Open Collections

UBC Theses and Dissertations

UBC Theses Logo

UBC Theses and Dissertations

The southwest zone breccia-centered silica-undersaturated alkalic porphyry Cu-Au deposit, Galore Creek,… Byrne, Kevin 2009

Your browser doesn't seem to have a PDF viewer, please download the PDF to view this item.

Notice for Google Chrome users:
If you are having trouble viewing or searching the PDF with Google Chrome, please download it here instead.

Item Metadata

Download

Media
24-ubc_2009_spring_byrne_kevin.pdf [ 7.04MB ]
Metadata
JSON: 24-1.0052382.json
JSON-LD: 24-1.0052382-ld.json
RDF/XML (Pretty): 24-1.0052382-rdf.xml
RDF/JSON: 24-1.0052382-rdf.json
Turtle: 24-1.0052382-turtle.txt
N-Triples: 24-1.0052382-rdf-ntriples.txt
Original Record: 24-1.0052382-source.json
Full Text
24-1.0052382-fulltext.txt
Citation
24-1.0052382.ris

Full Text

The Southwest Zone breccia-centered silica-undersaturated alkalic porphyry Cu-Au deposit, Galore Creek, B.C: Magmatic-hydrothermal evolution and zonation, and a hydrothermal biotite perspective by Kevin Byrne BA. (Hons.) Mod Geology, Trinity College, Dublin, 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 UNIVERSITY OF BRITISH COLUMBIA (Vancouver) April, 2009 © Kevin Byrne 2009 ABSTRACT Situated in northwest B.C Canada, the Southwest Zone Cu-Au breccia-centered deposit is one of twelve mineralized centers in the Galore Creek alkalic porphyry district. Formed in an island arc-setting outboard of ancestral North America in the lateTriassic, deposits in the Galore Creek district have a combined measured and indicated resource of 785.7 Mt at 0.52 per cent Cu, 0.29 g/t Au and 4.87g/t Ag. Mineralisation in the Southwest Zone is centered in narrow hydrothermally cemented breccias. Composite dikes of megacrystic orthoclase-phyric syenite and megacrystic orthoclase and plagioclase-phyric monzonite are cut by polylithic, poorly sorted pebble- cobble clast size matrix-rich breccias. Hydrothermally cemented breccias, characterised by the cement (infill) assemblage phlogopite ± K-feldspar ± magnetite ± anhydrite ± diopside ± suiphide, overprint the western contact between matrix-bearing breccias and megacrystic porphyries. Coeval with cemented breccia formation, intrusions of biotite phyric monzodiorite occur at the matrix-bearing breccia-wall rock contact and in the matrix-bearing breccias. Biotite-phyric monzodiorite and the principal cemented breccia domains are co-spatial and syn-Cu-Au. Drilling has outlined a zone of >0.3% hypogene Cu approximately 20-lOOm thick, 500m wide and 400m in length that strikes l0O, dips 45-60°S, and has a semi- ellipsoidal morphology. This mineralisation is coincident with potassic (stage D) alteration and infill. Cu-poor, diopside-dominated (calc-[potassic]) alteration formed contemporaneously with, and locally flanks, potassic-D infill. Sulphide minerals are zoned from a core of chalcopyrite-bomite, to chalcopyrite>pyrite, to pyrite>chalcopyrite out to pyrite only. Garnet-bearing peripheral propylitic alteration overlaps with a pyrite and Au-halo and locally overprints potassic and calc-Q,otassic) assemblages. Based on electron microprobe analysis, systematic spatial variations in Ti-content and Fe3/F2of infill biotite are evident. Increases in Ti-contents and Fe317Fe2overlap with positive gradients in Cu concentration, taken with interpreted alteration reactions, this suggest Cu-deposition is caused by decreasing 102 coupled with an increase in pH at 420-475°C. Low logQH2O/JHF), determined from infill biotite, distinguish potassic fluids in the Southwest Zone, and other alkalic porphyry deposits, from fluids in calc alkalic systems and reflects the contrasting magmatic composition. 11 TABLE OF CONTENTS ABSTRACT.I TABLE OF CONTENTS iii LIST OF TABLES vi LIST OF FIGURES vii ACKNOWLEDGMENTS xi CHAPTER 1: OVERVIEW I 1.1 Rationale for Study I 1.2 Thesis organization 2 1.3 Regional Geological Setting 3 1.4 Galore Creek District 5 1.4.1 Geography 5 1.4.2 Exploration History 5 1.4.3 Supracrustal Rocks 5 1.4.4 Igneous Rocks 8 1.4.5 Hydrothermally Altered and Mineralized Centers 11 1.4.6 Structural History 11 1.4.7 Southwest Zone 12 1.5 Breccias 15 1.5.1 Approach and Nomenclature 15 1.5.2 Fragmentation Mechanisms and Breccia Classification 15 1.5.3 Principal Breccia Fades Characteristics 17 1.6 Biotite 22 1.6.1 Crystal-chemistry 22 1.6.2 Normalization Schemes for Microprobe Data 24 1.6.3 Mechanisms of Ti Al and Fe3+ Incorporation 25 1.7 Biotite in the Porphyry Environment 27 1.7.1 Example Studies 27 1.7.2 Geothermometry 29 1.7.3 Oxygen Fugacity Estimates 31 1.7.4 Halogen and Water Fugacity Estimates 31 1.8 Research Objectives 32 1.9 Study Methodology 33 1.10 References 35 111 CHAPTER 2: MAGMATIC-HYDROTHERMAL EVOLUTION AND ZONATION OF A BRECCIA CENTERED CU-AU ALKALIC PORPHYRY: SOUTHWEST ZONE, GALORE CREEK 43 2.1 Introduction 43 2.2 Exploration History 46 2.3 Regional Geological Setting 46 2.4 District Geology 47 2.5 Rock types of the Southwest Zone 51 2.5.1 Coherent rocks 52 2.5.2 Clastic rocks 60 2.5.3 Rock paragenesis 67 2.6 Structural controls on rock distribution 68 2.7 Alteration 69 2.7.1 K-feldspar ± biotite ± hematite dusting (potassic-A and-B)-Stage-1 70 2.7.2 Phlogopite ± chlorite ± magnetite (potassic-C)-Stage-1 70 2.7.3 Phlogopite ± magnetite ± K-feldspar (potassic-D)-Stage-2 70 2.7.4 Diopside ± magnetite ± phiogopite (calcic-fpotassicD-Stage-2 72 2.7.5 K-feldspar-anhydrite (waning potassic)-Stage-2 76 2.7.6 Sericite ± anhydrite (phyllic)-Stage-3 76 2.7.7 Garnet ± chlorite (calcic)-Stage-3 77 2.7.8 Chlorite ± epidote ± calcite ± pyrite (propyiltic)-Stage-4 79 2.7.9 K-feldspar ± Fe-carbonate (carbonate-potassic)-Stage-4 80 2.7.10 Quartz (quartz veins)-Stage-4 81 2.8 Sulphide minerals 82 2.8.1 Potassic-D and calcic-fpotassicJ-Stage-2 82 2.8.2 Late-stage K-feldspar ± Fe-carbonate-Stage-4 84 2.9 Structural controls on alteration and mineralisation 84 2.9.1 Pre to syn Stage 2 and 3 84 2.9.2 Post-Stage 2 and 3 85 2.10 Oxygen and hydrogen isotope data 86 2.11 Metals Zoning 88 2.11.1 Copper, Gold and Silver 88 2.11.2 Lead and Zinc 90 2.11.3 Molybdenum 91 2.12 Discussion and Genetic Interpretation 92 2.12.1 Clastic rocks 92 2.12.2 Isotopic composition of hydrothermal fluids 95 2.12.3 Alteration zoning (Stages 2 and 3) 98 2.12.4 Metal transport, deposition and zoning 103 2.12.5 Palaeo-geometry and deposit model 111 2.13 Conclusions 116 iv 2.14 References .118 CHAPTER 3: COMPOSITION OF BIOTITE FROM THE SOUTHWEST ZONE ALKALIC PORPHYRY CU-AU DEPOSIT, GALORE CREEK, BC, CANADA: EVALUATION OF HYDROTHERMAL FLUID CHEMISTRY 124 3.1 Introduction 124 3.2 Geological framework 126 Regional: 126 District: 127 Deposit: 129 3.3 Biotite types classification 131 3.4 Analytical procedures and sample methodology 137 3.5 Normalization of microprobe data and estimation of Fe3 139 3.6 Biotite composition by type 142 3.7 Spatial variation in Ti and Fe3 145 3.8 Biotite halogen chemistry 150 3.9 Hydrothermal fluid halogen fugacity ratio estimates 153 3.10 Comparison of fugacity ratios with other porphyry systems 155 3.11 Discussion 158 3.11.1 Significance of gradients in infil biotite chemistry 158 3.11.2 Halogen chemistry 160 3.12 Conclusions 161 3.13 References 162 CHAPTER 4: CONCLUSIONS 167 4.1 Recommendations for future work (and other conjectures) 168 4.2 References 170 APPENDICIES A-D on disc I in sleeve and is also available at http:llwww.mdru.ubc.ca V LIST OF TABLES TABLE 1.1 Summary of Galore Creek intrusive rocks (Enns, et. al., 1995) 10 TABLE 1.2 Average conserved element ratios from Galore Creek intrusive rocks (Enns, et. al., 1995) 10 TABLE 1.3 Genetic classes and associated fragmentation processes, adapted from Sillitoe (1985), Davies (2002) and Davies et al (2008b) 20 TABLE 1.4 Common porphyry-system breccia facies and their descriptive criteria adapted from Seedorif et al. 2005, Sillitoe (1985), and Davies (2002) 21 TABLE 1.5 Substitution mechanisms for Ti4,Al3 and Fe3 in biotite 26 TABLE 2.1 List of abbreviations used in figures and tables 45 TABLE 2.2 Summary of Southwest Zone coherent rocks 53 TABLE 2.3 Summary of Southwest Zone clastic rocks 61 TABLE 2.4 Paragenetic stages of coherent and clastic rocks in the Southwest Zone 67 TABLE 2.5 Stable isotope results for Stage 2 infill phlogopite and phyllic stage sericite 86 TABLE 3.1 Hydrothermal apatite compositions from vein samples in the Southwest Zone 137 TABLE 3.2 Southwest Zone biotite data distribution 138 TABLE 3.3 Substitution mechanisms for Ti4,Al3 and Fe3 in biotite 140 TABLE 3.4 Biotite Thomson-space components 141 TABLE 3.5 Representative biotite compositions from the Southwest Zone by textural type 143 TABLE 3.6 Titanium, Fe3and Fe3/Fe2 values for infill biotite 146 vi LIST OF FIGURES FIG. 1.1 Map of British Columbia showing the location of the accreted Quesnellia and Stikinia ocean arc terranes, major alkalic Cu-Au porphyry deposits and alkalic intrusive centers, morphogeological belts and Galore Creek 4 FIG. 1.2 A. Major tectono-stratigraphic elements and the location of the Galore Creek district, Eskay Creek, Red Chris, Schaft Creek and Copper Canyon in the Stikinia terrane 6 FIG. 1.3 Simplified geological and alteration-mineralisation map of the Galore Creek district 14 FIG. 1.4 Ferromagnesian biotite with labeled structural sites 23 FIG. 1.5 Biotite composition trends in an idealized porphyry alteration model 29 FIG. 2.1 Map of British Columbia showing the location of the accreted Quesnellia and Stikinia ocean arc terranes, major alkalic Cu-Au porphyry deposits and alkalic intrusive centers, morphogeological belts and Galore Creek 44 FIG. 2.2 A. Major tectono-stratigraphic elements and the location of the Galore Creek district, Eskay Creek, Red Chris, Schaft Creek and Copper Canyon in the Stikinia terrane 47 FIG. 2.3 Simplified geological, and alteration and mineralisation map of the Galore Creek district illustrating the location of mineralized centers and major structural features 49 FIG. 2.4 Photographs of the Galore Creek valley, looking A south, and B north with labeled mineralized centers 50 FIG. 2.5 Simplified bedrock geology plan map and Cu-grade distribution in the Southwest Zone. Location of section line A-A’ (6333650N) and B-B’ (350030E) indicated with black lines 52 FIG. 2.6 Sequence of coherent and clastic rock emplacement in the Southwest Zone 54 FIG. 2.7 Photographs of coherent rocks in the Southwest Zone 56 FIG. 2.8 Distribution of clastic facies and simplified coherent units along cross section 6333650N, line A-A’ 57 FIG. 2.9 Distribution of clastic facies and simplified coherent units along cross section 350030E along line B-B’ 58 FIG. 2.10 Photographs of diagnostic features in biotite-phyric monzodiorite dike and dikelet facies 59 FIG. 2.11 Photographs of diagnostic features in matrix-bearing breccias 63 FIG. 2.12 Photographs of cemented breccia facies in the Southwest Zone: MC-BX and CM-BX 65 FIG. 2.13 Photographs of monolith in-situ cement dominated breccias (C-BX) 66 FIG. 2.14 Post-cemented breccia facies coherent units 68 vii FIG. 2.15 Paragenetic stages of coherent and clastic rocks, and alteration and mineralisation facies in the Southwest Zone 69 FIG. 2.16 Photographs of Stage 2 and phyllic alteration and infill 73 FIG. 2.17 Distribution and abundances of potassic-D and calcic-(potassic) and waning potassic hydrothermal minerals (Stage 1) and the> 0.3% Cu shell along section lines A-A’ and B-B’ 75 FIG. 2.18 Photographs and photomicrographs of Stage 3 and 4 alteration facies and veins 78 FIG. 2.19 Distribution and abundances of Stage 3 calcic and propylitic alteration minerals on section lines A-A’ and B-B’ 79 FIG. 2.20 Distribution of late-stage carbonate-potassic veins and alteration, hydrothermal quartz and Au intercepts> 3 g/t along section B-B’ 81 FIG. 2.21 Photomicrographs of sulphide and oxide mineralisation in the Southwest Zone 82 FIG. 2.22 Pattern of Stage 2 and 3 sulphide mineral distribution along section lines A-A’ and B-B’. 83 FIG. 2.23 Mineral 6018 and 6D values from the Southwest Zone 87 FIG. 2.24 A. Copper grade shells and hydrothermal diopside along sections A-A’ and B-B’ 89 FIG. 2.25 A. Lead, > 0.3 % Cu and > 0.1 g/t Au grade shells and distribution of moderately intense propylitic alteration along sections A-A’ and B-B’ 91 FIG. 2.26 A. Calculated isotopic compositions of water in equilibrium with infill phlogopite and, B. sericite from the Southwest Zone, with fluid compositions associated with potassic and phyllic alteration in other porphyry systems also shown 96 FIG. 2.27 A-C. Activity-activity diagrams showing stability of silicate minerals at 350°C, 500 bars pressure, modified from Beane and Titley, (1981) and Beane, (1982) 99 FIG. 2.28 Log f(o2) — pH diagram at 350°C showing solubility contours for Cu and Au and the stability fields of minerals in the Cu-Fe-C-System (redrafted and modified from Huston et al., 1993) 108 FIG. 2.29 Log f(02) — temperature diagram for solubility of Au as chloride and bisulfide complexes in the porphyry Cu environment (redrafted and modified from Huston and Large, 1989 and Jones, 1992) 110 FIG. 2.30 Stylized perspective projections of joined sections A-A’ and B-B’ in approximated palaeo-geometry (-60° tilt). A. Outline of cemented breccia domains projected beyond current erosion surfaces. B. Domains of intense potassic-D alteration comprising and calcic-(potassic) veins and alteration 112 FIG. 2.31 A-E. Time-integrated cartoons illustrating the evolution of the Southwest Zone breccia complex and paragenesis of Stage 1- early Stage 2 alteration events 113 vii’ FIG. 2.32 Time-integrated alteration and mineralisation schematic in estimated palaeo-geometry. A. Incipient stage 2 alteration; potassic-D and calcic-(potassic) alteration and infill form contemporaneously. B. Alteration and Cu distribution at the end of stage 2 and start of stage 3. Minor phyllic alteration develops in fractures zones at upper levels in the system. C. Distribution of propylitic alteration facies, Au-halo, and Cu, Zn and Pb contours at the end of stage 3 114 FIG. 3.1 Map of British Columbia and the Canadian Cordillera showing the location of the accreted Quesnellia and Stikinia ocean arc terranes, morphogeological belts and Galore Creek 126 FIG. 3.2 Regional scale geology of Galore Creek showing the location of Copper Canyon alkalic porphyry Cu-Au occurrence 128 FIG. 3.3 Simplified geological map and location of mineralised centers in the Galore Creek District 129 FIG. 3.4 Simplified geology along cross sections 6333650N (A-A’) and 350030E (B-B’) 131 FIG. 3.5 Photomicrographs least altered igneous and secondary biotite 133 FIG. 3.6 Infill biotite vein in megacrystic orthoclase-phyric syenites 134 FIG. 3.7 A. Plane light and B. polarized photomicrographs of one side of an infill biotite vein... 135 FIG. 3.8 Back scattered electron images of infill biotite 136 FIG. 3.9 XMg versus atoms per formulae unit AI(total), Si, Fe(total) Fe3IFe2, Ti, Cl, XF, Mn, and Ba for least-altered, secondary and infill biotite in polished thin sections from the Southwest Zone. 142 FIG. 3.10 Compositional variation of the cluster of infill biotite grains indicated in figure 3.6A. .145 FIG. 3.11 Ti % (norm) values of infill biotite, distribution of calc-silicate alteration and infill, and Cu concentrations on cross sections 6333650N (A-A’) and 350030E (B-B’) in the Southwest Zone. 147 FIG. 3.12 Fe3 I Fe2 of infill biotite, distribution of calc-silicate alteration and infill, and Cu concentrations on cross sections 6333650N (A-A’) and 350030E (B-B’) in the Southwest Zone. 148 FIG. 3.13 A. Ti vs. Fe3IFe2 and B. Ti vs. Si. All data in atoms per formula unit and normalized to scheme-B 149 FIG. 3.14 A. XFe versus log (XF/XOH) of infill biotite grains with 4 or more spot analysis per sample 151 FIG. 3.15 Reliable infill biotite data and the slope of a line (in red) calculated from equation (24) in Zhu and Sverjensky (1992) at 420°C 153 FIG. 3.16 A. XFe versus lo XF/XoH) and calculated log (fH2oIfHF) fluid B. XMg versus log (XFIXCI) and calculated log (fHFIfHCI) UI of repesentitive infill biotite anlayses 154 FIG. 3.17 A. Log(fH2O/fHF) versus log(fH2O/fHCI), and B. log(fHF/fHCl) versus log(fH2O/fHCl) ratios for potassic fluids in the Southwest Zone. Also shown are the fugacity values determined for other porphyry Cu systems (Selby and Nesbit, 2000; Twelker, 2007; Kroll et al., 2002) 155 ix FIG. 3.18 Comparison of secondary and least altered igneous biotite from the Southwest Zone and from Mo, W and Cu porphyry deposit types, modified after Brimhall and Ague (1988) 157 FIG. 3.19 Log f(o2) — pH diagram at 350°C showing solubility contours for Cu and the stability fields of minerals in the Cu-Fe-O-System, modified from Huston et al. (1993) 159 x ACKNOWLEDGMENTS This work forms part of a larger MDRU-CODES research project investigating shallow and deep-level alkalic mineral deposits, in collaboration with the GSC. Funding and field support is provided by Amarc Resources Ltd., Anglogold-Ashanti, Barrick Gold Corp., Imperial Metals Corp., Lysander Minerals Corp., Newcrest Mining Ltd., Newmont Mining Corp., Novagold Resources Inc., Teck Ltd., NSERC-CRD and Geoscience BC. Logistical and field support in 2007 from Novagold Resources Inc. and in 2008 the Galore Creek Mining Company is also acknowledged. Additional funding for the Southwest Zone project was supplied through SEG student research grants. I am particularly grateful to Geoscience BC and the Society of Economic Geologist student scholarships which facilitated my research and studying here in Canada. I would like to thank Dr. Richard Tosdal for guidance and supervision throughout this project. Additional thanks goes to my other committee member Dr. Greg Dipple and my external examiner Dr. Craig Hart. Dr. Claire Chamberlain, Dr. Kirstie Simpson, and Dr. Thomas Bissig were an integral part of the alkalic group and inexhaustible resources to the students involved in the project, thank you. Additional thanks to Dr. David Cooke (CODES) for his guidance and insight full suggestions and to Arne Toma (MDRU resource centre coordinator) for his help with practicalities. Editorial input from Dr. Claire Chamberlain, Dr. Kirstie Simpson and Dr. Andrew Davies, and their breccia discussion meetings greatly improved this thesis. Cheers to my peers for their support and encouragement during my stay at UBC, and for your technical discussions and assistance. I have learned from all of you. Particular thanks to my project and field partner Janina Micko, Dan McNeil for his helpful discussions in the pub, Chris Leslie for teaching me the ropes on the polisher (a little too well) and Curtis Brett for his trouble shooting and rubber arm. xi Chapter 1: Overview 1.1 Rationale for Study Characteristics of caic-alkalic porphyry deposits and the processes involved in their formation have been well documented and summarised succinctly by Titley and Beane (1981), Seedorff et al. (2005) and Holiday and Cooke (2007). In contrast, alkalic porphyry deposits are relatively under reported and no unifying model for alkalic systems exits. Work by Lang et al. (1995a and b), Jensen and Barton (2000) and Cooke et al. (2007) highlights the economic significance of alkalic porphyry Cu-Au deposits and note the subtle but significant differences from calc-alkalic systems, and variations within the porphyry class itself. Alkalic porphyry Cu-Au deposits are known in only a few metallogenic terranes, notably the Triassic and Jurassic Stikinia-Quesnellia marine volcanic arcs of British Columbia (Fig. 1.1), Canada (Barr et al., 1976; Lang et al., l995a), the Ordovician and early Silurian Lachlan Fold Belt in New South Wales, Australia (Cooke et al., 2007), Early Oligocene continental arc in the Chalkidiki Peninsula, Greece (Kroll et al., 2002), and in the Caraballo Mountains, North Luzon, Philippines (Wolfe et al., 1999). In order to construct models that better our understanding of alkalic porphyry systems, and facilitate exploration programs, integration of detailed structural, paragenetic, alteration zonation and geochemical information from individual deposits is essential. The Galore Creek district is hosted in the Stikinia island arc terrane. Twelve alkalic porphyry Cu-Au centers are currently recognized in the district, five of which have been delineated as deposits. Additionally, these alkalic porphyry Cu-Au centers represent a silica-undersaturated end member of this porphyry class (Lang et al., 1 995b). Combined, deposits in the Galore Creek district compose the largest undeveloped Cu-Au resource in the Stikinia-Quesnellia arc; total measured and indicated resources are 785.7 Mt at 0.52% Cu, 0.29g/t Au. The Central Zone deposit host more than half of the resource in the district with the remainder made up of smaller peripheral deposits. The 1 Southwest Zone deposit, the subject of this M.Sc, is one of the smaller deposits and provides an important case study in the alteration and sulfide mineral zonation, and the role of breccias in alkalic porphyry systems. Complicated zones of quartz-deficient alteration and an Au-halo characterize the deposit (Enns, et al., 2005; Schwab, et a!., 2008). Research aims are to characterize and map the three-dimensional distribution of coherent and clastic rocks, alteration, mineralization, and metal assemblages, and their paragenesis. Traditional studies of porphyry Cu deposit igneous and hydrothermal evolution rely heavily on deciphering host rock and intrusions and cross-cutting quartz vein relationships. Study of silica-undersaturated alkalic systems are challenged for several reasons: they lack quartz-phyric rocks; are characterised by complicated sedimentary and hydrothermal breccia facies; and host quartz-deficient veins and replacement alteration. The absence of quartz in the Galore Creek hydrothermal systems (Enns et al., 1995; Lang et al., 1 995c, Schwab et al., 2008), therefore, inhibits traditional methods of investigating porphyry system hydrothermal fluids, i.e. salinity, P and T, ö’80 and ÔD isotope composition from fluid inclusions in quartz. Alternative avenues of research and techniques are required to comment on the nature and evolution of hydrothermal fluids in silica deficient systems. Biotite monitors the physiochemical environment in which it grows (Beane, 1974; Guidotti, 1984; Speer, 1984). The chemical composition of biotites in the Southwest Zone is evaluated with respect to the geological framework and used to characterize and describe the magmatic-hydrothermal evolution of the Southwest Zone. Spatial variation in vein and cement biotite is used to elucidate processes involved in potassic alteration and Cu deposition. 1.2 Thesis organization Chapter 1 is a literature based review intending to provide the reader with additional background information on the Galore Creek district, breccias and biotite in the porphyry environment. Research objectives and study methodologies conclude this opening chapter. Chapter 2 is written as a stand alone manuscript intended for publication. In Chapter 2 the character, distribution, and timing of emplacement of coherent units and 2 elastic rocks is described with representative photos and illustrations. Alteration and suiphide mineral types, style, intensity and distribution are characterized and mapped, with an emphasis on paragenesis. Deposit scale zoning of Cu, Au, Ag, Pb, Zn and Mo is described and reviewed with respect to alteration and mineralisation assemblages and paragenesis. Chapter 2 culminates with a discussion of diagnostic features of the Southwest Zone and a proposed deposits model. The analytical component of this research is presented in Chapter 3. This chapter presents an evaluation of electron-microprobe analysis of biotite from the Southwest Zone and is also written as an independent document for publication. Spatial variations of vein and cement biotite composition are investigated and related to the distribution and concentration of Cu. These compositional variations and inferred wall-rock-fluid reactions are taken together to discuss and propose processes of Cu deposition. Additionally, fluid halogen fugacity estimates, based on biotite composition, are used to characterize potassic fluids in the Southwest Zone and also to distinguish alkalic Cu-Au porphyry systems from their caic-alkalic brethren. Conclusions drawing from data and interpretations presented in preceding chapters, and areas for future research are presented in Chapter 4. 1.3 Regional Geological Setting A collage of allochthonous oceanic and proximal to distal pericratonic terranes were accreted to the western margin of the North American craton during the Late Paleozoic through to Late Mesozoic (Monger, 1977; Monger and Irving, 1980; Monger et al., 1982; Coney, 1998). From east to west, the Foreland, Omineca, Intermontane, Coast, and Insular belts are the five morphogeological belts that compose the Canadian Cordillera (Gabrielse, et al., 1991) (Fig. 1.1). McMillan (1991) groups Stikinia, Cache Creek, Slide Mountain and parts of Quesnellia and Yukon-Tanana terranes into the Intermontane Superterrane. Similarities in rock type and geologic history between Stikinia and Quesnellia have led workers to believe that they are segments of the same Triassic arc (Wernicke and Klepacki, 1988; Nelson and Mihalynuk, 1993; Mihalynuk et al., 1994). The Galore Creek district is one of seven major mineralized alkalic porphyry systems in the Stikinia-Quesnellia arc: Afton/Ajax, Copper Mt./Ingerbelle, Mt. Milligan, 3 Lorraine, Mt. Polley, and possibly Red Chris. Theses B.C. alkalic deposits formed in two pulses between 210 and 180 million years ago (Mortensen et al., 1995) within the conjoined Quesnellia-Stikinia intraoceanic island arc. At that time, Atlantic rifting and spreading was underway causing ancestral North America to be displaced westward toward the offshore terrane complex (McMillan, 1991). Convergence and collision of the island arc with the western continental margin formed central B.C. and the Cordillera. The tectonic complexities therein resulted in increased (and petrologically unique) plutonic and volcanic activity of alkalic affinity. FIG 1.1 Map of British Columbia showing the location of the accreted Quesnellia and Stikinia ocean arc terranes, major alkalic Cu-Au porphyry deposits and alkalic intrusive centers, and morphogeological belts. Galore Creek is located at the western margin of the Intermontane Belt about 60km east of the US border. Data sourced from BC MapPlace (http://www.mapplace.caf). 1. Red Chris 2. Rugged Mountain 3. Ten Mile Creek 4. Butterfly Pluton 5. Zippa Mountain 6. Hogem Bathoith 7. Chuchi 8. Peach 9. Rayfield River 10. Kamloops Syenite 11. White Rocks Mountain 12. Averill Pluton 13. Copper Mountain 14. Kruger Syenite 4 North 200km Terranes cEQuesnellia Stikinia • Alkalic porphyry deposits • Alkalic intrusive centres • Cities / towns 4 1.4 Galore Creek District 14.1 Geography The Galore Creek district is situated at the western margin of the Intermontane Belt in the Stikinia terrane approximately 75 km northwest of the Eskay Creek Au-Ag mine and about 97 km north-northeast of Wrangell, Alaska (Fig. 1 .2A). The district is situated in a remote part of the northwestern British Columbia, 1,030 km northwest of Vancouver, in mountainous terrain. High elevations (>900 m) are covered by permanent snow and ice fields. Multiple alkalic porphyry Cu-Au centers cluster in the steep-sided Galore Creek valley and compose the Galore Creek district (Schwab et a!., 2008). The head of valley is at approximately 730 m elevation and marked by the West Fork and South Col Glaciers. 1.4.2 Exploration History The Galore Creek district was discovered in 1955, by testing anomalies in a regional aero-magnetic survey. It has been explored by several companies since (Enns et al., 1995). Early district field evaluations of magnetic-high geophysical anomalies led to discovery of Southwest Zone mineralisation in addition to other zones (Barr, 1966). Initial work in the Galore Creek district was carried out by Kennecott Corporation. Since 2003, NovaGold Resources Inc. has discovered the Bountiful and West Fork zones which have added considerable resources. In 2007, NovaGold Resources Inc. and Teck formed a joint-venture and the property is now managed through the Galore Creek Mining Corporation. 1.4.3 Supracrustal Rocks Stikinia consists of four, early Devonian to middle Jurassic, sequences separated by unconformities. The sequences are composed of arc-related mafic to felsic volcanic and sedimentary rocks and contain coeval plutons (Monger 1977 and Monger and Nokleberg 1996). The major tectonostratigraphic elements in the northwest Stikinia (Fig. 1 .2A) are (Allen et al. 1976 and Enns et al. 1995): 5 1) Late Paleozoic to Middle Jurassic island arc volcanic-plutonic and sedimentary rocks of the Stikine assemblage, the Stuhini Group and Hazelton Group; 2) Middle Jurassic to early Upper Cretaceous basin sedimentary rocks of the Bowser Lake Group; 3) Upper Cretaceous to Tertiary continental arc volcanic rocks of the Sloko Group and; 4) Late Tertiary to Recent post-orogenic plateau basalts of Edziza and Spectrum Ranges. Intrusive rocks Volcanic and Sedlmentray rocks • Coast plutonic complex Stuhini Group (Eacene) • (Mid to Upper Triassic) ,— Caic-alkalic plutonic suite Shale and argilite .__) (Early Jurassic) • (Middle Triassic) Galore Creek alkalic suite Stikirie assemabige(Upper Triassic) • (Devonian to Permian) Hickman Batholith • (Mid Triassic) • Limestone (Permian) 0 5 km FIG. 1.2 A. Major tectono-stratigraphic elements in northwest Stikinia and the location of the Galore Creek district and Copper Canyon (modified from Wheeler and McFeely 1991; Gabrielse et al. 1991; Logan and Koyanagi, 1994). B. Regional scale geology of Galore Creek showing the location of Copper Canyon alkalic porphyry Cu-Au occurrence (modified after Logan and Koyanagi, 1994 and Enns et al., 1995). BT = Butte Thrust Fault; CCT = Copper Canyon Thrust Fault. 132 13O NAD 83 JTM Zone SN Stikinia (J) Cret-Tertiary CD Jurassic • Tnassic Other deposits • Palaeozoic 1. Copper Canyon Other Terranes (E Undifferentiated o 50km 6 1.4.3.1 Paleozoic Stratigraphy Rocks of the Stikine assemblage are the oldest rocks in the Galore Creek area (Fig 1 .2B); the assemblage consist of Devonian, Carboniferous and Permian bimodal flows and volcaniclastics, interbedded carbonate and minor shale and chert (Monger 1977). Logan and Koyanagi (1994) outline five main sub-divisions for the Stikine assemblage: 1) Devonian-Carboniferous penetratively foliated limestone, phyllite, mafic and felsic flows and tuff are over lain by; 2) distinctive Lower to Middle Carboniferous limestone; 3) Upper Carboniferous to Permian thick-bedded conglomerate and siltstone overlie the limestone unconformably to conformably; 4) the limestone is in turn overlain by felsic to mafic volcanic and volcaniclastic rocks; the volcanic sequences are conformably capped by a thick Lower Permian succession of fossiliferous limestone. Stinine assemblage stratigraphy and facies change rapidly and laterally from conformable in one place to nonconformable in another. The nature of lithotypes and facies relationships indicates an island arc depositional environment in Paleozoic Stikinia (Logan and Koyanagi 1994). The Tahltanian Orogeny is invoked to explain the Permian Triassic disconformity in western Canada; the sedimentary and likely tectonic break has been recognized in many terranes at the P-T boundary. 1.4.3.2 Mesozoic Stratigraphy In the Galore Creek district Permian limestone is mostly overlain para unconformably by Lower to Middle Triassic sediments and elsewhere dis-conformably by Upper Triassic rocks (Logan and Koyanagi, 1994). The Stikine assemblage forms the floor to the Middle to Upper Triassic Stuhini Group. A variety of flows, tuffs and volcanic breccias and sedimentary rocks, of the Stuhini Group, define a volcanic edifice centered on the Galore Creek district (Monger, 1977). These rocks compose a volcano- sedimentary rock package deposited in a sub-aqueous to subaerial environment (Logan and Koyanagi, 1994). Panteleyev (1976) subdivided the volcanic rocks at Galore Creek into a lower unit of submarine basaltic to andesitic flows and breccias typical for the Stuhini Group, and an upper unit of partially subaerial, compositionally distinct alkali enriched flows and pyroclastic rocks (Eims et a!., 1995). Volcanic, volcaniclastic and volcanogenic sedimentary rocks constitute the bulk of the Upper Triassic stratigraphy in Galore Creek district, forming three compositionally distinct suites (Allen et al., 1976; Enns et al., 1995; Logan, 2004): 7 1) Subalkalic homblende-bearing basaltic andesite, 2) Subalkalic to alkalic augite-phyric basalt 3) An uppermost orthoclase and pseudoleucite-bearing alkali basalt. The Galore Creek alkalic intrusive suite is hosted in Stuhini Group rocks (Fig. 1 .2B). 1.4.4 Igneous Rocks Three major periods of intrusive activity are recognized in the Galore Creek region (Enns et a!. 1995): 1) Upper Triassic to Lower Jurassic intrusion that include the caic-alkaline Hickman pluton and the Galore Creek alkalic suite. 2) The Upper Cretaceous to Paleocene Coast plutonic complex that occurs as several granite bodies west of Galore Creek. 3) Tertiary quartz monzonite, diorite stocks and mafic to felsic dikes occur within west and north striking extensional structures. The Galore Creek alkalic suite is a multiphase complex of alkalic silica-undersaturated metaluminous intrusions (Lueck and Russel, 1994; Lang et al., 1995a; Enns et a!., 1995). Syenite, monzonite and monzodiorite dikes and stocks are hosted in the Stuhini group rocks (Fig. 1 .2B). These intrusions are sub-volcanic and interpreted to be coeval and comagmatic with the youngest parts of the Stuhini Group (Logan and Koyanagi, 1994; Enns et al. 1995; Mortensen et al., 1995). The highest concentration of intrusions occurs in the Galore Creek valley consisting of composite dikes and stocks 2 km long by 1 km wide elongated to the north (Fig. 1 .2B). Kennecott Corporation established a paragenetic sequence of twelve (11-112) units for the Galore Creek alkalic suite and these were further characterised by Enns et al. (1995) (Table 1.1). The intrusions are named and referred to by their unit number (Table 1.1). This study notes deviations in the paragenetic sequence outlined by Enns et al. (1995). Modal change in primary mineral assemblages of Galore Creek intrusions from syenite and alkali syenite (Ii through 15), to 8 monzonite (16 and 17) and monzodiorite intrusions (18), to syenite intrusions (19 through Ill) and, finally to quartz intrusions (112) suggest differences in the compositions of the parent melts (Enns et al., 1995). Stanley (1992) and Enns et al. (1995) identified distinct intrusive lineages from field, petrographic and petrochemical data. Conserved element ratios and Pearce element ratios analysis (Pearce, 1968; Russell et al., 1990) casts the twelve igneous phases into four lineages characterised by constant and distinctive TiO2IZr, Ti02/Hf, Ti02/Nb and Ti02/Th ratios (Table 1.2). Common sets of conserved element ratios include: Lineage A: Ii, 13, I4aJb, I9aJb and Ill intrusions Lineage B: 13’, 15 and 19b intrusions Lineage C: 16 and 18 intrusions Lineage D: 112 intrusion Changes from chemically more-evolved to less-evolved magma compositions associated with the intrusion of the 16, 17 and 18 units (Table 1 .1) may indicate introduction of new, less-evolved magma (lineage C) into the sub-volcanic environment (Enns et al., 1995). Uranium-lead dating of the Galore Creek alkalic suite complex was carried out by Mortensen et a!. (1995). Temporal grouping of intrusive phases by Mortensen et al. (1995) is with respect to mineralisation and is based on the paragenetic sequence described by Enns et al (1995). Pseudoleucite-orthoclase-phyric syenite porphyry that is pre-Cu, or potentially syn-Cu, returns a U-Pb age on zircon of 2 10±1 Ma. Two megacrystic orthoclase-phyric syenite intrusions that are late to post-Cu in some of the mineralized centers and pre-mineral in others are dated 205±2.3 Ma (U-Pb zircon) and 200.1±2.2 Ma (U-Pb titanite). 9 TABLE 1.1 Summary of Galore Creek intrusive rocks (Enns, et. al., 1995) , I Texture I Textureame -distinguishing feature fli ame -distinguishing feature Ii Pseudoleucite porphyritic 17 Orthoclase porphyry porphyry porphyry alkali syenite - pseudoleucite- alkali syenite bearing 12 Pseudoleucite megaporphyritic 18 Monzodiorite equigranular I senate megaporphyry alkali - pseudoleucite- -large aerial extent in syenite bearing the region 13 Gray orthoclase porphyritic 19a Hornblende-bearing megaporphyritic! porphyry syenite orthoclase hiatal; megaporphyry syenite post-mineral in the Central Zone 13’ Gray orthoclase porphyritic 19b Hornblende-bearing megaporphyriticl porphyry syenite - biotite-bearing orthoclase hiatal megaporphyry syenite - plagioclase-phyric; 14a Dark orthoclase porphyritic 1gb’ Orthoclase megaporphyritic porphyry syenite - pseudoleucite- megaporphyry syenite bearing 14b Dark orthoclase porphyritic 110 Biotite-bearing equigranular / hiatal porphyry alkali syenite - pseudoleucite- orthoclase porphyry bearing alkali syenite 14b’ Melanite porphyry porphyritic Ill Orthoclase porphyry [orphyritic alkali syenite -minor pyoxene; syenite melanite-beaing 15 Biotite-bearirig porphyritic Ill a Clinopyroxene-biotite porphyritic orthoclase porphyry porphyry syenite syenite 16 Monzonite equigranular / senate 112 Lavander quartz syenite equigranular I senate -syn-mineral in the Southwest Zone TABLE 1.2 Average conserved element ratios from Galore Creek intrusive rocks (Enns, et. al., 1995) Ratio Lineage A Lineage B Lineage C Lineage D Ti02/Zr 65 30 100 30 Ti02/Hf 2200 1100 4000 1600 Ti02/Nb 180 90 500 130 Ti02/Th 800 375 2500 725 10 1.4.5 Hydrothermally Altered and Mineralized Centers Hydrothermal alteration is developed in the multiphase complex of alkalic intrusion and encapsulating shoshonitic volcano-sedimentary rocks. The district includes approximately twelve zones of Cu-Au mineralisation (Fig. 1.3). The largest deposit, and the best documented, is the northerly-elongated Central Zone (Lang et al., 1995b, Micko, et al., 2007). Divided into three domains, each with characteristic alteration style and metal tenor (Schwab et al., 2008), the Central Zone is comprised of the North Gold Lens, Central Replacement Zone and the South Gold Lens. Smaller deposits peripheral to the Central Zone are also known, these include: Southwest Zone, Junction, Butte, West Rim, Westfork and the Saddle zones (Enns et a!. 1995; Schwab et al., 2008). The Southwest Zone is distinguished by the abundance and variety of breccias and its Au tenor. Timing relationships of mineralization to the various intrusive units suggests at least two distinct periods of hydrothermal activity punctuated by intrusion of voluminous megacrystic orthoclase-phyric syenite and monzonite dikes (Table 1.1 and Fig. 1.3). Early mineralisation occurs in the Central Zone and is hosted mostly in supercrustal rocks with subordinate Cu-Au in intrusions and hydrothermally cemented breccias. Sulfide mineralization in the Central Zone occurs as a disseminated and pervasive replacement and is best developed in the North Gold Lens and South Gold Lens. In contrast, sulfide mineralization is centered within hydrothermal cemented breccias in the Central Replacement Zone (Schwab et a!., 2008; Micko et a!., 2008). No obvious intrusive centre for mineralisation in the Central Zone has been recognized. Mineralisation in the Central Zone is truncated to the west by post-mineral megacrystic orthoclase-phyric syenite and monzonite dikes (Enns et al., 1995; and Schwab et al., 2008). A second stage of mineralisation occurs in megacrystic orthoclase-phyric syenite and monzonite in the Southwest Zone and West Fork (Schwab et al., 2008). 1.4,6 Structural History Three phases of deformation are recognized for the oldest Paleozoic rocks and one phase for Upper Triassic strata within the Stikine-Iskut River region (Panteleyev, 1976; Logan and Koyanagi, 1994; Logan, 2004). The supracrustal rocks at Galore Creek are interpreted to have undergone early and broad-scale post-Triassic north-south compression followed by post-early Jurassic open to tight folding and related faulting 11 from east-west compression. East-west compression led to the formation of northerly trending folds and thrust faults (Fig. 2b) (Logan and Koyanagi, 1994). The Galore Creek alkalic intrusions are proposed to have been intruded within an extensional stress-field possibly related to north-south trending vertical faults during the Late Triassic (Logan and Koyanagi, 1994; Mortensen et al., 1995). The most visible structures within the Galore Creek district are the post-mineral west-dipping Butte Thrust Fault (Schwab, et al., 2008) and the east-dipping Copper Canyon Thrust Fault (Fig. 1 .2B) (Bottomer and Leary, 1995). Weakly deformed epiclastic conglomerates and sandstones are thrust over undeformed Stuhini group rocks and alkalic intrusions along the Butte Thrust (Figs. 1 .2B and 1.3). The Copper Canyon Thrust places Permian limestone over the Stuhini Group rocks and the Copper Canyon porphyry Cu deposit (Fig. 1 .2B). The Central Zone’s local structural architecture is not well defined. However, the spatial distribution and orientation of lithological units as well as associated hydrothermal alteration halos imply tilting of the Central Zone 45-60° east (Micko et al., 2008). Tilting is interpreted to have occurred at the end of hydrothermal activity due to N-trending, upright folding (Logan and Koyanagi, 1989). The refolding and thrust structures in the district are interpreted to have formed in response to westerly directed compression some time in the Jurassic (Logan and Koyanagi, 1994). Emis et al. (1995) and Schwab et al. (2008) describe numerous smaller scale faults and lineaments: easterly trending Junction Fault; northwest-trending Southwest Zone Fault; north-trending faults such as the East, Central and West Fork Faults (Fig. 1.3). 1.41 Southwest Zone The Southwest Zone is situated 600 m southwest of the South Gold Lens (Fig. 1.3) and hosts some of the highest grade near-surface gold mineralization in the Galore Creek district. Megacrystic orthoclase-phyric syenite (19a) and monzonite (19b) host numerous breccia facies, younger porphyry dikes and significant Cu-Au. Enns et a!. (1995) and Sillitoe (1991) describe a diatreme breccia that is composed of unsorted polylithic clasts supported by rock-flour matrix, in the Southwest Zone. Clasts of “19a” in the breccia were also recognized and along with the presence of significant Cu-Au allowed interpretation of multiple stages of mineralization in the district. Enns et al., (1995) document K-silicate alteration with coincident disseminated chalcopyrite 12 mineralisation and also recognize mineralisation focused at the western diatreme porphyry wall rock contact. Following the last 5 years of exploration by NovaGold Inc., including nearly 720 drill-holes, the new core exposures facilitate an opportunity to improve our understanding of the deposit. Features recognized by Enns et al. (1995) and Schwab et al. (2008), but which have not placed in the context the deposit genesis, include: inferred post —Central Zone timing; numerous breccia facies; high tenor Au mineralisation; and a Au-halo. These specific features are addressed chapter 2 and help elucidate the processes involved in the formation of the Southwest Zone deposit. 