 2006 and references cited therein). These results are important in that they extend the range of porphyry-style mineralization in the district back by about 16 m.y. to the Triassic and thus increase the exploration potential for many older plutons in the district. The spatial coincidence of plutons (Fin monzogranite, Pine quartz monzonite, and Pine granodiorite) and dykes in the Pine-Fin-Mex areas suggest a possible structural control on the emplacement of felsic magmas. Candela and Piccoli (2005) describe the importance of local structures that control the emplacement of magmatic bodies and focus the upward transport of magmatic fluids. No large-scale structures were mapped in the limited outcrop or interpreted from aeromagnetic data in the 103 Pine-Fin-Mex area, however, this does not preclude the existence of a fault or the intersection of faults at depth. Structures are important in the Pine area in that they control the emplacement of NE and NW- trending felsic dykes and appear to channel late-stage fluids near the Pine quartz monzonite. The presence of the Black fault to the north suggests regional scale faults may be present at depth, thereby focussing magrnas and fluids in the Pine-Fin-Mex area, as proposed on a regional scale in the Toodoggone district by Diakow et al. (1991; 1993). The Mex system is approximately coeval with Pine and exhibits many of the same features, such as quartz-magnetite-pyrite-chalcopyrite veins and associated potassic alteration. The defined extent of the Mex porphyry systems is much smaller than the Pine system, possibly because it has been dissected by the Mex fault, which juxtaposes the Mex monzonite against the unmineralized Giegerich granodiorite. In general, the focussing of several mineralized magmatic phases in a multiphase porphyry system is considered to be beneficial to the size of the mineralizing system (c.f., Barra et a!. 2003). This process has not occurred at Pine. Rather, the emplacement of the unmineralized granodiorite and numerous syenite dykes has resulted in the volumetric dilution of the Pine quartz monzonite porphyry ore body. Possible reasons for the lack of mineralization associated with these late magmas may include a lowering of their oxidation state and potential to carry Au and Cu. The Pine porphyry and other porphyries at Kemess North and Kemess South appear to have formed at relatively deep positions in the crust (i.e., at least 6.5 kin) (McKinley 2006, Duuring writ. comm.). These porphyries are located in the southern portions of the Toodoggone district and are associated with exposed deeper country rock, such as Takla and Asitka Group rocks (Fig. 2.1). Pine demonstrates that, in the absence of country rock, other geobarometry tools (e.g., fluid inclusion and Al in-homblende studies) might be important for selecting eroded plutons that are likely to host porphyry systems. 104 CONCLUSIONS The Pine-Fin-Mex area exhibits several episodes of arc-related magmatism and porphyry-style mineralization that spans a 24 million year interval, from the Late Triassic to Early Jurassic periods. U-Pb dating of the Fin monzogranite, combined with Re-Os dating of genetically associated main-stage quartz pyrite-chalcopyrite ± molybdenite ± sphalerite veins, indicates that porphyry-style mineralization began in the Pine-Fin-Mex areas at ca. 218 Ma. Lead isotope values for suiphide minerals from Fin main-stage veins suggest that magmatic-derived fluids interacted with country rocks and possibly with other fluids. The Fin porphyry system represents the oldest known plutonic and porphyry Au-Cu-Mo system in the Toodoggone district. Duncan Member volcaniclastic rocks of the Toodoggone Formation overlie the Fin pluton and have a depositional age of 200.9 ± 0.4 Ma. The Duncan member rocks contain low-grade Au and Cu mineralization, suggesting that they have been mineralized by a later porphyry-style system (i.e., the Pine quartz monzonite). The Pine porphyry system occurs to the SW and borders the Fin porphyry system. The 197.6 ± 0.5 Ma Pine quartz monzonite displays a central potassic alteration zone that is flanked by phyllic and distal propylitic alteration zones. The younger age of the Pine quartz monzonite and related porphyry- style Au-Cu-Mo mineralization confirms that Pine and Fin are two distinct porphyry systems separated in time, although overlapping in space. Main-stage Au-Cu mineralization is related to quartz-magnetite chalcopyrite-pyrite veins, which are magnetite-rich and sulphide-poor, and associated with potassic (K feldspar-magnetite) alteration. The veins and alteration formed from a magmatic-related, high- temperature (430 to 550 °C) fluid at a depth of about 5.5 km. Late-stage anhydrite-pyrite ± specular hematite ± chalcopyrite, quartz-pyrite ± chalcopyrite and pyrite ± chalcopyrite veins and associated phyllic alteration in the Pine quartz monzonite coincided with the emplacement of the nearby Pine granodiorite stock. Late-stage fluids responsible for these veins were transported away from the Pine quartz monzonite via faults, causing weak Au and Cu enrichment in the Pin quartz monzonite and granodiorite. Metals were deposited from the late-stage fluid at initially high temperatures of 460 to 430 105 °C, falling to moderate-temperatures of around 340 °C for later veins (at consistent depths of about 5.0 to 5.5 km). Lead isotope values for suiphide minerals from late-stage veins show a metal contribution from Takia Group country rocks, suggesting that magmatic-derived fluids probably interacted with country rock and perhaps other fluids. The Mex system is approximately coeval with Pine and exhibits many of the same features, such as quartz-magnetite-pyrite-chalcopyrite veins and associated potassic alteration. The defined extent of the Mex porphyry systems is much smaller than the Pine system, possibly because it has been dissected by the Mex fault, which juxtaposes the Mex monzonite against the unmineralized Giegerich granodiorite. Intermediate-felsic dykes cross-cut the Fin, Pine, and Mex stocks. The earliest syenite dykes are weakly mineralized and intruded the Pine quartz monzonite at 193.8 ± 0.5 Ma. The age of these dykes provide a minimum age for mineralization at Pine. The timing of the rhyolite dykes constrains the end of magmatism in the Pine-Fin-Mex area to 193.6 ± 0.4 Ma. The geochemistry and S and Pb isotopic signatures of intrusive and volcanic rocks in the Pine Fin-Mex area suggest that the parent magmas for these stocks had a substantial contribution from a continental crustal source. 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The main intrusive phases in the Pine-Fin-Mex area are, in chronological order, the Fin monzogranite (217.8 ± 0.6 Ma), the Pine quartz monzonite (197.6 ± 0.5 Ma), the Mex monzonite, and the Pine granodiorite, followed by a series of three intermediate-felsic dykes. The Fin monzogranite is interpreted to have intruded Takla Group volcanic rocks, was subsequently uplifted and eroded, and overlain by the Duncan member andesite tuff at 200.9 ± 0.4 Ma. The Black Lake suite plutons then intruded the above units. All rocks display typical arc geochemical signatures, indicating the involvement of subducted sediments/oceanic lithosphere in the source magmas. The Pine system is centred in the potassically-altered quartz monzonite stock where the best grades of mineralization occur with high-temperature, intermediate sulphidation state quartz-magnetite chalcopyrite-pyrite veins. Coeval propylitic alteration occurs in surrounding country rocks. Late-stage fluids, with an initially high sulphidation state caused widespread phyllic alteration that replaces the potassic and propylitic zones. Mixing of a fluid derived from a crustal-contaminated magma is inferred to be the source of the late-stage fluids, which then mixed with a meteoric-derived fluid that scavenged Pb from Takia Group volcanic country rocks. The Fin porphyry system is an older, Late Triassic system which has been overprinted by the Pine hydrothermal system. The Mex hydrothermal system is interpreted to be contemporaneous and similar in style to the Pine system. 114 FURTHER WORK Suiphide minerals from main-stage veins from the Pine porphyry system were not analyzed for S and Pb isotopes. Suiphides in these veins are fine-grained and present in low concentrations. Bulk sampling and crushing was not possible because the samples collected during mapping are all significantly replaced by later veins and alteration zones and contain suiphides related to these events. It is likely that isolated main-stage veins exist at Pine; future S and Pb isotope studies should include these samples. The significance of the Fin porphyry system as a discrete entity from the Pine system only became apparent after laboratory analysis of the Fin samples (i.e., through geochemistry and geochronology). Consequently, the numbers of S and Pb isotope samples were weighted towards the Pine system rather than the Fin system. Hence, more samples from Fin might be useful in providing a more comprehensive data set for comparison with Pine data (e.g., samples of late-stage veins at Fin). The preliminary Re-Os molybdenite age obtained from a main-stage vein at Fin is older than the Fin monzogranite. This sample is one of two currently being analyzed at the University of Alberta. It is hoped that the second sample will better constrain the main-stage mineralization event at Fin. Fluid inclusion studies at Pine and Fin were largely unsuccessful due to the intense recrystallization of quartz veins. The absence of a penetrative foliation in the rocks suggests that stress was not as important as the effect of wide-spread thermal recrystallization due to the intrusion of multiple dykes in the Pine-Fin-Mex area. A total of 34 thin sections of veins were examined from throughout Pine and Fin. The veins were sampled from central and peripheral zones, deep and shallow levels within the systems, and at varying distances to (known) dykes. From these 34 thin sections, 11 were used for fluid inclusion microthermometry. Microthermometry was performed on 20 chips from 11 veins, with 203 individual fluid inclusions measured. From these data, only four inclusions gave meaningful interpretations. It is possible that further sampling might obtain less deformed samples for fluid inclusion analysis. However, this might take significant effort since vein samples did not appear significantly recrystallized until they were examined using cross-polarized, transmitted-light microscopy. 