13 349000 E 352000 E North Rim - 0 - 0 -;.. :. ‘ megacrystic orthoclase-phyric : syenite and monzonite (19a and b) dark orthoclase - : porphyry syenite; . -. * + * ± pseudoleucite- . — bearing (l4a and b) -: - —:-. -. -- Saddle FIG. 1.3 Simplified geological and alteration-mineralisation map of the Galore Creek district illustrating the location of mineralized centers and major structural features. Base map is courtesy of NovaGold Resources and has been modified from the original version. Bed rock type and distribution is compiled from drill-hole and outcrop data. Alteration-mineralisation assemblages and distribution is modified from Schwab et al. (2008) and are projected from the subsurface (from midway through each center). NGL = North Gold Lens; CRZ = Central Replacement Zone; SGL South Gold Lens; SWZF = Southwest Zone Fault; WFF = West Fork Fault; BF = Bountiful Fault; EF East Fault. Alteration. mineralization — bornite -chalcopyrite chalcopyrite -pyrite L___i pyrite — bornite -magnetite diopside — -chalcopyrite K-feldspar -biotite garnet-anhydrite -biotite sericite-anhydrite -carbonate chlorite-epidote -carbonate Rock Types hydrothermal breccia monzodiorite GED sedimentary -- augite and feldpar-bearing volcanic! volcaniclastic pseudoleucite-bearing volcanic! volcaniclastic North 0 km 14 1.5 Breccias 1.51 Approach and Nomenclature Clastic facies are any fragmental rock and are described in a non-genetic manner following schemes outlined in McPhie et al. (1993) and Davies (2002). The systematic description and classification of breccias associated with sedimentary, hydrothermal and magmatic-hydrothermal systems is a challenging and controversial problem facing ore deposit researchers and explorationists. Key features of clastic rocks are (adapted from Seedorffet a!., 2005; Davies et al., 2008a and b): 1. Nature of the material between fragments, i.e. matrix vs. cement 2. Characteristics of the matrix 3. Characteristics of the cement 4. Nature of the fragments and their internal organization 5. Geometry and dimensions of the breccia body Matrix is defined as a fine-grained elastic component that occurs between larger clasts. Matrix comprises comminuted wall-rock (rock-flour), lithic, and phenocryst fragments of sand to granule size (<0.5mm — 4mm). Cement is defined as the crystalline component precipitated from an aqueous fluid. In his wide-ranging overview of breccias in volcano-plutonic arcs, Sillitoe (1985) provided a useful, broadly applicable and flexible genetic classification of ore-related breccias. It emphasised the overlap between volcanic, intrusive and hydrothermal processes and incorporated several levels of classification. This classification is process- oriented and hierarchical. 1.5.2 Fragmentation Mechanisms and Breccia Classification Breccia classification is primarily based on the overall fragmentation mechanism and the inferred involvement of water, magma or tectonics in the fragmentation process (Sillitoe, 1985; Davies, 2002). Table 1.3, adapted from Sillitoe (1985) and Davies (2002), lists breccia genetic classes and associated fragmentation processes. Primary breccia features record the energetics responsible for fracturing and rearranging the rock mass 15 (Laznicka, 1988). The components used to distinguish the various breccia facies (Table 1.3) are diagnostic of their genetic origin and are used for classification. This classification is process-oriented and hierarchical. Assignment of a breccia to the appropriate genetic class category also considers the environment of formation and the source of fluids involved, if it can be established. Sillitoe (1985) further subdivides breccia classes by ore deposit type. Magmatic-hydrothermal breccias are the products of violent and rapid exsolution of magmatic-hydrothermal fluids from cooling magma stocks (second boiling) coupled with decompression and liquid-vapor separation of already exsolved aqueous phases (Bumham, 1985; Sillitoe, 1985; Foumier, 1999). Open space filling, hydrothermally cemented breccias, with variable amounts of matrix, are interpreted as magmatic hydrothermal breccias. Crystallization-induced volatile release from hydrous magma is termed second boiling (Bumham, 1979). Second boiling can release sufficient energy to cause fracture failure of wall rock. Fractures permit decompression and expansion of already exsolved fluids, and further exsolution of fluids from the melt, releasing additional mechanical energy to drive fragmentation (Bumham, 1985). This mechanism of fragmentation is a key process associated with fracture systems and breccias related to porphyry deposits (Burnham, 1985; Sillitoe, 1985) and can be explosive to non-explosive (Davies, 2002). Magmatic-hydrothermal breccia pipes are associated with both unmineralized and porphyry-type intrusions (Sillitoe, 1985). Subsurface phreatic and phreatomagmatic explosions (Sheridan and Wolhetz, 1981 and 1983) produce fragmental rocks classified as hydromagmatic breccias (Sillitoe, 1985) or as hydroclastic breccias (Davies et al., 2008a). Fragmentation to produce hydromagmatic breccias is caused by mechanical energy released during flashing of water to steam (Hedenquist and Henley, 1985). Flashing of hydrothermal fluids is caused by depressurization. Depressurization can occur if 1) confining pressure (lithostatic or hydrostatic, tensile strength of the rock) is reduced to values below hydrothermal fluid pressure; or 2) an increase in hydrothermal fluid pressure in excess of confining pressure (Bumham, 1985). Direct magma-water interaction causes steam expansion and subordinate quench-fragmentation that drive explosive fragmentation producing phreatomagmatic breccias (Sheridan and Woheltz, 1981; Hanson, 1991). Polylithic, matrix-supported breccias with subordinate juvenile component can be produced by phreatomagmatic explosions. Products of phreatomagmatic eruptions include water, 16 steam, fragmented country rock and by definition juvenile magmatic material (Woheltz and Sheridan, 1983). Phreatic eruptions are steam-driven explosions with no direct involvement of magma required and no production of juvenile magmatic fragments regardless of fluid source, temperature or composition (Cas and Wright, 1987; Barberi et a!., 1992). Poorly cemented breccias in matrix are the products of phreatic explosions. Hydraulic fracturing (Nelson and Giles, 1985; Jebrak, 1997) may also occur in phreatic explosions. A key distinction between breccias of hydromagmatic and magmatic hydrothermal origin is that fluid involved in pbreatic and phreatomagmatic explosions is sourced externally. Tabular hydrothermally-cemented breccias with angular wall rock fragments are interpreted as non-explosive hydraulic breccias. Hydraulic fracture can occur if fluid pressure is greater than the sum of the confining pressure and tensile strength of the rock (Philips, 1972; Jebrak, 1997). There is a potential for a continuum between non- explosive hydraulic fracture (no fluid flashing), to flashing that release sufficient energy to cause explosive fragmentation (phreatic brecciation) or decompression (magmatic hydrothermal brecciation) (Nelson and Giles, 1985; Davies, 2002). Nelson and Giles (1985) interpret the transition between these fragmentation mechanisms and breccia class end-members as a function of energy in the hydrothermal fluid. Davies et a!. (2008a and b) also suggest that multiple fragmentation mechanisms occur in conjunction to produce breccias subfacies that have characteristics. 1.53 Principal Breccia Fades Characteristics Breccias are fragmental rocks, with an enormous variety of characteristics, and occur across all classes of porphyry deposits (Sillitoe, 1985; Seedorff et al., 2005) and other geological environments. Seven principal descriptive breccia types, or facies, are recognized to occur in the subsurface of porphyry systems and are distinguished by sets of one or more of the five key features outlined above (Table 1.4; modified from Seedorffet al., 2005; Davies, 2002): • Igneous-cemented breccias • Open space filling, hydrothermally-cemented breccias • Poorly cemented breccias in matrix 17 • Polylithic, matrix-supported breccias with subordinate juvenile component • Tabular hydrothermally-cemented breccias with angular wall rock fragments • Clast-rich breccias with predominantly juvenile fragments • Tabular breccias bodies with angular wall rock fragments in a matrix Breccia facies based on these descriptive criteria can be related to genetic classes (Table 3), however, it is important to note that there is likely a continuum between certain fragmentation processes. Thus, the principal breccia facies (Table 1.4) are not necessarily unique to a single genetic class (Table 1.3). Open-space filling, hydrothermally cemented breccias are widespread in porphyry systems and subfacies are distinguished based by cement mineralology (Sillitoe, 1985; Seedorff et al., 2005). Single or multiple hydrothermally cemented breccias are generally lensoid, ovoid, or circular pipe-like bodies with steep to vertical dip (Sillitoe, 1985; Seedorff et al., 2005). Ore minerals commonly form cement in these breccias and along with gangue minerals, and associated alteration, can be zoned vertically and laterally (Simmons and Sawkins, 1983; Seedorf et. al., 2005). Sillitoe (1985) and Seedorff et al. (2005) report hydrothermally cemented breccias both with and without matrix infill leading to further facies distinctions. Furthermore, Davies et al. (2008b) and Jackson et al., (2008) document a continuum between cement and matrix infihled breccias at the Kelian epithermal gold deposit, Indonesia, and the Northeast Zone at Mt. Polley alkalic porphyry Cu deposit, BC, respectively. At these deposits, matrix formation and cement precipitation are inferred to have formed contemporaneously or very close together in time. Polylithic, matrix-supported breccias with subordinate juvenile component (Table 1.3) contain juvenile material as matrix or clasts. Juvenile material in the matrix can consists of whole or broken crystals of biotite, quartz, or feldspar. This breccia facies is typically matrix-rich, poorly sorted, polylithic, with clasts that exhibit rounding. Bodies of this breccia facies are also generally discordant to surrounding rocks, circular to ovoid in plan, subvertical to ftmnel-shaped, and have abrupt contacts with wall rock (Sillitoe, 1985; Seedorff et al., 2005). In the cases where clasts are rounded these clastic rocks are by definition conglomerates rather than breccias, however this carries sedimentary connotations and is not implemented here. 18 An additional breccia facies not included by Seedorff et al. (2005) has been adapted from Davies (2002) and Davies et a!. (2008a); tabular hydrothermally-cemented breccias with angular wall rock fragments are distinguished by their geometry, clast angularity and internal organization and absences of matrix. Breccias of this facies can be localized by faults and manifest as vein breccias. There is likely a continuum between open space filling, hydrothermally-cemented breccias and tabular hydrothermally-cemented breccias with angular wall rock fragments. 19 TABLE 1.3 Genetic classes and associated fragmentation processes, adapted from Sillitoe (1985), Davies (2002) and Davies et al (2008b) Breccia genetic class / Specific Breccia genetic class General Pertinent features germane to Examples and referencesfragmentation classification - nomenclature and process referencesclassification fragmentation (primary) mechanism(secondary) mechanism (31 Phreatic breccias - pebble dikes Magmatic volatile exsolution during second-boiling, fluid decompression ± gravitational collapse ± (hydraulic fracturing?) Fluids are sourced directly from crystallizing magma; products contain various amounts of matrix; generally not as spatially extensive as phreatomagmatic breccias No direct interaction with magma or production of juvenile fragments; externally derived fluid regardless of source, composition and temperature Direct interaction of magma with an external fluid; regardless of that fluids source, composition and temperature; by definition contains juvenile material Agua Rica (igneous breccia), Argentina (Landtwing et al., 2002); Los Pelambres (igneous breccia), Chile (Atkinson et al., 1996). - Wright and Bowes, 1963. Copper Mountain-Ingerbelle, BC (Stanley et al., 1995); Mount Polley (Canboo-Bell), BC (Fraser et al., 1994 and 1995; Jackson et al., 2008); Galore Creek (Dendritic breccia in Central Zone), BC (Allen et al., 1976; Enns et al., 1995; Micko et al., 2008). - Phillips, 1973; Norton and Cathles, 1973; Burnham, 1985. Island Copper, BC (Perellô et al., 1995) Oyu Tolgoi, Mongolia (Perello et al., 2001 Toquepala, PerO (Zweng and Clark, 1995) - Gustafson and Hunt, 1975; Wallace et al., 1978; Sheridan and Wolhetz, 1981. Batu Hijau (Santong diatreme); Southwest Zone at Galore Creek (Enns et al., 1995); El Teniente (Braden pipe), Chile (Skewes et al., 2002) - Sheridan and Wohletz, 1983; Wolfe, 1980; Sillitoe and Bonham,1984; Sillitoe, 1985; Houghton and Nairn, 1991; Morrisseyetal., 2000. (51 Hydraulic breccias (6) Magmatic (volcanic) breccias (7) Fault breccias Hydraulic ± steam expansion, implosion (wet) Magmatic volatile release (first-boiling) Wear abrasion, distributed crushing, implosion (dry) Hydrothermal breccias/ Tectonic breccias (Hydrocaistic subsurface breccias) Magmatic breccias Tectonic breccias nonxplosive Strong structural control; there is (wet) likely a continuum between non- explosive hydraulic and explosive fracturing driven by phase change (phreatic) or decompression (magmatic-hydrothermal) Explosive (dry) Non explosive (wet or dry) Kelian Epithermal Gold Mine (CMX breccia facies) (Davies et al., 2008a); Agua Rica, Argentina (Landtwing et al., 2002) - Nelson and Giles, 1985; Davies, 2002; Davies et al., 2008a and b Casino, BC (Godwin, 1976); Cave Peak, Texas (Sharp, 1979); Redwell Basin, Colorado (Thomas and Galey, 1982) - Sillitoe, 1985 Agua Rica (fault breccias), Argentina (Landtwing et al., 2002) - Davis and Reynolds, 1996; Passchier and Trouw, 2005 (11 Intrusion breccia (21 Magmatic-hydrothermal breccias Stoping and passive intrusion Magmatic (subsurface) breccias Hydrothermal breccias (Hydrocalstic subsurface breccias) Hydrothermal breccias (Hydrocaistic subsurface breccias) Hydrothermal breccias (Hydrocaistlc subsurface breccias) Hydro-magmatic breccias Hydro-magmatic breccias Non- explosive (dry) Explosive to non-explosive (wet) Explosive to non-explosive (wet) Explosive to non-explosive (wet) (4J Phreato magmatic breccias Steam expansion (phase change) ± hydraulic fracture Steam expansion ± hydraulic fracture L’) TABLE 1.4 Common porphyry-system breccia facies and their descriptive criteria adapted from Seedorff et a!. 2005, Sillitoe (1985), and Davies (2002) Breccia facies last characteristics Breccia bodyAlterationIgenetic class correlationi Matrix Cement and organization Geometry Size Relationship to Porphyry System - sub-fades modifiers No clastic material or Crystallized, Angular to subrounded Variable but Irregular patches, Up to Not unique to porphyry systems; voids uncomminuted clasts generally <10 m; general minor generally around the margin of —100 m near the walls and roofs of intrusive clast- to infill-supported an intrusion subvolcanic stocks; commonly pre igneous mineral; generally barren rock Originally voids with pen-space Commonly monolithic; Ovoid to irregular; variable quantities of filling angular to rounded clasts pipe-like; sheeted hydrothermal >0.01 to <10 m; contact with wall-rock; breccia gangue and dissipates rapidly into overlying ore minerals rocks >100 to Form during a variety of <1000 m mineralizing stages; can be good ore hosts; strong deposit-scale vertical zoning in some deposits, from shallower, lower temperature assemblages to deeper, higher temperature assemblages Poorly cemented Variable rock flour Uncommon Poly- to monolithic; In some cases Irregular or dikes, sills, and <500 m Commonly form relatively late in breccias in matrix (3) and open space; flow hydrothermal angular to rounded clasts unaltered but pipes; evolution of porphyry systems and — continuous bodies, banding minerals <1 m usually sericite, can envelop intrusive bodies or related to late barren dikes; polyllthic; irregular bodies (sulfides, advanced be controlled by faults generally dilute preexisting generally monolithic quartz) argillic, or hypogene grade; decrease in propylitic abundance downward in system Tabular hydrothermally- Absent or minor Open-space Monolithic angular wall varriable, Planar to locally massive; -10’s m Form in preexisting dilatational cemented breccias with (unless present in filling rock fragments >0.01 to dependent on veins, vein breccias; wide by zones or faults; potential host to ore angular wall rock overprinted hydrothermal <10 m, in-situ clast reactivity of elongation along faults; 100’s m fragments (5) protolith?) gangue and organization, cement- to fluids and wall transitional to stockwork veins long — diverse cement ore minerals clast supported rock mineralology Clast-rich breccias with Rock flour and predominantlyjuvenile abundant juvenile fragments (61 component, pumice Generally clast dominated; polylithic; intrusive fragments common; clast- to infill-supported Variable, Ovoid in plan, >500 to < Forms at all stages of development including subvertical to funnel shaped 1 000m of porphyry system; may be good potassic and bodies; a ore host sericitic diatreme Tabular breccias bodies Crushed rock, <30% Present if Angular clasts of wall; Variable but Tabular bodies, with angular wall rock matrix; may have breccia rock gouge -0.5 m clast generally absent commonly steep fragments in a matrix of slickensides was conduit size, clast- to infill- or minor crushed rock (71 subsequent supported fluid <50 m Not unique to porphyry systems; pre-, inter-, and post-mineral, locally may host ore. Important for locating faults; may influence distribution of late alteration or supergene features Igneous-cemented breccias (1) Open space filling, hydrothermally cemented breccias [2] — potassic; calcic, sodic; calc-potassic; sericitic ± sillcic ± tourmaline; advanced argilllc assemblages Polylithic, matrix- Fine grained rock Minimal open Polylithic, matrix-rich and Minor sericite, Circular to ovoid in supported breccias with flour and juvenile space supported; subangular to chlorite and clay plan, subvertical to subordinate juvenile material (broken and rounded wall rock clasts funnel-shaped bodies; component (4) unbroken quartz, variable size generally abrupt contact with biotite, feldspar <10 m; ± juvenile clasts, wall rock; a diatreme crystals) weakly vesiculated magmatic Lateral Commonly form relatively late, near extent the periphery of some orebodies; >1km; generally contain altered and vertical mineralized clasts; less widespread extent than magmatic-hydrothemial and >100 phreatic breccias; shown to root in to >1km post-mineral intrusions 1.6 Biotite Biotite has great importance for metallogenic and petrological studies by earth scientists because of their variable chemical composition and their ability to incorporate a variety of major, minor, and trace elements. Analogous with examples in metamorphic and igneous petrology, alteration minerals (assemblages) contain geochemical information that, in some instances, may help elucidate the conditions under which they formed. An outline of the essential elements of biotite chemistry, microprobe analysis normalization schemes and the mechanisms by which Ti4, Fe3 and Al3 are incorporated in the biotite lattice are presented and discussed below. 1.6.1 Crystal-chemistry Biotite is a true trioctahedral phyllosilicate with the simplified formula: IM23L112T4O0A The site I (alkali or interlayer site) is commonly K, Na, Ca, Ba; M (octahedral site) is usually Fe2,Mg2,Fe3,Ti4, Al, Mn, Cr; D represents a vacancy; T (tetrahedral site) is usually Si, Al, Fe3 and; A is OH, F, Cl, 0 (oxy-biotites) (Rieder Ct al., 1998). A solid 2+ . . 2+ solution series exists between KFe3A1SiO10(OH) (ammte) and KMg3A1SiO10(OH) (phlogopite) and occurs by Fe Mg substitution (Deer, Howie and Zussman, 1992). A variety of cation and anion substitutions in the alkali, octahedral, tetrahedral and hydroxyl sites affect the concentration of elements in biotite (Fig. 1.4). 22 Tetrahedral FIG. 1.4 Ferromagnesian biotite with labeled structural sites (http://staff.aist.go.jp). The most significant concentration variations are of the elements Fe, Mg, Al, Ti, Si, F, and Cl (Tischendorf et al., 2001). Biotite classification based to the ratio Mg/(Mg/Fetotaj) [Mg#] defines phiogopite as Mg#> 0.5 and annite as Mg# <0.5 (Tischendorf et al., 2004). Proton-deficient oxyannite, KFe3AlSii2(H )is a hypothetical end member in which Fe3 and 02. substitute for Fe2 and OW (Wones, 1963; Beane, 1974). Annite phiogopite-oxyannite form an additive ternary system in which the three cations (Fe2, Fe3,Mg2) substitute in the trioctahedral lattice sites of biotite. Biotite composition is typically described in mole fractions determined on the basis of all octahedral ions in the minerals crystal-chemical formula: phlogopite (Mg), annite (Fe2), Ti-biotite (Ti), P Doxy (Fe3), and Al-biotite (Al). Biotite analyses are obtained by different analytical methods (wet chemical, X-ray fluorescence, electron- and ion-microprobe) and require normalization schemes to calculate crystallo-chemical formulae. Alkali Hydroxi Octahedral 0 0 23 1.6.2 Normalization Schemes for Microprobe Data Formulae recalculation in not a trivial task in biotites as several normalization schemes can be implemented. Each scheme may introduce artifacts in the formulae, such as underestimation of octahedral and tetrahedral cations (Feldstein et a!., 1996; Waters and Charnely, 2002). Furthermore, electron microprobe analysis cannot distinguish between Fe2 and Fe3, and in the case ofH20, H and 0. Consequently, calculations of formula proportions of ions from analyses of a hydrous, Fe3tbearing mineral, such as biotite, require certain assumptions. Two normalization schemes are commonly applied. The first is expressed by the relationship total cations — (K +Na + Ca + Ba) = 7, which assumes complete occupancy of octahedral and tetrahedral sites (Ludington and Munoz, 1975). Such cation-based normalization procedures are unaffected by the valance state of iron and variations in H20-content. Assuming a full complement of 10 0 and 2 (OH, F, Cl) per biotite formula unit gives a charge of 22. Ferric iron would be equal to the difference between 22.0 and the calculated total positive charge; excess negative charge is compensated by replacing Fe2 with Fe3. Similarly, excess positive charge is compensated by replacing OH with excess oxygen, thus giving a minimum estimate of Fe3. If octahedral vacancies are present however, as indicated by Holdaway (1980), Dymek (1983) and Rebert et a!. (1995), the formula proportions will be overestimated in proportion to the number of vacancies, resulting in a corresponding charge excess, making Fe3 estimations impossible. Alternatively, a charge excess may not indicate vacancies, but could also be compensated by an 0 excess related to the occurrence of oxycomponents (Dymek and Albee, 1977). The second scheme, which assumes all iron occurs as Fe2, involves normalization to 11 oxygen atoms equivalents (neutrally charged mineral), and is expressed by the relationship: total cation charge = 22.0 (Deer, Howie and Zussman, 1992). By normalizing biotites to a set charge, the number of cations obtained may be less than eight, resulting in apparent vacancies that are often assigned to the interlayer (A) and br octahedral sites. This procedure is unaffected by the presence of vacancies, but causes cation abundances to be overestimated in direct proportion to the number of Fe3 present, which remains an unknown quantity. 24 Dymek (1983) and Waters and Charnley (2002) review the substitution mechanisms occurring in biotites from high-grade gneisses in West Greenland and in metapelites of the Namaqualand Metamorphic Complex, South Africa, respectively. Both studies evaluate the mechanisms of Ti incorporation and propose crystal-chemistry models that govern their biotite sample suites. Their results also led to tailored iterative normalization procedures for microprobe analyses of biotite that eliminate charge discrepancies, and therefore permits Fe3 estimation. 163 Mechanisms of Ti, Al and Fe3+ Incorporation Chemical variation in biotite is accommodated via Tschermak exchange (coupled substitution), vacancy substitutions, oxy-substitution (deprotonation / dehydrogenation reactions) or coupled interlayer-vacancy substitutions in the biotite lattice (Foster, 1960; Brigatti and Guggenheim, 2002; Waters and Chamley, 2002). Common substitutions involving Ti4,Fe3 and Al3 are presented in Table 1.5. 25 34 5 6 7 8 9 10 11 12 13 Ti-spineT Ti-vacancy (R) Ti-vacancy (Al) Ti-vacancy (Mg-Al) Ti-deprotonation (Mg) Ti-deprotonation (Al) Ti-deprotonation (Mg-Al) Al-Tschermak Al-vacancy Al-interlayer vacancy Coupled substitution Fe-vacancy oxybiotite exfoliation by loss of interlaver K A13 and Fe3 in biotite Reference Robert, 1976; Dyar et al, 1993 Dymek, 1983; Brigatti and Guggenheim, 2002 Dymek,1 983 DahI, 1969; Dymek, 1983 Brigatti and Guggenheim, 2002 Holdaway, 1980. Dyar et aL, 1993, Brigatti and Guggenheim, 2002 Waters and Charnley, 2002 Foster, 1960 Foster, 1960 Dymek, 1983 Rebert et al., 1995; Brigatti and Guggenheim, 2002 Reberteta., 1995 Wones and Eugster, 1965 Rebert et al., 1995; Scordari et al.. 2006 Divalent cations in the octahedral site (such as Fe2 and Mg2) are denoted as ‘(R) and cations in tetrahedral coordination are indicated with the IV superscript. A = alkali or interlayer site. fl represents a vacancy. Si4 and Al3are assumed in all reactions. Common substitutions involve replacement of Mg2 and Fe2 by Al3, Fe3, and Ti4 on octahedral sites, and Si4 by additional Al3 on tetrahedral sites. Biotite chemistry is further complicated by the apparent occurrence of vacancies, initially demonstrated by Foster (1960). By this mechanism, Ti4 and Fe3 are balanced within the octahedral sheet by means of formation of vacancies for space and charge balance. The most common .4+. . . . . mechanism of incorporation of Ti in biotite is by Ti-Tschermak and variations of Ti- Vacancy substitutions (substitutions 1 and 4, Table 1.5) (Dyar et al., 1993). These substitutions act to stabilize the mineral at particular physiochemical conditions. A dominant substitution mechanism may operate at specific T-P conditions (Guidotti et al., 1977) with multiple subordinate and concomitant mechanisms (Mesto et al., 2006). TABLE 1.5 Substitution mechanisms for Ti4, Substitution Name 1 “1Ti+ 2”Al = VIMg + Ti-Tschermak 2 VITi + IvAl = viAl + iV Coupled substitution VITi + “1Mg = 2”1A1 V1Ti + “1D = 2”1R I .5”1Ti + 0.5 vID = 2”1Al VITi + O.67”1L1 = VIMg ÷ 0.67”Al VITi + 202_+ H2 = VIMg + 20W V1TI + I .6702_ + 0.83H2 = 0.67”1Mg + 0.33”1Al + 1.670W VIAl + IVAI = 3V1R + iVSi 2”1Al + VID = 3VIR IVAI + K = AD + iV VIFe3+ + ivA13+= VIFe2++ IVSI 14 2”Fe3 + VID = 3V1Fe2+ 15 VIFe3+ + +O.5H2= VIFe2+÷0W 16 VIFe3+ + AD.... K + VIFe2 26 1.7 Biotite in the Porphyry Environment Two major genetic types of biotite occur in the porphyry environment; primary igneous and hydrothermal. Hydrothermal biotite is further subdivided into: metasomatised primary igneous minerals; metasomatised and recrystallized igneous minerals (secondary or shreddy biotite), and infill crystals precipitated directly from a fluid (vein and cement) (adapted from Jacobs and Parry, 1979). The majority of the previous studies of biotite in porphyry Cu deposits have focused on the determination of F and Cl contents, with the objective of distinguishing between mineralized and barren plutons (Mason, 1978; Hendry et al., 1988). In contrast, only a few studies characterize biotite composition with respect to alteration assemblages (Fig. 1.5) and estimate intensive parameters of the related hydrothermal fluids (Beane, 1974; Jacobs and Parry, 1979; Gunow, et al., 1980; Selby and Nesbit, 2000). There is a general paucity of work investigating major, minor and trace element composition of biotite in the porphyry environment particularly in the calc-alkalic Au rich and alkalic Cu-Au deposits. However, no examples of studies documenting spatial variation in cations of infill hydrothermal biotite composition are documented. Studies of biotite chemistry also allow an investigation of activity-composition models and their application in geothermometry and oxygen, water, halogen fugacimetry. Pertinent results of biotite studies from Santa Rita, Casino, Copper Canyon and Galore Creek are presented below. 1.7.1 Example Studies Santa Rita porphyry Cu deposit, New Mexico: Jacobs and Parry (1979) report variations in biotite composition from potassic, phyllic, propylitic and argillic alteration zones of the Santa Rita Cu-porphyry deposit. Based on this data, they interpreted HF, HC1, and H20 fugacities of the hydrothermal fluids. Mole fraction phlogopite, F and Si increase progressively from igneous to secondary to vein biotite types; the reverse relationship holds for Ti, Cl and Al. In each alteration zone type, with the exception of phyllic, mole fraction phlogopite correlates positively with F and Si and negatively with 27 Ti, Cl and Al (Fig. 5). Jacobs and Parry (1979) postulate that the variable but systematically related content of Fe, Mg, Ti and Al, in relic biotite in the phyllic and argillic alteration zones of the stock, resulted mainly from reequilibration with a solution that had variable ratios of Fe, Mg, and Ti to pH. They also conclude that the trend in biotite compositions across the phyllic-potassic transition zone of alteration, into the argillic zone, indicates a progressive increase in pH and in aMg2/aFe ratio of a prograding aqueous phase. Beane (1974) comments (on data from several study sites including Santa Rita) that igneous biotites generally have molecular ratios of Mg / Fe <1, and alteration biotites appear to be characterized by Mg / Fe> 1.5 and Fe3 I Fe of < 0.3 (Fig. 5). Casino porphyry Cu-Au-Mo occurrence, Yukon: Selby and Nesbitt (2000) took a similar approach as Jacobs and Parry (1979) to describe biotite in the Casino porphyry Cu-Au-Mo occurrence and describe broadly similar trends in biotite composition (Fig. 1.5). Both Selby and Nesbitt (2000) and Jacobs and Parry (1979) note that Ba concentration in igneous biotite decreases significantly in response to interaction with hydrothermal fluid. Based onJH20/ F,JH20/ C1 and fHF/HCL estimations they conclude that fluids associated with potassic and phyllic alteration were similar. Furthermore, they remark thatJH2O/HF values vary greatly among porphyry deposits and possible reflect the composition of source magmas. Copper Canyon alkalic porphyry Cu-Au occurrence, BC: Twelker (2007) noted that hydrothermal vein biotite, from Copper Canyon, is significantly depleted in Ti02 and has moderately high Fe3/F2ratios relative to typical igneous biotite and even hydrothermal biotite from other porphyry deposits. Comparison of Copper Canyon biotites with biotite intercept values from other prospects shows that estimated F activity values are typical of Cu porphyry deposits, but are lower than those of typical porphyry Mo deposits. Furthermore, Copper Canyon hydrothermal biotite has Cl activity typical of igneous rocks but slightly lower than average potassic fluids in porphyry Cu deposits (Munoz, 1984; Twelker, 2007). Galore Creek, Central Zone alkalic porphyry Cu-Au deposit, BC: Allen (1967) used the relative intensities of peaks (004 and 005) in XRD analysis to determine the composition of the octahedral layer in biotite. In this case, the biotite is strongly pleochroic tan to brown or tan to green with the amount and intensity of green biotite being the main variable. He notes that biotite is greenish in colour when Fe3 is present in 28 3+ . 3+ 2+ 3+ 2+proportions exceeding Fe /Ti = 1.0 and Fe /Fe = 0.17. Where Fe /Fe > 0.3 the mineral is pleochroic brown to green and possibly corresponds to least altered igneous biotite. Reddish brown biotite were estimated to have Fe3/Ti = 0.54-0.84 and Fe3/F2 0.07-0.09. The results of Allen (1976) conform to assertions made later by Beane (1974) that secondary biotites, characterized by Fe3/F2< 0.3, occur in proximity to intrusive centers and mineralization in the Central Zone. Beane (1974) reports that five biotite samples from the Central Zone show temperature of formation from 430° to 600° C and that a general increase in temperature is related to the proximity of the samples to the major intrusive syenite mass. The increase in temperature corresponds to increasing Fe3 content and Fe3/Fe2ratio at essentially constant Fe/Mg ratio (Beane, 1974). / / zmaynetiteericiezannydrite FIG. 1.5 Biotite composition trends in an idealized porphyry alteration model. Schematic porphyry alteration model modified after Guilbert and Lowell (1974). Biotite composition variations compiled from Beane (1974), Jacobs and Parry (1979), and Nesbit and Selby (2000). 1.7.2 Geothermometry Wones and Eugster (1965) took a simplistic approach to the biotite system and experimentally modeled biotite stability between annite and phlogopite. Their work demonstrated that increasing phlogopite content correlates with higher temperature and 102, following a given 102 buffer. Additionally, in the assemblage biotite-K-feldspar magnetite, mole fraction annite provides rough estimates offO2-Tconditions at constant fH20. For more reliable temperature and 102 estimates values for aKspar and aMag and Relic biotite composition trending to: Higher XM9 (phiogopitic) Lower FeO0i,Ti2 A1203,Si02,and F BaO and MnO Increase in Decrease in Mg/Fe ratio Fe3/F2ratio towards the porphyry center DEPTH (Km) Biotite type compositional trend: Igneous alteration vein Increase In XMg, 5102. and F Decrease in Ti02,A1203and Cl / P / —2 \ \ ‘p A=ArgiIIic; chlorite-kaolinte-adulaira -quartz P=Propylitic; chlorite—epidote-carbonate -anhydrite S=Phyllic; quartz-sericite-pyrite K=Potassic; quartzK:feIdsparbiotite — —4p/ / —6 Multiple porphyry stocks 29 required. Wones and Eugster (1965 acknowledge Fe3+ substitution in the octahedral site (PDoxy) but did not measure its abundance and its effects are disregarded. This geothermometer primarily considered biotite stability in P-T conditions of metamorphic and igneous crystallization. 2+ 3+ 2+Evaluation of mixing parameters for the Mg — Fe — Fe substitution in biotite provides a geothermometer for the potassic alteration zone in porphyry systems (Beane, 1974). Beane’ s (1974) geothermometer requires coexistence of biotite with magnetite (Fe304) and K-feldspar (KA1Si3O8)and builds on existing models for the annite-phiogopite geothermometer of Wones and Eugster, (1965) by including Fe3 substitution. This model expresses biotite as the ternary system annite-phlogopite-proton deficient oxyannite (PDoxy). In his thermodynamic calculations Beane (1974) assumes pressure is constant and its effect negligible, and that K-feldspar and magnetite are stoichiometric (aKS = aMag =1). Beane (1974) calculates equilibrium curves, from thermodynamic data, in which fixed PDoxy molar fractions versus molar fraction phlogopite can be used to calculate the temperature at which Fe and Mg partitioning took place. The validity of Beane’s geothermometer requires that Fe3 is accommodated as a component in hypothetical PDoxy and cogenetic K-feldsapr-biotite-magnetite. The geothermometer of Stormer and Charmicheal (1971) is based on fluorine hydroxyl exchange between biotite and apatite. Erratic F contents of igneous apatite, and to lesser extent hydrothermal apatite, often hinders application of this technique. Intergrown hydrothermal biotite and apatite however, is usually homogenous with respect to F (Jacobs and Parry, 1979). Ludington (1978) and Zhu and Sverjensky (1992) revised the biotite-apatite geothermometer to account for deviations from stoichiometric apatite, and octahedral site variations (Fe-Mg solution) known to affect the F-OH composition (Munoz and Ludington, 1974; Zhu and Sverjensky, 1991), making it more reliable. Robert (1976) experimentally determined that the solubility of Ti in phlogopite was a prograde function of temperature and retrograde function of pressure. The absolute abundance of Ti at a given P-T condition is strongly affected by the presence of a Ti saturating phase, like rutile and ilmenite (and perhaps sphene) (Guidotti et al., 1977). Furthermore, the distribution of Ti between silicate phases (biotite-muscovite hornblende) is not well established, thus, Ti content in biotite has not been successfully 30 applied as a geothermometer. Nevertheless, Ti-content in biotite may prove useful at distinguishing gradients in temperature. 1.7.3 Oxygen Fugacity Estimates Redox conditions of rocks and their coexisting fluids is of particular importance in the study of ore deposits and the proportion of Fe3 to Fe2 in biotite, in part, reflects these conditions (Lentz et al., 1994). Experiments indicate that the Fe3 content of biotite increase with 102 (Wones and Eugster, 1965; Deer, Howie and Zussman, 1992). In the reaction biotite + O.5oxygen = K-feldspar + magnetite + water, an oxidative transfer reaction results in Fe2, in the annite component, oxidizing to magnetite and is accompanied by Mg enrichment but no net increase in biotite (Wones and Eugster, 1965). However, Wones (1981) indicates that the calculatedf02,from the reaction above, is reliable between the nickel-nickel oxide and wustite-magnetite buffer but less so between nickel-nickel oxide and magnetite-hematite. By assuming 102 is buffered during cooling, Jacobs and Pary (1979) approximated J02T curves based on the relative positions of their biotite compositions to the nickel-nickel oxide and magnetite hematite buffers (Wones and Eugster, 1965), permitting 102 estimate at a known temperature. Other substitutions in biotite respond to variations in 102. F-OH substitution can affect the stability of biotite with F-rich biotites stable to higher 102 and lower thermal stability (Munoz and Ludington, 1974). At a minimum, variation in Fe3/ Fe2 ratio in biotite indicate changing 102 in the environment of formation; metamorphic equilibration; crystallizing magma; precipitation from magmatic-hydrothermal fluid. 1.7.4 Halogen and Water Fugacity Estimates The behavior of halogens is important in ore-forming processes as F and Cl are significant in the metal transporting process (Munoz, 1984). The substitution of F and Cl into the OW site of biotite is governed by a number of independent factors, principally: 1) the activity of these elements 2) the cation population of the octahedral layer in biotite (particularly Fe and Mg); 3) the temperature of hydroxyl-halogen exchange; and 4) the effects of post-crystallization leaching or enrichment due to hydrothermal fluids or groundwater (Munoz, 1984; Zhu and Sverjensky, 1991). Thus, this aspect of biotite composition usefully records an indirect measure of F and Cl activities of the 31 hydrothermal fluids that formed the biotite. Substitution of F for OH is thermodynamically favored in Mg-rich biotites and is governed by the Fe-F avoidance principle (Munoz, 1984). Similarly, there is a preferential association between Fe2-Cl compared to Mg-Cl referred to as Mg-Cl avoidance. Munoz (1984) tries to account for these effects in fugacity ratio calculation by correcting for each component affecting halogens in the biotite analysis (Mg, Fe, Al), allowing comparative discussion of biotite halogen content independent of general composition. Using experimental derived equilibrium constants for the reactions OH-biotite + HF F-biotite + H20 and OH biotite + HC1 Ci-biotite + H20 (Ludington and Munoz, 1975) a single numerical value can be defined to express relative F or Cl enrichment that considers both Mg/Fe and Al in the octahedral layer. This value is referred to as the fluorine or chlorine intercept value (IV [F] and IV [Cl]) (Munoz, 1984). For a known temperature, the intercept value equations can be combined with equilibrium constant equations to calculate the ratio of the fugacities; f1120/HF,JH20/ETC1 and JHF/HCL (Munoz, 1992; Zhu and Sverjensky, 1991, 1992). Conversely, intercept values can be used to calculate temperature if fugacities are known. 1.8 Research Objectives In order to develop holistic models that improve our understanding of alkalic porphyry systems integration of detailed structural, paragenetic, alteration zonation and geochemical information from individual deposits is essential. Furthermore, to place whole-rock or mineral chemistry studies in a meaningful context a coherent geological framework in vital. The principal objectives of this research are intended to contribute to the general characterization and understanding of alkalic porphyry systems, specific features of the Southwest Zone identified by Enns et al. (1995) are also addressed: • Establish the igneous stratigraphy for the Southwest Zone with an emphasis on emplacement sequence with respect to fragmentation, alteration, mineralisation and structural events. • Characterize elastic rocks and map the breccia facies architecture • Map and characterize hydrothermal alteration and suiphide minerals in time and space, and define assemblages and paragenesis. 32 • Identify mechanisms of Cu-Au localization: lithological and I or structural? • Describe metal halos with respect to alteration and mineralisation paragnesis: what is the nature of the Au-halo? • Characterise fluids associated with Cu-Au and hydrothermal alteration halos: can biotite chemistry be utilized to 1) assess the physicochemical conditions attending Cu-Au deposition? and 2) track the evolution of hydrothermal fluids?. • Develop a genetic model for the Southwest Zone deposit. 