115 An SEM-microprobe study might be employed to determine the oxidation state of the various intrusions in the Pine area by comparing the composition of magnetite and titanite in the intrusions. The oxidation states of mineralized versus unmineralized plutons might be an important indicator of the potential of an unknown pluton to host porphyry-style Au-Cu-Mo mineralization. Considerably more fundamental geological mapping should be completed at Mex. During the course of fieldwork at Pine and Fin, access to the Mex ridge and the Mex porphyry system were severely limited since access depended on helicopter-support. Detailed mapping at Mex, combined with more drilling and laboratory work (S and Pb isotopes) might lead to a more tightly constrained genetic model for Mex. Lastly, the detailed geological mapping could be extended outwards from the in the Pine-Fin Mex, where relationships are well studied. In particular, mapping could test the geographical extent of the Fin monzogranite as well as delineating alteration patterns associated with the Fin porphyry system (remembering that the propylitic alteration in the Fin monzogranite is interpreted to be the distal alteration of the younger Pine porphyry system). 116 Appendix A: Table A.1 Whole-rock major- and trace-elemenggeochemistr t of igneous rocks from the Pine-Fin-Mex area Volcanic units Intn,sive rocks Rock nodesite luff oodesite luff docile Wff docile wfl docile to& rnonzogeanite rnonzogronite rnonzograoite rnoozogrnnile ntoozogranite Field No. JD05-279 JDOS-217 31)04-100 31)04-101 JDO4-102 1005-269 JD05-272 11)05-281 31)05-285 P05-06 09 LabNo. 14139659 N 139661 Nl39651 N139652 Nl39653 N 139664 N 139662 N 139656 N 139663 N 139657 Property Pine/Moo Pine Tiger Notch Tiger Notch Tiger Notch Fin Fin Fin Fin Fin Long. (UTM) 640316 639497 640493 640510 640213 640295 640295 Lot. (UTM) 6343220 6343273 6345263 6345201) 6345103 6344705 6344705 Alteration epi+py epi+py (ncr) (ncr) (set) kern hem epi-chl kern hem minerals’ Si02 59.76 64.20 61.62 62.04 62.29 68.50 69.43 61.54 69.50 6g.9o Ti01 0.53 14.86 15.08 15.33 5.86 0.31 14.93 17.28 14.97 14.63 A1503 16.70 4.93 4.84 5.07 5.28 13.27 2.08 3.74 2.51 2.41 Fe50 5.06 4.46 3.70 3.35 3,95 2,96 1.53 4.38 1.43 1.74 MoO 0,26 .32 1.69 .75 .96 0,13 0.59 1.87 0.82 0.71 MgO 1.09 3.31 3.01 2.05 3.72 0.01 4.29 2.17 4.29 4.60 CoO 4.39 2.53 3.55 3.67 3.12 1.49 3.74 4.62 3.31 2.76 Nu0 2.71 0.01 0.01 0.01 0.01 4.23 0.01 0.01 0.01 0.01 LO 4.24 0.50 0.50 0.50 0.52 3.39 0.36 0.73 0.30 029 P205 0.24 0.25 0.10 0.10 0.12 0.13 0.12 0.22 0.12 0.10 Cr,03 coot 0.24 0.21 0.21 0.21 0.01 0.10 0.20 0.13 OH SeO 0.07 0.07 0.05 0.05 0.07 0.06 0.04 0.04 0.05 0.06 BoO 0.28 0.17 0.15 0.16 0.15 0.15 0.16 0.18 0.14 0.10 LOt 2.00 1.32 3.69 3.23 t.62 .06 0.87 1.94 1.03 2.11 Total 90.19 92,18 90.29 90,32 90.93 90.45 90.25 98.9 90.59 90,50 Ag <1 <I <1 <1 <1 <I <I <1 <1 <I Ao <0,001 <0.001 0.002 0.001 <0.001 <0,001 <0.001 0.002 <0,001 0.003 Ba 2660 1590 1540 1720 1500 1565 1660 1680 1400 1870 Ce 34.6 30 36 40.3 43 22.9 23.8 30.2 19.2 21.7 Co 4,7 7.4 9 II 10.4 6 6,6 9.8 4.7 5.3 Cr 20 10 20 50 20 40 40 30 20 20 Cs 0.9 0 8 9 2 I I I I Cu 15 <5 30 23 10 327 17 210 463 172 Dy 3 4 3 3 3 2 I 3 I Er 2 2 2 2 2 I I 2 1 Eu I I I I I I I I I Go t9 tO 17 20 8 21 19 19 19 20 Gd 3 4 3 4 4 2 2 3 2 Hf 3 3 4 4 3 4 4 4 4 4 Ho I I I I I 0 0 1 0 0 La 18 19 21 23 23 12 13 14 II II Lu 0 0 0 0 0 0 0 0 0 0 Mo <2 <2 4 2 <2 35 <2 31 20 12 Nb 6 7 7 8 7 4 5 5 4 4 Nd 17 20 16 19 20 It tO 15 9 9 Ni II 10 tO II 12 9 12 II II 12 Pb 24 34 8 II 7 12 8 9 8 8 Pr 4 5 4 5 5 3 3 3 2 2 Rb 95 52 115 134 96 110 96 124 99 83 Sm 4 4 3 3 4 2 2 3 2 2 Sn I 2 I I I I I 2 1 Sr 600 596 460 555 663 525 400 407 418 549 Ta <0.5 I I I I <0.5 <0.5 <0.5 <0.5 <0,5 Th I I 0 I 1 0 0 1 0 0 Th 5 4 7 9 8 2 4 4 3 3 TI 1 <11.5 I I <0.5 <0.5 1 I <0.5 Tm 0 0 0 0 0 0 0 0 0 0 U 2 2 3 4 3 I 2 I I V 152 118 162 191 188 61 56 126 56 61 W 3 2 13 10 6 5 3 5 4 3 V 20 23 19 22 22 10 tO 20 8 10 Yb 2 2 2 2 3 I I 2 I I Zn 156 159 77 95 79 104 05 87 04 239 Zr 114 119 131 143 122 122 123 137 119 119 117 Table A.1 Cont. Field No. F05-06 195.5 JD05-200 92-38 162.5 92-40 166.0 5005-134 Intrusive rocks Rock rn000ogrnnite gtz rnonzonite gtz moozonite gtz monzonite rnonzonite granodiorite granodiorite granodiorite granodiorite granodionte N 139658 N139697 N 139700 N 139699 N139671 