1.9 Study Methodology This M.Sc. study incorporates two field season of drill-core logging focused on description of rock types, alteration and mineralisation in the Southwest Zone alkalic porphyry Cu-Au deposit. A graphic logging scheme was implemented to record rock types, alteration and mineralization facies and textures, and paragenetic relationships encountered (Appendix, A). In all, core available from eighteen drill-holes was logged with additional information sourced from previous years drill-core data provided by NovaGold Resources (Appendix, B ) Orthogonal cross-sections were chosen that best transect mineralization and geological contacts; E-W section along 6333650N (A-A’) and N-S section along 350035E (B-B’). The volume of rock investigated is approximately 450 x 600 x 700 m. Representative samples of coherent and clastic rocks, alteration and mineralisation assemblages were taken for further investigation at the University of British Columbia. Data is compiled from core logging, petrography, spectroscopy, SEM, microprobe analysis and whole rock assays. The pattern of alteration and metal zonation and breccia facies architecture is used to interpret fluid-flow pathways and to track the physiochemical evolution of the hydrothermal system. Finally, descriptive, analytical and interpretative elements are combined to propose an ore deposit model for the Southwest Zone. Major and minor element contents, determined form electron microprobe analysis, of igneous and hydrothermal biotites are reported and discussed in Chapter 3. A crystal chemistry model for hydrothermal biotite is proposed and permits an estimate of Fe3 contents from microprobe analysis. Forty nine infihl biotite samples from different vertical and lateral position in the system were analyzed. A systematic deposit-scale 33 variation in vein and cement biotite chemistry is evident and described with respect to alteration assemblages and Cu concentration. Biotite analyses and activity-composition models are used to evaluate temperature, water and halogen fugacity, and aqueous solution characteristics of hydrothermal fluids attending formation potassic alteration and Cu deposition (Wones and Eugester, 1965; Beane, 1974; Munoz and Ludington, 1974; Munoz, 1992; and Zhu and Sverjensky, 1991 and 1992). Based on the variations in vein and cement (infill) biotite chemistry mechanisms of Cu transport and deposition are proposed. Additionally, halogen fugacity estimates from the Southwest Zone are compared with other alkalic porphyry systems, and to caic-alkalic deposits. 34 1.10 References Allen, D. G., 1967, Research investigation of Galore Creek potash feldspar and biotite; unpublished report for Stilcine Copper ltd., p. 5-8. Allen, D.G., Panteleyev, A. and Armstrong, A.T., 1976, Porphyry copper deposits of the alkalic suite: Galore Creek: In Sutherland Brown, A., ed., Porphyry Deposits of The Canadian Cordillera: Canadian Institute of Mining and Metallurgy Special Volume 15, p. 402-414. Atkinson, W.W., Jr., Souviron, A., Vehrs, T.I., and Faunes G., A., 1996, Geology and mineral zoning of the Los Pelabres porphyry deposit, Chile: Society of Economic Geology Special Publication 5, p. 1105-1119. Barr, D.A., 1966, The Galore Creek copper deposits: Canadian Institute of Mining and Metallurgy, Bulletin, 59, p. 841-853. Beane, R.E., 1974, Biotite stability in the porphyry copper environment: ECONOMIC GEOLOGY, v. 69, p. 241-256. Bottomer, L.R., and Leary, G.M., 1995, Copper Canyon porphyry copper-gold deposit, Galore Creek area, northwestern British Columbia: In Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Northern Cordillera: Canadian Institute of Mining and Metallurgy Special Volume 46, p. 645-649. Bumham, C.W., 1985, Energy release in subvolcanic environments: Implications for breccia formation: ECONOMIC GEOLOGY, v. 80, p. 15 15-1522. Bumham, C.W., 1997, Magmas and hydrothermal fluids, in Barnes, H.L., ed., Geochemistry of Hydrothermal Ore Deposits, third edition: New York, John Wiley and Sons, p. 63-123. Brigatti, M.F. and Guggenheim, 5., 2002, Mica crystal chemistry and the influence of pressure, temperature, and solid solution on atomistic models: In Mottana, A., Sassi F.P., Thompson, J.B. Jr., and Guggenheim, S., ed., Micas: Crystal Chemistry and Metamorphic Petrology: Reviews in Mineralogy & Geochemistry, 46, Mineralogical Society of America and The Geochemical Society, Washington, D.C. p. 1-97. Cas, R.A.F., and Wright, J.V., 1987, Volcanic successions: Modem and ancient. A geological approach to processes, products and successions: London, Allen and Unwin, 528 p. Cooke, D.R., Wilson, A.J., House, M.J., Wolfe, R.C., Walshe, J.L., Lickfold, V., and Crawford, A.J., 2007, Alkalic porphyry Au-Cu and associated mineral deposits of the Ordovician to Early Silurian Macquarie arc, NSW: Australian Journal of Earth Sciences, v. 54, p. 445-463. Coney, P.3., Jones, D.L. and Monger, J.W.H., 1980, Cordilleran suspect terranes: Nature, v. 288, p. 329- 333. Coney, P.J., 1989, Structural aspects of suspect terranes and accretionary tectonics in western North America: Journal of Structural geology, v. 11, p 117-125. Davies, A.G.S., 2002, Geology and genesis of the Kelian gold deposit, East Kalimantan, Indonesia (unpublished PhD thesis), University of Tasmania. Davies, A.G.S., Cooke, D.R., Gemmell, J.B., and Simpson, K.A., 2008a, Diatreme breccias at the Kelian gold mine, Kalimantan, Indonesia: Precursors to epithermal gold mineralization: ECONOMIC GEOLOGY, v. 103, p. 689-716. 35 Davies, A.G.S., Cooke, D.R., Gemmell, J.B., Leeuwen, T.V., Cesare, P., and Hartshorne, G., 2008b, Hydrothermal breccias and veins at the Kelian Gold Mine, Kalimantan, Indonesia: Genesis of a large epithermal gold deposit: ECONOMIC GEOLOGY, v. 103, p. 717-757. Dahi, 0., 1969, Irregular distribution of iron and magnesium among coexisting biotite and garnet: Lithos, v.2,p. 311-322. Davis, G.H., and Reynolds, S. J., 1996, Structural Geology of Rocks and Regions: John Wiley and Sons, New York, 366-371. Deer, W.A., R.A. Howie, Zussman, J., 1992, An Introduction to the Rock-Forming Minerals: 2nd ed. Pearson-Prentice Hall, p. 696. Dyar, M.D., Guidotti, C.V., Holdaway, M.J., and Colucci, M., 1993, Nonstoichiometric hydrogen contents in common rock-forming hydroxyl silicates: Geochimica et Cosmochimica Acta, v. 57 p. 2913- 2918. Dymek, R.F., 1983, Titanium, aluminum and interlayer cation substitutions in biotite from high-grade gneisses, west Greenland: American Mineralogist v. 68, p. 880-899. Dymek, R. F., and Albee, A.L, 1977, Titanium and aluminum in biotite from high-grade Archaean gneisses, Langq, West Greenland. Transactions, American Geophysical Union, EOS, v. 58, p. 525. Buns, S.G., Thompson, J.F.H., Stanley, C.R. and Yarrow, E.W., 1995. The Galore Creek porphyiy copper- gold deposits, Northwestern British Columbia: In Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Northern Cordillera: Canadian Institute of Mining and Metallurgy Special Volume 46, p. 630-644. Feldstein, S.N., Lang, R.A., Vennemann, T., and O’Neil, J.R., 1996, Ferric-ferrous ratios, H20 contents and D/H ratios of phlogopite and biotite from lavas of different tectonic regimes: Contributions to Mineralogy and Petrology, v. 126, p. 5 1-66. Foster, M.D., 1960, Interpretation of the composition of trioctahedral micas: United States Geological Survey Professional Paper 354-B, p. 11-49. Fournier R.0., 1999, Hydrothermal processes related to movement of fluid from plastic into brittle rock in the magmatic-epithermal environment: ECONOMIC GEOLOGY, v. 94, p. 1193-1211. Fraser, T.M., 1994, Hydrothermal breccias and associated alteration of the Mount Polley copper-gold deposit (93A112): In Geological Field work 1993, BC Ministry of Energy and Mines and Petroleum Re sources, Paper 1994-1, p. 259-267. Fraser, T.M., Stanley, C.R., Nikic, Z.T., Pesalj, R. and Gorc, D., 1995, The Mount Polley alkalic porphyry copper-gold deposit, south-central British Columbia: In Schroeter, T., ed., Porphyry Copper (± Au) deposits of the northwestern Cordillera of North America, Canadian Institute of Mining and Metallurgy Special Volume, no. 46, p. 609-622. Gabrielse, H., Monger, J. W. H., Wheeler, J. 0., and Yorath, C. J., 1991, Morphogeological belts, tectonic assemblages, and terranes: In Gabrielse, H., and Yorath, C. J., eds., Chapter 2 of Geology of the Cordilleran Orogen in Canada: Geological Survey of Canada, Geology of Canada, no. 4, p. 15-28. Garwin, S. L., 2002, The geological setting of intrusion-related hydrothermal systems near Bath Hijau porphyry copper-gold deposit, Sumbawa, Indonesia: Society of Economic Geologist Special Publication 9, p. 333-366 Gunow, A.J., Ludington, S., Munoz, J.L., 1980, Fluorine in micas from the Henderson molybdenite deposit, Colorado: ECONOMIC GEOLOGY, v. 75, p. 1127-1137. 36 Guidotti, C.V., Cheney, J.T., and Guggenheim, S., 1977, Distribution of titanium between coexisting muscovite and biotite in pelitic schists from northwestern Maine: American Mineralogist, v. 62, p. 438-448. Guidotti C.V., 1984, Micas in metamorphic rocks. In: S.W. Bailey, ed., Micas, Reviews in Mineralogy vol. 13, Mineralogical Society of America, p. 357-467. Gustafson, L.B., and Hunt, J.P., 1975, The porphyry copper deposit at El Salvador, Chile: ECONOMIC GEOLOGY, v. 70, p. 857-9 12. Godwin, C.I., 1976, Porphyry copper deposits of the calc-alkalic suite: Casino: in Sutherland Brown, A., ed., Porphyry Deposits of the Canadian Cordillera: Canadian Institute of Mining and Metallurgy, Special Volume, 15, p. 344-358. Hanson, R.E., 1991, Quenching and hydroclastic disruption of andesitic to rhyolitic intrusions in a submarine island-arc sequence, northern Sierra Nevada, California: Geological Society of America, Bulletin 103, p. 804-816. Hendry D.A.F., Chivas, A. R J., Long V. P. and Reed, S. J. B., 1985, Chemical differences between minerals from mineralizing and barren intrusions from some North American porphyry copper deposits: Contributions to Mineralogy and Petrology, v. 89, p. 317-329. Hedenquist, J.W., and Henley, R.W., 1985, Hydrothermal eruptions in the Waiotapu geothermal system, New Zealand: Their origin, associated breccias, and relation to precious metal mineralization: ECONOMIC GEOLOGY, v.80, p. 1640-1668. Holdaway, M.J., 1980, Chemical formulae and activity models for biotite, muscovite, and chlorite applicable to pelitic metamorphic rocks: American Mineralogist, v. 65, p. 711-719. Holliday, JR., and Cooke, D.R., 2007, Advances in geological models and exploration methods for copper±gold porphyry deposits; Proceedings of Exploration 07: Fifth Decennial International Conference on Mineral Exploration, p. 79 1-809. Houghton, B.F. and Naim, LA.. 1991. The 1976-1982 strombolian and phreatomagmatic eruptions of White Island, New Zealand: eruptive and depositional mechanisms at a ‘wet’ volcano. Bulletin of Volcanology, v. 54, p. 2549. Jacobs, D.C. and Parry, W.T., 1979, Geochemistry of biotite in the Santa Rita porphyry copper deposit, New Mexico: ECONOMIC GEOLOGY, v. 74, p. 860-887. Jackson, M., Chamberlain, C., Tosdal., R., Rees, C., Ferreira, L., and Taylor, C., 2008, Breccia evolution and distribution of associated alteration and sulfide minerals on sections 29 and 18, Northeast zone, Mount Polley alkalic porphyry Cu-Au deposit, British Columbia: In: MDRU-CODES: Shallow and deep alkalic deposits: Porphyry module -Year 3 meeting. Jébrak, M., 1997, Hydrothermal breccias in vein-type ore deposits: A review of mechanisms, morphology, and size distribution: Ore Geology Reviews, v. 12, p. 111-134. Jensen, E.P., and Barton, M.D., 2000, Gold deposits related to alkaline magmatism: Reviews in ECONOMIC GEOLOGY, v. 13, p. 279-314. Landtwing, M.R., Dillenbeck, E.D., Leake, M.H., and Heinrich, C.H., 2002, Evolution of the breccia hosted porphyry Cu-Mo-Au deposit at Agua Rica, Argentina: Progressive unroofmg of a magmatic hydrothermal system: ECONOMIC GEOLOGY, v. 97, p. 1273-1292. Landtwing MR, Pettke T, Halter WE, Heinrich CA, Redmond, PB, Einaudi, MT., Kunze, K., 2005, Copper deposition during quartz dissolution by magmatic-hydrothermal fluids: The Bingham porphyry: Earth and Planetary Science Letters, 235, p. 229-243. 37 Kroll, T., D., MUller, T., Seifert, PM., Herzig, A., 2002, Petrology and geochemistry of the shoshonite hosted Skouries porphyry Cu-Au deposit, Chalkidiki, Greece: Mineralium Deposita, v. 37, p. 137- 144. Lang, J.R., Stanley, C.R. and Thompson, J.F.H., 1995a, Porphyry copper deposits related to alkalic igneous rocks in the Triassic-Jurassic arc terranes of British Columbia: Arizona Geological Society, Digest 20, p. 2 19-236. Lang, J.R., Thompson, J.F.H., Stanley, C.R., 1995b, Na-K-Ca magmatic hydrothermal alteration associated with alkalic porphyry Cu-Au deposits, British Columbia: In: Thompson, J.F.H., ed., Magmas, fluids and ore deposits, Mineralogical Association of Canada Short Course, v. 23, p. 339-366. Lang, J.R., Lueck, B., Mortensen, J.K., Russell, J.K., Stanley, C.R. and Thompson, J.F.H., 1995c, Triassic Jurassic silica-undersaturated and silica-saturated ailcalic intrusions in the Cordillera of British Columbia: Implications for arc magmatism: Geology, v. 23, p. 451-454. Laznicka, P., 1988, Breccias and coarse fragmentites: Petrology, environments, associations, ores: Amsterdam, Elsevier, p. 832. Logan, J.M. and Koyanagi, V.M., 1994, Geology and mineral deposits of the Galore Creek area, northwestern British Columbia (1 04G/3 and 4): BC Ministry of Energy, Mines and Petroleum Re sources, Bulletin 92. Logan, J.M., 2004. Alkaline Magmatism and Porphyry Cu-Au Deposits at Galore Creek, North-western British Columbia: BC Ministry of Energy, Mines and Petroleum Resources, Geological Fieldwork 2004, Paper 2005-1, p. 237-248. Lentz, D.R. 1994, Exchange reactions in hydrothermally altered rocks: examples from biotite-bearing assemblages. In Lentz, D.R., ed., Alteration and Alteration Processes Associated with Ore-forming Systems: Geological Association of Canada, Short Course Notes 11, p. 69-99. Ludington, S.D., 1978, The biotite-apatite geothermometer revisited: American Mineralogist, v. 63, p. 551- 553. Ludington, S.D and Munoz, J.L., 1975, Application of fluorine-hydroxyl exchanged data to natural micas: Geological Society of America, Abstracts, Programs 7, p 1179. Lueck, B.A., and Russell, J.K., 1994, Silica-undersaturated, zoned, alkaline intrusions within the British Columbia Cordillera; in Geological Fieldwork 1993, paper 1994-1, Mineral Deposits Research unit contribution 038, British Columbia Geological Survey, p. 311-315. Mason, D.R., 1978, Compositional variations in ferromagnesian minerals from porphyry copper-generating and barren intrusions of the Western Highlands, Papua New Guinea: Economic Geology, v. 73, p. 878-890. Mesto, E., Schingaro, E., Scordari, F., Ottolini L., 2006, Electron microprobe analysis, secondary ion mass spectrometry and single crystal X-ray diffraction study of phlogopites from Mt. Vulture, Potenza, Italy: Consideration of cation partitioning. American Mineralogist, v. 91, p. 182-190. McMillan, W.J., 1991, Tectonic Evolution and Setting of Mineral Deposits in the Canadian Cordillera: BC Ministry of Energy, Mines and Petroleum Resources, Ore Deposits, Tectonics and Metallogeny in the Canadian Cordillera, Paper 1991-4. McPhie, J., Doyle. M. and Allen, R., 1993, Volcanic textures: a guide to the interpretation of textures in volcanic rocks; published by Centre for Ore Deposit and Exploration Studies, University of Tasmania, p. 198. Monger, J.W.H., 1977, Upper Paleozoic rocks of the western Canadian Cordillera and their bearing on Cordilleran evolution: Canadian Journal of Earth Sciences, v. 14, p. 1832-1859. 38 Monger, J.W.H., Price, R.A., and Tempelman-Kluit, D.J. 1982, Tectonic accretion and the origin of two major metamorphic and plutonic welts in the Canadian Cordillera: Geology, v. lOp. 70-75. Monger, J.W.H, and Irving, E., 1980, Northward displacement of north-central British Columbia: Nature (London), v. 285, p. 289-294. Mortensen, J.K., Ghosh, D. and Fern, F., 1995, U-Pb age constraints of intrusive rocks associated with copper-gold porphyry deposits in the Canadian Cordillera: In: Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Alkalic Suite, Canadian Institute of Mining and Metallurgy, Special Volume 46, p. 142-158. Micko, J., Tosdal, R.M., Chamberlain, C.M., Simpson, K, Schwab, D., 2007, Distribution of alteration and sulfide mineralization in the Central Zone of Galore Creek, northwestern British Columiba: Arizona Geological Society meeting, Ores and Orogenesis, Program with Abstracts, p. 175 Micko, J., Tosdal, R., Simpson, K., Chamberlain, C., and Schwab, D., 2008, Controls and hydrothermal zonation of Central Zone mineralization at the Galore Creek alkalic Cu-Au porphyry deposit, northwestern British Columbia: In: MDRU-CODES: Shallow and deep alkalic deposits: Porphyry module, Year 3 meeting. Mihalynuk, M.G., Nelson, J.L. and Diakow, L.J., 1994, Cache Creek terrane: oroclinal paradox within the Canadian Cordillera: Tectonics, v. 13, p. 575-595. Munoz, J.L., Ludington, S.D., 1974, Fluoride-hydroxyl exchange in biotite. American Journal of Science, v.274, p. 396-413. Morrisey et al., 2000. Morrisey, M., Zimanowski, B., Wohletz, K., Buettner, R., 2000, Phreatomagmatic fragmentation. In: Sigurdsson, H., Houghton, B., McNutt, S.R., Rymer, H., Stix, J., ed., Encyclopedia of Volcanoes. Academic Press, New York, pp. 431-445. Munoz, J.L., Swenson, A., 1981, Chloride-hydroxyl exchange in biotite and estimation of relative HCIIHF activities in hydrothermal fluids. ECONOMIC GEOLOGY, v. 76, p. 2212-2221. Munoz, J.L., 1984, F—OH and Cl—OH exchange in micas with applications to hydrothermal ore deposits: In: Bailey S.W., ed., Micas: Reviews in Mineralogy vol. 13, Mineralogical Society of America, p. 469-494. Munoz, J.L., 1992, Calculation of HF and HCL fugacities from biotite compositions: revised equations: Geological Society of America. Abstract Programs no. 24, A22 1. Nelson, J.L., and Mihalynuk, M. 1993, Cache Creek ocean: closure or enclosure: Geology, v. 21, p. 173- 176. Nelson, C.E., and Giles, D.L., 1985, Hydrothermal eruption mechanisms and hot spring gold deposits: ECONOMIC GEOLOGY, v. 80, p. 1633-1639. Norton, D. L., and Cathles, L. M., 1973, Breccia pipes - products of exsolved vapor from magmas: ECONOMIC GEOLOGY, v. 68, p. 540-546. Panteleyev, A., 1976, Galore Creek Map-Area: B.C. Ministry of Mines and Petroleum Resources, Geological Field Work, 1975, Paper 1976-1, p. 79-8 1. Passchier, C.W., and Trouw, R.A.J., 2005, Microtectonics; Springer, Berlin. Pearce, T.H., 1968, A contribution to the theory of variation diagrams: Contributions to Mineralogy and Petrology, vol.19, no.2, pp.142-57 39 Perello 3., Cox D., Garamjav D., 2001, Oyu Tolgoi, Mongolia; Siluro-Devonian porphyry Cu-Au-(Mo) and high-sulfidation Cu mineralisation with a Cretaceous chalcocite blanket: ECONOMIC GEOLOGY, V. 96,p. 1407-1428. Perello, J.A., Fleming, J.A., O’Kane, K.P., Burt, P.D., Clarke, G.A., Himes, M.D., and Reeves, A.T., 1995, Porphyry copper-gold-molybdenum deposits in the Island Copper cluster, northern Vancouver Island, British Columbia: In: Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Northwestern Cordillera of North America, Canadian Institute of Mining and Metallurgy, Special Volume 46, p. 2 14-238. Philips, R., 1972, Hydraulic fracturing and mineralization: Journal of Geological Society of London, v. 128, p. 337-359. Philips, W. J., 1973, Mechanical effects of retrograde boiling and its probable importance in the formation of some porphyry ore deposits. Transactions, Canadian Institute of Mining and Metallurgy, B82, p. 90-98 Robert, J.L., 1976, Titanium solubility in synthetic phiogopite solid solutions: Chemical Geology, v. 17, p. 213-227. Rieder, M., Cavazzini, G., Dtyakonov, Y.S., Frank-Kamenetskii, V.A., Gottardi, G., Guggenheim, S., Koval, P.V., Mueller, G., Neiva, A.M.R, Radoslovich, E.W., Robert, J., Sassi, F.P., Talceda, H., Weiss, Z., and Wones, D, 1998, Nomenclature of the micas. The Canadian Mineralogist, 36, Part 3, p. 905-9 12 Rebbert, C.A., Partin, B., Hewitt, D.A., 1995, Synthetic biotite oxidation under hydrothermal conditions: American Mineralogist, v. 80, p. 345-3 54. Russell, J.K., Nicholls, J., Stanley, C.R. and Pearce, T.H., 1990, Pearce element ratios: A paradigm for testing petrologic hypotheses: Transactions, American Geophysical Union, EOS, v.71, p. 234-341. Scordari, F., Ventruti, G., Sabato, A., Bellatreccia, F., Della Ventura, G., and Pedrazzi, G., 2006, Ti-rich phlogopite from Mt. Vulture (Potenza, Italy) investigated by a multianalytical approach: substitutional mechanisms and orientation of the OH dipoles: European Journal of Mineralogy, v. 18, p. 379-39 1. Schwab, D.L, Petsel, S., Otto, B.R., Morris, S.K., Workman, E. and Tosdal, R.M., 2008, Overview of the Late Triassic Galore Creek Copper-Gold-Silver Porphyry System: Arizona Geological Society meeting, Ores & Orogenesis, v. 22, p. 1-14. Schroeter, T., Pardy, J., and Cathro, M., 2004, Significant British Columbia porphyry Cu-Au resources: Min. En. Mines, Geo-file 2004-11, 7 p. Seedorff, E., Dilles, J.H., Proffett, J.M., Einaudi, M. T., Zurcher, L., Stavast, W.J.A., Johnson, D.A., and Barton, M.D., 2005, Porphyry Deposits: Characteristics and Origin of Hypogene Features: ECONOMIC GEOLOGY, 100th Anniversary Volume, p. 251-298. Selby, D., and Nesbitt, B.E., 2000, Biotite Chemistry of the Casino Porphyry Cu-Au-Mo Mineralization, Yukon, Canada: Evaluation of Magmatic and Hydrothermal Fluid Chemistry: Chemical Geology, v. 171, p. 77-93. Sharp, J.E., 1979, Cave Peak, a molybdenum-mineralized breccia pipe complex in Culberson County, Texas: ECONOMIC GEOLOGY, v. 74, p. 5 17-534. Sheridan, M.F., and Wohletz, K.H., 1981, Hydrovolcanic explosions: The systematics of water-pyroclast equilibration: Science, v. 212, p. 1387—1389. Sheridan MF and Wohletz KH, 1983, Hydrovolcanism: basic considerations. In: Explosive Volcanism (MF Sheridan and F Barberi, Eds) Journal of Volcanology and Geothermal Research v. 17, p. 1-29. 40 Sillitoe, R.H., 1991, Geological Reassessment of the Galore Creek Porphyry Copper-Gold Deposit, British Columbia: A report prepared for Kennecott Canada Inc., 8 p. Sillitoe, R.H. 1985, Ore-related breccias in vulcanoplutonic arcs. ECONOMIC GEOLOGY, v. 80, p. 1467- 15 15. Simmons, S. F., and Sawkins, F. J., 1983, Mineralologic and fluid inclusion studies of the Washington Cu Mo-W-Bearing breccia pipe, Sonora, Mexico: ECONOMIC GEOLOGY, v. 78, p. 52 1-526. Sillitoe, R.H., Baker, E.M., Brook, W.A. 1984, Gold deposits and hydrothermal eruption breccias associated with a maar-volcano at Wau, Papua New Guinea. ECONOMIC GEOLOGY, v. 79, p. 638-655. Sillitoe, R.H., and Bonham, H.F.J., 1984, Volcanic landforms and ore deposits: ECONOMIC GEOLOGY, V. 79, p. 1286-1298. Speer J.A., 1984, Micas in igneous rocks: In: Bailey S.W., ed., Micas: Reviews in Mineralogy vol. 13, Mineralogical Society of America, p 299-356. Skews, M. A., and Stren, C.R., 1996, Late Miocence mineralized breccias in the Andes of Central Chile: Sr- and Nd-isotopic evidence for multiple magmatic sources: Society of ECONOMIC GEOLOGY Special Publication 5, p. 33-41. Stormer, J.C., Jr., Carmichael, I.S.E., 1971, Fluorine-hydroxyl exchange in apatite and biotite: a potential igneous geothermometer: Contributions to Mineralogy and Petrology, v. 31, p. 121-131. Stanley, C.R., 1992, The Igneous Petrogenesis of the Galore Creek Intrusive Suite. In: Porphyry Copper Gold Systems of British Columbia. Annual Technical Report Year 1. Stanley, C.R., Holbek, P.M., Huyck, H.O., Lang, J.R., Preto, V.A.G., Blower, S., Bottaro, J.C. (1996): Geology of the Copper Mountain Alkalic Cu-Au Porphyry Deposit, Princeton, British Columbia. in ‘Porphyry Copper Deposits of the Northern Cordillera. ed. by Schroeter, T., Canadian Institute of Mining and Metallurgy, Special Vol. No. 46, Paper No. 43, p. 537-564. Thomas, J.A., and Galey, J.T., Jr., 1982, Exploration and geology of the Mt. Emmons molybdenite deposits, Gunnison County, Colorado: ECONOMIC GEOLOGY, v. 77, p. 1085-1104. Titley, S.R., and Beane, R.E., 1981, Porphyry copper deposits, Part 1. Geologic settings, petrology, and tectogenesis: Economic Geology 75th Anniversary Volume, p. 214-235. Tischendorf, G., Förster, H.-J. and Gottesmann, B., 2001, Minor- and trace-element composition of trioctahedral micas: a review: Mineralogical Magazine, v. 65, p. 249-276. Twelker, E., 2007, A breccia-centered ore and alteration model for the Copper Canyon alkalic Cu-Au deposit, British Columbia: M.Sc. thesis, University of Alaska Fairbanks, p. 75-83. Wallace, S.R., MacKenzie, W.B., Blair, R.G., and Muncaster, N.K., 1978, Geology of the Urad and Henderson molybdenite deposits, Clear Creek County, Colorado, with a section on a comparison of these deposits with those at Climax, Colorado: Economic Geology, v. 73, p. 325-368. Wernicke, B., and Klepacki, D.W. 1988, Escape hypothesis for the Stikine block: Geology, v.16, p. 461- 464. Wheeler, 3.0. and McFeely, P. 1991, Tectonic assemblage map of the Canadian Cordillera and adjacent parts of the United States of America; Geological Survey of Canada, Map 17 12A, scale 1:2 000 000. Wolfe, J.A., 1980, Fluidization versus phreatomagmatic explosions in breccia pipes; ECONOMIC GEOLOGY, volume 75, p. 1105-1111. 41 Wolfe, R. C., Cooke, D. R., and Joyce, P., 1999, Genesis of the alkaline Dinkidi Cu-Au porphyry deposit, North Luzon, Philippines, in PACRIM 1999 (Proceedings, 1999 PACRIM Congress): Melbourne Australia, Australasian Institute of Mining and Metallurgy. Waters, D.J. and Charnley, N.R., 2002, Local equilibrium in polymetamorphic gneiss and the titanium substitution in biotite: American Mineralogist, v. 87, p. 383-396. Wones, D. R., 1963, Physical properties of synthetic biotites on the join phlogopite-annite: American Mineralogist, v. 48, p. 1300-1321. Wones, D.R., and Eugster H.P., 1965, Stability of biotite: experiment theory and application: American Mineralogist, v. 50, p 1228-1272. Wones, D.R., 1972. Stability of biotite: a reply. American Mineralogist, v. 57, p. 316-317. Wones, D.R., 1981, Mafic silicates as indicators of intensive variables in granitic magmas: Mining Geology (Japan), v. 31, p. 191-212 Wohletz KH and Sheridan MF, 1983, Hydrovolcanic explosions II. Evolution of basaltic tuff rings and tuff cones: American Journal of Science, v. 283, p. 385-413. Wright, A. and Bowes D. R., 1963, Classification of volcanic breccias: a discussion: Bulletin Geological Society of America, v.74, p. 76-86. Zhu, C., Sverjensky, D.A., 1991, Partitioning of F—Cl—OH between minerals and hydrothermal fluids: Geochimica et Cosmochimica Acta v. 55, p. 1837-1858. Zhu, C., Sverjensky, D.A., 1992, F—Cl—OH partitioning between biotite and apatite: Geochimica et Cosmochimica Acta, v. 56, p. 3435-3467. Zweng, P.L. and Clark, A.H., 1995, Hypogene evolution of the Toquepala porphyry copper-molybdenum deposit, Moquegua, southeastern Peru, in Pierce, F.W., and Bolm, J.G., eds., Porphyry copper deposits of the American Cordillera: Arizona Geological Society Digest 20, p. 566-612. 42 Chapter 2: Magmatic-hydrothermal evolution and zonation of a breccia centered Cu-Au alkalic porphyry: Southwest Zone, Galore Creek. 2.1 Introduction Characteristics of caic-alkalic porphyry deposits and the processes involved in their formation have been well documented and summarised succinctly by Titley and Beane (1981), Seedorff et al. (2005) and Holiday and Cooke (2007). In contrast, alkalic porphyry deposits are relatively under reported and no unifying model for alkalic systems exits. Work by Lang et al. (1995b and c), Jensen and Barton (2000) and Cooke et a!. (2007) highlights the economic significance of alkalic porphyry Cu-Au deposits and notes the subtle but significant differences from caic-alkalic systems, and variations within the porphyry class itself. Additionally, alkalic porphyry deposits are exclusively of Cu—Au character (Cooke et al., 2007) and distinguished by high tenor Au mineralisation, making them attractive exploration targets. In the alkalic environment complicated K—Ca—Na metasomatic zones are typically developed (Lang et al., 1995), and both potassic and caic-potassic alteration assemblages can be intimately associated with mineralisation (Cooke et al., 2007 and references therein). Alkalic porphyry Cu-Au deposits are known in only a few metallogenic terranes, notably the Triassic and Jurassic marine volcanic arcs of British Columbia (Fig. 1), Canada (Barr et al., 1976; Lang et al., 1995a) and the Ordovician and early Silurian Lachlan Fold Belt in New South Wales, Australia (Cooke et al., 2007). The Galore Creek alkalic Cu-Au-Ag porphyry is located within the Stikinia terrane at the western margin of the Intermontane Belt in the Canadian Cordillera, northwestern British Columbia (Fig. 2.1). Host rocks and intrusions in the district are silica-undersaturated and hydrothermal alteration is distinguished by deficiency of quartz gangue (Enns et aL, 1995). A version of this chapter will be submitted for publication. Byrne, K., and Tosdal, R., Magmatic hydrothermal evolution and zonation of a breccia centered Cu-Au alkalic porphyry: Southwest Zone, Galore Creek. 43 The Southwest Zone Cu-Au breccia-centered deposit is one of twelve mineralized centers in the Galore Creek alkalic porphyry district and provides an important case study in alteration and mineralisation zonation, and the role of breccias in alkalic porphyry systems. Furthermore, the deposit is notable for its high tenor Cu-Au and a Au- halo (Enns, et al., 1995; Schwab et a!., 2008). FIG. 2.1 Map of British Columbia showing the location of the accreted Quesnellia and Stikinia ocean arc terranes, major alkalic Cu-Au porphyry deposits and alkalic intrusive centers, and morphogeological belts. Galore Creek is located at the western margin of the Intermontane Belt about 70km east of Wrangell, Alaska. Data sourced from BC MapPlace (http://www.mapplace.ca/). Research aims are to characterize and map the three-dimensional distribution of coherent and elastic rocks, alteration and mineralization assemblages, and their 44 paragenesis. This paper constitutes the first detailed description of the geology, host rocks, alteration systematics, and structural evolution of the Southwest Zone. The data presented is based on detailed dill-core logging of bore holes and a compilation of NovaGold Inc. data on two cross-sections. We emphasize controls on Cu-Au localization, the distribution of alteration, mineralisation and metal assemblages, and the processes involved in their zonation. Furthermore, we submit that selectively pervasive garnet alteration in the peripheries of the deposit is equivalent to propylitic alteration in calc-alkalic porphyry systems. Descriptive and interpretive elements are combined to propose an ore deposit model. Hosted in brecciated intrusive rocks that are distinctly younger than Central Zone Cu-Au mineralisation, the Au-rich Southwest Zone preserves late stages of magmatic hydrothermal activity in the Galore Creek district (Enns et al., 1995; Schwab et al., 2008). With a measured and indicated resource of 69 Mt grading 0.42% Cu, 0.63g/t Au, 2.67g/t Ag, the Southwest Zone contributes considerably to the total measured and indicated resources in the Galore Creek district, which are 785.7 Mt at 0.52% Cu, 0.29g/t Au and 4.87g/t Ag (NovaGold Resources Inc., February, 2008). Abbreviation used in the test and figures are listed in Table 2.1. TABLE 2.1 List of abbreviations used in figures and tables. act actinolite (ferro-) diop diopside (ferroan-) orth orthoclase alb albite diss disseminated plag plagioclase alt alteration epi epidote phi phiogopite amph amphibole fsp feldspar py pyrite anh anhydrite goe goethite pyro pyrophyllite and andalucite gnt garnet (grandite) pyx pyroxene ank ankerite gm groundmass qtz quartz ap apatite hbl homblende rt rutile bio biotite kaol kaolinite ser sericite bn bomite K-Spar K-Feldspar sph sphene carb carbonate mar margarite tal talc chi chlorite musc muscovite wol wolastinite cpy chalcopyrite mt magnetite 45 2.2 Exploration History The Galore Creek district was discovered in 1955, by testing anomalies in a regional aero-magnetic survey. It has been explored by several companies since (Enns et al., 1995). Early district field evaluations of magnetic-high geophysical anomalies led to discovery of Southwest Zone mineralisation in addition to other zones (Barr, 1966). Since 2003, NovaGold Resources Inc. has discovered the Bountiful and West Fork zones which have added considerable resources (SEDAR report #). In 2007, NovaGold Resources Inc. and Teck formed a joint-venture and the property is now managed through the Galore Creek Mining Corporation. 2.3 Regional Geological Setting The Canadian Cordillera is one of several orogenic belts, marginal to the Pacific Ocean, that form the Circum-pacific orogenic system extending from Alaska to Mexico. A collage of allochthonous oceanic and proximal to distal pericratonic terranes were accreted to the western margin of the North American craton during the Late Paleozoic through to Late Mesozoic (Monger, 1977; Monger and Irving, 1980; Monger et al., 1982; Coney, 1989). From east to west, the Foreland, Omineca, Intermontane, Coast, and Insular belts are the five morphogeological belts that compose the Canadian Cordillera (Fig. 1). Each belt is characterized by a distinctive combination of land forms, rock types, metamorphic grade and structural style (Coney, 1989; Gabrielse et al., 1991). McMillan (1991) groups Stikinia, Cache Creek, Slide Mountain and parts of Quesnellia and Yukon Tanana terranes into the Intermontane Superterrane. Quesnellia and Stikinia arcs host several similar aged alkalic intrusive centers and Cu-Au deposits (Fig. 2.1) and caic alkalic Cu-(Mo-Au) porphyry deposits. Similarities in rock type and geologic history between Stikinia and Quesnellia, including the presence of the silica-undersaturated alkalic porphyry deposits, have led workers to believe that they are segments of the same Triassic arc (Wemicke and Kiepacki, 1988; Nelson and Mihalynuk, 1993; Mihalynuk et al., 1994). The alkalic Cu-Au deposits in both Stikinia and Quesnellia terranes are products of two discrete alkaline magmatic events at the end of the Triassic and early Jurassic (Mortensen et al., 1995) and are interpreted to have formed outboard of ancestral North America in island-arcs tectonic settings (McMillan, 1991). Figure 2.2A shows the 46 northwest part of the Intermontane belt, Paleozoic and Mesozoic rocks in of the Stikinia terrane, and the location of Galore Creek and other base and precious metal deposits in the region: Eskay Creek, Red Chris, Shaft Creek and Copper Canyon. Intrusive rocks Volcanic and Sedimentray rocks Coast Plutonic Complex Stuhini Group (Eocene) (Mid to Upper Triassic) Caic-alkalic Plutonic suite Shale and argilite (Early Jurassic) (Middle Triassic) Galore Creek alkalic suite Stikirie assemabige(Upper Triassic) (Devonian to Permian) • Hickman Batholith • Limestone(Permian) FIG. 2.2 A. Major tectono-stratigraphic elements and the location of the Galore Creek district, Eskay Creek, Red Chris, Schaft Creek and Copper Canyon in the Stikinia terrane (modified from Wheeler and McFeely 1991; Gabrielse et al. 1991; Logan and Koyanagi, 1994). B. Regional scale geology of Galore Creek showing the location of Copper Canyon alkalic porphyry Cu-Au occurrence (modified after Logan and Koyanagi, 1994 and Enns et al., 1995). BT = Butte Thrust Fault; CCT = Copper Canyon Thrust Fault. 2.4 District Geology The Galore Creek alkalic intrusive suite is one of the largest and most silica undersaturated complexes to host porphyry Cu-Au deposits (Fig. 2.2B) and is therefore an end member example of the alkalic porphyry Cu-Au class (Lang et al., 1995). Galore Creek alkalic intrusions were emplaced between 2 10±1 and 200.1±2.2 Ma (Mortensen et al. 1995) and are hosted in supracrustal Upper Triassic Stuhini group rocks (Fig. 2B). Syenite, monzonite and monzodiorite dikes are hosted in rocks of shoshonitic affinity consisting of augite-phyric intermediate volcanic rocks, pseudoleucite-bearing phonolites 132 13O Other deposits 1. Copper Canyon 2. Schaft Creek 3. Red Chriss 4. Eskay Creek 47 and associated volcaniclastic and volcanogenic rocks (Lang et al., 1995a and b). Kennecott Corporation established a sequence of twelve (11-112) intrusions for the Galore Creek alkalic intrusive suite and these were further characterised by Enns et a!. (1995). This study however, notes deviations in the intrusion sequence outlined by Emis et al. (1995). Nearby, in the Copper Canyon area porphyry Cu-Au mineralisation is also hosted in alkalic intrusions (Bottomer, 1995). Hydrothermal alteration and mineralization is developed in the multiphase complex of alkalic intrusion and host shoshonitic volcano-sedimentary rocks. Twelve zones of Cu-Au mineralisation have been defined to data in the district (Fig. 2.3). The largest deposit, and best-documented, is the northerly-elongated Central Zone (Lang et al., 1995b; Micko et al., 2007 and 2008). Divided into three domains, each with characteristic alteration style and metal tenor (Schwab et al., 2008), the Central Zone is comprised of the North Gold Lens, Central Replacement Zone and the South Gold Lens. Smaller deposits peripheral to the Central Zone are also known, and include: Southwest Zone, Junction and North Junction, Butte, West Rim, Westfork and the Saddle zones (Enns et al. 1995; Schwab et al., 2008). Figures 2.4A and 4B are photographs illustrating the rugged terrane and the location of the mineralised centers in the Galore Creek district. Timing relationships of mineralization to the various intrusive units suggests at least two distinct periods of hydrothermal activity punctuated by intrusion of voluminous megacrystic orthoclase-phyric syenite and monzonite dikes. Early mineralisation occurs in the Central Zone and is hosted mostly in supracrustal rocks with subordinate Cu-Au in intrusions and hydrothermally cemented breccias. No obvious intrusive centre for mineralisation in the Central Zone has been recognized. Mineralisation in the Central Zone is truncated to the west by post-mineral megacrystic orthoclase-phyric syenite and monzonite dikes (Enns et al., 1995; and Schwab et a!., 2008). A second stage of mineralisation occurs in the same megacrystic orthoclase-phyric syenite and monzonite in the Southwest Zone (Fig. 2.3) and Middle Creek and Junction (Schwab et al., 2008). 48 FIG. 2.3 Simplified geological, and alteration and mineralisation map of the Galore Creek district illustrating the location of mineralized centers and major structural features. Alteration and mineralisation assemblages are projected from a depth approximately mid-way though each center. Base map is courtesy of NovaGold Resources and has been modified from the original version. Rock type and distribution is from drill-hole and outcrop data. Alteration-mineralisation assemblages and distribution is modified from Schwab et al. (2008). NGL = North Gold Lens; CRZ = Central Replacement Zone; SGL = South Gold Lens; SWZF = Southwest Zone Fault; WFF = West Fork Fault; BF = Bountiful Fault; EF = East Fault. Alteration- mineralization bornite -chalcopyrite chalcopyrite - pyrfte pyrite bornite -magnetite diopside — -chalcopyrite K-feldspar -biotite garnet-anhydrite -biotite sericite-arthydrite -carbonate chlorite-epidote -carbonate Rock Types hydrothermal breccia monzodiorite z 0 m ‘0 349000E 352000E megacrystic() orthoclase-phyric 0: syenite and monzonite (19a and b) dark orthoclase() porphyry syenite; * + + ± pseudoleudte bearing (14a and b) sedimentary augite and feld par-bearing volcanic / volcaniclastic pseudoleucite-bearing o volcanic / volcaniclastic -i North 0 km 49 FIG. 2.4 Photographs of the Galore Creek valley, looking A south, and B north. Outline of mineralised centers in yellow. Saddle prospect is not shown, but is to the upper mid-left and bottom right parts of frames A and B, respectively Three phases of deformation are recognized for the oldest Paleozoic rocks and one phase for Upper Triassic strata within the Stikine-Iskut River region (Panteleyev, 1976; Logan and Koyanagi, 1994; Logan, 2004). The supracrustal rocks at Galore Creek are interpreted to have undergone early and broad-scale post-Triassic north-south compression followed by post-early Jurassic open to tight folding and related faulting from east-west compression. East-west compression led to the formation of northerly trending folds and thrust faults (Logan and Koyanagi, 1994). The Central Zone’s local structural architecture is not well defined. However, the spatial distribution and orientation of lithological units as well as associated hydrothermal alteration halos imply tilting of the Central Zone 45-60° east (Micko et al., 2008). Tilting is interpreted to have occurred at the end of hydrothermal activity due to N-trending, upright folding in the Jurassic (Logan and Koyanagi, 1994). The most visible structures within the Galore Creek district are the post-mineral west-dipping Butte Thrust Fault (Schwab, et al., 2008) and the east-dipping Copper Canyon Thrust Fault (Bottomer and Leary, 1995), these are shown in Figure 2.2B. 50 2.5 Rock types of the Southwest Zone The Southwest Zone is situated 6OO m southwest of the South Gold Lens in the southern part of the Galore Creek intrusive-volcanic complex (Fig. 2.3), and is masked by glacial cover (Fig. 2.4). Description of lithofacies is based on re-logging of drill core and petrographic examination. In this contribution, rocks of the Southwest Zone are divided into two principal groups; coherent and clastic. Coherent facies are rocks formed from the cooling and solidification of magma and are characterised by aphanitic or phaneritic textures. Clastic facies are any fragmental rock and their descriptions follow the scheme outlined in McPhie et al. (1993). Additionally, characterization of clastic facies on the basis of infihl types (matrix and cement) is modified from Davies et al. (2008a). Matrix is defined as a fine-grained clastic component that occurs between larger clasts. Matrix comprises comminuted wall rock (rock-flour), lithic, and crystal fragments of sand to granule size (<O.5m — 4mm). Cement is defined as a crystalline component within the clastic rock precipitated from an aqueous fluid. Matrix-bearing breccia refers to clastic rocks with a component of matrix, with or without any additional cement. In the Southwest Zone, megacrystic orthoclase-phyric syenite and monzonite coherent facies are cross-cut matrix-bearing breccias and together these rocks host Cu Au (Fig. 2.5). 51 349700 E 345100 C 345500 E ‘% / ll — ll RockUnits % II - ••%•• 11% IN Lamprophyre % Matrix-bearing breccia Copper Distribution 0 200 Megacrystic orthoclase 0.1 -0.5% Cu > 0.5% CuI and plagioclase-phyric Near-surface Cu-ore; porphyries ..-.. bottom surafce and top —.. surface with dip arrow Southwest Zone • Cu-Ore at depth projectedFault down-dip undercover FIG. 2.5 Simplified bedrock geology plan map and Cu-grade distribution in the Southwest Zone. Location of section line A-A’ (6333650N) and B-B’ (350030E) indicated with black lines. Fire assay data was provided by NovaGold Resources Inc. 2.5.1 Coherent rocks Of the twelve intrusive rock types defined Enns, et a!. (1995) to define the Galore Creek Alkalic suite, eight are recognized in the Southwest Zone. Numerous dikes of intermediate to mafic to in composition are the youngest rocks in the Southwest Zone and are unrelated to Galore Creek suite magmatism (Enns, et a!., 1995). Textural characteristics and sequence of emplacement of coherent unit belonging to the Galore Creek alkalic suite are summarized in Table 2.2 and illustrated schematically in figure 2.6. It 52 TABLE 2.2 Summary of Southwest Zone coherent rocks. Unit’ # Coherent facies Name2 Lithology description Timing3 1 Feldspar-phyric 17 Phenocrysts: 1-2% lath and equant feldspars 0.5-1 cm. Pre syenite Groundmass: medium-grained equigranular. Distinguishing features: Forms thick bodies with unclear geometries. Appears to have variable phenocryst population containing some megacrysts. 2 Megacrystic 19A Phenocrysts: 10-20% 1-6 cm tabular orthoclase and -5% 0.2- Pre orthoclase- 1.0 cm lath and equant orthoclase Groundmass: medium phyric syenite grained salt and pepper texture, 7-15% mafics, 60% Fsp 3 Megacrystic 19B Phenocrysts: 5-20% 1-6 cm tabular Ortho, 5-10% 0.2-1.0 cm Pre orthoclase and lath and equant orthoclase, 20-35% 0.2-0.5 cm lath and plagioclase- subhedral plagioclase. Groundmass: medium-grained, 5% phyric mafics, 30-40% Fsp. monzonite 4 Megacrystic Ill Phenocrysts: 5-15% 0.5-2cm, lath to tabular orthoclase. Pre orthoclase- Groundmass: medium-grained equigranular. Distinguishing phyric features: trachytic orthoclase phenocrysts monzonite 5 Acicular- WFP Phenocrysts: 5-25% 0.2-1cm acicular feldspar Groundmass: Pre-syn? feldspar-phyric Dark fine-grained. Distinguishing features: Sometimes syenite displays irregular clast margins occurs throughout the matrix- bearing breccia body in minor amounts. 6 Biotite-phyric 16, Phenocrysts: 15-25% 1-2 mm biotite, 3% 2-3 mm Post monzodiorite 18B homblende needles and 60-70% feldspar. Distinguishing features: centered in the upper cemented breccia domain. Typically mineralised by fine-grained chalcopyrite. 7 Pyroxene, Vl Phenocrysts: 5-10% 0.2-0.8 cm hexagonal-shaped pyroxene Post homblende and related 5-10% 0.1-0.5 cm biotite (booklets), >3% 0.2-1 cm Hbl and biotite-phyric dikes? 3-5% 0.2-0.4 cm equant Fsp. Groundmass: dark green fine diorite grained. Distinguishing features: Host to weak Cu mineralisation. 8 Orthoclase and 19(C) Phenocrysts: 5-15% 0.5-1.5 cm Ortho (rare megacrysts) and Post plagioclase- 10-25% 0.2-0.6 cm anhedral plagioclase grains. Groundmass: phyric Dark fine-grained with 5-20% mafics. Distinguishing monzonite features: Glomerophyric feldspars phenocrysts. 9* Xenolith- Dl Phenocrysts: 10-20% 0.1-0.2 cm Bio and 5-10% 0.3-0.5 cm Post bearing Hbl. Groundmass: dark green fine-grained. lamprophyre 10* Mafic to D2-D3 Phenocrysts: 10% 0.2-0.3 cm lath shaped feldspars. Post Intermediate Groundmass: fine-grained to aphanitic dikes ‘Oldest to youngest units listed from top to bottom. Sequence of coherent rocks is based on observed and inferred cross-cutting relationships. 2Name refers to intrusive rock nomenclature used in Enns et al., 1995 and NovaGold Resources Inc. (and references therein). 3Timing is with respect to the formation of matrix-bearing breccia. *Not part of the Galore Creek Alkalic suite. 53 Coherent rocks 1 Orthoclase and plagioclase -phyric monzonite E9 Pyroxene,hornblende and biotite-phyric diorite Biotite-phyric monzodiorite El Acicular-feldspar-phyric syenite j Megacrystic orthoclase phyric monzonite Megacrystic orthodase and plagioclase-phyric monzonite Megacrystic orthoclase phyric syenite W Feldspar-phyric syenite Clastk rocks ® Cemented breccias (iJ Matrix-bearing breccia FIG. 2.6 Sequence of coherent and clastic rock emplacement in the Southwest Zone. Coherent units are grouped as pre- or post-matrix bearing breccia based on observed and inferred cross-cutting relationships. The six most significant coherent units are discussed here in detail, and these are: Megacrystic orthoclase-phyric syenite is characterized by 1-6 cm tabular orthoclase and 5% 0.2-1.0 cm lath and equant orthoclase phenocrysts, and a salt and pepper textured groundmass of K-feldspar, hornblende and biotite (unit 2, Table 2.1; Fig. 2.7A). The unit occurs as thick (>lOOm) composite dikes that intrude feldspar-phyric syenite (unit 1, Table 2.2). Megacrystic orthoclase and plagioclase-phyric monzonite is distinct from other megacrystic units by due to the presence of plagioclase phenocryst (unit 3, Table 2.2; Fig. 2.7B). These megacrystic porphyries (units 2 and 3) host younger less voluminous intrusions, and are the dominant wall rock to younger clastic rocks (Fig. 2.6). Acicular feldspar-phyric syenite forms thin (0.5-3 m) dikes cross-cut the megacrystic porphyries (unit 5, Table 1; Fig 2.6.). Acicular-feldspar-phyric syenite also occurs as an accessory clast type in matrix-bearing breccias TIME 54 Post-matrix-bearing breccia coherent units are volumetrically minor and hosted in thick composite dikes of megacrystic orthoclase-phyric syenite and monzonite, and elastic rocks, these are; Biotite-phyric monzodiorite post-dates matrix-bearing breccia and is spatially coincident with much of the high-grade Cu-Au in the Southwest Zone (unit 6, Table 2.2; Fig. 2.7C). Pyroxene and hornblende and biotite-phyric diorite occur as dikes less than 2 m wide, cutting cemented breccias. This unit locally contains minor Cu at its margins as veins and disseminations (unit 7, Table 2.2; Fig. 2.7D). Orthoclase and plagioclase-phyric monzonite forms 1-5 m wide dikes, cross-cut all breccia facies, are distinguished by characteristic glomerophyric-feldspar phenocrysts (unit 8, Table 2.2; Fig. 2.7E) and are post-Cu mineralization. 55 FIG. 2.7 Photographs of coherent rocks in the Southwest Zone. A. Megacrystic orthoclase-phyric syenite characterised by salt-and-peppered-texture groundmass (unit 2). B. Megacrystic orthoclase and plagioclase-phyric monzonite dike (unit 3). C. Biotite-phyric monzodiorite (unit 6). D. Pyroxene, hornblende and biotite-phyric diorite. Groundmass is pervasively chlorite altered (unit 7). Epidote-gamet forms clots in the groundmass. E. Orthoclase and plagioclase-phyric monzonite dike characterised by glomerophyric orthoclase phenocrysts (unit 8). Mineral abbreviations in italics refer to primary igneous minerals as apposed to alteration minerals. I cm 56 Biotite phyric monzonite occurs as dikes and dikelets which are focused at the contact between the matrix-bearing breccia and porphyry wall-rocks. Where it intrudes the matrix-bearing breccia, biotite-phyric monzodiorite occurs as volumetrically minor interconnected dikelets. The distribution of the megacrystic porphyry syenite and monzonite, biotite phyric monzonite, and various elastic facies are shown in figures 2.8 and 2.9 along cross-sections A-A’ (633650N) and B-B’(350030E), respectively. WEST f A• (/‘ ,, Upper (main) fl, —TI I11 11. ...‘ h J - I ... ‘ ‘ d e ‘DDJDC750mc0J b’’ - ,‘ I I V gJclJ ,ll 1U /‘ll I I! c) o I . ,9 , % - .; Q If dd d ct f’ U contactou?ofsection % DnIl core logged: 90 m • onsection/offsection. n • % •. çp cQ Drill-hole logs reviewed; , core either not logged or unavailable. . ..•. meters —‘ o 150 Clastic: cement-bearing breccia — Monolithic in-situ cement-dominatedas ic. ma FIX- earing reccia C-BX breccia (no matrix present, C-BX) Pebble to cobble clast size (with — 2 o <10% cement M-BX) — — - Cement-dominated breccia (>40% Cobble to boulder clast size (with ‘. cement, CM-BX) ‘. <10% cement M-BX) I Matrix-dominated breccia with (10 ‘— —, to <40% cement, MC-BX)Coherent: r Megacrystic orthclase-phyric Coherent: syenite and monzon,te; thick Biotite-phyric monzodiorite. L composite dikes form wall-rock to Ok Domain ofmatrix-bearing breccias) i es irregular dikelets FIG. 2.8 Distribution of elastic facies and simplified coherent units along cross section 6333650N, line A- A’ as indicated in Figure 2.5. A 50-60 m drill-hole data threshold applied to section 6333650N (A-A’) in all figures. SWZF = Southwest Zone Fault. 57 FIG. 2.9 Distribution of clastic facies and simplified coherent units along cross section 350030E along line B-B’ in Figure 2.5 and 2.7. Approximately 40 m drill-hole data threshold applied to section 350030E (B B’) in all figures. SWZF = Southwest Zone Fault. Within the biotite-phyric monzodiorite dike facies fine-grained margins are observed; in contrast, the peripheral dikelet facies is characterised by irregular-shaped margins with local alignment of biotite phenocrysts (Fig. 2.1 OA). NORTH B Section A-A’ (projected from TOm west) GC61-003 GC61 -004 6333900 N 0C05-651 GC6I -005 %ll %‘ o:; SOUTH B’ G005-677 0 GCO4-502 0C05-642 _________ — \ flll % W ll . Zg% ,%fl %%% %II So 750m \ll l Dike %ll /;Lower cemented -‘ breccias - Inferred dip direction of contact out of section. Drill core logged: on section I off section. Drill-hole logs reviewed; \ core either not logged or unavailable. meters 600 m 150 Upper (main) cemented breccia Elevation (meters) 0 U C 0 z 0 0 “S ,=I% %fl,:- -J (.%fl% & 4% & • C) Clastic: cement-bearing breccia Monolithic in-situ cement-dominated — —, breccia (no matrix present, C-BX) — Cement-dominated breccia (>40% cement, CM-BX) Matrix-dominated breccia with (10 ‘— —, to <40% cement, MC-BX) Coherent: Biotite-phyric monzodiorite: Dikes Domain of V irregular dikelets 450 m -t 6333300 N Clastic: matrix-bearing breccia Pebble to cobble clast size (with <10% cement M-BX) Cobble to boulder clast size (with <10% cement M-BX) Coherent:[ Megacrystic orthclase-phyricsyenite and monzonite; thick composite dikes form wall-rock to matrix-bearing breccias) 58 FIG. 2.10 Photograph of diagnostic features biotite-phyric monzodiorite dike and dilcelet facies. A. Dikelet facies intrudes MC-BX and exhibits weak flow fabric near the margins (white-dashed line). B. Cement (possible miaroles) composed of quartz-K-feldspar-chalcopyrite-anhydrite in biotite-phyric monzodiorite. Quartz related to potassic-D stage alteration is extremely rare; this is one ofjust two examples. C. Cement (possible miaroles) composed of K-feldspar-phlogopite-anhydrite-chalcopyrite in biotite-phyric monzodiorite. The dikelet facies may reflect peripheral fingering margins of the dike facies where it becomes disaggregated. Biotite-phyric monzodiorite is further characterised by volumetrically minor irregularly shaped vugs concentrically filled by hydrothermal minerals (Figs. 2.1 OB and C). Numerous dikes of mafic (unit # 10 group, Table 2.2) to intermediate in composition are the youngest rocks in the Southwest Zone and not related to Galore Creek alkalic suite magmatism (Enns et al., 1995). 59 25.2 Clastic rocks There are three principal clastic facies in the Southwest Zone based on the relative abundance of cement to matrix infihl, and these are: 1) M-BX, matrix-bearing breccia with negligible cement; 2) MC-BX, matrix-bearing breccias with 10-40% cement (matrix-dominated) and CM-BX, matrix bearing breccia with> 40% cement (cement-dominated) 3) C-BX, monolith-in situ (jig-saw fit) cement-dominated breccias (matrix absent) The distribution and geometry of the principal clastic facies are described first. Description of breccia components, clasts, matrix and cement follow. The paragenetic relationships between matrix-bearing breccias, cemented breccias and coherent units are described in section 2.5.3. 2.5.2.1 Matrix-bearing breccias The matrix-bearing breccias are approximately 400m wide by 800m long, extend to at least 600m below the surface, and are discordant to the trend of surrounding pre fragmentation porphyry dikes. The western matrix-bearing breccia-wall rock contact, in its present geometry, is locally overhanging and broadly trends N and dipping 60-70°W (Figs. 2.5 and 2.8). In contrast, the eastern contact is less well defined but appears to have a similar strike and dip. In plan view the breccia body has an irregular oval shape elongate to the north (Fig. 5). Section B-B’(Fig. 9) cuts a steeply west dipping overhanging protrusion of matrix-bearing breccia resulting in the observed rollover of the wall rock contact from north to south. The northern and southern matrix-bearing breccia wall rock contacts are not as well defined, although available drilling suggests the breccia body may continue south. Matrix-bearing breccias with no appreciable cement component (M-BX) are the most abundant breccia facies. Matrix-bearing breccias are predominantly polylithic, unbedded, poorly sorted to chaotic (massive), matrix-rich and matrix-supported. Features of the matrix-bearing breccia facies are summarized in Table 2.3. 60 TABLE 2.3 Summary of Southwest Zone clastic rocks. Clastic facies Sub- Cement I vein Grouping nomenclature1 Infill proportion2 facies13 mineralogy Contacts Lithology description M-BX: matrix- < 10% cement, M-BX-M phi, K-Spar, diop, Sharp to gradational contact Predominantly polylithic, bearing breccia 90% matrix (pebble- act, anh, mt, py, cpy with wall-rock marked by unbedded, unsorted to poorly cobble) monolithic very coarse facies sorted, matrix-rich and matrix supported. Clasts are sub 8. M-BX-C angular to sub-rounded and (cobble- pebble to boulder sized. The . boulder) clast to matrix ratio averages —70:30 E MC-BX: matrix- 10 to < 40% MC-BX-P phi > K-Spar ± anh, Gradational from weak Fractures, irregularly shaped dominated breccia cement, mt, diop, cpy, bn stockwork veins and locally veins and vugs are filled with remainder is sharp. varying proportions of cement matrix MC-BX-K K-Spar > phi ± anh, minerals. Sub-facies are . mt, diop, cpy, bn distinguished by cement type and abundance. Ciast diameters commonly greater cu CM-BX: cement- > 40% cement, CM-BX-P phi> K-Spar ± anh, Over short distances CM-BX than the drill-core width.8 dominated breccia remainder is mt, diop, cpy, bn is transitional from MC-BX. matrix Sharp contact between CM CM-BX-K K-Spar> phi ± anh, BX and M-BX also noted. . mt, diop, cpy, bn Late dikes typically occur proximal to cemented facies .D) contacts. cu C-BX: monolithic > — 90% C-BX-P phi > K-Spar ± anh, C-BX facies is hosted in in situ or jig-saw fit breccias in-situ cement- cement, no mt, diop, cpy, bn intrusive wail rocks and is are typical very coarse (large >. domainted appreciable contiguous with CM-BX and boulders) breccia matrix MC-BX. C-BX is transitional component C-BX phi, K-Spar, diop, to weak stock-work veins ci) act, anh, mt, py, cpy e O.2 ‘Clastic facies nomenclature and sub-facies refers to the scheme applied for sub-facies classification by cement mineralogy. 2Infihl proportions refer to the amount of matrix and cement as proportions of total mfihl. 3Abbreviations: K = K-feldspar, P phiogopite. 61 2.5.2.2 Cemented breccias Two spatially distinct cemented breccia domains are present: an upper (main) and a composite lower cemented zone (Figs. 2.8 and 2.9). These cemented breccia domains are composed of MC-BX, CM-BX and C-BX facies. The upper cemented breccia has a semi-ellipsoidal morphology 20-lOOm thick, 500m wide and 400m in length, tapering towards the tips. Multiple discontinuous cemented facies, 1 0-30m thick, characterize the lower cemented breccia. Cemented breccia facies of the upper zone strike 1 000 and dip 45-60°S, and taper along strike. Lower cemented breccias are 10-30m thick and have a similar geometry to the upper zone, but poor continuity. Intrusive wall rocks contain monolithic in-situ cemented breccia (coarse crackle-fracture) that lacks the matrix characteristic of other breccia facies. Monolithic in-situ cement-dominated breccias (C BX) are both spatially and temporally contiguous with cement-dominated domains within the matrix-bearing breccia (Figs. 2.8 and 2.9). Breccia margins are generally gradational over short distances but also locally abrupt, or marked by dike intrusion. Biotite-phyric monzodiorite dikelet facies are centered in MC-BX and CM-BX facies (Figs. 2.8 and 2.9). 2.5.2.3 Clastic rock components Clasts: Clasts in matrix-bearing breccias are typically pebble- to cobble-sized with a predominantly subrounded morphology (Figs. 2.1 1A, B, C and D). Coarser (cobble to boulder), monomict facies is proximal to the observed and inferred wall rock contact. Sorting and stratification occur over short intervals (5-10 cm) but is uncommon. Clasts in the matrix-bearing breccia are derived exclusively from surrounding porphyry wall rocks. Acicular feldspar-phyric syenite occurs throughout the matrix breccia body in minor quantities and locally displays irregular clast margins (Fig. 2.11E). Numerous clasts in the matrix-bearing breccia record older alteration-mineralisation events as evinced by alteration halos (Fig. 2.11A) and truncated K-feldspar and fine-grained biotite veins. Clast confined sericite alteration is also very locally observed. Some of the pre fragmentation alteration cannot be directly correlated to surrounding wall rocks, suggesting transportation of fragments from unobserved portions of the system. 62 FIG. 2.11 Photographs of diagnostic features in matrix-bearing breccias. A. M-BX, cobble clast size matrix-bearing breccia <10% cement, polylithic with hematite-K-feldspar altered clasts. B. Thin-section scan of matrix-bearing breccia, some clasts are highlighted with white-dashed lines. Red box indicates the area expanded in frame C. Photomicrograph of matrix composed of sand and granule fragments of wall rock (partially altered to magnetite-biotite and chlorite) and rare biotite-crystals. Note biotite-crystal is not within a larger fragment. D. M-BX, clast with magnetite alteration rind (white-dashed line) in matrix-rich breccia. E. M-BX, outsized clast of acicular feldspar-phyric syenite with an irregular margin. Unit 5 = acicular feldspar-phyric syenite. 63 Pre-fragmentation porphyry dike contacts cannot be traced into the matrix-bearing breccia body; further demonstrating transportation of clasts away from the site of fragmentation. No convincing examples of pre-fragmentation Cu mineralisation are observed. However, pyrite fragments in the matrix are locally present. Clasts in monolithic in-situ cement-dominated breccias are typically very coarse and jig-saw fit. Clast morphology and organization in the polylithic matrix-bearing breccias suggests rotation and transport. In contrast, superposition of cemented breccia facies (MC-BX, CM-BX and C-BX) shows little evidence of clast rotation or transport. Matrix: Matrix material is composed mostly of sand and granule sized wall rock fragments (rock-flour) with subordinate igneous biotite (Fig. 2.11 C) and feldspar crystals that may be derived from wall rock or possibly juvenile material. Fine grained biotite (green-dark green pleochroism) ± chlorite ± magnetite is interstitial to granule size lithic fragments (matrix) and gives the matrix a dark colour (Fig. 2.11A and C). This matrix feature is discussed in the alteration section (2.6.2). Cement: A phlogopite ± K-feldspar ± anhydrite ± magnetite ± diopside ± sulphide assemblage occurs as both hydrothermal cement and veins that cut the matrix- bearing breccia and host intrusions. Phlogopite and magnetite are the most common cement minerals associated with sulphides. Cemented breccia facies are distinguished and mapped by the abundance and type of hydrothermal cement present (Table 2.3). Hydrothermal cement within the matrix-bearing breccias and older bordering wall rocks varies from negligible to abundant. Cement textures vary throughout MC-BX and CM BX facies: irregular and straight walled interconnected fracture fill that cuts both matrix and clasts (Fig. 2.12A); open-space fill between clast and matrix (Fig. 2.12B, C and D), and irregular shaped vugs. MC-BX facies comprises <40% cement infill with matrix making up the greater portion of infill (Figs. 2.12B and C). CM-BX facies is distinguished by cement that comprises> 40% of infill (Figs. 2.12D, E and E). Biotite phyric monzodiorite dikelet facies are locally affected by cement, manifest as irregular shaped vugs (Figs. 2.1OB and C). Similarly, manifestation of cement in monolithic in situ-cemented breccia (C-BX) facies is both irregular (Figs. 2.13A and B) and more straight walled (Fig. 2.1 3C). Additionally, C-BX facies is locally coincident with biotite phyric monzodiorite dikelets (Figs. 2.13A and B). 64 FIG. 2.12 Photographs of cemented breccia facies in the Southwest Zone. A. MC-BX, matrix-bearing breccia <40 % cement infihl, overprinted by coarse K-feldspar cement. Matrix is pervasively chlorite- garnet altered. B. MC-BX, phiogopite-chalcopyrite cemented, white dashed line highlights a clast. C. MC BX, phlogopite-chalcopyrite±bornite±magnetite cement cuts matrix infihl. D. CM-BX, cement-dominated breccia, >40 cement. Note pervasive magnetite alteration of matrix. White dashed lines highlight domains of cement. E. CM-BX, phlogopite-bomite-chalcopyrite-magnetite-anhydrite-K-feldspar cement sub angular cobble sized clasts. F. CM-BX, K-spar rich cement and alteration with bomite. Unit # 2 = megacrystic orthoclase-phyric syenite. Unit # 5 = acicular feldspar-phyric syenite. 65 FIG. 2.13 Photographs of monolithic in-situ cement-dominated breccias (C-BX). A. Monolith in-situ cemented breccia (no matrix present) (GCO5-677, portions of core between 295.5 and 298.5 m). B. Line traces illustrating pertinent features of drill-core shown in frame A: cemented domains are highlighted with red dashed lines; biotite-phyric dikelets marked with thick black lines) and contain disseminated chalcopyrite and phiogopite alteration. Minerals within brackets indicate dominant wall rock alteration. Unit # 3 = megacrystic orthoclase-phyric monzonite. 66 25.3 Rock paragenesis Coherent and clastic rocks in the Southwest Zone are grouped into four stages defined by their timing with respect to the fragmentation events associated with the formation of matrix bearing-breccia and cemented breccia facies (Table 2.5). TABLE 2.4 Paragenetic stages of coherent and clastic rocks in the Southwest Zone Paragenetic Stage I Stage 2 Stage 3 Stage 4 stages1 Clastic rocks2 Matrix-bearing breccia Cemented breccias Coherent rocks3 Units I - 5: Unit 6: Units 7-8: Units 9-10: feldspar-phyric syenite; biotite-phyric pyroxene, xenolith megacrystic orthoclase- monzodiorite hornblende and bearing phyric syenite; biotite-phyric lamprophyre; megacrystic orthoclase diorite; mafic to and plagioclase-phyric orthoclase and Intermediate monzonite; megacrystic plagioclase- dikes orthoclase-phyric phyric monzonite monzonite; acicular-feldspar-phyric syenite 1 Paragenetic stages are described in the text. 2 Summary of clastic rock features is in Table 3. 3 Coherent rock characteristics summarized in Table 2. * Syn-breccia timing, paragenesis discussed in the body text. Feldspar-phyric syenite and megacrystic syenite and monzonite porphyry dikes are cut by matrix-bearing breccias. Acicular feldspar porphyry also occurs as clasts throughout the matrix-bearing breccia. However, some of these clast display irregular margins suggesting the unit was incorporated pre-solidification giving the unit syn matrix-bearing breccia timing. The coherent units cross-cut by the matrix bearing breccia and acicular feldspar porphyry are grouped into stage 1 (Table 2.4). The matrix-bearing breccias are superimposed by cemented breccias, locally resulting in the formation of MC-BX and CM-BX breccia facies (Figs. 2.8 and 2.9). Concurrent with cement emplacement in matrix-bearing breccias, the cemented breccia facies C-BX formed in adjoining coherent units. Cement emplacement is centered on and transects the older N- trending contact between matrix-bearing breccia and wall rock. Textural features and positions of biotite-phyric monzodiorite dikes and dikelets indicate emplacement was coeval with cemented-breccia formation. Cemented breccia facies (MC-BX, CM-BX and C-BX) and biotite-phyric monzodiorite are grouped together as stage 2 rocks. Cemented breccia facies are cross-cut by multiple dikes composed of pyroxene, homblende and 67 biotite-phyric diorite, and orthoclase and plagioclase-phyric monzonite and compose stage 3 (Fig. 2.14). FIG. 2.14 Post-cemented breccia facies coherent units. Xenolith-bearing lamprophyre and several dikes of mafic to intermediate composition are not part of the Galore Creek alkalic suite, as defined by Enns et al. (1995), and are grouped as stage 4 rocks. 2.6 Structural controls on rock distribution Biotite-phyric monzodiorite dikes occur at the contact between the matrix-bearing breccia and porphyry wall-rocks; dikelet facies occur in two discrete planar zones parallel to portions of the upper cemented zone (Fig. 2.9). The planar geometry, along- strike and down-dip linearity of the cemented zones implies structural control on their distribution. Megacrystic porphyry units and matrix-bearing breccias are cut by a post-mineral fault, the Southwest Zone Fault (SWZF). The fault strikes 12O13O0 and dips 6O° SW and its geometry is similar with reverse faults mapped elsewhere in the district scale (Fig. 2.3). A lamprophyre dike is locally coincident with this fault (Fig. 2.5). The timing of lamprophyre emplacement, however, is ambiguous. The extent and direction of WEST porphyry 49 wall rock wall roy trix-bearing matx- porphyry j bccia braa9.COfltacontçt, v ma meters 0 150 Post-cemented breccia dikes Lamprophyre Equigranular intermediate and mafic coherent units Orthoclase and plagioclase-phyric monzonite Pyroxene, hornblende and biotite-phyric diorite 68 displacement of coherent and elastic rocks shows a reverse separation, however, the true displacement is not well constrained. 2.7 Alteration The term assemblage herein specifies a group of hydrothermal minerals which appear to be stable together at the microscopic scale and to have formed contemporaneously (adapted from Seedorf and Einaudi, 2004). The concept of an assemblage is applied to alteration, vein and cement minerals. Veins and breccia cements, formed by precipitation of minerals from hydrothermal fluid (presumably into open spaces), are collectively termed infihl. Alteration stages are grouped according to timing with respect to the formation of matrix-bearing breccia and the cemented breccia facies and the coherent units affected (Fig. 2.15). Stagel Stage2 Satge4 FIG. 2.15 Paragenetic stages of coherent and clastic rocks, and alteration and mineralisation facies. Mineraology determined from detailed core logging, petrography, SWIR spectral data, electron micro- probe and SEM analyses. Colored lines in the lower panel are schematic representations of alteration and mineralisation facies distribution in space and time, highlighting inferred and observed paragenetic overlap and cross-cutting relationships. “Position” principally indicates lateral distances away from the locus of 0.3% Cu shells, for concise illustration of alteration mineral distribution see figures, 2.17, 2.18 and 2.20. Paracienetk stage = EE1 Clastk I rock unit matrix-bearing brecia,,?—— Alteration and pota5sk-A ________ mineral,sation __________ fades (tassic-B pas5i Time El cz cemented breccia fades (poassic-Diii fWaññg lcic-[patassicj) potassJ Siage3 (propyIitit) EJ veins 4? —E ptil.mt-chl Pasltcn Proximal K-Spar-hem K-Spar-bio —. (5-lOm) — — —. — Intermediate (5Dm) Ir-ani R gnt-chl-anh (>150m) — epi-chl-carb-(gnt) KEY Peripheral I K-spar-anh-ser-alb dop-nit-phla I slb-gn I-act = t py-cpy=c— (sphj=d a = major mineral, listed in relative abundance: b minor minerals listed in relative abundance:c suiphide minerals: d = laccessory minerals] 69 2.7.1 K-feldspar ± biotite ± hematite dusting (potassic-A and- B)-Stage-I Numerous clasts in the matrix-bearing breccia have alteration halos (Fig. 2.1OA) as well as truncated K-feldspar and finely-crystalline biotite veins. Correlative alteration is observed in the adjoining, pre-matrix-bearing breccia, stage 1 porphyry unit’s # 1-2 (Table. 2.2). Biotite (fine grained) veins and intense, texturally destructive, white- coloured K-feldspar alteration is the most abundant pre-matrix-bearing alteration stage (potassic A). Subordinate, intensely hematite-dusted K-feldspar alteration (potassic B) is also observed and overprints potassic-A. No Cu-Fe sulphide fragments or minerals confined to clasts are observed. 2.7.2 Phlogopite ± chlorite ± magnetite (potassic-C)-Stage-I Fine grained pervasive phlogopite (shreddy biotite) ± chlorite ± magnetite replaces mafic minerals and clastic fines, and occurs as microcavity fill in the matrix- bearing breccias (Fig. 2.1 1C). Potassic-C alteration, which varies in intensity, occurs throughout the matrix-bearing breccias and is predates Cu mineralisation. Clasts with biotite-magnetite alteration rinds are observed very locally within the matrix-bearing breccia particularly where the matrix is more intensely altered (Fig. 2.11 D). Based on the characteristics outlined, this alteration is interpreted to have occurred syn-to post-matrix- bearing breccia formation. 2.7.3 Phlogopite ± magnetite ± K-feldspar (potassic-D)-Stage 2 The potassic-D alteration assemblage comprises phlogopite ± magnetite ± K- feldspar ± apatite ± bomite ± chalcopyrite ± pyrite ± accessory rutile and is associated with the bulk of the Cu-Au and cementation. Cross-cutting relationships provide evidence that coherent units 1-6 (Table 2.2) are affected by potassic-D alteration and cement. Potassic-D alteration and infill is cross-cut by numerous narrow sinuous dikes (stage 3 and 4 coherent rocks, Fig. 2.15). The proportions of each mineral in the assemblage vary throughout the altered and cemented breccias. Coarse phiogopite ± magnetite alteration and infill is intimately associated with Cu-Fe-suiphides (Fig. 2.12B, C, D and E). Sporadic K-feldspar cement with weak to moderate wall rock alteration halos accompanies phlogopite-magnetite with Cu-Fe-sulphides. Intense pink coloured K-feldspar alteration and cement, with 70 subordinate phiogopite and magnetite, is recognized locally within the core (chalcopyrite-bornite zone) of the deposit (Fig. 2.12F). Recrystallization of igneous K- feldspar occurs during potassic alteration resulting in a milky white to pink colour and destruction of cleavage (Fig. 2.13C). Fine grained phiogopite replaces mafic minerals in vein alteration halos accompanying coarse infihl. The widest zone of intense potassic-D alteration occurs where cemented-breccias are localized at the contact between matrix- bearing breccia and porphyry wall rocks. Hydrothermal magnetite manifests as irregularly shaped replacement clots and as coarse infill (Fig. 2.12D), and is less abundant in the deeper southern and in the shallower eastern portions of the upper cemented breccia. Potassic-D alteration intensity and abundance of infihl minerals decreases outwards from the upper and lower cemented breccia zones. Alteration gradients are more pronounced in the direction perpendicular to the dip of cemented breccia zones. Intense potassic-D alteration grades outwards to weak potassic-D alteration and a mixed zone of potassic-D and propylitic assemblages. More locally, potassic-D alteration zones to, and overprints, a calcic-dominated alteration and infihl assemblage (Fig. 2.15). Biotite-phyric monzodiorite is less affected by infihl associated with potassic-D alteration. Petrographic examination of biotite-phyric monzodiorite facies reveals groundmass K-feldspar alteration and pervasive phiogopite replacement of igneous ferromagnesian minerals (mostly biotite). However, where cement is locally observed in the unit, it comprises crystals of K-spar ± phiogopite ± anhydrite ± rare quartz that extend from K-feldspar altered groundmass and project into quasi-spherical vugs 0.2 to 2cm in size (Fig. 2.1 OB and C). These vugs are commonly filled with various amounts of chalcopyrite and may be infilled miarolitic cavities. The small cement filled vugs lend a mottled texture to the coherent unit, compose up to ‘—5% vol., and contrasts with the typically “straight-walled” infill cement style of the surrounding cemented breccia facies. Primary wall-rock textures are weakly to moderately modified during potassic-D alteration event. The preservation of primary rock texture contrasts to younger, texturally destructive, potassic alteration. 71 2.7.4 Diopside ± magnetite ± phiogopite (calcic-(potassicJ) Stage-2 Calcic-Q,otassic) alteration assemblages and related veins are characterized by ferroan diopside ± magnetite ± ferro-actinolite ± garnet (grandite) ± albite ± chalcopyrite-pyrite ± accessory sphene (Fig. 2.16A). Infihl phiogopite locally accompanies diopside in veins (Fig. 2 1 6B), however, these mineral phases do not appear to be in textural equilibrium. Thus, the alteration stage is tentatively referred to as calcic with a potassic modifier. Cross-cutting relationships indicate that coherent units in stages 1 (Fig. 2.15) are affected by calcic-Q,otassic) alteration and cement. Calcic-(potassic) assemblages are present as veins and cement with subordinate pervasive wall rock alteration. Infill textures vary from intergrown fine-grained mineral aggregates of diopside-magnetite-garnet to coarse (>0.5cm) euhedral diopside ± magnetite growths. Calcic-Qotassic) stage veins and cement are characterised by diffuse and irregularly shaped margins. A patchy transitional zone (1-10 m wide) occurs where potassic-D and diopside-dominated assemblages overlap. Within this zone, phiogopite occurs with diopside where it occludes veins and cavities and is typically associated with pyrite ± chalcopyrite (Fig. 2.16B). However, rare examples phlogopite preceding diopside vein fill are locally noted (Fig. 2.1 6C). Albite occurs in minor to moderate amounts as milky white cryptocrystalline infihl along with diopside and magnetite and is typically altered to epidote (Fig. 2.1 6A). Albite is more common in the transitional zone between potassic-D-dominated and diopside-dominated alteration domains. 72 FIG. 2.16 Photographs of Stage 2 and phyllic alteration and infihl. A. Calcic-(potassic), diopside magnetite-albite-garnet infihl, overprinted by chlorite-garnet and epidote alteration. B. Calcic-(potassic) vein cutting feldspar-phyric syenite. Phlogopite and chalcopyrite occlude the diopside vein. C. Calcic (potassic) vein cutting megacrystic orthoclase-phyric syenite, diopside occurs in the vein centre succeeding phlogopite. D. Waning potassic alteration, K-feldspar-phlogopite vein with alteration halo cuts M-BX. E. Coarse waning potassic infill. F. Waning potassic, pervasive K-feldspar-anhydrite mafic mineral destructive alteration in megacrystic porphyry. G and H. Sericite-pyrite veins cutting M-BX. I. Sericite vein with sericite-anhydrite alteration halo overprints chalcopyrite and potassic D alteration in M-BX. 73 Garnet, a minor component of the Stage 2 calcic-Q,otassic) assemblage is honey-dark brown in colour, and occurs as 0.2-0.5 cm grains commonly lining vugs and veins with anhydrite. West of intense potassic-D stage alteration, actinolite locally substitutes for diopside in stage 2 calcic-Q,otassic) altered domains. Alternatively, actinolite could have formed by hydrolytic alteration of pyroxene. Diopside is overprinted by chlorite in most localities. Calcic-dominated veins occur in restricted zones (2-50m thick) along the fringes outboard of and below domains of intense potassic-D alteration (Fig. 2.17). Hydrothermal diopside is most abundant in the deep western and eastern flanks and in the shallower northern portions of the Southwest Zone (Fig. 2.17). In present day geometry, calcic-Q,otassic) alteration and infill occur in a footwall position to abundant phiogopite ± magnetite. However, minor amounts of diopside occur in isolated veins within domains of intense potassic alteration (Fig 2.17 section A-A’). Three short intervals of diopside infill occur in the hanging wall of the upper cemented breccia. The calcic-(potassic) assemblage distribution has a predominantly asymmetric geometry relative to intense potassic-D alteration. Potassic-calcic zonation is more pronounced in the lower cemented breccia domain. 74 FIG. 2.17 Distribution and abundances of potassic-D and calcic-Qotassic) and waning potassic hydrothennal minerals (Stage 1) and the > 0.3% Cu shell along section lines A-A’ and B-B’. * Inclusive of infihl and alteration material. The spatial relationship between diopside-dominated and potassic-D infihl and alteration assemblages is well constrained. Overlapping potassic-D and calcic-potassic) assemblages form narrow patchy transitional zones. Within these transitional zones phlogopite ± K-Feldspar predominantly occur in the vein or cavity centers. The inference being that, at that location diopside ± magnetite formed initially followed by phiogopite ± K-Feldspar. Two examples record the reverse of this infill sequence (e.g. 16C). The absence of more reliable paragenetic textures (off-setting and cross-cutting veins) between these two mineralogically contrasting assemblages precludes definitive WEST A 349650E I 3604006EAST 900m A’ >0.3% Cu — 750 m NORTH contact— B 6333900 N >0.3% Cu 750 m Diopside: — >2 % vol. Phlogopite*: — >12 % and — 3-12% vol. Magnetite: — >3 % vol. Domains of intense (>10 %), texturally destructive, XX’< K-feldpar-anhydrite alteration ± coarse K-feldspar infill o Location of biotite stable isotope sample t’ Location of sericite stable isotope sample meters 0 150 90m Matrix-bearing Porphyrybreccia ___—“ wall rock contact 450 m -r 6333300 N 75 evidence. However, the tentative mutually cross-cutting relationship together with the more consistent spatial relationship provides indicators that calcic alteration and infill paragenetically overlapped with, but is locally succeeded by potassic-D infill (Fig. 2.15). 2.7.5 K-feldspar-anhydrite (waning potassic)-Stage-2 Intense, texturally destructive, K-feldspar-anhydrite alteration locally overprints earlier Stage 2 assemblages and wall rock (Figs. 2.15 and 2.16D). Coarse K-feldspar (> 0.5cm) is commonly contiguous with the intense texturally destructive K-feldspar anhydrite wall rock alteration (Fig. 2.16E). K-feldspar occurs as veins, breccia cement, and within irregular shaped vugs. The assemblage is associated with minor amounts of phiogopite. Magnetite and mafic mineral destructive alteration halos of K-feldspar anhydrite locally forms 5 cm to 5 m wide domains resulting in the wall rock having a bleached appearance and distinctive pinky-peach to white colour (Figs. 2.16D and F). Broad zones of intensely altered wall rocks may be the result of overlapping alteration halos. The assemblage is typically Cu-poor and possibly grade destructive. Where it is spatially coincident with Cu-mineralisation, minor amounts of younger chalcopyrite occluded vugs and veins with anhydrite. Coarse K-feldspar-anhydrite veins and alteration is most abundant down-dip of the highest Cu-grade domain and proximal to the N-trending contact between the matrix- bearing breccia and wall rocks, in the deeper and southerly parts of the Southwest Zone (Fig. 2.17). The assemblage also occurs in structurally controlled horizons over short (0.5—2m) intervals. Cross-cutting alteration halos provide evidence that the Cu-poor potassic alteration and infihl is younger than potassic-D and calcic-Qotassic) event. K-feldspar anhydrite alteration and associated infill is likely to have occurred during the waning period of Cu-deposition. This alteration assemblage is locally overprinted by sericite particularly in the narrow fracture controlled zones. 2.1.6 Sericite ± anhydrite (phyllic)-Stage-3 Weak to moderate selective pervasive alteration and fracture controlled fine grained sericite ± anhydrite characterize phyllic alteration assemblages. Sericite alteration is prominent in some orthoclase and plagioclase-phyric dikes weakly altering 76 plagioclase phenocrysts. Light tan to dark green coloured sericite also occurs as fracture fill with 1-10 cm alteration halos (Fig. 2.16G and H) and locally overprints Stage 2 alteration and mineralization (Fig. 2.16H and I). The most abundant sericite occurs in vein and alteration halos in the eastern parts of the Southwest Zone, towards the limit of drill coverage. Late-stage phyllic alteration is commonly coincident with and overprints waning-stage potassic alteration and veins, indicating reuse of the same permeable zones. Anhydrite±chalcopyrite veinlets and subordinate veins inferred to be paragenetically equivalent to phyllic alteration and occur in minor amounts throughout the deposit. 2.7.7 Garnet ± chlorite (calcic)-Stage-3 Selectively pervasive garnet alteration of calcium-bearing mineral phases occurs as fine-grained (0.02-1 cm) buff-tan coloured aggregates (Figs. 2.7B, 2.18A and B). This garnet is distinguished from the calc-Q,otassic) phase by its colour and habit. Chlorite alteration accompanies garnet alteration of hornblende and the minerals appear to be in textural equilibrium. Garnet and chlorite±pyrite alteration occurs over wide intervals (>20 m) from approximately 10 to 50 m outboard of early Stage 2 alteration (Figs. 2.15 and 2.19). Locally garnet has overprinted well mineralized zones and high-temperature assemblages. Thus, in the Southwest Zone garnet can be spatially associated with Cu but does not appear to be intrinsic to ore formation. Calcic alteration occurs in potassic-D alteration domains where it is restricted to Stage 3 dikes (Figs. 2.7D and 2.19). Selective pervasive garnet also occurs in the peripheral zones (Fig. 2.19) where it is commonly overprinted by epidote and carbonate. Epidote overprints garnet in many locations forming a mantle or armor around garnet grains (Figs. 2.18 C and D). 77 FIG. 2.18 Photographs and photomicrographs of Stage 3 and 4 alteration facies and veins. A. Selective pervasive garnet-chlorite alteration overprinted by minor amounts of epidote in megacrystic orthoclase phyric monzonite. B. Garnet-chlorite-pyrite alteration of matrix in M-BX. Note K-feldspar alerted clasts are less susceptible to younger alteration. C. Plane D polarized photomicrographs of chlorite-garnet altered hornblende phenocrysts. Note fme-grained amorphous epidote mantles garnet grains. E. Carbonate potassic alteration and vein and alteration halo cut megacrystic orthoclase-phyric syenite. Magnetite is altered to hematite within the alteration halo. F Quartz-pyrite vein reopens fractures associated with carbonate-potassic veins. I 78 WEST A FIG. 2.19 Distribution and abundances of Stage 3 calcic and propylitic alteration minerals on section lines A-A’ and B-B’. 2.7.8 Chlorite ± epidote ± calcite ± pyrite (propylitic)-Stage-4 The minerals chlorite ± epidote ± albite ±calcite ± pyrite ± magnetite ± hematite composes a propylitic alteration (Figs. 2.7B, 2.1 1E and 2.16A). Moderate to strong, selective pervasive carbonate and epidote alteration locally overprint calcium-bearing igneous and hydrothermal minerals. Calcite also occurs as hairline fracture infill. Pale- brown carbonate (ferroan dolomite-ankerite) may reflect alteration of garnet. Propylitic EAST K 3504006 jJOO m Porphyry 1’ wall rock ‘ —: contact Matrix-bearing breccia NORTH B 6333900 N r’ Porphyry wall rock 900 contact 750 __J .‘ S.’ ‘ ‘ Epidote:—>1%vol. 1% Garnet:.- >1 % vol. ‘ Chlorite: >10 % and —3-10% vol. 600 meters 0 150 450 m -r 6333300 N 79 alteration occurs in post-Stage 2 coherent and clastic rocks (Fig. 2.15) and is most prominent in the northern shallower parts and eastern periphery of the Southwest Zone (Fig. 2.19). 