Fin Pine Pine Floe Men 640295 637985 638098 638116 640749 6344705 6343355 6343583 6343565 6342706 kern ser+chl+bem ksp-rnog ksp-mag ser(+chl) 3D05-288 92-39 144.4 97-12 309.6 )D05-117a N 139674 N 139676 N 139677 N 139655 Pine 638570 6343468 ser+chl+py Pine Giegerich 638570 641446 6343468 6341842 ser+chl+py (ser+chl) 3005.204 Lob No. N 139696 Property Pine Pine Long. (UTM) 637998 638057 Lat. (UTM) 6343447 6343384 Alteration ser+hem ser+ttem minerals’ +chl+py +chl+py Si02 66.72 65.92 57.01 56.53 63.90 65.63 64.60 63.03 62.44 63.97 Ti05 14.68 0.35 15.66 15.39 0.59 14.86 14.99 15.26 14.92 15.88 AlaO, 2.45 14.78 7.28 7.34 16.03 4.82 4.91 6.06 4.70 4.80 Fe0, 2.21 5.59 2.97 3.59 4.58 0.59 3.29 0.92 5.33 4.54 MoO 0.85 0.09 0.89 0.71 0.16 1.53 1.73 1.93 1.81 1.39 MgO 4.74 0.57 4.26 4.60 1.67 3.55 3.08 4.03 3.23 3.55 CoO 2.51 0.51 4.04 2.51 2.41 5.03 2.63 3.36 3.08 2.90 Na20 0.01 3.55 0.01 0.02 4.13 0.01 0.02 0.02 0.01 0.01 K20 0.31 6.26 0.53 0.48 3.68 0.48 0.45 0.52 0.47 0.59 P205 0.07 0.16 0.06 0.09 0.20 0.12 0.16 0.14 0.04 0.13 Cr203 0.11 <0.01 0.fl 0.21 0.01 0.22 0.14 0.22 0.13 0.20 SeQ 0.05 0.04 0.05 0.04 0.05 0.02 0.06 0.02 0.05 0,07 BoO 0.14 0.29 0.21 014 0.18 0.29 0.21 0.19 0,18 0.19 LOl 3.74 0.72 5.01 7.21 1.24 1.60 2.35 2.40 4.58 0.66 Total 98.58 98.82 98.99 98.87 98.84 98.76 98.62 98.09 98.97 98.86 Ag I 1 <1 <I <I <I <1 <1 <I <I Au 0.006 0.214 0.335 0.41 <0.001 0.016 0.002 0.009 0.003 0.001 Ba 1465 2940 2090 1440 725 2900 2030 1845 1740 1885 Ce 9.1 26 31 52 29.5 36.2 29.6 40.8 27.9 42.5 Co 4.5 2.8 4.7 6.4 10.3 5.2 7.8 5.7 10.0 7.8 Cr 40 20 30 20 20 30 80 40 90 20 Cs I I I I I I I I 1 Cu 47 102 663 586 36 24 64 130 tO <5 Dy 1 2 3 5 3 3 2 3 2 4 Er I I 2 4 2 2 I 2 I 2 Eu I I I 2 1 I I I I Ga 18 16 17 23 19 17 16 18 16 20 Gd 2 2 3 5 3 3 3 4 2 4 Hf 3 3 3 5 5 4 3 3 3 4 Ho 0 0 1 I 1 1 0 I 0 La 10 14 16 26 15 18 17 22 16 23 Lu 0 0 0 I 0 0 0 0 0 0 Ma 3 24 29 9 <2 5 3 14 3 <2 Nb 4 7 6 8 6 7 5 8 5 7 Nd 9 12 16 25 16 17 12 18 II 20 Ni II 13 7 12 11 9 6 5 7 II Pb 9 20 21 5 8 23 20 6 6 6 Pr 2 3 4 6 4 4 3 5 3 5 Rb 73 149 113 92 90 124 56 81 73 81 Sm 2 2 3 5 3 3 3 4 2 4 Sn I 2 3 4 I 2 I I I Sr 505 361 447 569 437 218 629 177 481 667 Ta <0.5 I <0.5 1 <0.5 I I I <0.5 Th 0 0 0 I I I 0 I 0 Th 2 5 4 4 4 8 7 7 6 7 TI I I 1 <0.5 I 1 <0.5 <0.5 <0.5 <0.5 Tm 0 0 0 I 0 0 0 0 0 0 U 1 2 2 2 I 3 3 4 3 3 V 53 71 102 20 101 126 116 123 124 124 W 3 4 0 9 2 6 3 5 7 2 Y 9 12 17 42 17 18 14 21 13 25 Yb I I 2 4 2 2 2 2 I 3 Zn 90 117 154 42 232 262 123 144 42 63 Zr 109 119 96 236 145 119 87 106 83 127 118 Table A.1 Cont. lntntsive rocks dykes Rock grnttodiorite gtz monzonite diorite diorite gubben gnbbro gabbrn gubbro gnbbrn moezonite monzonite Field No. 3005-1 17b 3005-156 SMRO5-l10 92-40 167 3005-188 3005.284 92-40 168.5 98-04 35.7 98-04 48.3 31305-127 F05.06 65.5Lab No. N 139675 N 139654 N139673 N 139672 N 139685 N 139660 N 139686 N 139351 N 139684 N 139690 N 139693Property Girgrich Duncan plnton Pine Pine Pine Pine Pine Pine Pine Fin FinLottg.(UTM) 641438 618010 639644 638116 638727 638727 638727 638727 638727 641102 640295Lot. (UTM) 6341829 6346051 6343217 6343565 6343484 6343484 6343484 6343484 6343484 6342245 6344705 Alteration (ser*chl) (ncr) epi*py chl+ser chl+ser cbl+ser rht+ser cbl+ser epi+srt-Iohl epi+ser+chl minernlst SiO 63.68 62.97 58.33 49.03 49.43 52.57 41.69 43.48 51.65 65.58 63.23TiO, 16.01 15.70 0.72 16.30 17.95 18.51 18.37 17,76 17.54 14.93 16.03Al,03 4.57 5.33 15.40 11.45 10.11 8.14 11.77 11.82 8.62 4.92 4.81 FeO 4.33 4.47 8.05 3,07 5.31 3.97 13.24 0.32 5.28 1.34 3.45 MnO 1.37 1.93 0.40 4.38 3,93 307 3.91 439 2.96 1.63 1.71 MgO 3.51 3.33 3.30 4.78 4.64 5.71 2.86 2.92 5.17 3.42 3.08CoO 2.96 3.33 5.32 2.66 .73 1.40 0.16 1.47 2.49 3.16 4.19Na,O 0.01 0.01 1.08 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 KO 0.50 0.49 2.52 1.18 0.96 0.99 1.19 1.11 0.81 0.43 0.51 P0, 0.13 0.11 0.12 0.24 0.28 0.29 0.18 8.25 0.25 0.23 0.21 CrtO5 0.18 0.21 <0.01 0.37 0.35 0.42 0.32 0.29 0.32 0.05 0.22 SrO 0.06 0.07 0.06 0.06 0.05 0.04 0,01 0.02 0.02 0.04 0.06 BoO 0.17 0.17 0.15 0.32 0.16 0.07 0.02 0.09 0.12 0.17 0.24 LOl 0.89 0.62 2.73 5.09 3.37 3.55 4.61 4.62 3.04 2.86 1.58 Total 98.35 98.73 98.19 98.92 98.27 98.74 98.31 98.54 98.25 98.78 99.34 Ag <1 <I I <1 <1 <1 <I <I <I I <1 Au <0.001 <0.001 0.004 0.008 0.003 <0.001 0.003 0.001 0.001 0.007 <0.001 Ba 1560 1685 1520 3070 1560 793 215 825 1130 1865 2280 Cc 35,4 43.4 23.5 29.3 37.1 56.9 31.4 32 40.1 26.5 40.4 Co 5.9 11 36.9 27.4 23.1 15.8 31.4 28.6 17.8 9.1 8.8 Cr 60 40 40 20 <10 10 10 20 <10 20 20 Cs 1 2 5 0 0.2 1.1 <0.1 <0.1 0.1 I 1 Cu <5 10 355 37 17 14 46 132 16 454 16 Dy 3 3 4 4 4.4 6.6 4.1 4.8 4.4 2 3 Er 2 2 2 3 2.7 4.5 2.6 2.7 2,8 I 2 Eu I 1 I 1 1.5 2.3 1.5 1.4 1.5 I Go 17 18 20 7 18 27 24 24 19 19 18 Gd 4 4 4 4 4.8 7.1 4.7 5.1 4.8 2 4 Hf 3 3 3 2 2 6 2 2 3 4 4 Hn I I I I 0.9 1.5 0.8 0.7 0.9 0 Lu 19 23 II 12 17.4 25.7 13.8 13.7 19.4 15 21 Lu 0 0 0 0 0.4 0.7 0.3 0.1 0.4 0 0 Mo <2 <2 <2 <2 <2 3 <2 <2 <2 6 3 Nb 6 9 4 4 5 It 4 4 5 6 7 Nd 16 19 14 19 21.9 