2.7.9 K-feldspar ± Fe-carbonate (carbonate-potassic)-Stage-4 K-feldspar ± Fe-carbonate ± specular hematite ± pyrite ± minor chalcopyrite characterises a carbonate-potassic alteration stage. Strong, texturally destructive, K- feldspar halos surround Fe-carbonate ± specular hematite ± pyrite veins (Fig. 1 8E). Tightly spaced veins and veinlets occur as 0.5 to -40m wide fracture zones that cross-cut older Stage-2 and 3 alteration and Cu-Au mineralisation. Veins and alteration halos have low-angles to the vertical drill-core axis and occur proximal to steeply dipping fault zones. K-feldspar with this stage is pervasive and salmon-coloured. Within the K- feldspar alteration halos magnetite is strongly altered to hematite and pyrite replaces sulphide and mafic minerals (Fig. 1 8F). A moderately intense foliation is commonly spatially coincident with the assemblage; timing between fabric development and veining is inconsistent. Carbonate-potassic veins and alteration are associated with Cu-poor Au- rich (>0.5 glt) mineralisation and elevated pyrite contents. Fe-carbonate ± specular hematite veins and associated potassic alteration is the result of a younger steeply dipping event (Fig. 2.20). NORTH B 6333900 N SOUTH B meters 0 150 Elevation(meters) 450 m -r 6333300 N 80 FIG. 2.20 Distribution of late-stage carbonate-potassic veins and alteration, hydrothermal quartz and Au intercepts> 3 g/t along section B-B’. 2.7.10 Quartz (quartz veins)-Stage-4 Quartz ± minor pyrite ± chalcopyrite ± specularite veins with thin quartz alteration halos occur in some fault and shear zones. The veins occur as individual and fracture mesh veins 0.02 to 2 m wide. Quartz veins cross-cut and are locally concomitant with carbonate-potassic veins (Fig. 2.18F). Late fractures and post-mineral faults locally truncate and brecciate quartz. The erratic distribution of quartz has made it difficult to characterize; however, it consistently occurs with carbonate-potassic veins and alteration in fracture / fault zones. Quartz is commonly but not always coincident with anomalously high (>2 g/t up to 20 g/t) Au and weak pyrite mineralisation (Fig. 2.20). 81 2.8 Sulphide minerals 2.8.1 Potassic-D and calcic-(potassicJ-Stage-2 Sulphide mineralisation is focused on and zoned about discrete tabular domains of biotite-phyric monzodiorite and coincident cemented breccia facies. Chalcopyrite is the dominant Cu-Fe-suiphide associated with potassic-D stage alteration and infihl. It generally replaces mafic minerals, occurs with magnetite, and is predominantly fine grained (Figs. 2.7C, 2.12C and 2.21A and B). Coarser replacement and suiphide cement also occurs (Fig. 2.12E). Minor bornite (trace- 0.5 %), locally intergrown with chalcopyrite (Fig. 2.20C), occurs as cement and within mafic sites and microcavities in the core of the deposit. Copper-iron sulphides are - . - ç .7. K-Sparalt FIG. 2.21 Photomicrographs of suiphide and oxide mineralisation in the Southwest Zone. A. Plane and B. reflected light photomicrographs of replacement phlogopite and pyrite-magnetite intergrowths. Specularite locally overprints magnetite. Red rectangle in frame A indicates field of view in frame B. C. Reflected light photomicrograph of bomite-chalcopyrite replacement (left bleb) and possible microcavity fill (right bleb) and hydrothermal apatite grains. D. Reflected light photomicrograph of pyrite overprinting chalcopyrite in pervasively K-feldspar altered groundmass. 82 deposited late in the cement paragenesis and also commonly replace coarse hydrothermal magnetite. Specularite locally occurs with chalcopyrite and magnetite and is likely replacing magnetite (Fig. 2.21B). Pyrite manifests as late veinlets and also selectively replaces mafic minerals and chalcopyrite (Fig. 2.21 D). Extension to cemented breccia domains selective replacement style mineralization dominates the rock. In the upper and lower cemented breccia domains, suiphide minerals are zoned from a core of chalcopyrite-bomite, to chalcopyrite>pyrite, out to pyrite>chalcopyrite (Fig. 2.22). 450m ___________________ Porphyry if contact __— wall rock 6333300 N FIG. 2.22 Pattern of Stage 2 and 3 suiphide mineral distribution along section lines A-A’ and B-B’. This deposit zonation is also observed at a smaller scale (5m), e.g. the farthest east bornite occurrence on section B-B’ at 750m elevation (indicated with and X in Fig. 2.22). WEST A 750 m 600 m SOUTH 750 m meters 0 150 cpy+bn cpy> py py> cpy py only ________ Matrix-bearing breccia 600 m Biotite-phyric ‘monzodiorite fades 83 Lateral sulphide zonation gradients are more pronounced than the up-dip gradients in both the upper and lower ore zones. The most abundant chalcopyrite (‘‘>3%) is coincident with intense phiogopite ± magnetite dominated potassic-D alteration and cementation. Calcic-dominated alteration and veins, characterised by diopside, is not associated with abundant chalcopyrite. Diopside is spatially associated with Cu in the transitional zone between calcic-dominated and intense potassic alteration and cement. Chalcopyrite is ubiquitous in biotite-phyric monzodiorite with subordinate microscopic bornite. Within the chalcopyrite-bomite sulphide zone, cemented and altered breccias contain more bornite than biotite-phyric monzodiorite dikes and dikelets, implying the bulk of bornite mineralization preceded intrusion. Strong, pervasive texturally destructive K-feldspar ± anhydrite alteration (waning potassic) is prominent in areas with less abundant chalcopyrite, suggesting this stage may be grade destructive. Additional Cu mineralisation is associated with minor late chalcopyrite-anhydrite veinlets. 2.8.2 Late-stage K-feldspar ± Fe-carbonate-Stage-4 Late-stage Fe-carbonate ± specular hematite ± pyrite veins, with K-feldspar alteration halos, is associated with pyrite and minor amounts of chalcopyrite. Late stage pyrite replaces sulphides as well as primary and hydrothermal mafic minerals. In drill hole GCO5-665, -48O-21Om, bornite and chalcopyrite are overprinted chalcopyrite and pyrite, respectively. 2.9 Structural controls on alteration and mineralisation 2.9.1 Pre to syn Stage 2 and 3 The upper (main) Cu zone has along strike and own dip continuity, in contrast, the lower Cu zone is discontinuous composing multiple domains of zoned alteration and mineralisation (Fig. 2.17, section B-B’). Mineralised zones have a roughly tabular and stacked geometry (Figs. 2.17 and 2.22) and represent at least two corridors of hydrothermal activity. These tabular corridors of> 3% Cu strike 1OO and dip 45 to 60° south, and intersect the north-trending matrix-bearing breccia-wall rock contact. The chalcopyrite-bomite zone dips steeper where localized at this intersection (Fig. 2.22, section A-A’). Gradients in suiphide and alteration assemblages and Cu-Au tenor are 84 most pronounced in the direction perpendicular to the dip of the Cu zones. Zonation about a planar geometry, and along strike and dip continuity, implies a structural influence on the distribution of alteration and mineralisation. Furthermore, to support this interpretation, the somewhat ellipsoidal geometry of the upper Cu zone is analogous to elliptical fault geometries described by Peacock (2002) and Walsh et al. (2003). Following this theme, the similarities in geometry and zonation patterns of all the Cu zones can perhaps be related to a series of small nucleating faults that propagated into the matrix-bearing breccia and hydrothermal system, nucleating on the breccia - wall rock contact. The geometry of the Au-rich zones and coincident carbonate-potassic alteration is not clear from the evidence. However, angle to drill-core axis and correlation between closely spaced drill-holes suggest the ‘-l to 5 m Au-rich zones trend 1100 and dip 70°- 80° south (Fig. 2.20). 2.9.2 Post-Stage 2 and 3 Post-mineral reverse movement along the Southwest Zone Fault locally truncates Cu-Au mineralisation and alteration assemblages. Along the fault, calcic-(potassic) alteration facies is superposed with potassic-D and propylitic assemblages (Fig. 2.17). Similarly, suiphide zonation patterns and Cu-grade change dramatically across the fault (Figs. 2.17 and 2.20). There is between 200 and 250 m of reverse separation. 85 2.10 Oxygen and hydrogen isotope data Potassic-D and calcic-Qotassic) phlogopite from four locations in the Southwest Zone (Fig. 2.17) was analyzed for oxygen and hydrogen isotope composition. A single sample of vein sericite associated with phyllic stage alteration was also analyzed to deduce the sources of hydrothermal fluids. Sample methodology and analytical procedures are presented in Appendix C. Two of the four phiogopite samples were taken from diopside-bearing veins to test for deviation from potassic-dominated zone isotopic composition. The mineral ö’8Oand D composition of phiogopite in both alteration zones is similar: ö’80 is between 5.6 and 6.O% and öD ranges from -78.5 to -84.9% (Table. 2.5 and Fig. 2.23). TABLE 2.5 Stable isotope results for Stage 2 infill phiogopite and phyllic stage sericite. Mineral Drill Depth Alteration Measured (%o) Tm1 (°C) Calculated fluid hole # (m) assemblage VSMOW compositions2 8O’ D 6O’ (%o) 6D (%o) phi GCO4-502 420 K-Spar-phl- 5.9 -78.5 400—475— 7.9—8.3— (-28.6)—(-37.6)— anh-bn-cpy 550 8.4 (-44.5) phl GCO5-651 221 phl-diop-mt- 5.6 -81.6 400—475— 7.7—8.0— (-31.8)—(-40.7)— anh-py 550 8.1 (-47.4) phl GCO5-665 259 phl-diop-mt- 6.0 -84.9 400—475— 8.0—8.3— (35.OH-44.0)- py 550 8.5 (-50.7) phl GCO5-642 182 phl-mt-anh- 5.8 -81.5 400—475— 7.9—8.2— (-31.7)—(-40.6)— cpy-py 550 8.3 (-47.3) ser GCO5-696 214 ser-anh-py 12.5 -99 400—350 11.1— (-69.0)—(-61.2) 10.3 300—250 9.1—7.7 (-50.8)--(-37.5) ‘Modeled temperatures of formation of biotite are estimated form biotite-halogen geothermometery (Chapter 3, section 3). Modeled temperatures for fluids in equilibrium with sericite are based on a range of temperatures at which sericite most commonly forms at in the porphyry environment (Beane and Titely, 1881). 2Calculated 6018 fluid composition from biotite and sericite are determined using mineral-fluid fractionation factors of Zheng (1993) and O’Neil and Taylor (1969), respectively, while 6D for both minerals is determined using Suzuoki and Epstein (1976). Fractionation factors for D/H in Suzuoki and Epstein (1976) are less reliable under 350°C. 86 -50 -60 -70 >40 0 -100 * -110 3.5 5.5 7.5 9.5 11.5 13.5 8180 (%oVSMOW) • infill phiogopite (- diopsicie) - , , ‘j’ sericiteo infihl phiogopite (i- dopside) FIG. 2.23 Mineral 8018 and 6D values from the Southwest Zone. Sample location indicated in figure 17. Sericite associated with phyllic alteration has O18 and 6D composition of 12.5% and - 99%, respectively (Table 25 and Fig. 2.23). The isotopic composition of fluids in equilibrium with phiogopite and sericite, at a range of temperatures, is presented in Table 2.5. The single sericite sample data displays markedly different composition relative to calcic-potassic and potassic-D samples, as illustrated in figure 2.23. 87 2.11 Metals Zoning 2.11.1 Copper, Gold and Silver The Southwest Zone has a Cu-rich core (Cu>Au) that zones to a Au-enriched halo (Au>Cu). The widest ore grade Cu intercept is along the matrix-bearing breccia wall rock contact in cemented breccia domains. Transecting this contact, Cu is present as two south-dipping corridors with characteristic metal tenor. Gold is concentrated in the upper (main) and lower Cu zones (Fig. 2.24A). Intense potassic-D alteration hosts the bulk of >0.3% Cu mineralisation. Lower Cu-grades between 0.1 — 0.3 % are associated with less intense potassic, calcic and some propylitic alteration. The presence of diopside-bearing alteration and infill is marked by a sharp decrease in Cu-grade (Fig. 2.24A). The Cu-core of the upper (main) ore zone has Au/Cu ratios less than or equal unity. Ore in the lower Cu-ore zones has higher Au/Cu ratios and is relatively more Au enriched than the upper (main) ore zone. Both zones exhibit an up-dip trend towards Au/Cu >1 (Figs. 2.24B and C). Bomite predominantly occurs within the <1 to 1.5 Au/Cu ratio range and correlates with greater than 3 g/t Au within the >1 % Cu shell. Up-dip increase in Au/Cu ratio within the >1% Cu shell of the upper ore zone are attributed to an increase in Au content at relatively constant Cu grades. Shallow portions of the 0.3 % Cu shells with high Au/Cu ratios overlap with elevated pyrite content and! or propylitic alteration (Figs. 2.24B and C). In these domains propylitic alteration overprints potassic assemblages. Outwards and downwards perpendicularly away from the> 0.3 % Cu shell, the Au/Cu ratio increases to values > 1.5. In contrast, Au/Cu ratios are < 1.5 and relatively constant above the upper surface of the upper (main) Cu-ore zone in sections A-A’ and B-B’. These across-dip or lateral trends to higher Au/Cu ratios outboard of the > 1 % Cu shell are due to sharp decrease Cu content at relatively constant Au grade. In the lower mineralised zones, an opposite trend is locally observed. There, Au/Cu values decreases with diminishing Cu grade. A second domain of Au-enrichment without appreciable Cu is also present in the Southwest zone. Here, Au forms a halo peripheral to the Cu zones and is characterised by Au grades of 0.1 - 0.3 g!t and Au/Cu ratios> 1.5 (Fig. 2.24). 88 epidote + garnet s Aproximate zone Domain of chlorite + wall rock porphyry + pyre of I - 3 % pyriteporphyry wall rock /matrixbearinjtf matrix-bearing Doma First pervasivein of diopsidebreccia ,contact breccia veins and alteration chlonte / ,_ contact\ FIG. 2.24 A. Copper grade shells and hydrothermal diopside along sections A-Al and B-B’. B. Gold grade shells and distribution of moderately intense propylitic alteration along sections A-Al and B-B’. C. Gold- copper ratio shells and zones of approximate 1 - 3 % pyrite along sections A-Al and B-B’; Au/Cu ratio shells are clipped where Au or Cu values are below 0.1. The best developed Au-halo is in the eastern parts of section A-A’, and the shallow western parts of section B-B’. These broad domains of Cu-poor, Au mineralisation are A. Copper Section A-A WEST A C.GtioT 89 spatially coincident with: 1) chlorite alteration; and 2) moderately intense chlorite epidote-garnet alteration; and 3) pyrite (Fig. 2.24). Silver grades between 1 and 2 ppm and Pb-Zn mineralization correlate with portions of the Au-halo. Silver is concentrated with Au in the Cu-ore zones. Silver grade in excess of 5 ppm correlate well with >0.3% Cu and >0.3g/t Au domains. Some of the highest Ag grade is coincident with bornite in the core of the deposit. Silver content decreases laterally outwards from the> 0.3% Cu domains and intense potassic alteration. Weak to moderate Ag mineralization (1 - 2 ppm) persists into propylitic alteration with sporadic intervals of 10 m at 2 - 5 ppm. Carbonate-potassic assemblage with pyrite contains elevated in Au ( >0.7 up to 8 g/t) but is generally Cu-poor. This alteration is also coincident with high Ag contents. Carbonate-potassic alteration and mineralisation overprints potassic-D stage metal zonation and is focused in steeply dipping mineralised zones. 2.11.2 Lead and Zinc Within the 0.3% Cu shell of the upper (main) ore zone Pb concentrations rarely exceed 5 ppm. In comparison, Pb values range between 5 and 10 ppm in the lower Cu- ore mineralized zones (Fig. 2.25A). The widest intervals of >10 ppm Pb overlaps with >90 ppm Zn in the northern and eastern extent of the deposit (Figs. 2.25A and B). Short intervals of higher grade Pb (> 20 ppm) are associated with sporadic galena-anhydrite veins that cut older Cu-mineralisation. Late stage potassic-carbonate alteration and veins are also coincident with elevated Pb grades. Lead-zinc-silver mineralisation is developed peripheral to high grade Cu-Au and overlaps with portions of the Au-pyrite-halo. An increase in Zn concentartion correlates positively with an increase in Au/Cu ratio. The upper Cu ore zone is mostly characterized by Zn grades between 45 and 90 ppm (Fig. 2.25B). Zinc grades increase up-dip within the 0.3 % Cu shell to values in excess of 90 ppm, portions of this high tenor Zn are coincident with propylitic alteration overprinting potassic assemblages. Outboard of the Cu-ore zones, the broadest domains of> 90 ppm Zn mimic the distribution of moderately intense propylitic alteration (Fig. 2.25B). As with Au and Pb, the lower Cu-ore zones are also relatively enriched in Zn, in comparison to the upper Cu-ore zone. 90 2.11.3 Molybdenum Molybdenite occurs as late veins with narrow (<0.5cm) quartz envelopes and as fracture patina. These fractures and veins are commonly reopened and filled with anhydrite ± pyrite and trace chalcopyrite. Based on this timing of Mo mineralisation is inferred as late stage 2. The highest density of elevated Mo (>l0ppm) values occurs in hanging wall rocks of the Southwest Zone Fault (Fig. 2.25B). B’ C. Molybdenum porphy — f . . . waftroc7 matrix-beanng 4 breccia / 1CfltC 1 4 matrix-bearing I contact\/ breccia J Domain of chlorite - >0.3% Cu epidote t garnet Mo (ppm) - pyrite > 0.1 g/tAu Highest density 6 0 meters 150 -P . FIG. 2.25 A. Lead,> 0.3 % Cu and> 0.1 g/t Au grade shells and distribution of moderately intense propylitic alteration along sections A-Al and B-B’. B. Zinc grade shells, distribution of peripheral-stage moderately intense propylitic alteration, and> 0.3 % Cu and> 0.1 g/t Au grade shells and along sections A-Al and B-B’. C Distribution of the domains of highest density of anomalously high Mo along section B- B’. Zn (ppm) I • > 125 45- •90-125 (1-45 91 212Discussion and Genetic Interpretation 2.12.1 Clastic rocks 2.12.1.1 Matrix-bearing breccia Key features of matrix-bearing breccia are: 1) abundant fine grained phiogopite magnetite-chlorite alteration and microcavity fill in the matrix; 2) dominantly matrix-rich and supported; 3) monotonous and poorly sorted nature; 4) polylithic, rounded to sub rounded, clasts of exclusively intrusive units; 5) moderately large areal extent (inferred from drill-core); 6) hosted in porphyry wall rocks. Hydrothermal cement (Stage 2) in the matrix-bearing breccia is the result of a younger superimposed event, and is not directly linked to brecciation processes. Matrix in the matrix-bearing breccia is pervaded by phlogopite-magnetite-chlorite (potassic-C) as alteration and microcavity fill at most localities. The distribution, mineralology and textural features of this alteration and microcavity infill suggest magmatic-derived hydrothermal fluids dominated the environment synchronous with and post-fragmentation. Fragment rounding and mixing, matrix (rock-flour) generation, and differential vertical displacement of fragments is considered compatible with fluidization as a transport mechanism during the formation of subsurface breccias (McCallum, 1985; Sillitoe 1985). Whole and unbroken phenocrysts in rock flour matrix breccias are interrupted by Seedorff et al. (2005) as juvenile (tuffaceous) material. By analogy with Seedorff et al. (2005), biotite crystals in the matrix (Fig. 1 6C), which do not appear to be part of a clast, are tentatively interpreted as a juvenile component. Moreover, delicate fluidal clasts are not unequivocally present, however, acicular feldspar-phyric syenite very locally displays irregular clast margins. These irregular margins suggest the unit is incorporated into the breccia before solidification at some localities. Juvenile fragments and clasts with irregular margins suggest the presence of magma during fragmentation. Based on the previously described characteristics and features, matrix-bearing breccias in the Southwest Zone are interpreted to be the product of explosive fragmentation and primarily classified as hydrothermal (Sillitoe, 1985) or hydroclastic sub surface (Davies et al., 2008a) breccias. Explosive fragmentation is inferred to be the result of rapid expulsion of magmatic-hydrothermal fluids from cooling magma stocks (second boiling) coupled with decompression and liquid-vapor separation of already 92 exsolved aqueous phases (Bumham, 1985; Sillitoe, 1985; Fournier, 1999). Breccia facies generated by this fragmentation mechanism are termed magmatic-hydrothermal breccias (Sillitoe, 1985). Juvenile material in the breccia suggests the presence of magma during fragmentation. In addition to fragmentation caused by expulsion of fluids from crystallizing stocks, fragmentation caused by steam expansion due to magma-fluid interaction (Sheridan and Woheltz, 1981; Hedenquist and Henley, 1985) is inferred to have occurred and produce juvenile material in the matrix. Subsequently, matrix-bearing breccias were produced by a hybrid of fragmentation processes and classified as magmatic-hydrothermal breccias with a subordinate phreatomagmatic component. Top and bottom terminations of the matrix-bearing breccia are not exposed in present day geometry and current erosion level. Therefore, it remains an unknown as to whether the magmatic-hydrothermal-phreatomagmatic explosions lead to the disruption of rocks through to the paleosurface. 2.12.1.2 Cemented breccias Cemented breccia facies (MC-BX, CM-BX and C-BX), veins, and their associated alteration, account for much of the Cu-Au budget in the Southwest Zone. Similar open-space filling, hydrothermally cemented breccias are widespread in porphyry systems and can be spatially associated with increased abundance of Cu-Au (Seedorff et al., 2005). Generally, hydrothermally cemented breccias form single or multiple lensoid, ovoid, or circular pipe-like bodies with steep to vertical dip (Sillitoe, 1985; Seedorff et al., 2005). At the Southwest Zone, the most abundant cement occurs in two sub parallel horizons, the upper and lower cemented breccia domains. Potassic-D stage cement minerals indicate moderately high temperature fluids of a dominantly magmatic source (Ulrich et al., 2001; Seedorff et a!., 2005), and indicated by stable isotope results (discussed below). Superimposition of hydrothermally cemented breccia facies on matrix-bearing breccias and porphyry wall rock is not associated with significant clast rotation or transport; implying fragmentation was non-explosive. Cement textures indicate emplacement by infill of: 1) old and new fractures; 2) original open space between fragments in the matrix-bearing breccias; and by 3) open-space generated by possible chemical corrosion or winnowing of matrix fines. Based on cement 93 mineralology and the environment of formation cemented breccias are interpreted to be the results of non-explosive fragmentation caused the by migration of magmatic hydrothermal fluids. In this scenario, fragmentation of wall rock is the result of mechanical energy released during second boiling, decompression (Philips, 1972; Burnham, 1985) and subsequent hydraulic fracturing (Jebrak, 1997). Based on the inferred fragmentation mechanisms and criteria presented by Sillitoe (1985), Davies (2002) and Davies et al. (2008b), hydrothermally cemented breccias in the Southwest Zone are interpreted as hybrid magmatic-hydrothermal-hydraulic breccias. Root zones to the cemented breccias were not directly observed, however, some magmatic-hydrothermal breccias are known to root in porphyry intrusions (Zweng and Clarke, 1995; Jackson, 2008). Infihl minerals, alteration and mineralisation are centered on, and zoned about, cemented breccias and biotite-phyric monzodiorite facies. Distribution of infill and alteration facies, mineralisation, and metals are discussed below. 2.12.1.3 Evolution of the Southwest Zone breccia complex Overpressure at the top of a hydrous magma chamber, due to second boiling, caused a rupture of magmatic-hydrothermal fluids and initiated explosive fragmentation. Fragmentation continued with decompression and was accompanied by subordinate phreatomagmatic explosions caused by the interaction of acicular feldspar-phyric syenite and water. Explosions propagated into megacrystic porphyry wall rocks causing differential clast displacement, mixing and comminution which resulted in the formation matrix-bearing breccias. Ambient magmatic-hydrothermal fluids cemented microcavities ore spaces) and altered clastic fines in the matrix-bearing breccias, greatly reducing porosity and permeability in matrix-rich facies. Permeability regimes established post-matrix-bearing breccia focused subsequent magmatic-hydrothermal fluids and strongly influenced the distribution of cemented breccia facies. The geometry of the upper and lower cemented breccia domains suggest structures played an important role in their genesis. An array of faults or fracture zones is interpreted to have intersected the matrix-bearing breccia-wall rock contact. The faults may have been dilatational features related to palaeo-stress fields. Magmatic volatiles are inferred to have accumulated again in cupola underlying the matrix-bearing breccias to 94 be eventually expelled. The energy released during magmatic-hydrothermal fluid expulsion was sufficient to fracture roof rocks but non-explosive. Decompression and hydraulic fracturing formed fluid channels permitting access to the overlying rocks. Metal-bearing potassic-D fluids were preferentially channeled by the matrix-bearing breccia-wall rock contact and multiple intersecting fractures zones, with additional permeability generated by hydraulic fracturing. Fluid migration through newly formed fracture networks and the preexisting permeability architecture resulted in formation of the cemented breccia facies and Cu-Au zones. Cartoons illustrating the evolution of the breccia complex are presented below in the section 2.12.5 Q,alaeo-geometry and deposit model). 2.12.2 Isotopic composition of hydrothermal fluids The composition of fluid in equilibrium with potassic-D and calcic-(potassic) infill phlogopite is calculated for a range of temperatures based on biotite-fluid halogen partitioning geothermometery that utilizes thermodynamic models of Zhe and Sverjensky (1991, 1992). The modeled temperatures are consistent with the temperature range over which potassic alteration is generally accepted to occur (Seedorff et al., 2005 and references therein). Waters for potassic alteration associated with the formation of phlogopite, calculated at 475°C, have ö’80 and 8D of 8.0-8.3 and -37.6 to -40.7, respectively, and overlap with residual magmatic water (Fig. 2.26A). There is no isotopic distinction between phiogopite that formed in the presence of hydrothermal diopside versus phlogopite sampled from potassic-dominated alteration zones. The similarity in values is tentatively interpreted that fluids involved in the formation of both alteration facies are of magmatic origin. A caveat with this inference, however, is that the isotopic composition of diopside, actinolite or garnet was not directly measured; the isotopic composition of the fluids that formed phiogopite in the presence of diopside is inferred to be representative of calcic alteration and veins. At 475°C, the isotopic composition of Southwest Zone potassic fluid overlaps with the field of porphyry-Cu biotites (Sheppard et al., 1971), but is enriched in 8D compared to the potassic fluids in the British Columbia field (Fig. 2.26A). The later may represent contrasting magma source composition. If phiogopite formed at 400°C or less, calculated fluid isotopic 95 compositions shift to higher D values of ‘- > -32% (Fig. 2.26A). In this lower temperature scenario, the relatively high D values could be due to D-enriched magmatic-hydrothermal fluids produced by Rayleigh fractionation during degassing of the magma (Brigham and O’Neil, 1985). 0 -20 -40 -60 ‘0 -80 ____________ Th Bingham -100 -100 V -4 -2 0 2 4 6 8 10 12 4 -2 0 2 4 6 8 10 12 88O (%o, VSMOWI 6180 (%o, VSMOW) • infill phlogopitea ( diopside)b L.J porphyry Cu biotites, 650Cc * sericitect o infihl phiogopite (+ diopside) D potassic alteration (Z phyllic alteration FIG. 2.26 A. Calculated isotopic compositions of water in equilibrium with infihl phlogopite and, B. sericite from the Southwest Zone, with fluid compositions associated with potassic and phyllic alteration in other porphyry systems also shown (modified after Harris et al., 2002; Dilles et al., 1992 and references therein). The Southwest Zone potassic field is based on a range of calculated temperatures. Note the similarity of isotopic fluid values of infill biotite irrespective of the presence of diopside. Refer to text for discussion of marked points a, b and c along the black arrow in B. afluid composition of infill phlogopite is based on a range of temperature estimates from biotite halogen-chemistry geothermometry (Chapter 3, section 3.5). bdenotes the presence (+) or absence (-) of diopside in the vein sample. °“porphyry Cu biotites” is the field of calculated water in equilibrium with hydrothermal biotite at 65 0°C from Santa Rita, Bingham, and Ely porphyry deposits (Sheppard et al., 1971). dfluid composition are calculated for a range of possibilities as the temperature of sericite formation is unknown. FSE = Far Southeast Cu-Au deposit, Philippines; E26N = Endeavour 26 North porphyry Cu-Au deposit, eastern Australia. Potassic alteration field for British Columbia porphyry systems is modified after Harris et al. (2002). Ranges of residual magmatic water, i.e., that remaining in intrusion after degassing and crystallization, are from Taylor (1974). Meteoric water and SMOW values are from Rollinson (1993). Unfortunately, due to the absence of quartz veins, temperature estimates for sericite formation associated with phyllic alteration was not attempted. Therefore, modeled temperatures for fluids in equilibrium with sericite are calculated in 50°C increments over a range in which sericite commonly forms in porphyry systems; - 400- 0 .“ •SMOW Meteroric water -20’ •SMOW LA Meteroric water Southwest El Salvador, FSE ( Zone\,• 4OOC 47C residual magmatic V755OC Yerington- British Cloumbia” Southwest Zone -80 400t residual magmatic — water 96 250°C (Seedorifet al., 2005; Harris et al., 2002). Isotopic values for phyllic alteration are shown in figure 2.26B with phyllic alteration from other porphyry deposits (field of Southwest Zone potassic fluids is also provided as reference). Sericite alteration in the eastern parts of the Southwest Zone was caused by öO’8-enriched fluids in comparison to phyllic alteration in other porphyry deposits, irrespective of the modeled temperatures (Fig. 26B). Furthermore, between 320 and 250°C, calculated compositions overlap with the residual magmatic field and parts of calculated potassic waters field (Fig. 2.26B). The widely used genetic model of phyllic alteration invokes mixing between magmatic and meteoric fluids and is supported by isotopic evidence (Taylor, 1997). Other studies, however, have demonstrated phyllic alteration to have formed from moderately high temperature fluids of magmatic origin (Hendenquist et al, 1998; Harris et al., 2002). Also shown in the figure is a modeled phase separation evolution line for an exsolved magmatic fluid, indicated with black arrows with points labeled a, b and c (from Harris et al., 2002). If a magmatic fluid, of starting isotopic composition (a), reaches its solvus the fluid becomes immiscible and vapor-liquid separation occurs (Heinrich, 2005). Sericite formed through reaction with the vapor phase will be depleted in 018 and enriched D (Slumovich et al., 1999) and proceed from point (a) towards point (b), in figure 2.26B (Harris et al, 2002). Reactions with the separated brine have the opposite effect and produce an 018 enriched assemblages. Cooling of this brine will result in further depletion in D and 016 (Slumovich et al., 1999) resulting in sericite isotopic compositions that are shifted down the modeled line, similar to point (c). For moderately high temperatures (‘—S >270°C), phyllic alteration in the Southwest zone is depleted in D and 016, relative to potassic fluids, consistent with formation from a magmatic brine or a cooled magmatic brine. For phyllic stage sericite to be enriched in D and depleted in 018 relative to a potassic starting composition lower temperatures of formation are required (<250°C). Without accurate temperature estimates for phyllic alteration a single model (cooling-brine or vapor) cannot be resolved. Nevertheless, the values O18% and D%o, over the range of geological reasonable temperatures of formation, indicate that the fluids responsible for phyllic alteration were of magmatic origin or were equilibrated with magmatic rocks. 97 2.12.3 Alteration zoning (Stages 2 and 3) Caic-silicate alteration occurs in several alkalic porphyry Cu deposits, particularly silica undersaturated systems, in BC (Lang et al., 1995a and c). Lang Ct al., (1995c) and Arancibia and Clark and (1996) advocate that the composition and petrogenetic evolution of the associated magmas is the ultimate cause of fluids that form Na and Ca-rich alteration. In contrast, Dilles and Einaudi (1992) propose that Na-Ca alteration at the Yerington deposit formed, synchronously with potassic alteration, from hypersaline formation waters that were heated during flow into the cuplola regions. Changes in fluid chemistry because of mineral precipitation or wall rock reactions may also be able to generate Ca-rich alteration (Lang et al., 1995c). This study does not include an investigation of the isotopic composition of calc-silicate alteration minerals and so the potential for non-magmatic fluid involvement in their formation cannot be completely accessed. Nevertheless, models of formation of caic-silicate alteration and veins are proposed which may explain the style and distribution of this alteration facies in the Southwest Zone. Finally, some thoughts on what is required to test these hypotheses are presented. 2.12.3.1 Early Stage 2 alteration and infill: Potassic assemblages in the core of many porphyry Cu deposits are generally considered to have formed from moderately high-temperature (350-550°C), saline fluids of magmatic character (Ulrich et al., 2001; Seedorff et al., 2005). Potassic-D stage alteration and infill minerals are interpreted to have formed from magmatic fluids under similar conditions. The silicate portion of main stage potassic assemblage can be graphically shown on an aK/aH vs. aMg2/aHphase diagram Q,oint (1), Fig. 2.27A). 98 6 9 + z 0(U 0) 0 L (1) k-spar phi ,ø- - - musc / chi , 350°C I +qtz + I C”\(U +\ C.)(U Potassic-D alteration is characterised by abundant infihl phlogopite-magnetite-K feldspar. Replacement of wall rock ferromagnesian minerals (hornblende, pyroxene and biotite) and plagioclase by hydrothermal phlogopite and K-feldspar, respectively, also + (U 0) 0 4 9 4 5 log aMg>2 6 6 8 7 + C”\(U C.)(U 0) 0 6 5 log aMg,.2 T(C) FIG. 2.27 A-C. Activity-activity diagrams showing stabilities of silicate minerals at 350°C, 500 bars pressure, modified from Beane and Titley, (1981) and Beane, (1982). D. Stability relations of minerals as a function of log aCa2/aHand temperature of a coexisting aqueous phase. All systems conserve A1203 among solids and contain excess Si02 and H20. Positions (1), (2) and (3) in phase the phase diagrams are referred to in the text. Ideally these phase diagrams would have been constructed in the absence of quartz to best represent Southwest Zone alteration assemblages, however, at lower values of aH4SiO positions of mineral boundaries shift but the topology of the diagrams does not change significantly (compare figure 9 of Lang et al., 1995). aK/aW = 1 (K-feldspar stable) in panel B and 103.25 (muscovite stable) in panel C. Composition of biotite, pyroxene and garnet used in these calculations are similar to Southwest Zone hydrothermal minerals, which are: phlogopite = K(Mg065F3AlSiO1o(OH)2; pyroxene = Ca(MgogFe02)Si26;garnet = Ca3(Al03Feo7)2SiO12(Beane, 1982). Green line in frames B and C highlights calcic phases. The stability of chlorite is dashed because it is partly based on empirical observations (Beane and Titley, 1981). 99 occurs. Furthermore, magnetite alteration commonly occurs with potassic alteration. K metasomatism of plagioclase and hornblende, and magnetite alteration consumes H and reduces hydrothermal fluids and can be represented by the reaction (balanced, on Al): 5Ca02N8A1Si38(igneous plagioclase) + 1.5Ca2(Mg,Fe)4Si7Al02OH) (igneous hornblende) + 0.6Fe0(k) + 0.102 + 9K(aq) + 3H(aq) + 2.5H4S04(aq) = 7KA1Si3O8 (hydrothermal K-spar) + 2K(Mg,Fe)3AlSiOio(OH (hydrothermal phlogopite) + 4Ca2(a+ 4Na(aq) + 6H20 + 0.2Fe34 (1) The absence of quartz in the potassic alteration zone indicates silica does not become sufficiently saturated in the fluid to precipitate the mineral. Hydrothermal fluids sourced from silica-undersaturated magmas may be inherently silica poor (Lang et al., 1 995b). The silicic acid consumed by reaction 1 could buffer fluid silica activity further inhibiting quartz precipitation. K-metasomatism of igneous orthoclase is suggested by recrystallization and cleavage destruction commonly observed. Analysis of secondary biotite (section 3.3) suggests replacement of homblende may be accompanied by oxidation of Fe2 to Fe3, and thus further lower fluid 102. The spatial distribution and textural relationships indicate that potassic-D alteration paragenetically overlapped with, but locally succeeded, calc-silicate veins (diopside-magnetite±garnet). Einaudi (1982) demonstrates that the temperature range of pyroxene and garnet stability overlaps with K-silicate alteration; supporting a high-temperature tie between these two early stage 2 assemblages. If solutions initially compatibility with K-feldspar and biotite (point (1), Fig. 2.27A), in Al-rich environments, undergo significant increase in aCa2/aW, their composition is elevated to the region of Fe-bearing garnet stability. The chemical range of calc-silicate veins in the Southwest Zone is further constrained by the presence of diopside; point (2) on figure 2.27B. A sharp increase in aCa2/aHcan occur if fluids generating potassic alteration move into Ca-rich sedimentary rocks. However, diopside and garnet are not associated with Ca-rich protolith, and therefore Ca-saturation in the hydrothermal fluid occurred by a different means. As K in the fluid is consumed during potassic alteration (reaction 1), Ca and subordinate Na is liberated. It is proposed that, due to extensive K-metasomatism, hydrothermal fluids become progressively enriched in Ca causing localised Ca saturation to produce diopside veins: 100 Ca2(aq) + (Fe02, Mgo 8)2+ + 2H4SO(a = Ca(Feo.2,Mgo.s)Si206(hydrothermal diopside) + 2H0+4W (2) Increasing aSiO2 shifts reaction (2) to the right producing more diopside buffering total fluidH4SiO.Reaction 2 suggests formation of hydrothermal diopside is sensitive to the activity of aqueous silica. Cooling could perhaps be attributed to increase in aH4SiO. Moreover, reaction (2) produces H causing a localised decrease in pH that perhaps increases the solubility of Cu-Cl complexes in the fluid (see section 2.12.4). Hydrothermal albite is also interpreted to form in response to Na saturation in mechanisms similar to reaction 2. If Ca-rich magmatic or non-magmatic fluids, drawn in towards the magmatic-hydrothermal plume, are the cause of alteration perhaps more widespread veining and pervasive alteration would occur. The proposed model may explain why diopside veins and alteration form, sporadic and volumetrically minor domains, outboard of intense potassic-D alteration and the observed paragenetic relationship. 2.12.3.2 Late Stage 3 alteration Propylitic alteration is generated by the addition of volatiles (1420, H, and C02) with only minor cation metasomatism (Titley and Beane, 1981). Epidote-Na-feldspar calcite-chlorite-pyrite (propylitic) alteration assemblages have been interpreted to form at 250-350°C by Reed (1997). Mixing of inwardly convecting, thermally prograding, meteoric-derived fluids and magmatic fluids on the periphery of the hydrothermal system is commonly inferred to form a zone of low temperature propylitic alteration (Titley and Beane, 1981; Seedorff et al., 2005). Consequently, pervasive selective chlorite-epidote calcite alteration and disseminated pyrite in peripheral-stage propylitic alteration zone of the Southwest Zone is considered to have formed at similar temperatures and circumstances. The encroachment of propylitic alteration on early central higher temperature potassic-D and calcic-Qotassic) alteration may be affected by decreased + + 2+2+ aK /aH and aMg /a H cooling. Selective pervasive garnet, accompanied by chlorite alteration of ferromagnesian minerals, composes stage 3 calcic alteration. Point (3) on figure 2.27C is an approximate representation of the alteration minerals in this assemblage. Garnet-chlorite alteration is 101 K-feldspar stable suggesting aK/aH and aMg2/aH decreased sufficiently to form chlorite after hornblende but not muscovite after K-feldspar (aK/aH values < 4 at 350°C). As with calcic-Q,otassic) alteration, the role of non-magmatic fluids in the formation Ca-silicate minerals has not been accessed. For this reason garnet formation due to Ca-metasomatism as result of interaction with an external fluid Ca-rich cannot be affirmed or discredited. However, the location, timing and selective pervasive style of Stage 3 calcic garnet alteration is similar to propylitic alteration commonly found in calc alkalic porphyry systems in which little or no Ca-metasomatism need occur. In this scenario, much of the garnet in the Southwest zone could have formed by the reaction: Ca02N8Al12Si808(igneous plagioclase) + 1.4Ca2(Mgo.Feo6Si7Al02OH) (igneous homblende) + K(Mgo.6Feo.43A1SiO10(OH)2(biotite) +6.4H + 9.7H20 + 1.1502 =Ca3FeAl(Si04)(grandite) + (Mgo.Feo5Ali3Oio(OH)8(chlorite) + KA1Si3O8+ 0.4Mg2+ 0.8Na + 1 .6Fe3+4.H4SO4(aq) (3) Calcium liberated from chloritization of igneous hornblende or hydrothermal diopside can form hydrothermal garnet. Potassium released by chloritization of biotite is inferred to have formed accessory K-feldspar. The Na liberated could subsequently be consumed by epidote. Ferrous iron produced in reaction 3 could form alteration hematite or substitute in epidote. A simplified version of reaction (3) can be written: 2.5Ca(MgoFeo.)4Si7Al02OH)(igneous homblende) + Mg2 + 6H + 802 = Ca3FeA1(Si0 (grandite) + 2(Mgo.6Feo.45AlSi3Oio(OH)8 (chlorite) + 6H4SO4(aq) + 2Fe + 2Ca (4) Reaction (4) is consistent with alteration mineraology shown in figure 2.1 8C and D. In this reaction, Fe produced by can be consumed to form pyrite. It is clear from figure 2.27C that garnet formation is favored by high values of aCa2/aH. The formation of carbonate or anhydrite would buffer aCa2/aHto lower values hindering garnet formation. Similarly, anhydrite formation would reduce fluid aCa2/aH. Anhydrite stability is enhanced at temperatures in excess of 3 00°C (Titley and Beane, 1981). Furthermore, the bulk of anhydrite vein formation succeeds garnet alteration. Thus, widespread anhydrite precipitation did not occur until sufficiently high 102 temperatures prevailed and excess oxidized S was available. The dashed blue lines in figure 2.27D indicate an upper limit on aCa2/aH imposed by the saturation of the solution with respect to calcite at a(C02)= 1 (upper line) and a(C02)= 40 (lower line). Since calcite buffers aCa2/aH, the formation of garnet is favored by low CO2 contents in the hydrothermal fluid. Garnet is commonly overprinted by epidote (Fig. 2.18 C and D) and carbonate in the propylitic zones. Figure 2.27D (black arrow) shows a 2+2+hypothetical path in which cooling and small decrease in aCa /a H (due to increase in + 2+H or decrease in aCa ) leads to epidote and or carbonate overprint of garnet. Although less spatially extensive, garnet-chlorite alteration at the Southwest Zone can perhaps be correlated to inner propylitic alteration described at Ridgeway (Wilson et al., 2003). In summary, magmatic, or external non-magmatic, derived Ca-rich fluids is not necessarily required to form high temperature caic-silicate veins; extensive K metasomatism may lower fluid K/Ca ratio sufficiently to promote diopside precipitation. The isotopic composition of Sr(87Sr/6r), if present, in minerals can be used as a tracer to indicate the source fluid from which the mineral formed, thus Sr-isotope analysis of diopside and Sr-chemostratigraphy could help determine the minerals genesis. Broad zones of propylitic alteration that overprint high temperature assemblages in the Southwest Zone are largely interpreted to be in response to thermally retrograding magmatic-dominated hydrothermal fluids with an unimown degree of entrainment of thermally prograding non-magmatic (meteoric or formation) fluids. Selective pervasive garnet alteration can form without Ca-metasomatism inline with models for propylitic alteration. Alternatively, Ca for garnet formation was sourced externally and added to the rock. As with methods to assess the genesis of hydrothermal diopside, Sr-isotope analysis can be used to explore the role of external fluid involvement in peripheral-stage garnet formation. Similarly, 6018 and 6D isotope analysis of cogenetic chlorite could elucidate the presence of non-magmatic fluids. 2.12.4 Metal transport, deposition and zoning There is an intimate relationship between alteration, suiphide minerals and metal tenor in the Southwest zone. Specific features include: (1) gold-rich Cu mineralisation (2) extensive magnetite alteration; (3) systematic variations in Au/Cu ratio; and (4) halo of weak to moderate Au-mineralization coincident with propylitic alteration. Changes in fluid cation and anion concentration, temperature, pH and f(o2) are just some of the 103 intensive parameters that can control mineral stabilities and sulphide deposition. These features will be discussed and models for metal transport and deposition, and metal zoning proposed. 2.12.4.1 Gold-rich Cu mineralization The Southwest Zone has average Au contents >0.4 g/t Au and can be defined as, albeit arbitrarily, Au-rich porphyry Cu deposit (Sillitoe, 2000). A first order control on the metal budget of porphyry deposits is the process of metal partitioning from the magma to a volatile phase (Candela and Piccoli, 2005). Therefore, any crystallizing phase that depletes the parent melt in Au prior to magmatic-volatile phase saturation can potentially reduce the Au available for deposition (Simon et al., 2008). Gold concentration in magnetite is 10 to 100 times greater than silicates (Simon et al., 2003). In magmas where magnetite remains above the solidus until a late stage, loss of Au will be minimized and the potential to form an Au-rich deposit body via interaction with hydrothermal fluids increases (Simon et al, 2003). Gammons and Williams-Jones (1997) report that Au solubility as Cl complexes at Cu-porphyry temperatures (-500°C) is highest for a fluid that is oxidized, highly saline, and K-rich. Sillitoe, (2000) notes that Au-rich porphyry deposits contain alteration magnetite averaging 3-10 % vol. manifest as veins, irregular clots, disseminated grains and grain aggregates. In the Southwest Zone magnetite is associated with potassic alteration and breccia cement and is coincident with much of the Cu-Au. Late stage crystallization of magnetite in the melt may have facilitated Au-enrichment of the magmatic-volatile phase. Similarly, higher Fe contents in the exsolved magmatic-hydrothermal fluids may also have been promoted by late magmatic magnetite formation. In summary, late magmatic magnetite formation is a mechanism by which fluids responsible for potassic alteration and Cu-deposition can become Au-enriched and capable of extensive magnetite alteration. Furthermore, the presence of an appreciable volume of alteration and infill magnetite suggests oxidized fluids were reduced by interaction with wall rock Fe2 (Wilson et al., 2006). 2.12.4.2 Metal transport and deposition Transport of Cu and Au (and Fe) in magmatic brines is widely accepted to be in the form of chloride complexes (Hayashi and Ohmoto, 1991). Hydrogen sulphide (H2S) and SO2 are the predominant sulfur species that occur in high temperature (>400°C) 104 magmatic fluids typical of porphyry copper deposits (Rye, 1993). Cu and Au transport is favored by oxidizing conditions with sulfur transported as primarily SO2 (Barnes and Czamanske, 1967; Gammons and Williams Jones, 1997). Deposition of metals in hydrothermal systems occurs where changes such as cooling, decrease in Cl, pH increase due to rock alteration, phase separation, or fluid mixing cause the aqueous metal concentration to exceed saturation (Hemley and Hunt, 1992; Gammons and Williams Jones, 1997). Fluid cooling is a widely accepted model for ore formation in porphyry Cu deposits (e.g. Redmond et al., 2004). However, some workers (Cooke et al., 2004) contend that cooling alone is not capable of causing efficient sulphide deposition and that extreme temperature gradients are required to produce appreciable concentrations of metals. Such conditions could be achieved in the subsurface by fluid mixing. Mixing of magmatic fluids with lower-salinity meteoric waters, however, was not considered a dominant factor driving sulfide deposition at Bingham (Redmond et al., 2004). Consideration of sulfur speciation and S-isotope studies provides some insight in metal transport and deposition. Oxidizing conditions favors the stability S042 over H2S in magmatic-hydrothermal fluids. However the availability of reduced sulfur is of major importance to the problem of deposition or non-deposition of metals (Hemley and Hunt, 1992). Copper-iron-sulphide deposition therefore requires either: (1) a S042 reduction mechanism at the trap site (e.g., fluid mixing or water-rock interaction); (2) a supply of external H2S that the Cu-Au-bearing fluids interact with or; (3) a huge excess of sulfur flushing through the system with oxidized sulfur failing to precipitate and sulfides scavenging the smaller proportion of reduced sulfur from mineralizing fluids (Cooke et al., 2004). Vertical and lateral S-isotope zonation characterized by low ö34S values in the deposit cores, and a gradual increase in ö3’1S values outward to the system periphery where ö34S values merge with near-zero background values has been described at several alkalic porphyry deposits of British Columbia (Deyell and Tosdal, 2005) and New South Wales (Lickfold, 2001; Wilson et al., 2006). A pilot study in the Southwest Zone indicates a similar pattern of634S variation (Deyell and Tosdal, 2005). One of the most effective methods of producing large variations in 34S values is the reduction of S042 to H2S (Ohomoto, 1972). Thus, numerous workers (Lickfold, 2001; Deyell and Tosdal, 2005; Wilson et al., 2006) suggest that relatively oxidized Cu-bearing fluids become progressively reduced producing large variations in the of Cu-Fe sulphide minerals. Wilson et al. (2006) propose a particular physio-chemical process wherein 34S- 105 enriched aqueous S042 species are reduced through interaction with primary Fe2 bearing minerals in the host rock. Abundant Fe3 bearing alteration minerals (magnetite, hematite, and epidote) in potassic and propylitic zones are diagnostic of this inorganic sulfate reduction. The interaction of oxidized metal-bearing fluids with wall rock is invoked to reduce S042 to H2S at the trap site facilitating suiphide deposition in the Southwest Zone, similar to scenario 1 in the above list (Cooke et al., 2004). At temperatures below 400—350°C, aqueous sulfate can be generated according to the following hydrolysis reaction of SO2 in magmatic fluids (Holland and Mallinin, 1979): 4S02(aq) + 4H20(l) = H2S + 3S042(aq) + 6H(aq) (5) Sulphate and H2S is then available for interaction with the wall rock. Aqueous sulfate is reduced to sulfide at the same time that wall rock Fe2 (in hornblende, pyroxene, biotite and magnetite is oxidized to Fe3 and consume H by the reaction: l2FeO(rock) + S042(aq) + 2H(aq) = 4Fe3O()+ H2S(aq) (6) The abundant magnetite associated with potassic-D alteration is interpreted to be the result of fluid-rock reactions such as 6. Conceivably the Fe3 produced could also be transported to open spaces to form vein and cement magnetite. Similarly, vein and cement phiogopite can incorporate Fe3 (Chapter 3). Alteration of magnetite to hematite locally occurs in minor amounts during main stage potassic alteration. This reaction consumes 02 and causes fluids to become more reduced, as follows: 4Fe3O+02= 6Fe2O3 (7) Reaction 7 consumes 02 shifting reaction 8 to the right causing reduction of S042 to H2S by the reaction (Huston et al., 1993): S042 + 2H = H2S + 202 (8) 106 There are only minor amounts of hematite associated with potassic-D alteration and Cu. Reactions such as 6 are, therefore, interpreted to be the dominant mechanism to reduce S042 to H2S, and responsible for the spatial634Ssulfide zonation observed in the Southwest Zone. It is proposed that reactions 1 and 6 occur in tandem on initiation and during potassic-D alteration. Wall-rock reactions 1, 6, 7 and 8 have the net affect of increasing fluid pH, lowering the oxidation state, and generating reduced S for Cu-Fe-suiphide deposition. Metals zoning background information: Figure 2.28 is a J02-pH diagram constructed at 3 50°C showing Cu-Fe-S-O mineral stabilities, Au solubility as AuC12 and Cu solubility as CuCl° for salinities, S and Cl contents similar to the porphyry environment (redrafted and modified from Huston et al., 1993). 107 - 20 - 25 (‘1 e4 o .2 .2 - 30 0 2 4 -40 FIG. 2.28 Log f(02) — pH diagram at 350°C showing solubility contours for Cu and Au and the stability fields of minerals in the Cu-Fe-O-System (redrafted and modified from Huston et al., 1993). Inset shows an expanded view of the area indicated with the light grey box. P1, P2 and the green circles labeled a-d indicate hypothetical fluid chemical evolution paths which can result in contrasting Au/Cu values through time and space. Gold solubility as AuCl2 and Cu solubility as CuCl°. The 30 ppm Cu solubility contour in the inset is an approximate position. Although constructed at a fixed temperature, this figure demonstrates the effects of changing pH and 102 on fluid Cu and Au carrying capacity. Figure 2.28 illustrates fluid evolution paths that can cause Au to precipitate prior to and more rapidly than Cu (P 1), others that cause them to deposit at a similar rate (P2), and paths that can cause Au to precipitate more slowly than Cu (points (a)-(d), inset), in response to changes inJO2-pH at a fixed temperature. Hemley and Hunt (1992) demonstrate that cooling decreases the solubility Cu and Au in Cl complexes. A decrease in temperature will therefore enhance metal deposition caused by favorable changes in pH and 102 conditions. Furthermore, cooling may play a more significant role in mineralisation in late stages of fluid evolution and peripheral parts of the system. pH 108 Potassic-D suiphide assemblages indicate Cu deposition within the bornite chalcopyrite and chalcopyrite-pyrite stability fields constraining the pertinent part of figure 2.24 to the area within the grey box. Magnetite is commonly associated with Cu- Fe suiphides in the Southwest Zone. Hematite, however, is shown to be stable within the window of decreasing Cu-solubility in figure 2.28. So before proposing a metal zoning model for the Southwest Zone some comments on mineral stabilities are required to better relate observed mineral assemblages to the fields shown in figure 2.28. Beane and Titley (1981) review the stability relations for oxides and sulphides of Fe and Cu as a function of activities of 02 and S2 at 300°C and 550°C (Fig. 2.24 in Beane and Titley, 1981). They show that increasing temperature moves stability relations to higher activities of 02 and S2 but mineral compatibilities do not change significantly. Thus, the magnetite field expands, and is stable at higher activities of 02, with increasing temperature. Accordingly, an increase of temperature in figure 2.28 would stabilize magnetite to higher values of log 102 such that magnetite dominates over hematite and coexists with chalcopyrite ± bornite through the decreasing fluid Cu-solubility window. 2.12.4.3 Metals zoning model A hypothetical pathway of Cu-Au deposition in the Southwest Zone is proposed to account for the divergence in Cu and Au zonation patterns in the Southwest Zone (points (a)-(d), inset, Fig. 2.28). Position (a) represents the initial conditions of potassic D magmatic-hydrothermal fluid in which both Cu and Au are undersaturated. Potassic alteration (reaction 1) occurs in tandem with oxidation of wall rock Fe2 to magnetite (reaction 6) consuming 02 and H. Reactions 1 and 6 shift fluids to higher pH and less oxidized conditions lowering the solubility of Cu and Au (Fig. 2.28). Inversely, reactions decreasing pH may be able to increase Cu and Au solubility inhibiting deposition, i.e. formation of hydrothermal diopside (reaction 2). Copper and Au deposition commence together as point (b) is approached and bornite-chalcopyrite mineralization occurs. Au precipitation does not occur in excess of Cu (i.e. Au/Cu 1) between points a and b because of two possible reasons: (1) fluids initially evolved obliquely to Au solubility contours; and (2) Au is above 10 ppb and perhaps still undersaturated in the magmatic hydrothermal fluid. Nevertheless, coprecipitation with chalcopyrite is a viable means of Au deposition from undersaturated solutions (Gammons and Williams-Jones, 1997). 109 Continuing K-metasomatism is interpreted to decrease 102 and increase pH during fluid evolution. Between points (b) and (c), Cu and Au continue to coprecipitate and the suiphide assemblage shifts from bornite-chalcopyrite to chalcopyrite-pyrite stable. Along this path segment fluid evolution is less oblique to the Au contours (inset, Fig 2.28.) and hence more efficient Au deposition can occur towards point (c). A sharper decrease in the solubility of Au in comparison to Cu is interpreted to cause an increase in AuJCu values within the Cu zone during potassic-D stage alteration. From point (c) to (d) the capacity of the fluid to carry Cu is diminished and minor Cu is being deposited, however, Au is still relatively soluble (-40 ppb Au; Jones, 1992). Gold deposition, from a Cl- complex, between points (c) and (d) in figure 2.24, is enhanced by cooling to temperatures compatible with propylitic alteration (<320°C, Fig. 2.29). FIG. 2.29 Log f(02) — temperature diagram for solubility of Au as chloride and bisulfide complexes in the porphyry Cu environment (redrafted and modified from Huston and Large, 1989 and Jones, 1992). Activities of Cl, total S and alkali cations are similar to those modeled in figure 2.28. Note the solubility minimum of chloride complexed Au is at about 300°C. During this stage of fluid evolution cooling and fluid buffering caused by chlorite alteration of ferromagnesian minerals (reactions 3 and 4) results in weak to moderate (0.1-0.3 g/t) Au precipitation without significant Cu. The grey arrow on figure 2.29 illustrates the evolution from high temperature Cu-Fe-sulphide-bearing potassic assemblages (similar to points a-c, Fig. 2.28), to lower temperature pyrite-bearing 150 200 250 300 350 T(C) 110 propylitic alteration with coincident weak to moderate Au deposition. Pyrite-bearing propylitic alteration locally overprints potassic zone Cu-Au mineralization, introducing more Au, and modifying original Au/Cu to higher values. To summarize, the trend towards elevated Au/Cu values within potassic alteration zones is interpreted to be a depositional zoning feature which reflects a change in metal assemblage over an interval of space at a given time. These variations in Au/Cu values are evoked to be the result of a fluid evolution path that caused Au and Cu to be deposited at different rates largely in response to changes infO2 and pH. Copper-poor Au mineralization associated with pyrite-bearing propylitic alteration likely occurred by cooling and fluid neutralization outboard of potassic alteration. Downward collapse of isotherms, accompanied by waning of magmatic fluids, is invoked to cause retrograde flow of propylitic fluids and local overprinting of earlier temperature potassic alteration and mineralization. 2.12.5 Palaeo-geometry and deposit model 2.12.5.1 Palaeo-geometry Westerly directed compression some time in the Jurassic formed north-trending upright folds and thrust structures in the Galore Creek region (Logan and Koyanagi, 1994). North-trending fold axis traces are illustrated in figure 2.2B. Reviews by Beane and Titely (1982), Seedorffet al. (2005) and Cooke et al. (2007) indicate that the primary geometry of porphyry stocks and hypogene mineralisation is typically vertical to subvertical. Furthermore, hydrothermally cemented breccias generally have pipe-like bodies with steep to vertical dip (Sillitoe, 1985; Seedorff et al., 2005). Taken with the regional tectonic model (Logan and Koyanagi, 1994), the shallow dip of cemented breccia facies and Cu-Au, and the position of alteration footprints suggest the Southwest Zone magmatic-hydrothermal system is tilted 50-60° northeast. Figure 2.30 illustrates the palaeo-geometry of matrix-bearing breccias, cemented breccias facies, biotite-phyric monzodiorite dikes (A), and the distribution of potassic-D and calcic-potassic alteration and infill (B). 111 FIG. 2.30 Stylized perspective projections of joined sections A-A’ and B-B’ in approximated palaeo geometry (6O0 tilt). A. Outline of cemented breccia domains projected beyond current erosion surfaces. B. Domains of intense potassic-D alteration comprising phlogopite±magnetite±K-feldspar and calcic (potassic) diopside-dominated veins and alteration. Lines of sight for frames A and B are from different positions. Post-stage 4 faults and offsets are removed for clarity. 2.12.5.2 Deposit model The events, sequence and interpreted processes leading to the formation of the Southwest Zone deposit are presented with a series of time-integrated schematic figures illustrated in palaeo-geometry. Evolution of the Southwest Zone breccia complex and stage 1 to early stage alteration events are depicted in figure 2.31: 112 ‘L’A V’ L1 %1L’I • % — , ‘I FIG. 2.31 A-E. Time-integrated cartoons illustrating the evolution of the Southwest Zone breccia complex and paragenesis of Stage 1- early Stage 2 alteration events. Porphyry wall rock is removed in frames D and E for clarity. Numbers in squares and circles refers to coherent and clastic rock units (Tables 2 and 3; figure 2.15). Note, bifurcation and shallow-dipping margins occur in the Braden matrix-bearing breccia pipe at El Teniente, Chile (D. Cooke, pers. com). 1) Multiple phases of megacrystic porphyry intrude Stuhini group supracrustal country rocks. Pre-breccia stage 1 coherent units are affected by potassic A and B alteration (Fig. 2.31 A); 2) A hydrous melt is inferred to intrude the composite porphyry stock at depth, leading to accumulation of magmatic volatiles in the roof! carapace region (Fig. 2.31B); 3) Explosive fragmentation ensues when rapid expulsion of magmatic-hydrothermal fluids occurs. Hybrid magmatic-hydrothermal and phreatomagmatic explosions form the matrix-bearing breccia (Fig. 2.3 OC-D); 4) Post-fragmentation, the magmatic-hydrothermal fluids responsible for brecciation alter and fill microcavities in the matrix (potassic-C). An array of faults cut the breccia extending into porphyry wall rocks in response to reestablishment of the local stress-field (Fig. 2.31D); c;i 7Stuhinigroup LV’ £— - 7 A 7 A .,_IsI. 0 II’ . 8 — 113 5) After matrix-bearing breccia emplacement, a cupola is inferred to have formed due to magma replenishment at depth. Cement textures in biotite-phyric monzodiorite (Figs. 2.1OC and 2.D) are similar to miaroles described by Candela and Blevin (1995). Miarolitic cavities are good evidence for magmatic volatile phase exsolution (Candela, 1997). Suggesting biotite-phyric monzodiorite is closely linked to the magma from which fluids responsible for potassic cementation and Cu-Au mineralisation are sourced. 6) Magmatic-hydrothermal fluids, released from a crystallizing cupola (monzodiorite), ascend along the matrix-bearing breccia-wall rock contact and intersecting fault zones leading to the formation of cemented breccia facies (Fig. 2.3 1E). The ellipsoidal geometry of the cemented breccia and Cu-Au domains is inherited from these pre mineral faults. Figure 2.32 is a series of time-integrated frames illustrating stage 2 and 3 alteration and mineralisation paragenesis and zoning in interpreted palaeo-geometry. Explanation Domains of Wanin9 stage Propylitic dlopside infill \ potassic infill alteration and alteration and alteraion (chl-epi-carb- .inferred inferredlJ (K-Spar-anh) K-spar-py) LI t f Ph ir j \. Limitof >‘ °, Domatnsof s >9OppmZn >lOppmPb lt I rite V First chlorite e039U or,e ç S)f N) FIG. 2.32 Time-integrated alteration and mineralisation schematic in estimated palaeo-geometry. A. Incipient stage 2 alteration; potassic-D and calcic-(potassic) alteration and infill form contemporaneously. B. Alteration and Cu distribution at the end of stage 2 and start of stage 3. Minor phyllic alteration develops in fractures zones at upper levels in the system. C. Distribution of propylitic alteration facies, Au- halo, and Cu, Zn and Pb contours at the end of stage 3. The affects of the post-mineral Southwest Zone fault have been removed in all frames. 114 At the deposit-scale, the matrix-bearing breccia-wall rock contact and intersecting faults zones focused ascending fluids (Figs. 2.31E and 2.32A). However, at a smaller scale pre existing permeability and new fractures, created by hydraulic fracturing, act to locally de focus metal-bearing fluids. Partial fluid defocusing increases surface area, the extent of fluid wall rock interaction, and available reaction time promoting mineral deposition (Heinrich, et al., 1996). Furthermore, fracture networks allow drainage of wall-rock- equilibrated fluid out of the main channel ways. Alteration and mineralisation processes and paragenesis are summarized from the point of initial cemented-breccia fragmentation and magmatic-hydrothermal fluid migration (Figs. 2.32A and 2.31E): 2+.1) Magnetite alteration of wall rock Fe , in tandem with K-metasomatism (at 45 0°C), reduces oxidized S providing H2S for sulphide deposition. Continued input of fresh potassic fluids from below drive spent fluids (wall rock-fluid equilibrated), containing Ca liberated by potassic alteration, upwards and outwards through the cemented-breccia permeability network. Fluids become locally saturated with Ca, at sufficiently high aSiO2, to form diopside infill and alteration on the flanks of Cu mineralisation (Figs. 2.32A and B) and coincident potassic-D alteration (Fig. 2.30B) Biotite-phyric monzodiorite dike and dikelet facies are emplaced syn-Cu (Fig. 2.32A and B). TheJO2and pH hydrothermal fluid evolution path causes Cu-Au depositional zoning and a trend towards higher Au/Cu values upwards and outwards away from the focus of mineralisation. 2) Waning stage Cu-poor potassic alteration overprints early stage 2 alteration in the deeper parts of the hydrothermal system (section B-B’, Fig. 2.17B). This potassic alteration also occurs at shallower levels in the systems in fracture zones (Fig. 2.32B). Fracture controlled Mo mineralisation is spatially offset from the Cu-zones and inferred to occur early in stage 3. Phyllic alteration, best developed in the upper parts of the system, is likely to have formed from cooling magmatic brines related to waning potassic alteration (Fig. 2.32B); 3) Selective pervasive garnet-bearing propylitic alteration commences during thermal decline of the magmatic-hydrothermal system at temperatures between 250 and 350°C (Fig. 2.32C). Ingress of non-magmatic (thermally prograding) external fluids is likely to occur at this stage. A Cu-poor weak-Au (Pb) halo forms around the 0.3% Cu shell (Fig. 2.32C). This Au-halo overlaps with pyrite-bearing propylitic alteration and 115 encroaches upon higher temperature alteration facies. Gold deposition in the propylitic alteration zone is interpreted to occur between 300 and 350°C by cooling and fluid pH buffering via chlorite formation. Epidote and carbonate overprint garnet alteration due to continued thermal decline; 4) Post-hydrothermal system folding and thrust structures tilt the deposit 60° northeast resulting in shallowly dipping Cu zones. Tilting, displacement along the Southwest Zone fault and erosion remove upper and lateral parts of the hydrothermal system disrupting patterns of alteration, mineralisation and metal zonation. 2.13 Conclusions Detailed drill core logging, and analyses of samples, on two cross-sections, has characterized the coherent and clastic rocks, alteration, mineralisation and metals distribution and paragenesis in the Southwest Zone. On the basis of these features and a model of their spatial and temporal interplay, conclusions specific to the Southwest Zone can be deduced. Matrix-bearing breccias and megacrystic porphyry units host Cu and Au centered in potassic hydrothermally cemented breccias. The contact between matrix-bearing breccias and porphyry wall rocks served as the principal conduit and trap for ascending metal-bearing magmatic-hydrothermal fluids and biotite-phyric monzodiorite dikes. An array of faults intersecting this contact strongly influenced fluid-flow and the geometry cemented breccias. The same structures resulted in multiple domains of zoned alteration and mineralisation. Cu-poor, diopside-dominated infill and alteration formed contemporaneously with potassic infihl locally flanking high-tenor Cu and domains of intense K-metasomatism. Within the Cu zones, variations in Au/Cu can be attributed to changes in fluid 102 and pH and CuCl and AuC1 solubility. Weak potassic and calcic (potassic) alteration zones to and is locally overprinted by selective pervasive propylitic style garnet alteration. Garnet-bearing propylitic alteration extends to peripheral parts of the Southwest Zone silica-undersaturated hydrothermal system. A Au-(Pb)-halo encircles high temperature assemblages and the Cu-core of the deposit. This Au mineralisation is coincident with pyrite-bearing propylitic alteration and in places overprints potassic 116 alteration and Cu-Au. Post-hydrothermal system, regional folding and thrusting tilted the Southwest Zone deposit 5O6O0 northeast. In addition to conclusions specific to the Southwest Zone, several broader implications germane to the Galore Creek district and other alkalic porphyry systems are evident. Matrix-bearing breccias are commonly miss-interpreted as post-mineral, hydrothermal-system-terminating, events in porphyry environments. This study demonstrates that matrix-bearing breccias can host porphyry related hydrothermal alteration and affiliated Cu-Au and the importance of disparate rock packages to localize ore fluids. Hydrothermal garnet is described in central positions with Cu in the Central Zone (Micko et al., 2008) and in several locations with Cu at Mt. Polley (Pass et al., 2008) However, the nature and significance of peripheral selective pervasive garnet alteration has not been previously reported. Recognition of garnet-bearing propylitic alteration may provide a useful vector to alkalic porphyry centers. 117 2.14 References Arancibia, 0. N. and Clarke A.H., 1996, Early magnetite-amphibole-plagioclase alteration-mineralization in the Island copper porphyry copper-gold-molybdenum deposit, British Columbia: ECONOMIC GEOLOGY, v. 91, p. 402-438. Barr, D. A., 1966, The Galore Creek copper deposits: Canadian Institute of Mining and Metallurgy, Bulletin, 59, p. 84 1-853. Barnes, H.L., Czamanske, G.K., 1967, Solubilities and transport of ore minerals: In Geochemistry of Hydrothermal Ore Deposits, ed. by H.L. Barnes, Holt (Rinehart and Winston), New York, p. 334- 381. Beane, R. E., 1974, Biotite stability in the porphyry copper environment: ECONOMIC GEOLOGY, v. 69, p. 241-256. Beane, R.E., 1982, Hydrothermal alteration in silicate rocks, Southwestern North America in Titley, S. R., ed., Advances in Geology of the Porphyry Copper Deposits, Southwestern North America: Tucson, Univ. Ariz. Press, Chapter 6, p. 125. Bottomer, L.R., and Leary, G.M., 1995, Copper Canyon porphyry copper-gold deposit, Galore Creek area, northwestern British Columbia: In Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Northern Cordillera: Canadian Institute of Mining and Metallurgy Special Volume 46, p. 645-649. Bumham, C.W. and Ohmoto, H., 1980, Late stage processes of felsic Magmatism in Inshiara, S. and Takenouchi, S. ed., Granitic Magmatism and Related Mineralization, Special Issue No. 8, Society of Mining Geologist, Japan, p. 1-11. Bumham, C.W., 1985, Energy release in subvolcanic environments: Implications for breccia formation: ECONOMIC GEOLOGY, v. 80, p. 15 15—1522. Brigham, R. H., and O’Neil, J. R., 1985, Genesis and evolution of water in a two-mica pluton: A hydrogen isotope study: Chemical Geology, v. 49, p. 159-177 Candela, P.A., and Blevin, P.L., 1995, Do some miarolitic granites preserve evidence of magmatic volatile phase permeability: ECONOMIC GEOLOGY, v. 90, p. 2310-2316. Candela, P.A., 1997, A review of shallow, ore-related granites: textures, volatiles, and ore metals, Journal of Petrology, v. 38, p. 1619-1633. Candela, P.A., and Piccoli, P.M, 2005, Magmatic processes in the development of porphyry-type ore systems: 100th Anniversary Volume, p. 25-37 Cooke, D.R., Wilson, A.J. and Davies, A.G.S, 2004 Characterictics and genesis of porphyry copper-gold deposits: 24 ct Au workshop, ed. by D. R. Cooke, C. Deyell and 3. Pongratz, Hobart, Tasmania, Centre for Ore Deposit Research, University of Tasmania, 2004. Cooke, D. R., Wilson, A. 3., House, M. J., Wolfe, R. C., Walshe, 3. L., Lickfold, V. and Crawford, A. J., 2007, Alkalic porphyry Au - Cu and associated mineral deposits of the Ordovician to Early Silurian Macquarie Arc, New South Wales: Australian Journal of Earth Sciences, v. 54:2, p. 445- 463. Coney, P.J., 1989, Structural aspects of suspect terranes and accretionary tectonics in western North America: Journal of Structural Geology, v. 11, p 117-125. 118 Davies, A.G.S., 2002, Geology and genesis of the Kelian gold deposit, East Kalimantan, Indonesia (unpublished PhD thesis), University of Tasmania, 4O4p. Davies, A. G., Cooke, D. R., Gemmell B. J., and Simpson K. A., 2008a, Diatreme breccias at the Kelian Gold Mine, Kalimantan, Indonesia: precursors to epithermal gold mineralisation: ECONOMIC GEOLOGY, v. 103, p. 689-7 16. Davies, A.G.S., Cooke, D.R., Gemmell, J.B., Leeuwen, T.V., Cesare, P., and Hartshorne, G., 2008b, Hydrothermal breccias and veins at the Kelian Gold Mine, Kalimantan, Indonesia: Genesis of a large epithermal gold deposit: Economic Geology, v. 103, p. 717—757. Dilles, J.H., and Einaudi, M.T., 1992, Wall-rock alteration and hydrothermal flow paths about the Ann- Mason porphyry copper deposit, Nevada—a 6-km vertical reconstruction: ECONOMIC GEOLOGY, v. 87, P. 1963—2001. Deyell, C.L., and Tosdal, R.M., 2005, Alkalic Cu-Au deposits of British Columbia: sulfur isotope zonation as a guide to mineral exploration: British Columbia Ministry of Energy and Mines, paper no. 2005-1, p. 191-208. Drummond, S.E., Ohomoto, H., 1985, Chemical evolution and mineral deposition in boiling hydrothermal systems: Economic Geology, v. 80, p. 126-147. Einaudi, M.T., 1982, General features and origin of skarns associated with porphyry copper plutons, southwestern North America: In Titley, S.R., ed., Advances in Geology of the Porphyry Copper Deposits, Southwestern U.S., University of Arizona Press, p.185-209. Enns, S. G., Thompson, J. F. H., Stanley, C. R. and Yarrow, E. W., 1995, The Galore Creek porphyry copper-gold deposits, Northwestern British Columbia: In Schroeter, T, ed., Porphyry Copper (± Au) Deposits of the Northern Cordillera: Canadian Institute of Mining and Metallurgy Special Volume 46, p. 630-644. Foumier R.O., 1999, Hydrothermal processes related to movement of fluid from plastic into brittle rock in the magmatic-epithermal environment: ECONOMIC GEOLOGY, v. 94, p. 1193-1211. Gabrielse, H., Monger, J. W. H., Wheeler, 3. 0., and Yorath, C. J., 1991, Morphogeological belts, tectonic assemblages, and terranes: In Gabrielse, H., and Yorath, C. 3., eds., Chapter 2 of Geology of the Cordilleran Orogen in Canada: Geological Survey Of Canada, Geology of Canada, no.4, p. 15-28. Gammons, C.H., and Williams-Jones, A.E., 1997, Chemical mobility of gold in the porphyry-epithermal environment: ECONOMIC GEOLOGY, v. 92, p. 45-59. Hayashi, K., and Ohmoto, H., 1991, Solubility of gold in NaC1- andH2S-bearing aqueous solutions at 250 - 350°C. Geochimica et Cosmochimica Acta, 55, 2111-2126. Harris, A.C., and Golding, S.D., 2002, New evidence of magmatic-fluid-related phyllic alteration: Implications for the genesis of porphyry Cu deposits: Geological Society of America, v. 30, p. 335-338. Holand, H. D. and Malinin, S. D., 1979, Solubility and occurrence of non-ore minerals. In Geochemistry of Hydrothermal Ore Deposits, ed. by H.L. Barnes, Holt (Rinehart and Winston), New York, p. 461- 508. Hedenquist, J.W., and Henley, R.W., 1985, Hydrothermal eruptions in the Waiotapu geothermal system, New Zealand: Their origin, associated breccias, and relation to precious metal mineralization: ECONOMIC GEOLOGY, v. 80, p. 1640—1668. Heinrich, C.A., Walshe, J.L., and Harold, B.P., 1996, Chemical mass transfer modeling of ore-forming hydrothermal systems: current practice and problems: Ore Geology Reviews, v.10, p. 319-338. 119 Hemley, J.J., and Hunt, J.P., 1992, Hydrothermal ore-forming processes in light of studies in rock buffered systems:II. Some general geologic applications: ECONOMIC GEOLOGY, v. 87, p. 23-43. Huston, D.L., and Large, R.R., 1989, A chemical model for distribution of gold in volcanogenic massive sulphide deposits: Ore Geology Reviews, v. 4, p. 17 1-200. Huston, D. L., Bolger, C., and Cozens, G, 1993, A comparison of mineral deposits at the Gecko and White Devil deposits; implications for ore genesis in the Tennant Creek District, Northern Territory, Australia: ECONOMIC GEOLOGY, v. 88, p. 1198-1225. Jacobs, D. C. and Parry, W. T., 1979, Geochemistry of biotite in the Santa Rita porphyry copper deposit, New Mexico. ECONOMIC GEOLOGY, V. 74, 860-887. Jackson, M.L., 2008, Evolution of the Northeast zone breccia body, Mount Polley mine, British Columbia: Unpublished MSc thesis, The University of British Columbia. Jago, P., 2008, Metal- and alteration-zoning, and hydrothermal flow paths at the moderately-tilted, silica- saturated Mt. Milligan Cu-Au alkalic porphyry deposit: Unpublished MSc thesis, The University of British Columbia. Jébrak, M., 1997, Hydrothermal breccias in vein-type ore deposits: A review of mechanisms, morphology, and size distribution: Ore Geology Reviews, v. 12, p. 111—134. Jones, B.K., 1992, Application of metals zoning to fold exploration in porphyry copper systems: Journal of Geochemical Exploration, v. 43, p. 127-155. Lang, J.R., Lueck, B., Mortensen, J.K., Russell, J.K., Stanley, C.R. and Thompson, J.F.H., 1995a, Triassic Jurassic silica-undersaturated and silica-saturated alkalic intrusions in the Cordillera of British Columbia: Implications for arc magmatism: Geology, v. 23, p. 45 1-454. Lang, J.R., Stanley, C.R. and Thompson, J.F.H., 1995b, Porphyry copper deposits related to alkalic igneous rocks in the Triassic-Jurassic arc terranes of British Columbia: Arizona Geological Society, Digest 20, p. 2 19-236. Lang, J.R., Thompson, J.F.H, Stanley, C.R., l995c, Na-K-Ca magmatic hydrothermal alteration associated with alkalic porphyry Cu-Au deposits, British Columbia: In Thompson, J.F.H., ed., Magmas, fluids and ore deposits, Mineralogical Association of Canada Short Course, v. 23, p. 339-366. Lickfold V., Cooke D. R., Smith S. G. and Ulirich T. D., 2003, Endeavour Cu—Au porphyry deposits, Northparkes, N.S.W.: intrusive history and fluid evolution: ECONOMIC GEOLOGY, v. 98, p. 1607- 1636. Logan, J.M. and Koyanagi, V.M., 1994, Geology and mineral deposits of the Galore Creek area, northwestern British Columbia (104G/3 and 4): BC Ministry of Energy, Mines and Petroleum Re sources, Bulletin 92. Logan, J. M., 2004, Alkaline Magmatism and Porphyry Cu-Au Deposits at Galore Creek, Northwestern British Columbia: BC Ministry of Energy, Mines and Petroleum Resources, Geological Fieldwork 2004, Paper 2004-1. Seedorif, E., Dilles, J.H., Proffett, J.M., Einaudi, M. T., Zurcher, L., Stavast, W.J.A., Johnson, D.A., and Barton, M.D., 2005, Porphyry Deposits: Characteristics and Origin of Hypogene Features: ECONOMIC GEOLOGY, 100th Anniversary Volume, p. 251-298. Sheridan, M.F., and Wohletz, K.H., 1981, Hydrovolcanic explosions: The systematics of water-pyroclast equilibration: Science, v. 212, p. 1387—1389. 120 Simon, A.C., Pettke, T., Candela, P.A., Piccoli, P.M. and Heinrich, C.A., 2003, Gold solubility in magnetite: preliminary implications for magmatic gold budgets: American Mineralogist, v. 88, p. 1644-165. Simon, A.C., Candela, P.A., Piccoli, P.M. and Englander, L., 2008, The effect of crystal — melt partitioning on the budgets of copper, gold and silver: American Mineralogist, v. 93, p. 1437-1448 Sillitoe, R.H., 1985, Ore-related breccias in volcanoplutonic arcs: ECONOMIC GEOLOGY, v. 80, p. 1467- 15 14. McCallum, M.E., 1985, Experimental evidence for fluidization processes in breccia pipe formation: Economic Geology, v. 80, p. 1523-1543. McPhie, J., Doyle, M., Allen, R., 1993, Volcanic Textures: A guide to interpretation of textures in volcanic rocks: CODES Key Centre, University of Tasmania Hobart. McMillan, W.J., 1991, Tectonic Evolution and Setting of Mineral Deposits in the Canadian Cordillera: BC Ministry of Energy, Mines and Petroleum Resources, Ore Deposits, Tectonics and Metallogeny in the Canadian Cordillera, Paper 199 1-4. Micko, 3., Tosdal, R. M., Chamberlain, C. M., Simpson, K., and Schwab, D., 2007, Distribution of alteration and sulfide mineralization in the Central Zone of Galore Creek, northwestem British Columiba: Arizona Geological Society meeting, Ores & Orogenesis, Program with Abstracts, p. 175. Micko, J., Tosdal, R., Simpson, K., Chamberlain, C., and Schwab, D., 2008, Controls and hydrothermal zonation of Central Zone mineralization at the Galore Creek alkalic Cu-Au porphyry deposit, northwestem British Columbia: In: MDRU-CODES: Shallow and deep alkalic deposits: Porphyry module, Year 3 Final meeting. Mihalynuk, M.G., Nelson, J.L. and Diakow, L.J., 1994, Cache Creek terrane: oroclinal paradox within the Canadian Cordillera: Tectonics, v. 13, p. 575-595. Mortensen, J. K., Ghosh, D. K. and Fern, F., 1995, U-Pb geochronology of intrusive rocks associated with copper-gold porphyry deposits in the Canadian Cordillera: In Schroeter, T, ed., Porphyry Copper (± Au) Deposits of the Northem Cordillera: Canadian Institute of Mining and Metallurgy Special Volume 46, p. 142-158. Munoz, J.L., 1984, F—OH and Cl—OH exchange in micas with applications to hydrothermal ore deposits: In Bailey, S.W. Ed., Micas. Rev. Mineral. v. 13 p. 469—494. Nelson, J., and Mthalynuk, M., 2004, Mega-terranes and deep structures: tectonics and the potential for major new mineral deposits in British Columbia: Mineral Exploration Roundup, Vancouver, Abstract with programs, p. 26. Ohmoto, H., 1972, Systematics of sulfur and carbon isotopes in hydrothermal ore deposits: ECONOMIC GEOLOGY, v. 67, p. 55 1-578. O’Neil, J.R., and Taylor, H.P. Jr. 1969, Oxygen isotope equilibrium between muscovite and water: Journal of Geophysics Research, v. 74, p. 60 12-6022. Panteleyev, A., 1976, Galore Creek Map-Area: B.C. Ministry of Mines and Petroleum Resources, Geological Field Work, 1975, Paper 1976-1, p. 79-81. Pass, H., 2007, Geology, mineralization, alteration and geochemistry of long section 17-34, NE Zone, Mt. Polley Cu-Au porphyry Deposit: MDRU-CODES: Shallow and deep-level ailcalic deposits: Porphyry module: year 2 meeting. 121 Philips, R., 1972, Hydraulic fracturing and mineralization: Journal of Geological Society if London, v. 128, p. 337-359. Rollinson, H.R., 1993, Using carbon isotopes: In Using Geochemical Data: Evaluation, Presentation, Interpretation: Longman Press. Reed, M.H., 1997, Hydrothermal alteraion and its releationship to ore flid composition: In Barnes, H.L., ed., Geochemistry of hydrothermal ore deposits, 3rd ed.: New York, John Wiley and Sons, p. 303- 365. Redmond, P.B., Einaudi, M.T., man, E.E., Landtwing, M.R., Heinrich, C.A, 2004, Copper deposition by fluid cooling in intrusion-centered systems: New insights from the Bingham porphyry ore deposit, Utah: Geology, v. 32, p. 217—220. Rye, R.O., 1993, The evolution of magmatic fluids in the epithermal environment: the stable isotope perspective: ECONOMIC GEOLOGY, v. 88, p. 733-553. Schwab, D. L, Petsel, S., Otto, B. R., Morris, S. K., Workman, E. and Tosdal, R. M., 2007, Overview of the Late Triassic Galore Creek copper-gold-silver porphyry system: Arizona Geological Society meeting, Ores & Orogenesis, v. 22, p. 1-14. Selby, D., Nesbitt, B.E., 2000, Chemical composition of biotite from Casino porphyry Cu—Au—Mo mineralization, Yukon, Canada: evaluation of magmatic and hydrothermal fluid chemistry: Chemical Geology, v. 171, p. 77—93. Seedorff, E.and Einaudi, M. T., 2004, Henderson Porphyry Molybdenum System, Colorado: I. Sequence and Abundance of Hydrothermal Mineral Assemblages, Flow Paths of Evoloving Fluids and Evoltionary Style: Economic Geology 99, p. 3-37. Seedorff, E., Dilles, J. H., Proffett, J. M., Einaudi, M. T., Zurcher, L., Stavast, W. J. A., Johnson, D. A., and Barton, M. D., 2005, Porphyry Deposits: Characteristics and Origin of Hypogene Features: Economic Geology, 100th Anniversary Volume, p. 251-298. Sharp, Z. D., 1990, Laser-based microanalytical method for the in situ determination of oxygen isotope ratios of silicates and oxides: Geochimica et Cosmochimica Acta, v. 54, p. 1353-1357. Sheppard, S.M.F., Nielsen, R.L., and Taylor, H.P., Jr., 1971, Hydrogen and oxygen isotopic ratios in minerals from porphyry copper deposits: ECONOMIC GEOLOGY, v. 66, p. 515-542. Shmulovich, K.I., Landwehr, D., Simon, K., and Heinrich, W., 1999, Stable isotope fractionation between liquid and vapor in water-salt systems up to 600°C: Chemical Geology, v. 157, p. 343—354. Sillitoe, R. H. 2000. Gold-rich porphyry deposits: descriptive and genetic models and their role in exploration and discovery. Reviews in ECONOMIC GEOLOGY, v.13, p. 315—345. Suzuoki, T. and Epstein, 5. (1976). Hydrogen isotope fractionation between OH-bearing minerals and water: Geochimica et Cosmochimica Acta, v. 40, p. 1229-1240. Titley, S.R. and Beane, R.E., 1981, Porphyry copper deposits: ECONOMIC GEOLOGY, 75th Anniversary Volume, p. 214-269. Twelker E., 2007, A breccia-centered ore and alteration model for the Copper Canyon alkalic Cu-Au deposit, BC: Unpublished MSc. Thesis, The University of Alaska Fairbanks. Taylor, H.P., 1974, The application of oxygen and hydrothermal isotope studies to problems of hydrothermal alteration and ore deposition: ECONOMIC GEOLOGY, v. 69, p. 843—883. 