32.2 19.9 21.7 21.9 II 18 Ni 7 11 14 13 <5 7 7 II <5 10 II Pb 6 8 73 <5 <5 7 <5 5 5 20 II Pr 4 5 3 4 5 7.5 4.4 4.4 5.3 3 5 Rb 68 112 112 67 31.8 40 2.6 26 53.2 76 102 Sm 3 4 3 4 5.1 7.2 4.7 5.2 4.8 2 4 Sn I 2 I 2 I 3 I <I I 2 2 Sr 566 698 505 480 465 336 109.5 216 222 483 561 Ta I I <0.5 <0.5 <0.5 0.7 <0.5 <0.5 <0.5 1 Th I I I I 0.7 1.1 0.7 0.6 0.7 0 1 Th 6 10 2 2 2 5 I 2 3 9 7 fl <0.5 <0.5 I <0,5 <0.5 <0.5 <0.5 <0,5 <0,5 1 Tm 0 0 0 0 0 I 0 0 0 0 0 U 3 3 I I I 2 I I I 4 3 V 94 155 187 318 241 182 384 409 214 116 127 W 1 4 3 4 I 3 4 3 2 8 7 Y 20 20 24 25 25 41 24 28 25 10 21 Yb 2 2 2 2 2 4 2 3 3 I 3 Zn 61 55 283 269 126 222 91 135 143 3540 148 Zr 112 110 98 74 84 234 61 73 96 131 115 119 Table A.1 Cont. dykes Rock monzonite moozonite monzonite syenite I syenite I syenite I syessite syenite 1 syenile I ayeoile I synnile I Field No. F05-06 122.0 F05-06 155.2 F05-06 163.4 3005-186 31)05-240 F05-05 166.9 31)05-120 3005.123 3005-130 3005-282 31)05.289 N 139691 N 139692 N 139666 N 139665 Fin Fin Fin Fin 640295 640295 63951! 640505 6344705 6344705 6343457 6344898 epi+ser+chl epi+ser+chl hem hem N 139669 N 139668 N 139670 N 139681 N 139602 Men Mex Men Pine Pine 641275 641279 640955 638294 638294 6342035 6342179 6342306 6343337 6343337 hem hem hem hem hem (“-nbI+ser) (+chl+ser) (+chl+ser) (+chl+ser) (+chl+ser) Lab Na. N 139694 Property Fin Long. (UTM) 640295 La!. (UTM) 6344705 AI!eeal,on epi+ser+chl minerals’ Si05 60.22 Ti05 15.78 Al,05 5.37 Fe503 4.92 MoO 1.82 Mg0 2.27 CaO 4.29 Na50 0.0! KO 0.69 P205 0.37 Ce,03 0.24 SeO 0.05 BoO 0.25 LOl 2.18 Total 98.45 Ag <I An 0.00! Ba 2380 Ce 40.8 Ca 10.6 Cr 50 Cs Cu <5 Dy 4 Er 2 Eu Ga 19 Gd 4 Hf 4 Ha I La 2! Lu 0 Mo 2 Nb 7 Nd 19 Ni 10 Pb 16 Pr 5 Rb 108 Sm 4 Sn Sr 478 Ta Tb I Tb 7 11 1 Tm 0 U 3 V 142 w 7 V 24 Yb 3 Zn 186 Zr 134 63.42 15.60 4.80 4.92 1.54 2.66 3.27 0.01 0.49 0.23 0.19 0.08 0.19 1.36 98.75 0.802 1895 37.8 7.8 20 0 25 18 4 20 0 2 17 10 14 4 61 738 112 Is 19 608 131 (1-chl’+ser) 58.37 65.32 16.11 14.82 6.77 3.78 5.60 2.7! 3.09 1.45 5.2! 4.12 2.50 3.08 0.01 0.01 0.66 0.38 0.34 0.14 0.28 0.14 0.09 0.03 0.20 0.17 1.87 1.94 99.09 90.00 0.00! <0.001 2060 1620 41.! 32.3 14.5 7.9 50 40 33 7 4 2 2 1 21 15 4 2 4 3 21 19 0 0 3 <2 7 5 20 14 25 8 22 6 5 4 59 60 4 2 866 249 0 6 0 <0.5 <0.5 0 0 2 3 196 102 10 2 24 15 3 2 280 119 121 104 N 139667 Fin 648295 6344705 hem (+chl+ser) 63.20 15.33 4.83 3.73 1.78 3.66 3.6! 0.01 0.52 0.15 0.17 0.05 0.20 1.12 98.34 <0.00! 1905 31.6 10.4 20 32 17 17 14 76 394 <0.5 136 16 84 124 (+cbl+ser) 62.31 15.27 5.29 4.74 1.74 3.22 3.55 0.0! 0.53 0.38 0.19 0.06 0.17 1.44 98.09 <0.001 1685 34.2 11.7 20 43 19 18 0 9 16 10 75 559 <0.5 44 18 183 134 63.31 15.10 5.17 3.27 1.86 3.48 3.03 0.01 0.51 0.14 0.16 0.07 0.21 1.93 98.22 <0.001 2130 34 12.7 l0 503 2 18 4 It 0 <2 6 14 10 8 4 101 611 8 0 3 149 16 2 241 129 64.44 14.84 4.95 2.66 1.84 3.84 3.23 0.01 0.47 0.15 0.16 0.05 0.22 1.32 98.17 <0.001 2120 32.9 10.6 30 44 16 18 0 <2 14 10 13 4 90 454 <0.5 <0.5 0 137 I5 1040 109 66.13 14.69 3.71 2.19 1.24 3.90 3.94 0.01 0.41 0.12 0.12 0.04 0,19 1,55 98.24 <0.001 2020 32.8 8.9 40 0 62 2 2 16 3 4 19 0 2 6 14 9 6 4 97 3 337 0 I0 0 4 109 18 2 256 110 65.70 64.46 14.53 15.17 3.56 4.32 2.78 5.20 1.29 1.49 3.54 3.47 3.51 3.53 0.0! 0.01 0.35 0.44 0.09 0.15 0.10 0.13 0.04 0.05 0.20 0.22 2.59 2.23 95-57 90.85 0.001 0.003 1930 2100 29.5 32.1 7.2 10.4 10 10 17 12 2 3 I 2 15 16 2 3 3 3 0 1 17 19 0 0 <2 3 5 5 12 14 5 5 7 8 3 4 96 86 2 3 456 482 0 0 7 7 <0.5 <0.5 0 0 3 4 84 103 3 2 12 19 I 2 49 174 93 109 120 Field No. 98-02 42.4 5005-210 JDOS-286 LnbNo. N 139683 N 139679 N 139678 Properly Pine Pine Pine Lang. (UTM) 630294 637965 638294 Lal. (IJTM) 6343337 6343293 6343337 Alteration hem hem+chl hem+chl minerals’ (+chl+ser) +ser ‘45cr Si02 65.42 65.09 64.86 Ti05 4.79 15.25 15.37 A150, 3.59 3.73 4.00 2.47 1.46 2.07 MnO .38 1.19 1.42 MgO 3.60 522 482 CaO 3.59 376 3.30 Nn20 0.01 0.02 0.01 1(20 0.48 0.38 0.50 P,05 0.07 0.10 0.11 Cr205 0.11 0.15 0.13 SeO 0.04 0.04 0.03 Ba0 0.19 0.20 0.17 LOl 2.84 0.95 1.85 Total 98.65 98.32 98.71 Ag <1 <1 <1 An 0.002 <0.001 0.001 Ba 1885 1985 1630 Ce 30.5 46.3 41.1 Co 8 7.1 8.6 Cr 10 50 50 Cs I I Ca 7 16 24 Dy 2 3 3 Er 1 2 2 Es I I Ga 15 16 17 Gd 2 4 4 Hf 3 4 4 Flo 0 1 I La 17 25 22 Lu 0 0 0 Mo <2 <2 3 Nb 5 7 6 Nd 12 20 18 Ni 7 6 7 Pb 6 7 <5 Pr 3 5 5 Rb 100 99 88 Sm 2 4 3 So 1 1 1 Sr 409 389 339 Ta I I Th 0 I Th 8 10 10 TI <0.5 <0.5 <0.5 Tm 0 0 0 U 3 5 5 V 94 78 99 W 3 2 3 Y 13 19 19 Yb I 2 2 Zn 49 94 61 Zr 96 148 142 JDO5-283 98-04 32.5 90-04 44.8 N 139687 N 139688 N 139689 Pine Pine Pine 638695 638727 638727 6343375 6343484 6343484 hem+chl bem+chl hem+ohl *ser +ser +ser 73,95 62.39 74.08 12.97 17.40 12.57 1.34 3.83 1.44 0.50 1.79 0.66 024 1.06 0.15 2.73 6.48 3.51 546 3.66 4.36 0.01 0.01 0.02 0.17 0.51 022 0.03 0.11 0.06 0.02 0.19 0.04 0.02 0.03 0.02 027 020 0.21 0.99 1.69 1.47 98.79 99.34 98.79 <1 <1 <1 <0.001 <0.001 0.001 2680 1970 2070 41.5 64.6 37.9 1.1 4.6 1.4 10 10 10 I I 6 <5 <5 2 6 2 2 I 2 22 13 3 6 2 4 5 4 0 I 0 26 32 23 0 I 0 <2 2 4 8 9 6 14 31 14 5 9 5 <5 6 0 4 8 4 141 73 ItO 2 6 2 I 1 253 443 192 I I 0 I 0 14 7 II <0.5 <0.5 0 1 0 5 3 4 16 66 17 3 3 5 14 41 12 2 4 2 27 104 28 134 227 129 Table A.1 Cont. dykes Rock syonite I syenite 2 syenite 2 syenite 2 rhyolite rhyolite ehyolile 97-12 214.6 N 139680 Pine 638570 6343468 hem+chl +ser 63.52 14.51 4,21 2.70 1.30 4.36 4.14 0.01 0.45 0.10 0.14 0.03 0.19 2.89 98.56 <0.001 1845 36.9 9.1 50 25 12 19 <2 16 104 296 <0.5 102 18 I 56 138 Major elements are in wI. %; trace elements ore in ppm ‘chl =chlonte, epi = epidote, hem = hematite, py = pyrite, ser aericite. Parenthesis denote minor presence. 