122 Ulrich, T., Gunther, D., and Heinrich, C.A., 2001, The evolution of a porphyry Cu-Au deposit, based on LA-ICP-MS analysis of fluid inclusions of fluid inclusions: Bajo de Ia Alumbrera, Argentina: ECONOMIC GEOLOGY, v. 96, p. 1743-1774. Wemicke, B., and Kiepacki, D.W. 1988, Escape hypothesis for the Stilcine block: Geology, v.16, p. 461— 464. Wilson, A. J., Cooke, D. R., and Harper, B., L., 2003, The Ridgeway gold—copper deposit: a high-grade alkalic porphyry deposit in the Lachian Fold Belt, New South Wales, Australia: ECONOMIC GEOLOGY v. 98, p. 1637-1666. Wilson, A.J., Cooke, D.R., Harper, B.J., Deyell, C.L., 2006, Sulfur isotope zonation in the Cadia district, southeastern Australia: exploration significance and implications for the genesis of alkalic porphyry gold-copper deposits: Mineralium Deposita, v. 42, p. 465-487. Zheng, Y.-F., 1993, Calculation of oxygen isotope fractionation in hydroxyl-bearing silicates. Earth and Planetary Science Letters, v. 120, p. 247-263. Zhu, C., Sverjensky, D.A., 1991, Partitioning of F—Cl—OH between minerals and hydrothermal fluids: Geochimica et Cosmochimica Acta v. 55, p. 1837—1 858. Zhu, C., Sverjensky, D.A., 1992, F—Cl—OH partitioning between biotite and apatite: Geochimica et Cosmochimica Acta, v. 56, p. 3435—3467. 123 Chapter 3: Composition of biotite from the Southwest Zone alkalic porphyry Cu-Au deposit, Galore Creek, BC, Canada: evaluation of hydrothermal fluid chemistry 3.1 Introduction Biotite monitors the physiochemical environment during growth (Guidotti, 1984; Speer, 1984). Many thermodynamic variables influence the complex chemistry of biotite, therefore, its composition is potentially useful in understanding some of the physical and chemical conditions of hydrothermal events leading to the formation of porphyry Cu deposits (Beane, 1974; Jacobs and Parry, 1979; Munoz, 1984); particularly temperature, oxidation state, pH, and HF, HC1, and H20 fugacities. The majority of the previous studies of biotite composition in porphyry Cu deposits have focused on distinguishing between mineralized and barren plutons (Parry and Jacobs, 1975; Mason, 1978; Hendry et al., 1988). In contrast, only a few studies characterize biotite composition with respect to alteration assemblages and porphyry centers: Santa Rita Cu deposit, New Mexico (Jacobs and Parry, 1979); Henderson Mo deposit, Colorado (Gunow, et al., 1980); Casino Cu-Au-Mo occurrence, Yukon (Selby and Nesbit, 2000). There is a paucity of work investigating maj or and minor element composition of biotite in the alkalic porphyry Cu- Au class. Furthermore, no examples of studies investigating spatial variations in vein biotite composition from within a porphyry Cu system are documented. This manuscript examines biotite in a spatial context based upon the Southwest Zone breccia-centered silica-understated alkalic porphyry Cu-Au deposit at Galore Creek. The Galore Creek alkalic porphyry district is located within the Stikinia terrane (Fig. 3.1) at the western margin of the Intermontane Belt in the Canadian Cordillera, NW British Columbia, Canada (Allen et al., 1976). The Southwest Zone deposit is one of twelve Cu-Au mineralized centers in the Galore Creek district (Enns et al., 1995; Schwab et al., 2008). A version of this chapter will be submitted for publication. Byrne, K., and Dipple, G.M., Composition of biotite from the Southwest Zone alkalic porphyry Cu-Au deposit, Galore Creek, BC, Canada: evaluation of hydrothermal fluid chemistry. 124 In the Southwest Zone, Cu-Au is intimately associated with quartz-deficient biotite- K feldspar-magnetite alteration and infill (vein and cement). Our analysis show that the biotite types, least altered igneous, secondary and infill form distinct compositional groups which reflect their genetic origins. Infill biotite composition varies systematically with respect to Cu-grade and proximity to calc-silicate veins and alteration, which locally occurs outboard of intense potassic alteration. Gradients in the composition of infill biotite are interpreted to be mostly in response to changingfO2, and represent a record of hydrothermal fluid evolution. Based on the variations in infill biotite chemistry mechanisms of Cu deposition are proposed. Estimates of HF, HC1, and H2O ftigacities ratios of the hydrothermal fluids, attending infill biotite formation suggest lowerH20/HF compared to most calc-alkalic porphyry Cu system. Furthermore, fugacity values from the Copper Canyon (B.C) and Skouries (Greece) alkalic porphyry Cu deposits are also distinct from caic-alkalic class. Contrasting fugacity values are interpreted to reflect magmatic source compositions and their influence on the composition and evolution of magmatic volatile phase. Electron microprobe analysis of biotite in polished thin-sections is used to characterise biotite types in the Southwest Zone and to test for chemical variation within individual infill-type grains. Ten linearly independent Thompson-space components are used to describe the composition of infihl biotite and the most common substitution mechanisms involving Al3, Fe3, and Ti4. This model accounts for several vacancy substitutions, permitting an estimation of Fe3 from electron microprobe analyses. Grains of infihl biotite were randomly selected from 45 individual drill-core samples taken from different locations in the Southwest Zone. Atoms per-formulae unit Ti and Fe3 / Fe (calculated) are plotted with Cu-concentration shells on cross-sections through the Southwest Zone. The relation between XF, Xci, XOH and XMg* are used to estimate HF, HC1 and H20 fugacity ratios (Munoz, 1992; Zhu and Sverjensky, 1991, 1992) of the hydrothermal fluids attending biotite formation in the chalcopyrite suiphide zone. Halogen data, and major and minor cation compositional features of infill biotite from Southwest Zone are compared to other porphyry systems, such as Bingham, Utah; Casino, Yukon; Los Pelambres, Chile; Santa Rita, New Mexcio; Bakircay, Turkey; Skouries, Greece; and Copper Canyon British Columbia. 125 Canadian Cordillera • Cities / towns • Alkalic porphyry deposits 200km FIG. 3.1 Map of British Columbia and the Canadian Cordillera showing the location of the accreted Quesnellia and Stikinia ocean arc terranes, morphogeological belts, Galore Creek, and major alkalic Cu-Au porphyry deposits. Data sourced from BC MapPlace (http://www.mapplace.caf). 3.2 Geological framework Regional: The Quesnell-Stikine arcs, in the Canadian Cordillera, are host to Triassic and Early Jurassic alkalic porphyry Cu-Au deposits (Fig. 3.1) formed outboard of ancestral North America in island-arc tectonic settings (McMillan, 1991). The majority of the alkalic porphyry Cu-Au deposits are products of a discrete magmatic event at the end of the Triassic, with a second and smaller event in the Early Jurassic (Mortensen et al., 1995). Of the Triassic complexes, the Galore Creek alkalic intrusive suite is one of the largest and most silica undersaturated intrusive complexes in the world Explanation Terranes Quesneiha Stikinia 126 (Lang et al., 1995a), and host to the most significant undeveloped alkalic porphyry Cu- Au resource in North America. District: The Galore Creek alkalic intrusive suite is hosted in Stuhini Group rocks (Fig. 3.2) composed of a volcano-sedimentary rock package deposited in a sub aqueous to subaerial environment (Logan and Koyanagi, 1994). The district contains numerous dikes and thick composite bodies, commonly characterized by megacrystic K- feldspar phenocrysts up to 10cm long. The Galore Creek alkalic intrusions are proposed to have been intruded during an extension possibly along north-south trending faults during the Late Triassic (Logan and Koyanagi, 1994). The most visible structures within the Galore Creek district are the post-mineral west-dipping Butte Thrust Fault (Schwab, et al., 2008) and the east-dipping Copper Canyon Thrust Fault (Bottomer and Leary, 1995) (Fig. 3.2). Weakly metamorphosed epiclastic conglomerates and sandstones are thrust over undeformed Stuhini group rocks and alkalic intrusions along the Butte Thrust. Hydrothermal alteration is developed in the multiphase complex of alkalic intrusion and encapsulating shoshonitic volcano-sedimentary rocks. The district includes approximately twelve zones of Cu-Au mineralization (Fig. 3.3). The largest deposit, and the best documented, is the northerly-elongated Central Zone (Enns, et al., 1995; Lang et al., 1995b; Micko, et al., 2008). Smaller deposits peripheral to the Central Zone are also known, and include: Southwest Zone, Junction, Butte, West Rim, Westfork and the Saddle zones (Emis et al. 1995; Schwab et al., 2008). Mineralisation in the Central Zone is truncated to the west by post-mineral megacrystic orthoclase-phyric syenite and monzonite dikes (Enns et al., 1995; and Schwab et al., 2008). A second stage of mineralisation occurs in megacrystic orthoclase-phyric syenite and monzonite in the Southwest Zone (Fig. 3.3). 127 z0 Intrusive rocks Volcanic and Sedimentray rocks _ Coast Plutonic Complex Stuhini Group (Eocene) (Mid to Upper Triassic) Calcaikalic Plutonic suite Shale and argilite (Early Jurassic) (Middle Triassic) , Galore Creek alkalic suite Stikine assemabigei (Upper Triassic) (Devonian to Permian) • Hickman thohith • Limestone (Permian) FIG. 3.2 Regional scale geology of Galore Creek showing the location of Copper Canyon alkalic porphyry Cu-Au occurrence (modified after Logan and Koyanagi, 1994 and Enns et al., 1995). BT = Butte Thrust Fault; CCT = Copper Canyon Thrust Fault. 128 352000 — a :::-addIe km 1 Rock Types dark orthoclase por fT hydrothermal ,——- phyry syenite; pseudoleucite : breccia .iitY ± pseudoleucite- -bearing bearing volcanic / volcaniclastic monzodiorite € sedimentary megacrystic orthoclase-phyric augite and feldpar syenite and G:D -bearing volcanic •‘ monzonite /volcaniclastic FIG. 3.3 Simplified geological map and location of mineralised centers in the Galore Creek District. Deposit: The Southwest Zone, the focus of this contribution, is a Cu-Au breccia-centered deposit with a measured and indicated resource of 69 Mt grading 0.42% Cu, 0.63g/t Au, 2.67g!t Ag. Numerous megacrystic orthoclase-phyric syenite and monzonite intrusions are cut by matrix-bearing breccias and cemented breccias (Fig. 3.4). The matrix-bearing breccias body is approximately 400m wide by 800m long that extends to at least 600m below the current surface (Figs. 3.3 and 3.4). Matrix-bearing breccias are predominantly polylithic, unbedded, poorly sorted, matrix-rich, and matrix supported. A biotite ± K-feldspar ± magnetite ± anhydrite ± diopside ± suiphide assemblage occurs as both hydrothermal cement and veins that cut the matrix-bearing breccia and host intrusions. Two cemented breccia domains transect the matrix-bearing 340050 F 129 breccia-porphyry wall rock contact, forming the upper (main) and lower composite cemented breccia zones (Fig. 3.4). The east-trending upper cemented breccia dips ‘4O0 south has a semi-ellipsoidal morphology 20-lOOm thick, 500m wide and 400m in length, and tapers towards its tips (Fig. 3.4). Mineralization is focused on the western matrix- bearing breccia-wall rock contact (Fig. 3.3) and is coincident with much of the cemented breccia domains. Biotite-phyric monzodiorite post-dates matrix-bearing breccia and occurs as dikes and dikelets focused at the contact between the matrix-bearing breccia and porphyry wall-rocks (Fig. 3.4). Where it intrudes the matrix-bearing breccia, biotite phyric monzodiorite occurs as volumetrically minor interconnected dikelets of the intrusion. Cemented breccia facies, mineralization and alteration is centered on and zoned about biotite-phyric monzodiorite dikes. Biotite ± magnetite ± K-feldspar ± apatite ± bornite ± chalcopyrite ± pyrite ± accessory rutile compose the potassic alteration assemblage in the Southwest Zone, and is associated with the bulk of the Cu-Au and cementation (section 2.7.3, Chapter 2). The abundance of potassic infill minerals decreases outwards from the upper and lower cemented breccia zones. Alteration gradients are more pronounced in the direction perpendicular to the dip of cemented breccia zones, reflected by change in alteration facies and sulphide assemblage over short distances of 5-1 Om. Intense potassic alteration grades outwards to weak potassic alteration and a mixed zone of potassic and propylitic assemblages (Figs. 2.17 and 2.19, Chapter 2). More locally, diopside-magnetite ± garnet (calcic) veins and alteration occurs outboard of intense potassic alteration. The spatial distribution and textural relationships indicate that calcic alteration and veins paragenetically overlapped with, but were locally succeeded by, potassic alteration (section 2.7.4, Chapter 2). Sulphide mineralisation is focused on and zoned about discrete tabular domains of biotite-phyric monzodiorite and coincident cemented breccia facies. The most abundant chalcopyrite (>3%) is coincident with intense phlogopite ± magnetite dominated potassic alteration and cementation. In the upper and lower cemented breccia domains, sulphide minerals are zoned from a core of chalcopyrite-bomite, to chalcopyrite>pyrite, out to pyrite>chalcopyrite (Fig. 2.22, Chapter. 2). The shallow dip of Cu-Fe sulphide mineralisation, metals zoning and the position peripheral alteration facies suggest the deposit is tilted 600 east, consistent with observations made elsewhere in the district by Micko et al. (2008). Tilting is interpreted 130 O Inferred dip direction of contact Out of section. Drill core: on section/off section. meters 0 150 3.3 Biotite types classification Two major genetic types of biotite occur in the porphyry environment; igneous and hydrothermal. Hydrothermal biotite is further subdivided by a scheme adapted from Jacobs and Parry (1979): metasomatised igneous minerals; metasomatised and recrystallized igneous minerals (secondary or shreddy biotite), and infihl crystals precipitated directly from a fluid (vein and cement). In the megacrystic orthoclase-phyric syenite and monzonite stocks least altered igneous biotite occurs predominantly as euhedral to subhedral phenocrysts between 0.5 and 2 mm (Fig. 3.5A). Some igneous biotite is ragged and splintered (Fig. 3. 5B). Least altered igneous biotite is distinguished by brown-dark brown pleochroism. In the potassic alteration zone, igneous biotite phenocrysts locally display recrystallized margins, anomalous pleochroism and diffuse cleavage (Fig. 3.5B). Anomalous pleochroism suggests portions of the phenocrysts are altered without recrystallization (Jacobs and Parry, 1979). Hydrothermal biotite occurs as aggregates of flakes, 0.05-0.2 mm in size, resulting from recrystallization of igneous to have occurred post hydrothermal activity due to north-trending, upright folding (Logan and Koyanagi, 1994). NORTH B 6333900 N EAST SOUTH P 9Om t,5 75Gm • ) /•1 t., 4’, f Clastic: matrix-bearing breccia I€ Pebble to cobble clast size (with clastic: cement-bearing breccia21 J’ <10%cementM-BX) Cobble to boulder clast size (with <10% cement M-BX) I I Coherent:s.oherent i [ Megacrystic orthclase-phyric I Biotite-phyric monzodiorite:syenite and monzonite; thick Dikes and dikelets ____________ composite dikes form wall-rock to matrix-bearing breccias) L FIG. 3.4 Simplified geology along cross sections 6333650N (A-A’) and 350030E (B-B’) 3 600 ro Elevation(meters) 450 m 6333300 N 131 biotite or replacement of igneous hornblende (Fig. 3.5C). Secondary biotite is characterised by felted textures and green-brown to dark green pleochroism (Figs. 3 .5C and D). Fine grained accessory rutile and sphene is locally intergrown with the aggregates of secondary biotite. Igneous groundmass consisting of fine grained K feldspar and plagioclase also appears to be locally replaced by secondary biotite in zones of intense potassic alteration (Fig. 3 .5D). 132 FIG. 3.5 A. Photomicrograph of three least altered igneous biotite phenocrysts in K-feldspar groundmass of megacrystic orthoclase-phyrie syenite. Opaque minerals are predominantly magnetite. B. Photomicrograph of partially recrystallized igneous biotite phenocrysts in potassicaly altered megacrystic orthoclase-phyric syenite. Portions of the primary biotite have been metasomatised resulting in diffuse cleavage boundaries and anomalous pleochroism. C. Plane light and D. polarized photomicrographs of secondary biotite replacing igneous homblende. Note the margins of some igneous biotite grains are recrystallized to secondary biotite in the bottom left side of frame D. E. Plane light and F. polarized photomicrographs of moderately pervasive secondary biotite alteration in megacrystic orthoclase and plagioclase-phyric monzonite. PPL = plane polarized transmitted light. XPL = cross polarized transmitted light. Abbreviations: Ap = apatite; alt = alteration; bio = biotite; cpy = chalcopyrite; gm = groundmass; hbl = hornblende; ig = igneous, K-spar = K-feldspar; ortho = orthoclase; pheno = phenocrysts. 133 The most abundant infihl biotite occurs in the cemented breccia domains, manifest as vein (Fig. 3.6A) and cement (Fig. 3.6B) material. Hydrothermal infihl biotite forms 0.2 to 3 cm coarse interlocking euhedral grains that are typically intergrown with variable proportions of hydrothermal K-feldspar, magnetite, anhydrite, and Cu-Fe suiphides (Figs. 3.6A and B). FIG. 3.6 A. Infihl biotite vein in megacrystic orthoclase-phyric syenite. Porphyry groundmass is pervasively K-feldspar altered and hematite dusted, making the rock a pink colour. B. Infihl biotite cement in cemented breccia. Biotite is intergrown with pyrite-chalcopyrite, hydrothermal K-feldspar and anhydrite. Abbreviations: anh = anhydrite; bio = biotite; cpy = chalcopyrite; K-spar = K-feldspar; mt = magnetite; ortho = orthoclase. Tan-brown pleocliroism further distinguishes infihl biotite from the other biotite types (Fig. 3.7A). Inclusions of magnetite and chalcopyrite commonly occur between cleavage planes of coarse infihl biotite. 134 FIG. 3.7 A. Plane light and B. polarized photomicrographs of one side of an infill biotite vein showing coarse interlocking grains of biotite intergrown with hydrothermal K-feldspar, apatite and anhydrite. Red box indicates the field of view in figure 7A. Red circles indicate the location of electron microprobe analyses shown in figure 3 .9A. Abbreviations: Ap = apatite; anh = anhydrite; bio = biotite; cpy = chalcopyrite; gut = garnet; K-spar = K-feldspar; rt = rutile. Accessory hydrothermal rutile locally occurs at grain intersections of infill biotite (Fig. 3.8A). Chlorite alteration along fracture and cleavage planes varies in intensity from obvious to cryptic. Lighter shades of gray in back scattered electron imaging, highlights incipient chlorite alteration not visible in transmitted light (Fig. 3.8A).These alteration halos correspond to a sharp decreases in K contents. Hydrothermal apatite is locally intergrown with infill biotite (Fig. 3 .8B). 135 FIG. 3.8 A. Back scattered electron image of infihl biotite exhibiting alteration (chlorite?) along fractures; lighter grey than surrounding biotite. Chalcopyrite occurs between cleavage planes and rutile at grain boundaries. B. Back scattered electron image of hydrothermal apatite (false coloured) intergrown with small aggregates of infill biotite in a vein with an irregular morphology. Abbreviations: Ap = apatite; anh = anhydrite; bio = biotite; cpy = chalcopyrite; gut = garnet; hem = hematite; K-spar = K-feldspar; mt = magnetite; it = rutile. Hydrothermal apatite in the Southwest zone is rich in F and S and Cl poor similar to igneous apatite at Galore Creek (Lighat and Tosdal, 2008). Electron microprobe analysis of apatite intergrown with infihl biotite is presented in Table 3.1. 136 TABLE 3.1 Hydrothermal apatite compositions from vein samples in the Southwest Zone DH-ID (GCO5) -678 -678 -678 -678 -678 -661 -661 G-661 Depth (m) 266.5 266.5 266.5 266.5 266.5 292 292 292 Sample-ID I 279-Al I 279-A4 I 279-A5 1 279-A7 I 279-A6 1 063-A4 1 063-2B 1 063-2C No. of spot analysis 11 4 6 6 8 9 8 10 Oxide wt% SiO2 0.84 0.91 0.94 0.79 0.66 0.38 0.43 0.31 FeO 0.05 0.05 0.05 0.09 0.09 0.03 0.08 0.04 CaO 54.23 54.22 54.14 54.15 53.89 54.41 54.09 54.56 Na20 0.17 0.18 0.17 0.12 0.21 0.16 0.21 0.19 S03 1.16 1.16 1.22 0.90 0.81 0.64 0.78 0.59 P205 39.53 39.24 39.30 39.52 39.92 40.28 40.33 40.72 F 3.68 3.73 3.75 3.76 3.70 3.64 3.71 3.64 CI 0.06 0.07 0.05 0.02 0.06 0.05 0.06 0.05 Sub total 99.06 98.88 98.86 98.74 98.94 99.34 99.43 99.98 -0 = F+ Cl 1.56 1.59 1.59 1.59 1.57 1.54 1.58 1.54 Total 97.50 97.29 97.27 97.15 97.37 97.80 97.85 98.43 Sic 0.072 0.078 0.080 0.068 0.056 0.033 0.037 0.026 Fe2+ 0.004 0.004 0.004 0.006 0.007 0.002 0.006 0.003 Ca 4.962 4.980 4.967 4.986 4.950 4.989 4.954 4.970 Na 0.028 0.030 0.028 0.020 0.035 0.027 0.035 0.031 S 0.075 0.075 0.078 0.058 0.052 0.041 0.050 0.038 F 0.994 1.012 1.016 1.023 1.002 0.984 1.004 0.979 Cl 0.008 0.011 0.007 0.003 0.008 0.007 0.009 0.007 P 2.858 2.848 2.849 2.875 2.897 2.918 2.918 2.931 C Atoms per 12 oxygen atoms, excluding F and Cl 3.4 Analytical procedures and sample methodology Paragenetically constrained wall rock, vein and cement samples were taken from drill-core for analysis by electron microprobe. Electron-probe micro-analyses of biotites, in polished thin section and grain separates in polymer pucks, were done on a fully automated CAMECA SX-50 instrument, operating in the wavelength-dispersion mode, at the University of British Columbia. Element determinations for Si, Al, Fetotai, Mg, Ti, Mn, Ba, Na, K, F and Cl in biotite were made with the following operating conditions: excitation voltage, 15 kV; beam current, 10 nA; peak count time, 20 s (40 s for F, Cl); background count-time, 10 s (20 s for F, Cl); spot diameter, 10 j..tm. Data reduction was completed using the ‘PAP’ (pZ) method (Pouchou and Pichoir 1985). For the elements considered, the following standards, X-ray lines and crystals were used: synthetic phiogopite, FKa, TAP; albite, NaKa, TAP; anorthite, A1Ka, TAP; synthetic phiogopite, MgKa, TAP; synthetic phiogopite, SiKa, TAP; scapolite, ClKa, PET; synthetic phiogopite, KKa, PET; diopside, CaKa, PET; rutile, TiKa, PET; synthetic magnesiochromite, CrKa, LIF; synthetic rhodonite, MnKa, LIF; synthetic fayalite, 137 FeKa, LIF. A microprobe analysis from polished thin section is defined as the arithmetic mean of normalized cations of at least three spot analyses per biotite grain. A second suite of infill biotite samples were taken from 45 different locations in the Southwest Zone. Grains of infihl biotite were taken randomly from each of the 45 individual drill- core samples, set in polymer puck, and polished in preparation for analysis. An analysis of grain separate biotite composes at least two probe spot analyses. Barium content of infill biotite grain separate samples was not analysed for. Analysis with totals, inclusive of calculated H20, between 98 and 101.5% are considered good. In addition to the oxide total check, stoichiometric criteria are applied to define good data: tetrahedral site = 4; octahedral site = 2.96 — 3.04; alkali site = 0.95 and 1.05; and calculated charge and oxygen equivalents within ± 0.1 of 22 and 11, respectively. The normalization scheme to determine site occupancies is described in the next section. Biotite data population by type, number of samples, grains and spot analysis is presented in Table 3.2. All data is compiled in Appendix D. TABLE 3.2 Southwest Zone biotite data distribution. No. of samples No. of grainsTextural type No. of spot analysis Leastalerted igneous 5 10 45 Secondary 5 7* 37 InfiN (thin-section) 9 20 132 Infill (grain separate) 45 45 99 * Grains consists of clusters of fine-grained of secondary biotite 138 3.5 Normalization of microprobe data and estimation of Fe3 The problems inherent in generating mineral formula for mica from microprobe data alone have long been recognized (Holdaway, 1980; Dymek, 1983; Feldstein, 1996). Ferric iron determinations by wet chemical analysis and Mössbauer spectroscopy imply significant amounts Fe3 in biotite in most geological environments (Jacobs and Parry, 3+1979; Guidotti and Dyar, 1991; Feldstein et al., 1996). Experiments indicate that the Fe content of biotite increase with 102 (Wones and Eugster, 1965; Deer, Howie and Zussman, 1992). Thus, the ratio of Fe3 to Fe2 in biotite can be used to evaluate magmatic or hydrothermal 102 conditions. This is of particular importance in the study of porphyry Cu magmatic-hydrothermal systems as the solubility of CuC1(aci) and AuCl(a are affected by variation infO2 and pH (Huston et al., 1993). Electron microprobe analysis cannot distinguish between Fe2 and Fe3, and H20 is likewise not determined. Consequently, calculations of formula proportions of ions from analyses of a hydrous, Fe3-bearing mineral, such as biotite, require certain assumptions. Two normalization schemes are commonly applied. The first is expressed by the relationship: total cations — (K+Na+Ca+Ba) 7 which assumes complete occupancy of octahedral and tetrahedral sites (Ludington and Munoz, 1975). Such cation-based normalization procedures are unaffected by the valance state of Fe and variations inH20-content. In biotite, assuming a full complement of 10 0 and 2 (OH, F, Cl) per formula unit, Fe3 would be equal to the difference between 22.0 (11 0) and the calculated total positive charge. If octahedral vacancies are present, however, the formula proportions will be overestimated in direct proportion to the number of vacancies, resulting in a corresponding charge excess, making Fe3 estimations impossible. The second scheme, assumes all iron occurs as Fe2, involves normalization to 11 0 atoms equivalents (neutral mineral), and is expressed by the relationship: total cation charge = 22.0 (Deer, Howie and Zussman, 1992). By normalizing biotites to a set charge, the number of cations obtained may be less than 8, resulting in apparent vacancies that are often assigned to the interlayer (A) and br octahedral sites (VI). This procedure is unaffected by the presence of vacancies, but causes cation abundances to be 139 overestimated in direct proportion to the number of Fe3 present, which remains an unknown quantity. Chemical variation in biotite can be described, cation exchange, coupled substitution (such as Tschermak exchange), vacancy substitutions, oxy-substitution (deprotonation / dehydrogenation reactions) or coupled interlayer-vacancy substitutions in the biotite lattice (Foster, 1960; Brigatti and Guggenheim, 2002; Waters and Charnley, 2002). By accounting for vacancies substitutions, a tailored cation-based normalization scheme can be applied that will eliminate charge discrepancies permitting estimation of Fe3 (Dymek, 1983; Waters and Chamley, 2002). Common substitutions involving Ti4, Fe3 and Al3, and those reviewed by Dymek (1983), are presented in Table 3. .4+ 3+ 3+.TABLE 3.3 Substitution mechanisms for Ti , Al and Fe in biotite Substitution Name Reference 1 “1Ti+ 2AI = VIMg ÷ 2IV Ti-Tschermak 2 VITi + VI D = 2”1R Ti-vacancy (R) 3 1 .5”Ti + 0.5 ‘ D = 2”1A1 Ti-vacancy (Al) 5 vj + O.67”Ll = vIM9 + 0.67”1Al Ti-vacancy (Mg-Al) 6 VIAl + IVAI = 3VIR + lV Al-Tschermak 7 2”1Al + VIE = 3VIR Al-vacancy 8 VIFe3+ + IVAI3+._ VIFe2++ IV Coupled substitution Robert, 1976; Dyaretal, 1993 DahI, 1969; Dymek, 1983 Brigatti and Guggenheim, 2002 Holdaway, 1980. Foster, 1960 Foster, 1960 Rebert et al., 1995; Brigatti and Guggenheim, 2002 9 2”Fe3+ VIE = 3VIFe2+ Rebert et al., 1995 10 “Fe3= “Al Foster, 1960; RebertetaL, 1995 Divalentcations in the octahedral site (such as Fe and Mg) are denoted as “(R) and cations in tetrahedral coordination are indicated with the IV superscript. A = alkali or interlayer site. The symbol LI represents a vacancy. Si4 and Al3are assumed in all reactions. Fe-vacancy Twelve linearly independent Thompson space components (Thomson, 1982) are defined that can describe all biotite analysis from the Southwest Zone and account for exchange mechanisms 1-8 in Table 3. The additive component phlogopite and the other eleven linearly independent exchange components are listed in column A Table 4. 140 TABLE 3.4 Biotite Thomson-space components A B PhI* K(Mg)3S1A2010(OH K(Mg)3S1AI2O10(OH I TIMgAL2 Fe2TiSL 2 AI2Mg1SL FeAlMg.iSi.1 3 FeMgi FeMgi 4 MnMgi MnMgi 5 Al2 L”Mg3 Ti [VIM9 6 CaMgi CaMgi 7 BaMgi BaMgi 8 S1LIAKAI S1L?’KA1 9 NaK.1 NaK1 10 F(OH).i F(OH).1 11 CI(OH) .i CI(OH)1 *Phl = phiogopite The advantage of adopting this approach is that a single normalization scheme (A) is used to provide an objective description of all data. The transformation involves the conversion of mineral analysis to cation formulae based on cation sum = 8 and does not initially impose site constraints. The value equal to the difference between calculated oxygen equivalent and 11 is attributed to Fe3. Normalization by these exchange components resulted in overestimation of octahedral occupancy, unrealistically low Fe3 estimates (< 0.02 apfu), and negative values of Al. The later is likely caused by insufficient cations to fill the tetrahedral site to the ideal value of 4. This approach was abandoned as results were non-stoichiometric. Ferric iron can also be accommodated in the tetrahedral site (Foster, 1960; Deer, Howie and Zussman, 1992; Rebert, et al., 1995; Rieder et al., 1998) most simply by vIAl exchange (substitution mechanism 10, Table 3). A second set of Thompson space components which constrain Fe3 to the tetrahedral site are listed in column B, Table 4. Cation normalization by transformation of this set of components (scheme B) produces stoichiometric site occupancies and oxygen equivalent values between ± 0.1 of 11. Ferric iron content estimated by the second set by scheme B yield more realistic vales comparable with wet chemical Fe3 determinations for biotite formed in similar 141 geological environments (Santa Rita, Jacobs and Parry, 1979; Copper Canyon, Tweilcer, 2007). All biotite analysis and calculated formula are presented in Appendix D. Values of OH are calculated by difference assuming hydroxyl site occupancy of 2 and water contents are calculated with the oxygen normalization factor (norm factor). Molar (X) Mg is defined as Mg / (Fetotai+Mg). The molar proportion of F is defined as a fraction of hydroxyl site anions; F / (F+Cl+OH). Similarly, Xci = Cl / (F+Cl+OH) and XOH = XOH / (F+Cl+OH). 3.6 Biotite composition by type Results of electron-microprobe analysis of least altered igneous, secondary and infill-type biotite in the Southwest Zone, normalized by scheme-B, are illustrated in major and minor element plots in Figure 9. Representative biotite gain compositions (section 3.4) from the Southwest Zone are presented in Table 3.5. 142 • . • 0 •.6*•. . . •. . . . .5 0.90 00 1.25 1.15 a 0.95 1.00 * 0.80 If 0A0 if 0.20 0.00 0.40 0.30 U. x 0.20 0.10 0.00 0.62 00 • A0 0 I. A I .0C . A 0 S fl7f . . .. . A•4A. • • •. .e . A.. •, . 0Io 0 on0 ° A . e..• • • A • •• A • A 2.9$ 785 2.75 2.65 0.24 020 .5 o.is I- 0.12 0.08 0.03 0.02 0.01 . •. : • . tO 4 A As 0 a o 0 . : .: •,,, •• 0.007 .50.005 U 0.003 0.010 0.005 0.000 a . : •• A A a.0 0’ a • . • ?4c.0.66 0.70 0.74 __________________________________ __________________________________ XM9 -. 0.62 0.66 0.70 0.74 0.62 0.66 0.70 0.74 Explanation - biotite types XMg XMg Least altered igneous A Secondary • Infilt FIG. 3.9 XM9 versus atoms per formulae unit A1<tO), Si, Fe(0)Fe3/Fe2, Ti, Cl, XF, Mn, and Ba for least-altered, secondary and infihl biotite in polished thin sections from the Southwest Zone. Each symbol represents an arithmetic mean of at least three or more normalized microprobe spot analyses per biotite 3+ 2+gram. XMg = Mg / (Fe10+Mg); Fetotai Fe + Fe ; X = F / (F+Cl+OH); apfu = atoms per formula umt. TABLE 3.5 Representative biotite compositions from the Southwest Zone by textural type B DH-ID -651 -661 -678 -655 -661 -678 -642 -651 Depth (m) 87.8 183 300.8 220 292 266.5 284 87.8 Biotite type Igneous Igneous nd ry Infill Infihl Infill Infill B B 5 5 7 6 8 7 5 5 B B B B B No. of spot analysis Oddewt% A*____ A A A A A A Si02 36.53 36.90 39.18 39.83 38.71 37.04 38.81 39.24 Ti02 3.79 3.94 2.89 1.86 2.10 2.66 2.45 2.25 A1203 14.28 14.52 12.04 13.56 12.35 13.40 13.08 12.64 FeO 14.37 14.38 13.46 12.60 12.58 14.20 12.75 13.07 MnO 0.56 0.44 0.32 0.20 0.55 0.36 0.38 0.56 MgO 16.33 15.16 17.16 16.94 18.05 16.55 18.21 18.12 GaO 0.08. 0.04 0.01 0.02 0.00 0.02 0.01 0.02 Na20 0.12 0.13 0.05 0.05 0.04 0.11 0.19 0.10 1(20 9.31 9.70 10.15 10.22 10.24 10.07 9.83 9.87 BaO 0.20 0.14 0.01 0.00 0.05 0.06 0.00 0.00 F 0.58 2.22 2.88 2.66 1.90 1.46 1.96 1.00 Cl 0.04 0.03 0.03 0.05 0.05 0.05 0.05 0.04 Subtotal 96.20 97.61 98.15 97.94 96.63 95.97 97.72 96.93 0 = F+Cl 0.25 0.94 1.22 1.13 0.81 0.62 0.84 0.43 (+H20) 4.09 4.07 2.68 2.81 3.12 3.30 3.17 3.62 Total 100.04 100.74 99.61 99.63 98.94 98.65 100.05 100.12 Si 2.675 0.050 2.717 0.043 2.899 0.061 2.924 0.073 2.879 0.052 2.770 0.012 2.828 0.011 2.866 0.012 Al(VI) 0.875 0.097 0.915 0.016 0.958 0.012 0.957 0.022 0.973 0.014 0.963 0.010 0.914 0.017 0.920 0.010 Fe3 0.450 0.143 0.368 0.053 0.142 0.067 0.119 0.085 0.148 0.052 0.266 0.014 0.258 0.024 0.214 0.014 SUM IV 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 Ti 0.209 0.027 0.218 0.037 0.161 0.032 0.103 0.017 0.118 0.032 0.150 0.005 0.134 0.011 0.124 0.005 Al(IV) 0.357 0.115 0.344 0.054 0.092 0.062 0.216 0.093 0.109 0.074 0.218 0.018 0.209 0.020 0.168 0.016 IV-Vacancy 0.162 0.037 0.206 0.027 0.135 0.022 0.151 0.035 0.098 0.017 0.126 0.008 0.109 0.010 0.100 0.005 Fe2 0.431 0.091 0.517 0.062 0.691 0.045 0.655 0.057 0.635 0.110 0.622 0.063 0.519 0.017 0.584 0.075 Mn 0.035 0.006 0.027 0.009 0.020 0.003 0.013 0.004 0.035 0.007 0.023 0.004 0.024 0.003 0.035 0.006 Mg 1.783 0.047 1.664 0.051 1.893 0.081 1.854 0.120 2.000 0.154 1.845 0.054 1.978 0.022 1.973 0.067 SUM VI 2.976 2.978 2.992 2.991 2.995 2.983 2.972 2.984 Ca 0.006 0.012 0.003 0.005 0.001 0.001 0.002 0.001 0.000 0.001 0.002 0.001 0.001 0.001 0.002 0.002 Na 0.017 0.005 0.019 0.006 0.007 0.004 0.008 0.003 0.005 0.003 0.015 0.005 0.027 0.006 0.014 0.004 K 0.870 0.096 0.911 0.017 0.958 0.012 0.957 0.022 0.972 0.013 0.961 0.010 0.914 0.017 0.920 0.010 Ba 0.006 0.002 0.004 0.003 0.000 0.000 0.000 0.000 0.002 0.001 0.002 0.001 0.000 0.000 0.000 0.000 A-Vacancy 0.125 0.097 0.085 0.016 0.042 0.012 0.043 0.022 0.027 0.014 0.037 0.010 0.086 0.017 0.080 0.010 SUMA 1.024 1.022 1.008 1.009 1.005 1.017 1.028 1.016 F 0.134 0.037 0.517 0.093 0.673 0.123 0.617 0.066 0.447 0.061 0.345 0.021 0.452 0.052 0.231 0.034 CI 0.005 0.001 0.004 0.002 0.004 0.002 0.006 0.002 0.006 0.003 0.006 0.002 0.007 0.002 0.005 0.001 OH 1.861 0.037 1.479 0.092 1.322 0.122 1.377 0.066 1.547 0.062 1.649 0.022 1.542 0.052 1.764 0.035 SUM 2.000 2.000 2.000 2.000 2.000 2.000 2.000 2.000 Norm factor 4.400 4.424 4.446 4.411 4.468 4.495 4.378 4.388 Oxygen 10.994 10.993 XM9 0.669 0.653 * Column A are averages of multiple spot analyses, column B is the standard deviation. 10.996 10.996 10.998 10.993 10.987 10.993 0.694 0.705 0.719 0.675 0.718 0.712 143 The compositions of biotite types are shown by plotting XMg against Altotai, Si, Fe(total), Fe3t’Fe2, Ti, XF, Cl, Mn and Ba atoms per formulae unit (Fig. 9). All analyses of biotite plot within the phlogopite compositional field, where phlogopite is equal to Mg # values> 0.5 (Tischendorf et al., 2004). All but one igneous biotite sample have XMg < 0.68. Hydrothermal biotite mostly plots in the > 0.68 XMg fields, though there is considerable variation (XMg of 0.67 to 0.74). Silica atoms per formulae unit correlate with XMg for secondary and infill biotite and less so for igneous biotite. Igneous biotite has the lowest Si contents compared to the other types. Secondary biotite contains the most Si, whereas infill biotite plots in between the two biotite types. Molar proportion Mg correlates negatively with Fe(total) for each biotite-type, this is part a reciprocal affect. Infill biotite exhibits the widest range in Fe(total), whereas igneous and secondary biotites cluster at 0.85 and 0.75-0.85, respectively. Igneous biotite has markedly higher Fe3/Fe2 3+ 2+ values in comparison to the other types. Secondary biotite has Fe I Fe values between 0.2 and 0.3. Infill biotite plot between igneous and secondary forming two arrays correlating positively with XMg. Titanium content does not correlate well with XMg for any biotite type. Highest Ti content occurs in least altered igneous biotite. Secondary biotite plot mostly below Ti values of 0.14, and infill biotite plots has Ti contents in between the other types. Molar proportion F has a weak positive correlation with XMg for secondary and infill biotite types. Secondary biotite has the highest XF values compared to the other biotite types. No obvious correlation between Cl contents and XMg is evident. Most infihl biotite contains > 0.005 atoms per formulae unit Cl. Whereas igneous and secondary have Cl values mostly below 0.006. Secondary and igneous biotites have similar values of Cl. Highest Mn contents occur in least altered igneous biotite. Secondary biotite has the lowest Mn values and infill biotite Mn content is mostly in between the other textural types. Igneous biotite generally has high Ba values compared to hydrothermal biotites. The four igneous samples with low Ba may have undergone metasomatism (Beane, 1974, Jacobs and Parry, 1979), due to interaction with hydrothermal fluids, without any significant recrystallization, or change in Ti and Mg contents. The trends in biotite composition with respect to textural type are generally consistent with observations made at Santa Rita (Beane, 1974, Jacobs and Parry, 1979) and Casino (Selby and Nesbit, 2000) porphyry Cu deposits. Southwest Zone infill biotite, 144 however, has higher Ti content in comparison to biotite vein samples from these deposits and the nearby Copper Canyon porphyry Cu-system (Twelker, 2007). Trends in Fe3/ Fe2 values from high with igneous to lower values for hydrothermal biotite is also reported at Santa Rita (Jacobs and Parry, 1979), however, the average Fe3 content of infill Southwest Zone biotite is 5% higher in comparison. Individual infill biotite grains were tested for compositional heterogeneity. Some grains exhibit varying Ti, F and Mg contents with respect to grain margins; Ti decrease and XMg and F contents increases towards the margins of the cluster of biotite grains (Figs. 3.7A and 3.10). F Ti 0.55 0.16 rim core rim XMg 05 0.15 Ti I’ 0.45 0.14 0.74 04 0.13 = F — •• / 0.35 0.12 ::: on 0.2 0.09 0.71 0.15 0.08 0,7 0.1 0.07 4 3 2 1 (apfu) microprobe spot analyses FIG. 3.10 Compositional variation of the cluster of infihl biotite grains indicated in figure 3 .7A with a red line. The cores of some infill biotite are enriched in Ti and depleted in F and Mg relative to the rims. Highest contents (apfu) of Mg and F occur closest to the grain margin. Energy-dispersive X-ray spectroscopy of infihl biotite suggests that these compositional heterogeneities are most pronounced within the region closest to the grain margin. However, most infill biotite grains appear to be relatively compositional homogenous. Furthermore, grains from different locations in the Southwest Zone exhibit larger variation in composition than that observed in single grains. 3.7 Spatial variation in Ti and Fe3 Titanium, Fe3and Fe2 analyses of infill biotite from 49 different locations in the Southwest Zone are plotted to test for systematic spatial variations in composition. 145 Sample coverage and distribution was limited by the availability of drill-core and good data. The average, maximum, minimum, median and standard deviation of plotted infihl biotite analyses are presented in Table 3.6. TABLE 3.6 Titanium, Fe3and Fe3/ Fe2 values for infihl biotite n = 49 (159 spot analysis) Ti Fe3+ Fe3 I Fe2 Average 0.147 0.233 0.436 Max 0.223 0.455 2.378 Mm 0.030 0.115 0.147 Median 0.147 0.239 0.414 Stdev. 0.032 0.051 0.283 Concentrations of Ti, Fe3 and Fe2 in infill biotite do not vary by a large degree (Table 6). However, relatively consistent trends in Ti and Fe3 /Fe2 with respect to Cu concentration, and to a lesser extent proximity to calc-silicate veins, are observed. Titanium content is normalized to the highest Ti contents measured (0.223 apfu) and plotted as a percentage. Titanium contents are generally high in biotite samples within> 1% and 0.3-1 % Cu shells (Fig. 3.11). 146 WEST A Ti % (norm) Q<58 58-70 • 71-80 •>80 I 1% norm =Ti (apfu) ___________________ normalized to Ti (apfu) maxof 0223 FIG. 3.11 Ti % (norm) values of infill biotite, distribution of caic-silicate alteration and infihl, and Cu concentrations on cross sections 6333650N (A-A’) and 350030E (B-B’) in the Southwest Zone. Blue coloured transparent polygons highlight areas discussed in text. On section A-A’ the highest Ti values are off-set to the west of the 1% Cu shell (Fig. 3.11). On section B-B’, highest biotite Ti-content is in the deepest part of the system and overlaps with> 0.3% Cu mineralisation (polygon W, Fig. 11). Titanium % (norm) values less than 58% are generally in Cu-poor zones (<0.1%) and I or overlap with caic-silicate veins and alteration (although there are exceptions). Ti-concentrations are typically higher in well mineralised zones compared to adjacent Cu-poor or weakly Cu mineralised rocks. The absolute Ti concentration varies between Cu domains, however, the relationship between Ti and Cu is fairly consistent; a sharp in Ti-contents corresponds to a decrease in Cu concentration (Fig. 3.11). This relationship can be seen Cu grade • >1 0.1-0.3 • 0.3-1 cZ <0.1 EAST meters 150 147 in some mineralised zones where steep gradients in Ti-content occur in the direction perpendicular to the dip of mineralisation (polygons X, Y, Z, Fig. 3.1 1). Biotite Ti- content also exhibit, shallower, up-dip gradients within Cu mineralised zones. This is most clear in areas highlighted with polygons X and Z (Fig. 3.11). Furthermore, high values in the deepest parts of section B-B’ (polygon W) zone to moderate and low values in the 0.3% shell in the shallower mineralized zone (polygon X) in figure 3.11. Ferric-ferrous iron ratios show more scatter with respect to the distribution of Cu than Ti (Fig. 3.12). WEST o <0.25 c 0.25 -0.33 0.33-0.37 • 0.37- 0.40 • 0.40 -0.45 _________________ • 0.45 -0.50 FIG. 3.12 Fe3/ Fe2 of infihl biotite, distribution of caic-silicate alteration and infill, and Cu concentrations on cross sections 6333650N (A-A’) and 350030E (B-B’) in the Southwest Zone. Infihl biotite from several Cu-poor rocks have similar Fe3’/ Fe2’ values compared to well mineralized domains (Fig. 3.12). However, some zonation patterns in Fe3 / Fe2 and Cu- Cu grade % • >1 ED 0.1-0.3 • 0.3-1 j <0.1 EAST A’ 7- . __- 0.35 .2 •0A3 •:.:) 00.13 NORTH meters 0 150 148 3+ 2+ concentrations are discernable. Ratios of Fe to Fe increase sharply across the 1 /o Cu contour on sections A-A’ and B-B’ (polygons Q and R, Fig. 3.12). In the area highlighted by polygon Y (Fig. 3.12), Fe3 I Fe2 decrease up-dip within the 1% Cu shell and also trend to lower values laterally into the 0.3% Cu shell. The lowest Fe3 / Fe2 in this area correlate with Cu-concentrations of 0.1-0.3%. The Fe3 / Fe2 also decreases up-dip within the 0.3% Cu shell in polygon S (Fig. 3.12), although, lateral changes in Fe3/Fe2 are not coupled with Cu-concentrations in this area. Infill biotite Ti-concentration has a broadly positive correlation Fe3 / Fe2, suggesting oxidation state affects the solubility of Ti in biotite (Fig. 3.1 3A). Additionally, Ti-concentration exhibits a negative correlation with Si (Fig. 3.13B). 