2From outside of map orea 121 A pp en di x B :T ab le B .1 M ic ro pr ob e an al ys es o f a m ph ib ol es in th e ig ne ou s ro ck s o f t he Pi ne -F in -M ex ar ea . Ro ck ty pe Sa m pl e n o . Po in t S i0 2 T i0 2 A 1 203 C r 203 Fe O T M nO M gO C aO N a 20 K20 F To ta l G ie ge ric h JD OS -1 17 a 1 45 .3 8 1.2 5 7. 29 0.0 1 15 .41 1.0 7 12 .37 11 .3 6 1.2 5 0. 73 0. 25 96 .3 7 gr an od io rit e 2 44 .9 9 1. 16 7. 10 0. 02 15 .21 1. 02 12 .2 8 11 .50 1.2 5 0. 78 0. 20 95 .5 2 3 48 .2 0 0. 43 4. 93 0. 03 14 .1 3 0. 97 13 .97 11 .56 0. 85 0. 45 0. 19 95 .71 4 44 .7 0 12 8 7. 17 0. 03 15 .23 1.0 4 12 .2 2 11 .4 8 1. 28 0. 73 0. 29 95 .4 6 5 44 .8 2 1.3 5 7. 49 0.0 1 15 .28 1. 09 12 .0 9 11 .4 7 1.2 8 0. 75 0. 27 95 .8 9 6 43 .4 0 1. 62 8.3 1 0. 00 15 .83 0. 87 11 .60 11 .50 1.3 7 1.0 5 0. 30 95 .8 7 7 43 .4 7 1.6 3 8. 37 0. 00 16 .7 9 0. 84 10 .96 11 .31 1. 40 1.0 3 0.3 1 96 .11 8 43 .6 7 1. 44 8.2 1 0. 00 15 .5 6 0. 73 12 .05 11 .37 1.3 1 0. 97 0. 25 95 .5 7 9 43 .8 1 1. 59 8. 17 0. 00 15 .8 4 0. 83 11 .7 3 11 .54 1.3 3 1.0 3 0. 39 96 .2 5 10 43 .7 3 1.6 3 8. 43 0. 00 16 .20 0.9 1 11 .56 11 .3 8 1.3 4 1. 04 0. 32 96 .5 3 11 49 .1 1 0. 59 4. 45 0.0 1 13 .23 0. 98 14 .67 11 .69 0. 76 0. 34 0. 23 96 .0 6 12 47 .2 3 0. 84 5. 98 0. 00 14 .8 8 1.1 1 13 .23 11 .71 1.0 3 0. 62 0. 24 96 .8 7 13 46 .8 5 0. 87 6. 09 0. 00 14 .61 1.1 5 13 .18 11 .60 1.0 5 0. 64 0. 38 96 .4 2 14 48 .6 9 0. 49 4. 62 0. 00 13 .83 1.0 5 14 .1 9 11 .7 3 0. 70 0. 35 0. 17 95 .8 2 15 45 .6 0 1. 20 6. 70 0. 00 14 .96 0. 93 12 .97 11 .4 0 1.0 5 0.7 1 0. 34 95 .8 6 16 45 .0 3 1. 32 7. 13 0.0 1 15 .33 1. 07 12 .1 5 11 .2 9 1.1 9 0. 72 0. 28 95 .5 0 17 44 .1 5 1.6 4 7. 87 0. 00 15 .73 0. 92 11 .8 2 11 .35 1.4 0 0. 82 0. 25 95 .9 4 18 44 .7 4 1. 50 7. 48 0.0 1 15 .01 1.0 2 12 .2 8 11 .4 6 1. 29 0. 82 0. 32 95 .9 2 19 45 .4 8 1. 07 7. 27 0. 00 15 .42 1. 10 12 .41 11 .4 7 1. 09 0. 72 0. 28 96 .3 2 20 44 .7 4 1. 44 7. 30 0. 00 15 .40 1. 06 12 .1 3 11 .56 1.2 7 0. 72 0. 26 95 .9 0 21 43 .7 9 1. 19 8. 68 0. 00 16 .00 0. 84 11 .7 3 11 .7 3 1. 40 0.7 1 0. 25 96 .3 2 22 43 .4 4 2. 04 8.9 1 0. 00 14 .7 4 0.8 1 12 .08 11 .55 1.6 5 0. 85 0. 26 96 .3 3 23 42 .0 8 2. 48 9.8 3 0. 00 15 .27 0. 85 11 .3 9 11 .1 6 1.7 1 1. 04 0.3 1 96 .11 24 42 .64 2. 03 9. 09 0. 00 16 .07 0. 94 11 .1 8 11 .2 0 1.6 5 0. 79 0. 30 95 .8 9 25 41 .4 7 1.7 3 10 .60 0. 02 16 .53 0. 92 10 .57 11 .52 1. 66 0. 98 0. 24 96 .2 5 M ex JD O5 -1 34 1 50 .7 7 0. 57 3. 29 0. 00 13 .78 0. 72 14 .63 11 .3 9 0. 68 0. 30 0.2 1 96 .3 4 m o n zo n ite 2 49 .0 5 0. 89 4. 35 0. 03 15 .07 0. 93 13 .91 11 .2 7 1.1 4 0. 45 0. 18 97 .2 9 3 50 .01 0.8 1 4. 02 0. 00 14 .7 8 0. 74 14 .1 4 11 .2 4 1.1 4 0. 35 0. 30 97 .5 3 4 50 .8 3 0. 57 3. 30 0. 00 14 .1 0 0. 72 14 .99 11 .35 0. 86 0. 30 0.2 1 97 .2 2 6 50 .61 0.6 1 3. 42 0. 00 14 .21 0. 75 14 .7 7 11 .30 0. 86 0. 33 0. 19 97 .0 3 7 45 .7 3 1.9 1 6. 88 0.0 1 15 .43 0. 68 12 .58 10 .85 1. 80 0.6 1 0.2 1 96 .6 9 8 49 .9 4 0. 75 3. 90 0.0 1 14 .58 0. 68 14 .11 11 .21 0. 98 0. 33 0. 24 96 .7 3 9 49 .3 0 0. 84 4. 29 0. 00 14 .7 6 0. 73 14 .1 2 11 .1 7 1. 06 0. 42 0. 24 96 .9 4 T ab le B. 1 M ic ro pr ob e an al ys es o f a m ph ib ol es in th e ig ne ou s ro ck s o f t he Pi ne -F in -M ex ar ea . R oc k ty pe Sa m pl e n o . Po in t S i0 2 T 10 2 A 1 203 C r 203 Fe O T M nO M gO C aO N a 20 K20 F To ta l G ie ge ric h JD 05 -1 17 a 1 45 .3 8 1.2 5 7. 29 0.0 1 15 .41 1.0 7 12 .3 7 11 .3 6 1.2 5 0. 73 0. 25 96 .3 7 gr an od io rit e 2 44 .9 9 1. 16 7. 10 0. 02 15 .21 1. 02 12 .2 8 11 .50 1.2 5 0. 78 0. 20 95 .5 2 3 48 .2 0 0. 43 4. 93 0. 03 14 .1 3 0. 97 13 .97 11 .5 6 0. 85 0. 45 0. 19 95 .71 4 44 .7 0 1. 28 7. 17 0. 03 15 .23 1. 04 12 .2 2 11 .4 8 1. 28 0. 73 0. 29 95 .4 6 5 44 .8 2 1. 35 7. 49 0. 01 15 .28 1. 09 12 .0 9 11 .4 7 1. 28 0. 75 0. 27 95 .8 9 6 43 .4 0 1. 62 8.3 1 0. 00 15 .83 0. 87 11 .60 11 .5 0 1. 37 1.0 5 0. 30 95 .8 7 7 43 .4 7 1.6 3 8. 37 0. 00 16 .79 0. 84 10 .96 11 .31 1. 40 1.0 3 0.3 1 96 .11 8 43 .6 7 1. 44 8.2 1 0. 00 15 .56 0. 73 12 .05 11 .3 7 1.3 1 0. 97 0. 25 95 .5 7 9 43 .8 1 1. 59 8. 17 0. 00 15 .84 0. 83 11 .7 3 11 .5 4 1.3 3 1.0 3 0. 39 96 .2 5 10 43 .7 3 1.6 3 8. 43 0. 00 16 .20 0.9 1 11 .5 6 11 .3 8 1. 34 1. 04 0. 32 96 .5 3 11 49 .1 1 0. 59 4. 45 0.0 1 13 .23 0. 98 14 .67 11 .69 0. 76 0. 34 0. 23 96 .0 6 12 47 .2 3 0. 84 5. 98 0. 00 14 .88 1.1 1 13 .23 11 .71 1.0 3 0. 62 0. 24 96 .8 7 13 46 .8 5 0. 87 6. 09 0. 00 14 .61 1.1 5 13 .18 11 .6 0 1.0 5 0. 64 0. 38 96 .4 2 14 48 .6 9 0. 49 4. 62 0. 00 13 .83 1.0 5 14 .1 9 11 .7 3 0. 70 0. 35 0. 17 95 .8 2 15 45 .6 0 1. 20 6. 70 0. 00 14 .96 0. 93 12 .97 11 .4 0 1. 05 0. 71 0. 34 95 .8 6 16 45 .0 3 1. 32 7. 13 0.0 1 15 .33 1.0 7 12 .1 5 11 .2 9 1. 19 0. 72 0. 28 95 .5 0 17 44 .1 5 1. 64 7. 87 0. 00 15 .73 0. 92 11 .8 2 11 .3 5 1. 40 0. 82 0. 25 95 .9 4 18 44 .7 4 1. 50 7. 48 0.0 1 15 .01 1.0 2 12 .2 8 11 .4 6 1. 29 0. 82 0. 32 95 .9 2 19 45 .4 8 1. 07 7. 27 0. 00 15 .42 1. 10 12 .41 11 .4 7 1.0 9 0. 72 0. 28 96 .3 2 20 44 .7 4 1. 44 7. 30 0. 00 15 .40 1. 06 12 .1 3 11 .56 1. 27 0. 72 0. 26 95 .9 0 21 43 .7 9 1. 19 8. 68 0. 00 16 .00 0. 84 11 .7 3 11 .7 3 1. 40 0.7 1 0. 25 96 .3 2 22 43 .4 4 2. 04 8.9 1 0.0 0 14 .74 0.8 1 12 .08 11 .55 1.6 5 0.8 5 0.2 6 96 .33 23 42 .0 8 2. 48 9.8 3 0. 00 15 .27 0.8 5 11 .39 11 .16 1.7 1 1.0 4 0.3 1 96 .11 24 42 .6 4 2.0 3 9.0 9 0. 00 16 .07 0.9 4 11 .18 11 .20 1.6 5 0. 79 0.3 0 95 .89 25 41 .4 7 1.7 3 10 .60 0.0 2 16 .53 0.9 2 10 .57 11 .52 1.6 6 0.9 8 0.2 4 96 .25 M ex JD O5 -1 34 1 50 .77 0. 57 3. 29 0. 00 13 .78 0.7 2 14 .63 11 .39 0.6 8 0.3 0 0.2 1 96 .34 m o n zo n ite 2 49 .05 0. 89 4. 35 0.0 3 15 .07 0.9 3 13 .91 11 .27 1.1 4 0.4 5 0.1 8 97 .29 3 50 .01 0.8 1 4. 02 0. 00 14 .78 0.7 4 14 .14 11 .24 1.1 4 0.3 5 0.3 0 97 .53 4 50 .83 0. 57 3. 30 0.0 0 14 .10 0.7 2 14 .99 11 .35 0. 86 0.3 0 0.2 1 97 .22 6 50 .61 0.6 1 3.4 2 0. 00 14 .21 0.7 5 14 .77 11 .30 0. 86 0.3 3 0.1 9 97 .03 7 45 .73 1.9 1 6.8 8 0.0 1 15 .43 0.6 8 12 .58 10 .85 1.8 0 0.6 1 0.2 1 96 .69 8 49 .9 4 0.7 5 3.9 0 0.0 1 14 .58 0.6 8 14 .11 11 .21 0. 98 0.3 3 0.2 4 96 .73 9 49 .3 0 0. 84 4. 29 0. 00 14 .76 0.7 3 14 .12 11 .17 1.0 6 0.4 2 0.2 4 96 .94 T ab le B .1 Co nt . Ro ck ty pe Sa m pl e n o . Po in t S i0 2 T i0 2 A 1 203 C r 203 Fe O T M nO M gO C aO N a 20 K20 F To ta l M ex JD O5 -1 34 10 49 .2 2 0. 93 4. 19 0. 00 15 .18 0. 84 13 .93 10 .97 1. 12 0. 37 0. 18 96 .9 3 m o n zo n ite 11 49 .4 5 0. 90 4. 01 0. 00 14 .65 0. 85 14 .4 7 10 .8 9 1.1 3 0. 38 0.2 1 96 .9 4 12 49 .6 8 0.7 1 3. 86 0.0 1 14 .4 4 0. 79 14 .35 11 .1 5 1. 04 0. 35 0. 33 96 .6 9 13 49 .9 3 0. 93 3. 85 0. 00 14 .7 0 0. 78 14 .1 5 11 .2 8 1.0 1 0. 33 0. 25 97 .2 2 14 49 .5 8 0. 93 4. 17 0.0 1 14 .8 2 0. 68 14 .4 3 11 .1 5 1. 09 0. 38 0. 20 97 .4 3 15 48 .6 7 0. 79 4. 12 0. 05 14 .3 8 0. 80 14 .2 4 11 .3 4 1.0 1 0.4 1 0. 27 96 .0 8 16 47 .7 5 1.6 5 5.3 5 0. 00 12 .3 9 0. 50 14 .81 11 .60 1.2 8 0. 49 0. 20 96 .0 2 17 45 .3 8 2. 10 7. 96 0. 00 12 .57 0. 47 14 .63 10 .92 1. 90 0. 57 0. 17 96 .6 7 18 46 .0 2 2. 05 7. 56 0. 00 12 .4 4 0. 43 14 .90 11 .1 9 1. 79 0. 58 0. 17 97 .1 2 19 45 .0 3 2. 16 7. 89 0. 00 12 .61 0. 48 14 .4 2 11 .0 6 1.9 3 0. 52 0. 25 96 .3 5 20 46 .1 6 1.7 8 6. 73 0. 00 12 .8 4 0. 50 14 .83 11 .1 8 1. 64 0.5 1 0. 25 96 .4 3 21 46 .8 8 1. 26 5. 77 0. 03 15 .28 0. 63 13 .4 0 10 .85 1.5 2 0. 58 0. 19 96 .3 9 22 47 .0 2 1.1 3 5.9 1 0.0 1 14 .67 0. 55 13 .8 6 10 .91 1.4 7 0. 62 0. 18 96 .3 3 23 46 .8 4 1.7 5 6. 33 0. 00 14 .3 8 0. 56 13 .72 10 .86 1.4 9 0. 60 0. 34 96 .8 5 24 46 .2 4 1.6 5 6.5 1 0. 01 14 .5 2 0. 45 13 .8 6 11 .05 1.6 1 0. 63 0. 23 96 .7 4 25 46 .8 0 1.2 4 5. 73 0.0 1 15 .93 0. 85 12 .8 6 11 .07 1.3 5 0. 57 0. 28 96 .6 8 M ex ID 12 OB 2 44 .6 5 1.8 1 7. 95 0. 00 14 .1 5 0. 47 13 .2 2 11 .51 1. 42 0.9 1 0. 15 96 .2 4 sy en ite dy ke 3 44 .3 0 1.7 8 8. 22 0.0 1 14 .2 4 0. 44 13 .1 6 11 .34 1. 44 1.0 1 0. 23 96 .1 7 4 44 .2 1 2. 01 8. 40 0. 00 14 .66 0.5 1 12 .98 11 .4 8 1.4 2 0. 94 0. 18 96 .7 9 5 45 .01 1. 57 7. 52 0. 00 14 .31 0. 54 13 .27 11 .4 9 1.2 8 0. 88 0. 09 95 .9 5 6 44 .9 2 1.7 3 7. 88 0. 02 13 .93 0. 45 13 .33 11 .2 9 1.3 6 0.9 1 0. 14 95 .9 5 7 43 .7 4 1. 83 8. 35 0.0 1 14 .5 0 0. 59 12 .71 11 .4 4 1. 40 0. 90 0. 17 95 .6 4 8 43 .5 9 1. 92 8.5 1 0. 05 15 .04 0. 49 12 .59 11 .2 2 1.3 8 0. 92 0. 10 95 .8 0 9 43 .7 2 1. 89 8. 40 0. 00 14 .93 0. 47 12 .57 11 .36 1.4 2 0. 88 0.2 1 95 .8 3 10 43 .7 1 1. 79 8. 40 0. 00 14 .7 8 0.6 1 12 .60 11 .41 1.4 1 0. 85 0. 12 95 .7 0 11 44 .8 4 1. 68 7. 75 0. 03 14 .1 9 0. 48 12 .90 11 .4 3 1.3 1 0. 92 0. 22 95 .7 5 12 45 .3 8 1. 48 7. 39 0. 00 14 .31 0. 51 13 .43 11 .33 1. 32 0. 86 0. 09 96 .1 0 13 45 .2 4 1. 54 7. 41 0. 00 14 .1 0 0. 47 13 .32 11 .41 1.3 2 0. 83 0.2 1 95 .8 6 14 45 .1 8 1. 46 7. 53 0. 00 14 .7 9 0. 63 13 .13 11 .2 9 1.1 9 0. 78 0. 18 96 .1 5 15 42 .6 4 1. 78 10 .01 0.0 1 15 .65 0. 55 11 .8 0 11 .68 1.6 5 0. 99 0. 18 96 .9 3 16 45 .3 9 1.5 1 7. 48 0. 00 14 .2 6 0. 43 13 .28 11 .62 1. 26 0. 80 0. 22 96 .2 6 17 44 .6 2 1.7 5 7. 94 0. 00 13 .61 0. 50 13 .4 7 11 .1 7 1. 40 0. 81 0. 06 95 .3 3 18 44 .8 9 1. 66 8. 06 0. 02 13 .4 9 0. 55 13 .55 11 .32 1.3 7 0. 83 0.2 1 95 .9 6 T ab le B. 1 C on t. R oc k ty pe Sa m pl e n o . Po in t S i0 2 T i0 2 A 1 203 C r 203 Fe O T M nO M gO C aO N a 20 1( 20 F To ta l M ex JD 12 OB 19 42 .6 1 2. 11 9. 79 0. 00 15 .10 0. 45 12 .0 6 11 .53 1. 62 1. 04 0.1 1 96 .4 3 sy en ite dy ke 20 45 .2 6 1. 64 7. 59 0. 02 13 .58 0. 55 13 .7 0 11 .4 4 1.3 8 0. 85 0. 18 96 .1 8 21 44 .6 7 1. 77 8. 12 0. 00 14 .4 9 0. 52 13 .2 9 11 .54 1. 40 0. 84 0. 16 96 .7 8 22 45 .1 6 1. 64 7. 82 0. 00 14 .36 0. 52 13 .03 11 .2 9 1.3 5 0. 84 0. 06 96 .0 8 23 46 .6 3 1.2 3 6. 39 0. 02 15 .27 0. 80 12 .1 7 11 .56 0. 95 0. 67 0. 06 95 .7 5 24 44 .6 9 1. 74 8. 18 0.0 1 14 .1 6 0. 49 13 .2 8 11 .38 1. 39 0. 77 0. 16 96 .2 6 25 45 .4 4 1. 64 7. 62 0. 02 14 .4 6 0. 50 13 .2 4 11 .3 4 1. 28 0. 82 0. 17 96 .5 4 26 45 .2 3 1.5 5 7. 65 0. 00 14 .07 0. 53 13 .3 9 11 .4 5 1. 34 0.8 1 0. 09 96 .1 0 Fi n F0 5- 05 16 6 1 43 .8 9 1. 94 8. 25 0.0 1 13 .7 0 0. 55 13 .1 3 11 .2 8 1.4 8 0. 88 0. 08 95 .21 sy en ite dy ke 2 44 .3 5 1.8 5 8.0 1 0. 03 13 .82 0. 55 13 .4 4 11 .4 5 1.4 3 0. 95 0. 28 96 .1 6 3 42 .9 3 2. 00 9.2 1 0.0 1 14 .3 0 0. 45 12 .6 9 11 .4 4 1. 60 1. 06 0. 23 95 .9 2 4 46 .2 5 1.3 5 6. 43 0. 00 13 .53 0. 57 13 .5 2 11 .7 4 1.3 5 0. 78 0. 26 95 .7 7 5 44 .7 6 1. 70 7. 76 0. 00 13 .42 0. 53 13 .43 11 .3 6 1. 44 0. 94 0. 23 95 .5 7 6 44 .7 9 1. 82 7. 77 0. 04 13 .25 0. 57 13 .67 11 .35 1. 42 0. 89 0. 15 95 .7 2 7 44 .9 2 1. 92 8. 20 0. 00 13 .63 0.5 1 13 .43 11 .4 3 1. 56 0. 84 0. 28 96 .7 2 8 45 .1 9 1. 78 7. 68 0. 00 13 .37 0. 47 13 .7 0 11 .3 9 1. 44 0. 87 0. 20 96 .0 8 9 45 .0 7 1. 77 7. 78 0. 00 13 .29 0.5 1 13 .62 11 .3 0 1. 40 0. 86 0. 19 95 .7 9 10 44 .3 2 1. 70 7. 53 0. 00 13 .66 0. 52 13 .73 11 .2 8 1. 36 0. 86 0. 22 95 .1 8 11 44 .0 9 2. 17 8. 53 0. 00 14 .1 7 0. 50 12 .8 9 11 .3 0 1.6 4 0.9 1 0. 19 96 .3 8 12 44 .2 7 1. 64 7. 56 0.0 1 13 .68 0.5 1 13 .7 7 11 .52 1. 40 0. 86 0. 26 95 .4 7 13 44 .6 2 1.8 3 7. 68 0.0 1 14 .06 0. 55 13 .4 0 11 .4 2 1.3 7 0.9 1 0. 20 96 .0 6 14 45 .3 0 1.6 3 7. 42 0. 00 13 .53 0. 56 13 .95 11 .4 8 1.3 5 0. 82 0. 25 96 .2 9 15 46 .1 7 1.7 1 7. 40 0. 03 13 .23 0.5 1 14 .0 0 11 .3 9 1.4 2 0. 77 0. 39 97 .0 2 16 45 .1 1 1. 68 7. 59 0. 04 13 .24 0. 50 13 .7 6 11 .41 1.4 3 0. 83 0. 18 95 .7 8 17 45 .1 0 1.7 3 7. 73 0. 00 13 .36 0. 49 13 .73 11 .33 1. 42 0. 85 0. 15 95 .8 9 18 45 .1 0 1. 67 7. 42 0.0 1 13 .21 0. 49 14 .0 7 11 .4 6 1.4 1 0. 78 0. 23 95 .8 5 19 47 .0 6 1. 27 5.7 1 0. 00 14 .59 0. 57 13 .31 11 .68 1. 00 0. 68 0. 14 96 .0 2 20 44 .7 2 1. 94 8. 07 0. 00 13 .7 6 0. 50 13 .4 9 11 .51 1. 49 0. 86 0. 32 96 .6 5 21 44 .5 3 1.9 1 8. 12 0. 03 13 .47 0. 47 13 .2 9 11 .52 1. 46 0. 89 0. 23 95 .9 2 22 44 .2 8 1. 94 8. 15 0. 00 13 .64 0. 55 13 .33 11 .3 7 1. 50 0. 87 0. 38 96 .0 0 23 44 .5 9 1. 82 7. 79 0. 00 13 .21 0.5 1 13 .71 11 .1 3 1. 49 0. 85 0. 16 95 .2 6 24 45 .3 7 1.8 1 7. 58 0. 02 13 .4 9 0. 46 13 .7 4 11 .3 8 1. 38 0.9 1 0. 29 96 .4 3 25 44 .9 1 1.7 5 7. 82 0. 02 13 .88 0. 49 13 .63 11 .4 7 1. 40 0. 94 0. 22 96 .5 4 T ab le B. 1 C on t. Ro ck ty pe Sa m pl e n o . Po in t S i0 2 T i0 2 A 1 203 C r 203 Fe O T M nO M gO C aO N a 20 K20 F To ta l Pi ne 92 -3 8 19 5 1 46 .9 1 1.7 3 7. 04 0. 04 12 .3 2 0. 54 14 .4 4 11 .0 8 1. 47 0. 76 0. 29 96 .6 3 sy en ite dy ke 2 46 .1 9 1.7 8 7. 09 0. 00 12 .62 0. 60 14 .59 11 .3 9 1. 47 0. 78 0. 19 96 .7 0 3 46 .0 9 1. 72 6.9 1 0. 03 12 .2 3 0. 52 14 .7 6 11 .11 1. 46 0. 65 0. 38 95 .8 6 4 46 .0 1 1. 74 6. 99 0. 00 12 .8 5 0. 52 14 .4 5 11 .0 2 1. 46 0. 68 0. 38 96 .1 0 5 44 .91 1.9 5 7. 43 0. 00 12 .61 0. 43 14 .31 11 .21 1. 60 0. 77 0. 33 95 .5 4 6 44 .0 6 2. 21 8. 48 0. 02 13 .23 0. 37 13 .8 4 11 .2 6 1. 69 0. 78 0. 26 96 .2 0 7 45 .2 7 1.9 1 7. 22 0. 00 12 .35 0. 40 14 .57 11 .1 6 1. 56 0. 67 0. 34 95 .4 5 8 43 .4 0 2. 17 8. 53 0. 02 13 .13 0. 45 13 .73 11 .1 0 1.8 4 0. 86 0.3 1 95 .5 4 9 45 .0 9 2. 00 7. 54 0. 00 12 .4 4 0. 39 14 .4 0 11 .06 1. 60 0.7 1 0. 33 95 .5 5 10 44 .1 8 2. 13 8. 05 0. 00 13 .2 3 0. 38 13 .99 11 .1 0 1.7 1 0. 84 0. 26 95 .8 7 11 44 .1 2 1.7 8 8. 58 0. 00 12 .6 6 0. 45 14 .3 4 11 .35 1.6 8 0. 72 0. 35 96 .0 2 12 46 .1 3 1.3 7 6.5 3 0.0 1 12 .8 7 0. 50 14 .68 11 .3 4 1. 29 0. 69 0. 18 95 .5 7 13 45 .3 6 1.2 6 6. 67 0. 00 12 .4 5 0. 48 14 .7 0 11 .4 7 1.3 3 0. 69 0. 28 94 .6 9 14 44 .5 2 1. 77 7. 64 0. 00 13 .48 0. 49 13 .79 11 .4 5 1. 49 0. 85 0. 34 95 .8 2 15 44 .0 9 1.8 4 8. 70 0. 00 12 .71 0. 47 13 .94 11 .4 2 1.7 3 0. 69 0. 35 95 .9 3 16 46 .4 9 1. 26 6. 40 0. 00 13 .43 0. 73 14 .05 11 .52 1. 30 0. 66 0. 27 96 .11 17 46 .0 7 1.2 3 6. 23 0. 00 13 .13 0.7 1 13 .94 11 .2 4 1.2 8 0. 68 0. 12 94 .6 3 18 46 .8 9 1.1 5 6. 25 0. 02 12 .7 7 0. 79 14 .4 9 11 .32 1.2 8 0. 62 0. 25 95 .8 3 19 46 .4 8 1. 59 6. 70 0. 00 12 .81 0. 55 14 .4 3 11 .3 6 1. 34 0. 74 0. 26 96 .2 7 20 46 .0 0 1.4 5 6.7 1 0. 02 12 .9 9 0. 69 14 .2 6 11 .4 8 1.4 2 0. 66 0.3 1 95 .9 8 21 45 .8 2 1.2 2 6. 53 0. 00 12 .97 0. 77 14 .63 11 .4 7 1.3 2 0. 62 0.3 1 95 .6 5 22 46 .0 8 1.2 3 7. 03 0. 02 13 .3 2 0. 68 13 .99 11 .50 1.4 2 0. 65 0.2 1 96 .1 4 23 45 .31 1.2 1 7. 55 0. 00 13 .4 7 0. 72 13 .94 11 .4 2 1. 46 0. 75 0. 22 96 .0 5 24 45 .3 6 2. 17 6. 77 0. 00 12 .92 0. 65 14 .35 11 .41 1. 49 0. 64 0. 40 96 .1 4 25 46 .8 2 1.3 3 6.5 1 0.0 1 12 .7 6 0. 56 14 .81 11 .2 2 1. 40 0. 69 0. 23 96 .3 3 C ’

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