0.20 O.15 0.10: value at 2,66 Si 0.00 2.70 FIG. 3.13 A. Ti vs. Fe3 I Fe2 and B. scheme-B. • •• • • •‘C ?. • •. • 2.75 2.80 2.85 2.90 2.95 Si(apfu) Ti vs. Si. All data in atoms per formula unit and normalized to The specific Ti-content and Fe3 / Fe2 vary between each of the Cu domains. However, gradients in biotite Ti-content and Fe3 / Fe2 and their relationship to Cu- concentration are relatively consistent In summary, Ti-content and Fe3 / Fe2 appear to 149 0.25 0.20 0.15 — 0.10 • •8 • • :..;‘: • T maximumI = standard I deviation - value at0.05 2.3 ratio 0.00 0 0.2 0.4 0.6 0,8 Fe3IFe24 (apfu) • • be coupled with Cu in three ways: 1) highest Ti-content overlaps with well mineralised zones; inversely, the lowest values are associated with Cu-poor rocks and calc-silicate alteration. However, there are several exceptions to this pattern. The correlation between highest and lowest Fe3 / Fe2 and Cu concentrations is not as clear; 2) Gradients in Ti- 3+ 2+content and Fe / Fe to lower values correlates with decreasmg Cu concentrations (particular lateral variations); 3) Although less pronounced, Ti-content and Fe3 / Fe2 also trend to lower values up-dip within Cu-mineralised zones. 3.8 Biotite halogen chemistry To evaluate halogen contents of Southwest Zone biotite and for comparison to data sets from other porphyry deposits the total cation charge = 22.0 normalization scheme is applied. This permits application of reciprocal VI(Mg Fe2, Al) (F, Cl, OH) mixing models of Zhu and Sverjensky (1991, 1992) that do not account for Fe3 and assume all Fe to be octahedral. Similarly, halogen fluid-biotite partitioning coefficient equations by Munoz (1992) do not account for octahedral vacancies, and assume all Fe to be Fe2 and octahedral. Subsequently, the values of XFe, XMg and XMg* in this section of the paper are calculated from data normalized to total cation charge = 22.0. Where XFe (determined from octahedral site cation fractions) is defined as Fe2 +‘Al / 2+VI *.Mg+Fe + Al (Zhu and Sveijensky, 1992), XMg = Mg / Fe(total) + Mg, and XMg 15 defined as Mg / sum of all octahedral cations (Munoz, 1992). Spot analyses of halogen content of infill biotite from the potassic zone are shown in figure 3.14. Infill biotite from the Southwest Zone possesses log (XFIXOH) values between -0.9 and -0.3, and log (XF/XcI) values between 1.4 and 2.9. Fluorine and Cl contents of biotite are controlled to some extent by XFe and XMg (Fig. 14). With increasing XMg and decreasing XFe the values of log (XF/XC1) and (XF/XOH) increase, respectively. This suggests that these biotites exhibit both Mg-Cl and Fe-F avoidance (Munoz, 1984, 1990). 150 045 3.0 2.8 2.6 2.4 ‘I 2.2 2.0 1.8 1.6 1.4 1.2 0.60 0.65 0.70 0.75 0.&) XM9 • biotite data considered reliable a biotite data suspect FIG. 3.14 A. XFe versus log (XF/XOH) of infill biotite grains with 4 or more spot analysis per sample (includes 2 grain separate data). See text for explanation of reliable and suspect biotite data. The good negative correlation between XFe and log (XF/XOH) indicates infill biotites conform to the Fe-F avoidance rule. B. XMg versus log (XF/XCI) of reliable infill biotite. Although there is considerable scatter, log (XF/XC1) has a broadly positive correlation with XMg suggesting that infill biotite mostly conform to Mg-Cl avoidance rule. Substitution of F and Cl into the OW site of biotite is governed by a number of independent factors; which are principally: 1) activity of these elements present during formation; 2) cation population of the octahedral layer (particularly Mg and Fe); 3) temperature of hydroxyl-halogen exchange (i.e. the temperature of precipitation from a hydrothermal fluid) and; 4) effects of subsolidus leaching or enrichment due to hydrothermal fluids or groundwater (Munoz, 1984). Zhu and Sverjensky (1991, 1992) 0.20 0,25 0.30 XFe 0.35 0.40 • a a a • .- . I • •‘ • •• •_ •• 151 developed an internally consistent thermodynamic model describing the reciprocal mixing properties of biotite. Their work allows the study of partitioning of F-Cl-OH between biotites and hydrothermal fluids involving biotite with variable Mg-Fe-Al composition. The thermodynamic formulations (equations 23 and 24) of Zhu and Sverj ensky (1992) demonstrate that the compositions of biotites, formed under the same T, P, and fluid composition, will form linear trends on XFe versus log (XF/XOH): log (XF/XOH) = O.5(AGR,b/2.303RT)(XMg) + 0.5Kf— log(aH2O/aHF°) and XMg versus log (XC1/XOH) plots: log (Xcl/X0H) = 0.5(AGR,c/2.303RT)(XMg) + 0.5Ka log(aH2O/aHCl°) The slopes of these linear trends are a function of T, but independent of fluid chemistry and P (Zhu and Sverjensky, 1992). Infihl biotite data in figure 3.14A is sorted into two groups; reliable (black circles) and suspect (open circles). Sample data is considered suspect if it forms near vertical or near horizontal linear trends as the slopes of these lines yield unrealistic temperatures of formation for hydrothermal biotite; <100 and >800°C. This technique is adapted from Coulson et al. (2001). Halogen composition of biotite is susceptible subsolidus re-equilibration (Dilles, 1987). The biotite samples in the suspect group may have undergone subsolidus re-equilibration with late hydrothermal fluids, a similar phenomena is described by Munoz and Swenson (1981). An alternative explanation is that mineral and fluid did not reach equilibrium (Zhu and Sverjensky, 1992). Removing suspect infihl biotite samples from the data population reduces the scatter in log (XF/XOH), as illustrated by comparing black filled circles and open circles in figure 3.14A, and highlights a strong negative correlation with XFe. There is still considerable scatter, however, in XMg versus log (XF/Xc1) values for reliable infill biotite samples (Fig. 3.14B). Therefore, it appears that the Mg-Cl avoidance rule is more often violated than Fe-F avoidance rule for infihl biotite in the Southwest Zone (similar observations are made by Munoz, 1990 and Selby and Nesbit, 2000). Molar proportion of Fe versus log (XF/XOH) of reliable infihl biotite data are plotted per sample in figure 3.15 (5 samples, 15 grains and 101 spot analyses). To estimate the temperature of formation, the slopes of linear trends of individual sample spot analyses is compared to slopes of 152 temperatures calculated from equation 24 of Zhu and Sverjensky (1992). As a group the data show some scatter, however, on a sample-by-sample basis spot analysis form linear trends that closely parallel the slope of the line calculated at 420°C (Fig. 3.15). Furthermore, the average of these slopes of linear trends of the five samples is - -1.6 corresponding to a calculated temperature of 420°C, plus or minus 50°C. 0.45 1279 I. 1276A D 1063 X 1401A + 1136 FIG. 3.15 Reliable infill biotite data (i.e. data that form shallow linear trends) and the slope of a line (in red) calculated from equation (24) in Zhu and Sverjensky (1992) at 420°C. Note that reliable biotite data form linear affays, per sample, that closely parallel the calculated 420°C slope. 3.9 Hydrothermal fluid halogen fugacity ratio estimates Fluorine and Cl analyses and XOH of reliable biotite samples are used to calculate the log (fIo/fHF), (fo/fHc1) and (tIF/fHc1) values for hydrothermal fluids associated with the precipitation of infill biotite (potassic alteration) and Cu-Au deposition in the Southwest Zone. The fugacity ratios are calculated using the equations of Munoz (1992), which are based on coefficients for F-Cl-OH partitioning between biotite and hydrothermal fluid (Zhu and Sverjensky, 1991, 1992). These equations are: log(JH2O/JHF)th”’= 1 000/T (2.37+ 1.1 (XMg*)b) + 0.43 - log(JH2O/JHCl)fh1d = l000/T (115+ 0.55(XMg*)”°) + 0.68 - log(Xc1/XoH)° 0.20 0.25 0.30 0.35 0.40 XFe 153 log(fHF/JHCl)’ = - 1 000/T (1.22+ 1 .65(XMg*)”°) + 0.25 + Figure 3.16 demonstrates that once the effects of coupled exchange on the R and hydroxyl sites are accounted for, halogen fugacity ratios do not change significantly with variable XMg and XFe at a fixed temperature. °AJ 365t -0.2 67SCf ... I’ 4,5 I’ -0.6 4.0 0 -1 35_i -1.2 0.20 0.25 0.30 0.35 0.40 XFe 35 . 0 . •., -0.5 a .. • . 3.0 •• 475t a 2.5 . • 365C 2,OV •. • 20 • • • • -2.5u. •aI••. •.‘ 3 • .•• 0 15 F1p•l . • - •• •• -4.0 1.0 ____________________________________________ 0.60 0,65 0.70 0.75 XM9 molar ratios of hydroxyl • fugacity ratios site components calcauted at 42OC FIG. 3.16 A. XFe versus log (XF/XOH) and calculated log (t2O/fHF) B. XMg versus log (XF/Xc1) and calculated log (fyIfHcj) fluid Correcting for the effects of coupled exchange on the R and hydroxyl sites (Zhu and Sverjensky, 1991, 1992; Munoz, 1992) show that there is essentially no change in fugacity ratios of halogens with XFe and XMg at fixed temperature. Red lines indicate the direction of shift in fugacity ratios corresponding to 55°C increase and decrease. Fluids associated with the formation of infill biotite and Cu-Au mineralisation in the Southwest Zone have log(fH20/JHF) values of 5.3—5.6 at 420°C (Fig. 3.16A). At the same temperature, log(JHF/JHC1) values exhibit a broader range; -0.13 to -1.54, however, most data cluster between -1 and -1.45, with only a few points extending to lower values 154 (Fig. 3.16B). LogQH2O!JHC1) values are 4.8—6.2 with most points clustering at 4.8-5.5. At higher temperatures log(jH20/JHF), log(jHO/JHC1) values decrease, whereas log(jHF/JHC1) values increase (Fig. 3.16). The similarity in fugacity ratios of analyses from the five reliable infihl biotite samples supports paragenetic observations that this biotite formed during a single hydrothermal event. 3.10 Comparison of fugacity ratios with other porphyry systems Figure 3.17 shows calculated fugacity ratios determined from infill biotite associated with potassic alteration in the Southwest Zone. II Log(fH2o/ HcI) FIG. 3.17 Log(fH2OIfHF) versus log(fH2O/fHC1) ratios for potassic fluids in the Southwest Zone. Also shown are the fugacity values determined for other porphyry Cu systems (modified after Selby and Nesbit, 2000). Southwest Zone fligacity ratios are calculated per-sample at 420°C (red polygon) and 475°C (gray circles) to illustrate the effect of an increase in temperature. The red polygons comprise 87% of Southwest Zone fugacity values. The dashed black line encompasses the remaining 13% of values that extend away from the main fugacity ratio population. Calc-alkalic deposits green text; alkalic systems = black text. Also included in figure 3.17 are the potassic alteration fields for Bingham, Santa Rita, Casino, Los Pelambres, Bakircay (Selby and Nesbit, 2000 and references therein), Copper Canyon (Twelker, 2007), and Skouries (Kroll et al., 2002) porphyry Cu-systems. Santa Rita, Los Pelambres, Bakircay porphyry Cu deposits and the Casino porphyry Cu occurrences are associated with caic-alkaline magmatism. In contrast, the Southwest Zone, Copper Canyon, and Skouries are associated alkalic magmatism. Bingham lies 3.0 3.5 4.0 4.5 5.0 5.5 6.0 155 between these two groups hosted in moderately high-K caic-alkaline rocks (Maughan et al., 2002). Furthermore, mineralisation at Bingham is linked to mafic alkalic magmas (Maughan et al., 2002). Fugacity values for Copper Canyon (B.C) and Skouries (Greece) biotite were recalculated at 550°C based on the geothermometry of Twelker (2007) and Frei (1995), respectively. Fugacity values for the other deposits were recalculated by Selby and Nesbit (2000). Southwest Zone has similar log(f1420/JHC1) values to calc-alkalic systems (Fig. 3.17). Copper Canyon is characterised by high logQH2O/j’HC1) values and overlaps with Bingham, whereas Skouries has log(JH201JF{Cl) lower then most calc-alkalic systems. With the exception of Casino, the caic-alkaline systems have log(J1120/JHF) values in the range of 5.8 to 6.4. In contrast, logQFI2O/JH ) values from the Southwest Zone and the other alkalic systems are considerably lower (Fig. 3.17). Moreover, taking the upper temperature range (475°C) for Southwest Zone biotite values shifts log(fH20/JHF) to lower values and closer to potassic fields from Copper Canyon and Skouries. The upper temperature estimate is consistent with hydrothermal magnetite and Mg-rich biotite stability (Hezarkhani et al., 1999). The fugacity ratios determined for potassic alteration (Selby and Nesbit, 2000) at Bingham show considerable scatter. The data cluster, however, to describe two populations with respect to log(fH2O/JHC1) and logQH2O/JFIF). Based on the data presented in figure 3.17 certain empirical observations can be made, alkalic porphyry systems are broadly characterised as having low log(fH20/JHF) values compared to caic-alkalic porphyry systems. Some analyses from the Southwest Zone and the field for Copper Canyon have logQ1120/JHC1) higher than the caic-alkalic systems. Potassic fluids in calc-alkalic porphyry class show less variation in fugacity ratios and are characterised by relatively high log(fH20/JHF) ratios. Fugacity estimates for Bingham overlap and sit between these two generalized populations. By considering biotite composition in terms of XF/XOH and log XMg/XFe Brimhall et al. (1988) demonstrated that igneous compositions can be grouped into three main classes correlated with a distinct type of base metal ore (Fig. 3.18). 156 1.0 0,5 — a z 0 . -0.5 o -i.o -1.5 contaminated F type / primative I-type melts -2.0 -1.0 -0.5 tO Southwest Zone Lorraine • igneous o re-equilIbrated A igneous 0 igrieous(class 1) • secondary FIG. 3.18 Comparison of secondary and least altered igneous biotite from the Southwest Zone and from Mo, W and Cu porphyry deposit types modified after Brimhall and Ague (1988). Also shown are the igneous biotite compositions from the Mt Milligan (De Long, 1996) and Lorraine (Bath, 2008) alkalic porphyries. Hydrothennal biotites typically form arrays with positive slopes extending away from the composition of local igneous biotite composition (as depicted by the gray arrows) The orange polygons represent the composition field of igneous biotite from Santa Rita, (bottom) and Butte, Montana (top) and are from Brimhall and Ague (1988). Least altered igneous biotite, with BaO not less than 0.1%, from the Southwest Zone and igneous biotite from the Mt Milligan and Lorraine alkalic deposits are plotted in Figure 3.18 for comparison. Southwest Zone igneous grains form a steep array extending to high XF/XOH values towards secondary biotite compositions (Fig. 3.18). These spurious igneous grains with halogen composition approaching the secondary biotite fields may have undergone re-equilibration of the hydroxyl site without recrystallization (Munoz, 1984). Only three igneous grains clusters near fresh igneous values from the other deposits. Hydrothermal biotites typically form arrays with positive slopes extending away from the composition of local igneous biotite composition (Brimhall and Ague, 1988), as depicted by the gray arrows in figure 18. Brimhall and Crerar, (1987) postulate that the ‘— 3/2 slope observed is not solely the result of Fe-F avoidance (Munoz, 1984) and is caused by wall rock oxidation at relatively constant temperature and aqueous activity ratios of a(log(XF/XoH)°/a(XMg/XFe)”° = 1.5. This relationship is not observed for the Southwest Zone igneous-secondary biotite array (Fig 3.18). The deviation from Mi Potassic alteraion 0 Log(XMgIX F,) Mt MIlhgan 157 trends typically observed between igneous and secondary biotite to high log(XF/XoH) values relative to log(XMg/XFe), suggests alteration by potassic fluids with low log(fH2O/JHF). The steep trend observed for Southwest Zone igneous-secondary biotite array is in support of fugacity ratios estimates for potassic fluids (Fig. 3.17). 3.11 Discussion 3.11.1 Significance of gradients in infihl biotite chemistry The variations in Ti-content and Fe3/ Fe2 of infihl biotite could be influenced by a number of parameters, acting alone or in unison, such as: 1) temperature; 2) oxidation state; 3) silica activity in the fluid attending biotite growth. The degree to which these factors influence infihl biotite Ti-concentration Fe3/ Fe2by are evaluated. Robert (1976) experimentally determined that the solubility of Ti in phiogopite was a prograde function of temperature and retrograde function of pressure. Titanium in biotite and Cu concentration appear to correlate, implying that Cu deposition is related to cooling. This poses some problems, however, as steep temperature gradients over narrow zones of mineralisation are unlikely (Cathles, 1977). Furthermore cooling alone is not an efficient mechanism of metal deposition (Cooke and Simmons, 2000; Drummond and Obmoto, 1985). Experiments indicate that the Fe3 content of biotite increase with 102 (Wones and Eugster, 1965; Deer, Howie and Zussman, 1992). The solubility of CuCl(a is strongly affected by variations in 102 and pH (Huston et al., 1993). A progressive decrease in fluid oxidation state, accompanied by pH increase, can lower Cu solubility causing deposition of Cu-Fe sulphides (gray arrow, Fig. 3.19). 158 -20 - 25 -30 - 35 -40 pH FIG. 3.19 Log f(02) — pH diagram at 350°C showing solubility contours for Cu and the stability fields of minerals in the Cu-Fe-O-System (modified from Huston et al., 1993). Gray arrow labeled P1 represents a hypothetical Cu deposition caused by a decrease in f(o2) and increase in pH. Thus, it is proposed that Cu deposition is caused by lowering fluid oxidation state which in turn is recorded by decrease in infihl biotite Fe3 / Fe2 across gradients in Cu- concentration. Intense potassic alteration characterises Cu-enriched rocks and can be related to oxidation state and pH. K-metasomatism of plagioclase and homblende and magnetite alteration of wall rock FeO consumes H and reduces hydrothermal fluids and is represented by the reaction (balanced on Al): 5Ca02Nao.8AlSi3O (igneous plagioclase) + 1.5Ca2(Mg,Fe)4Si7A1O2OH) (igneous hornblende) + 0.6FeO(k) + 0.102 + 9K(aq) + 3H(aq) + 2.5H4SO4(aq) = 7KA1Si3O8 (hydrothermal K-spar) + 2K(Mg,Fe)3A1SiO0(OH) (hydrothermal phiogopite) + 4Ca2(aq) + 4Na(aq) + 6H20 + 0.2Fe3O4 (1) Assuming relatively constant total concentrations of Ti in fluids, the positive correlation of Ti-concentration and Fe3 / Fe2 (Fig. 3.1 3A) suggests oxidation state of the fluid is a factor controlling Ti-solubility in biotite. However, Si and Ti have a negative correlation (Fig. 3.1 3B) suggesting a relationship with fluid aH4SiO.The distribution of most infill biotites with low Ti and high Si contents is outboard of the domains of intense potassic 159 alteration (and high tenor Cu). Thus, if fluids are derived from magmatic rocks, the freshest fluid have higher Fe3 and Ti, and lower aH4SiO Furthermore, an in increase aH4SiOin is favored by cooling (Foumier and Potter, 1982). Russell et al. (1999) speculate that aH4SiO impacts Ti contents in hydrothermal and magmatic gamets. 3.11.2 Halogen chemistry Southwest Zone potassic fluids are characterised by moderately low log(fH201J1-IF) values in comparison to most other calc-alkalic porphyry systems. The Skouries, Copper Canyon alkalic porphyry Cu-systems also exhibit low log(fH2OfHF), suggesting, halogen contents of exsolved magmatic-hydrothermal fluids is largely a function of the composition of the parent magma. Most of the deposits have similar logQH2O/HCl), though Skouries is distinguished by lower values. Timing of vapor saturation in a crystallizing melt also influences halogen composition of exsolved fluids (Candela, 1989). Early vapor saturation removes Cl and enriches residual melt and subsequent fluids in F (Candela, 1986). Accordingly, progressive loss of Cl-bearing phases and residual F-enrichment is attributed to decreasing HC1!HF values of successive intrusions in the Emerald Lake pluton (Coulson et al., 2001). It would seem fortuitous if this process occurred at each of the three alkalic systems. Moreover, Coulson et al. (2001) measure half a log variations in fugacity ratios between early and late intrusions, whereas values of potassic fluids from the Southwest Zone, Copper Canyon, and Skouries are closer to one log value different from typical caic-alkalic systems. Other factors affecting fugacity ratios, post-exsolution, are temperature, pH, and fluid composition. HF is the predominant F-bearing species at high temperatures (Zhu and Sverjensky, 1991), whereas HC1° is present in small concentrations and subject to fluid composition and pH (Zhu and Sverjensky, 1992). Alkali cations and metals tend to complex with Cl at the temperatures of interest (Brimhall and Crerar, 1987), potentially lowering fluid C1 concentrations. Hydrogen chloride tends to disassociate with decreasing temperate to H and cr (Helgeson, 1967). However, K-feldspar stability during high temperature potassic alteration constrains pH to mildly acidic solutions of pH = 5 (Beane, 1974; Helgeson, 1970). Fugacity ratio estimates involving HC1 are sensitive to Cl speciation, pH, and the presence of alkali-Cl and metal-Cl complexes and are not necessarily representative of total Cl. 160 Quantitative assessment of the effects Fe3 on the F-Cl-OH partitioning are incomplete, however, natural samples with high (>45% Fe(total) as Fe3) appear to influence partitioning (Zhu and Sverjensjy, 1991,1992). Infill biotite from the Southwest Zone and Copper Canyon contain an average of 27% and 20% Fe(totaj) as Fe3, respectively. Thus, although there is potential for Fe3tF (or even Fe3-Cl avoidance) to affect partitioning, the Fe3 contents of Southwest Zone and Copper Canyon biotite do not appear to be sufficiently high to influence calculated fugacity ratios. Low log(fH20/JHF) values from the Southwest Zone, and particularly the other alkalic deposits, imply F-enrichment in comparison to caic-alkalic magmatic hydrothermal systems. The presence of hydrothermal fluroapatite also indicates high activities of F in potassic fluids attending Cu deposition in the Southwest Zone. 3.12 Conclusions Based on our chemical characterization of infihl biotite from the Southwest Zone we propose that Fe3 / Fe2 variations reflect the oxidation state of fluids associated with potassic alteration and Cu deposition. Trends to high Fe3/ Fe2 and Ti-contents broadly overlap with positive gradients in Cu-concentration. On the basis of trends in mull biotite Fe3 / Fe2, we submit that decrease in 102 coupled with an increase in pH was the primary Cu-Fe sulphide deposition mechanism. Elevated Ti-concentrations of infill biotite in high tenor Cu-zones may be attributed to both oxidation state and low aH4SiO and represent the most primitive fluids. The spatial variations of infill biotite chemistry are a record of the physiochemical evolution of magmatic-hydrothermal fluids during wall rock-buffering and Cu-deposition. Halogen compositions, corrected for the effects of coupled exchange on the XMg site, show that the potassic fluids from the Southwest Zone, Copper Canyon and Skouries alkalic Cu-Au deposit contrast with those from calc-alkalic porphyry deposits. We submit that low log(JH2O!JHF) of exsolved magmatic-hydrothermal fluids reflects the magmatic composition. Variations in log(JH20/JHCJ) could also be attributed source compositions, however, these ffigacity estimates are more sensitive to post exsolution processes and other uncertainties. 161 Evaluating spatial zonation in infihl biotite chemistry is another method to assess the processes of porphyry Cu deposit genesis and is particularly applicable to systems characterised by the absence of hydrothermal quartz. 3.13 References Allen, D.G., Panteleyev, A. and Armstrong, A.T., 1976, Porphyry copper deposits of the alkalic suite: Galore Creek: In Porphyry Deposits of The Canadian Cordillera, A. Sutherland and Brown, ed., CIM Special Volume 15, p. 402-4 14. Bath, A., and Cooke, D.R., 2008, The importance of biotite for the deposition of sulfides at the Lorraine Cu-Au Porphyry deposit, north-central British Columbia: Shallow- and Deep-Level Alkalic Mineral Deposits: Developing an Integrated Exploration Model, Year 3, Final Report Porphyry Module. Beane, R.E., 1974, Biotite stability in the porphyry copper enviromnent: ECONOMIC GEOLOGY, v. 69, p. 241-256. Bottomer, L.R., and Leary, G.M., 1995, Copper Canyon porphyry copper-gold deposit, Galore Creek area, northwestern British Columbia: In: Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Northern Cordillera ed. by Canadian Institute of Mining and Metallurgy Special Volume 46, p. 645- 649. Brigatti, M.F. and Guggenheim, 5., 2002, Mica crystal chemistry and the influence of pressure, temperature, and solid solution on atomistic models: In Mottana, A., Sassi F.P., Thompson, J.B. Jr., and Guggenheim, S., ed., Micas: Crystal Chemistry and Metamorphic Petrology: Reviews in Mineralogy & Geochemistry, 46, Mineralogical Society of America and The Geochemical Society, Washington, D.C. p. 1-97. Brimhall, G.H. and Crerar, D.A., 1987, Ore fluids: magmatic to supergene. In Thermodynamic Modelling of Geological Material: Minerals, Fluids and Melts, ed. By I.S.E Carmicheal and H.P. Eugster, Reviews in Mineralology, Vol. 17, Mineralogical Society of America. Brimhall, G.H. and Ague, J.J., 1988, Granite systems: In Hydrothermal Processes-Applications to Ore Genesis. Eds. H.L. Barnes and H. Ohomoto, Riedel Publishers, Dordrecht, Holland, p. 33. Cooke, D. R., Wilson, A. J., House, M. J., Wolfe, R. C., Walshe, J. L., Lickfold, V. and Crawford, A. J., 2007, Alkalic porphyry Au - Cu and associated mineral deposits of the Ordovician to Early Silurian Macquarie Arc, New South Wales: Australian Journal of Earth Sciences, v. 54, p. 445-463. Cooke, D.R. and Simmons, S.F., 2000, Characteristics and genesis of epithermal gold deposits: Reviews in ECONOMIC GEOLOGY, v. 13, p. 221-224. Coulson, I. M., Dipple, G.M., Raudsepp, M., 2001, Evolution of HF and HCI activity in magmatic volatiles of the gold-mineralised Emerald Lake pluton, Yukon Territory, Canada: Mineralium Deposita, v. 36, p. 595-606. Cathies, L.M., 1977, Fluid flow and genesis of hydrothermal ore deposits: ECONOMIC GEOLOGY, 75th Anniversary Volume, p. 424-457. Candela P.A., 1989, Magmatic ore-forming fluids, thermodynamic and mass-transfer calculations of metal concentrations: In: Whitney J. A., and Naldrett A. J., eds., Ore Deposition Associated with Magmas, Reviews in Economic Geology, v. 4,p203—21. Candela, P.A., 1986, Toward a thermodynamic model for the halogens in magmatic systems: An application to melt-vapor-apatite equilibria: Chemical Geology, 57, 289-30 1. 162 Candela P. A., Piccoli P. M., 1995, Model ore-metal partitioning from melts into vapor and vapor/brine mixtures: In Thompson J. F. H., ed., Magmas, fluids, and ore deposits, Mineralogical Association of Canada Short Course, v. 23, P. 101—127. Dahi, 0., 1969, Irregular distribution of iron and magnesium among coexisting biotite and garnet: Lithos, v.2,p.311-22 Deer, W.A., R.A. Howie, Zussman, 3., 1992, An Introduction to the Rock-Forming Minerals: 2nd ed. Pearson-Prentice Hall, p. 301. Delong, R.C., 1996, Geology, alteration, mineralization and metal zonation of the Mt. Milligan porphyry copper-gold deposits: Unpublished M.Sc. thesis, University of British Columbia. Dilles, J.H., 1987, The petrology of the Yerington batholith, Nevada: Evidence for the evolution of porphyry copper ore fluids: ECONOMIC GEOLOGY, v. 82, p. 1750-1789 Dyar, M.D., Guidotti, C.V., Holdaway, M.J., and Colucci, M., 1993, Nonstoichiometric hydrogen contents in common rock-forming hydroxyl silicates: Geochimica et Cosmochimica Acta, v. 57 p. 2913— 2918. Dymek, R.F., 1983, Titanium, aluminum and interlayer cation substitutions in biotite from high-grade gneisses, west Greenland: American Mineralogist v. 68, p. 880—899. Drummond, S.E. and Ohmoto, H., 1985, Chemical evolution and mineral deposition in boiling hydrothermal systems: ECONOMIC GEOLOGY, v. 80, p. 126-147. Enns, S.G., Thompson, J.F.H., Stanley, C.R. and Yarrow, E.W., 1995, The Galore Creek porphyry copper- gold deposits, Northwestern British Columbia: In: Schroeter, T., ed., Porphyry Copper (± Au) Deposits of the Northern Cordillera ed. by Canadian Institute of Mining and Metallurgy Special Volume 46, p. 630-644. Fournier, R.O., Potter, R.W, 1982, An equation correlating the solubility of quartz in water from 25° to 900°C at pressures up to 10,000 bars: Geochimica et Cosmochimica Acta, v. 46, p. 1969-1973. Feldstein, S.N., Lang, R.A., Vennemann, T., and O’Neil, J.R., 1996, Ferric-ferrous ratios, H20 contents and D/H ratios of phiogopite and biotite from lavas of different tectonic regimes: Contributions to Mineralogy and Petrology, v. 126, p. 51—66. Frei, R., 1995, Evolution of mineralizing fluid in the porphyry copper system of the Skouries Deposit, Northeast Chalkidiki (Greece); evidence from combined Pb-Sr and stable isotope data: ECONOMIC GEOLOGY, v. 90, p. 746-762. Jacobs, D.C. and Parry, W.T., 1979, Geochemistry of biotite in the Santa Rita porphyry copper deposit, New Mexico: ECONOMIC GEOLOGY, v. 74, p. 860-887. Foster, M.D., 1960, Interpretation of the composition of trioctahedral micas: United States Geological Survey Professional Paper 354-B, p. 11—49. Guidotti, C.V., 1984, Micas in metamorphic rocks: In: Bailey S.W., ed., Micas, Reviews in Mineralogy, vol. 13, Mineralogical Society of America, p. 357—467. Guidotti, C.V., and Dyar, M.D., 1991, Ferric iron in metamorphic biotite and its petrological and crystallographic implications: American Mineralogist, v. 72, p. 161-175. Hendry D.A.F., Chivas, A. R J., Long V. P. and Reed, S. 3. B., 1985, Chemical differences between minerals from mineralizing and barren intrusions from some North American porphyry copper deposits: Contributions to Mineralogy and Petrology, v. 89, p. 317-329. 163 Holdaway, M.J., 1980, Chemical formulae and activity models for biotite, muscovite, and chlorite applicable to pelitic metamorphic rocks: American Mineralogist, v. 65, p. 711-719 Huston, D.L., Bolger, C., and Cozens, G, 1993, A comparison of mineral deposits at the Gecko and White Devil deposits; implications for ore genesis in the Tennant Creek District, Northern Territory, Australia: ECONOMIC GEOLOGY, v. 88, p. 1198-1225. Helgeson, H.C., 1967, Thermodynamics of complex dissociation in aqueous solution at elevated temperatures: Journal of Physical Chemistry, v. 71, p. 3121-3136. Helgeson, H.C., 1970, A chemical and thermodynamic model of ore deposition in hydrothermal systems: Mineralogical Society of America, Special publication, Paper 3, p. 155-186. Hezarkhani, A., Williams-Jones, A.E., and Gammons, G.H., 1999, Factors controlling copper solubility and chalcopyrite deposition in the Sungun porphyry deposit, Iran: Mineralium Deposita, v. 34, p. 770-783. Kroll, T., D., MUller, T., Seifert, PM., Herzig, A., 2002, Petrology and geochemistry of the shoshonite hosted Skouries porphyry Cu-Au deposit, Chailcidiki, Greece: Mineralium Deposita, v. 37, p. 137- 144. Lang, J.R., Stanley, C.R. and Thompson, J.F.H., 1995a, Porphyry copper deposits related to alkalic igneous rocks in the Triassic-Jurassic arc terranes of British Columbia: Arizona Geological Society, Digest 20, p. 2 19-236. Lang, J.R., Thompson, J.F.H., Stanley, C.R., 1995b, Na-K-Ca magmatic hydrothermal alteration associated with alkalic porphyry Cu-Au deposits, British Columbia, in Thompson, J.F.H., ed., Magmas, fluids and ore deposits, Mineralogical Association of Canada Short Course, v. 23, p. 339-366. Logan, J.M. and Koyanagi, V.M., 1994, Geology and mineral deposits of the Galore Creek area, northwestern British Columbia (104G/3 and 4): BC Ministry of Energy, Mines and Petroleum Re sources, Bulletin 92. Liaghat, S. and Tosdal, R, 2008, Apatite chemical composition and textures as a probe into magmatic conditions at Galore Creek porphyry copper-gold deposit, British Columbia: Geochimica et Cosmochimica Acta, Volume 72, Issue 12, p. A550. Ludington, S.D and Munoz, J.L., 1975, Application of fluorine-hydroxyl exchanged data to natural micas: Geological Society of America, Abstracts, Programs 7, p. 1179. Mason, D.R., 1978, Compositional variations in ferromagnesian minerals from porphyry copper-generating and barren intrusions of the Western Highlands, Papua New Guinea: ECONOMIC GEOLOGY, v. 73, p. 878—890. Maughan, D.T., Keith, J.D., Christiansen, E.H., Pulsipher, T., Hatori, K., and Evans, N.J., 2002, Contributions from mafic alkaline magmas to the Bingham porphyry Cu-Au-Mo deposit, Utah, USA: Mineralium Deposita, v. 37, p. 14-37. McMillan, W.J., 1991, Tectonic Evolution and Setting of Mineral Deposits in the Canadian Cordillera: BC Ministry of Energy, Mines and Petroleum Resources, Ore Deposits, Tectonics and Metallogeny in the Canadian Cordillera, Paper 1991-4. Mortensen, J.K., Ghosh, D.K. and Fern, F., 1995, U-Pb geochronology of intrusive rocks associated with copper-gold porphyry deposits in the Canadian Cordillera: Canadian Institute of Mining and Metallurgy, Special Vol. 46, p. 142-158. Micko, 3., Tosdal, R., Simpson, K., Chamberlain, C., and Schwab, D., 2008, Controls and hydrothermal zonation of Central Zone mineralization at the Galore Creek alkalic Cu-Au porphyry deposit, 164 northwestern British Columbia: In: MDRU-CODES: Shallow and deep alkalic deposits: Porphyry module, Year 3 Final meeting Munoz, J.L., 1984, F—OH and Cl—OH exchange in micas with applications to hydrothermal ore deposits: In: Bailey, S.W. ed., Micas. Reviews in Mineralology, vol. 13, p. 469-494. Munoz, J.L., 1990, F and Cl contents of hydrothermal biotites: a re-evaluation: Geological Society of America Abstract Programs no 22, p. A135. Munoz, J.L., 1992, Calculation of HF and HCL fugacities from biotite compositions: revised equations: Geological Society of America. Abstract Programs no. 24, p. A22 1. Munoz, J.L., Swenson, A., 1981, Chloride-hydroxyl exchange in biotite and estimation of relative HCI/HF activities in hydrothermal fluids: ECONOMIC GEOLOGY v. 76, p. 2212—2221. Parry, W.T., and Jacobs, D.C., 1975, Fluorine and chlorine in biotite from Basin and Range plutons: ECONOMIC GEOLOGY, v. 70, p. 24 1-256. Pouchou, J.L. & Pichoir, F., 1985, PAP (pZ) procedure for improved quantitative microanalysis: Microbeam Analysis, p. 104-106. Rebbert, C.A., Partin, E., Hewitt, D.A., 1995, Synthetic biotite oxidation under hydrothermal conditions: American Mineralogist, v. 80, p. 345-354. Rieder, M. and Cavazzini, G., D’yakonov, Yu.S., Frank-Kamenetskii, V.A., Gottardi, G., Guggenheim, S., Koval, P.V., Muller, G., Neiva, A.M.R., Radoslovich, E.W., Robert, J.-L., Sassi, F.P., Takeda, H., Weiss, Z. and Wones, D.R., 1998, Nomenclature of the micas: The Canadian Mineralogist, v. 36, p. 905-912. Robert, J.L., 1976, Titanium solubility in synthetic phlogopite solid solutions: Chemical Geology, v. 17, p. 213-227. Tischendorf, G., Forster, H.-J., Gottesmann, B., and Rieder, M., 2004, A new graphical presentation and subdivision of potassium micas: Mineralogical Magazine, v. 68, p. 649-667. Taylor, R.P., 1983, Comparison of biotite geochemistry of Bakircay, Turkey, and Los Pelambres, Chile, porphyry copper systems: Transactions, Institute of Mining and Metallurgy, Sect. B 92, p. B 16-22. Seedorff, E., Dilles, J.H., Proffett, J.M., Einaudi, M. T., Zurcher, L., Stavast, W.J.A., Johnson, D.A., and Barton, M.D., 2005, Porphyry Deposits: Characteristics and Origin of Hypogene Features: ECONOMIC GEOLOGY, 100th Anniversary Volume, p. 251-298. Selby, D., and Nesbitt, B.E., 2000, Biotite Chemistry of the Casino Porphyry Cu-Au-Mo Mineralization, Yukon, Canada: Evaluation of Magmatic and Hydrothermal Fluid Chemistry: Chemical Geology, v. 171, p. 77-93. Schwab, D.L, Petsel, S., Otto, B.R., Morris, S.K., Workman, E. and Tosdal, R.M., 2008, Overview of the Late Triassic Galore Creek Copper-Gold-Silver Porphyry System: Arizona Geological Society meeting, Ores & Orogenesis, v. 22, p. 1-14. Speer J.A., 1984, Micas in igneous rocks: In: Bailey S.W, ed., Micas, Reviews in Mineralogy, vol. 13, Mineralogical Society of America, p. 299-356. Thompson, J.B., 1982, Composition space: An algebraic and geometric approach in characterization of metamorphism through mineral equilibria, ed. Ferry, J.M., Reviews in Mineralogy, Mineralogical Society of America, v.10, p. 1-32. 165 Twelker, E., 2007, A breccia-centered ore and alteration model for the Copper Canyon alkalic Cu-Au deposit, British Columbia: M.Sc. thesis, University of Alaska Fairbanks, p. 75-83. Waters, D.J. and Charnley, N.R., 2002, Local equilibrium in polymetamorphic gneiss and the titanium substitution in biotite: American Mineralogist, v. 87, p. 383—396. Wones, D.R., and Eugster H.P., 1965, Stability of biotite: experiment theory and application: American Mineralogist, v. 50, p. 1228-1272. Zhu, C., Sverjensky, D.A., 1991, Partitioning of F—Cl—OH between minerals and hydrothermal fluids: Geochimica et Cosmochimica Acta v. 55, p. 1837—1858. Zhu, C., Sverjensky, D.A., 1992, F—Cl—OH partitioning between biotite and apatite: Geochimica et Cosmochimica Acta, v. 56, p. 3435—3467. 166 Chapter 4: Conclusions Interpretations of data collected during this study allow the key features of the Southwest Zone deposit to be established, and these are summarized below. They provide an insight into alteration zonation of a silica-undersaturated magmatic hydrothermal alkalic system and controlling factors of Cu-Au mineralisation and zonation. This chapter also addresses how the results have helped to resolve the original objectives. • Hydrothermally cemented breccias are localized at the contact between matrix- bearing breccias and porphyry wall rock and along an array of faults intersecting this contact. Biotite-phyric monzodiorite dikes and dikelets were also focused by these structural features and were emplaced syn-Cu. • Alteration, sulphide minerals and metals are zoned about cemented breccias and biotite-phyric monzodiorite. • Potassic alteration and cemented breccia facies formed at moderately high temperatures (420 ± 50C°) from magmatic fluids. Phlogopite-magnetite cement and alteration is associated with the most abundant Cu-Fe-sulphide. Cu-poor, diopside-dominated infill and alteration formed contemporaneously with potassic infill locally flanking high-tenor Cu and domains of intense K-metasomatism. Diopside infill and alteration is likely to have formed from Ca-saturated magmatic fluids. • Systematic spatial variations in Ti-content and Fe3/F2of infill biotite, together with interpreted alteration reactions, indicate Cu deposition was primarily caused by a decrease in magmatic-hydrothermal fluid oxidation state coupled with an increase in pH. The same mechanisms may also be attributed to the divergence in 167 Cu and Au zonation patterns within high tenor Cu domains. Cooling may have also contributed to a decrease in fluid Cu solubility. • Selectively pervasive garnet and chlorite alteration formed outboard of potassic and calcic-Q,otassic) alteration and infill and locally overprints higher temperature assemblages. Garnet-chlorite alteration in the Southwest Zone is interpreted to be equivalent to propylitic alteration and is not associated with significant Ca-metasomatism. • A Au-(Pb)-halo encircles high temperature assemblages and the Cu-core of the deposit. This Au mineralisation is coincident with pyrite-bearing propylitic alteration (250-350°) and interpreted to be caused by cooling and fluid neutralization by chlorite alteration. • Post-hydrothermal system, regional folding and thrusting tilted the Southwest Zone deposit 50-60° northeast. 4.1 Recommendations for future work (and other conjectures) Mafic rocks associated with a few porphyry deposits show evidence of the transfer of sulfur and other components to the mineralized felsic intrusions (Hattori & Keith, 2001). The timing of emplacement with respect to Cu and the composition of biotite phyric monzodiorite and pyroxene and homblende and biotite-phyric diorite suggest mafic magma may have played a role in the genesis of the Southwest Zone. Immobile element and Pb-isotope characterization could lead recognition of potentially productive intrusions in the Galore Creek area. This type of characterization may also shed light on the metallogenesis of island arc porphyries. The interplay between hydrothermal fluid F concentration, temperature and alkali cation concentrations is likely to influence the mobility of high field strength elements such as Ti and REE’s (Salvi, et al., 2000). The evaluation of infill biotite crystallo 168 chemistry indicates relatively high mineral Ti concentration and high fluid aHF. Additionally, hydrothermal gangue mineralogy are compatible with deposition from moderately hot ( 4200) silica-undersaturated and oxidized fluids. Investigation of the affects halogen fugacity and oxidation state on mineral stability relationships in alteration halos at the Southwest Zone may provide insight into some the unusual features of the alkalic magmatic-hydrothermal system. This study highlights the usefulness of estimating biotite Fe3/ Fe2 ratio in evaluating fluid oxidation state. Establishing the oxidation state of Fe in biotite is somewhat problematic and involves several assumptions, however, our Fe3 estimates are comparable to measured concentrations determined by more traditional methods. Thus, the normalization methods presented in this thesis could potentially be applied in other studies, but would require calibration to measured Fe3 concentrations. Additionally, shortwave infrared spectrometry could prove to be a quick means of evaluating biotite Fe3 contents if spectra can be correlated to either normalized probe data or wet chemical analysis. 169 4.2 References Hattori, K.H. and Keith J.D., 2001, Contribution of mafic melt to porphyry copper mineralization: Evidence from Mount Pinatubo, Philippines and Bingham deposit, Utah: Mineralium Deposita, v. 36, p 799-806. Salvi, S., Fontan, F., Monchoux, P., Williams-Jones, A.E., and Moine, B., 2000, Hydrothermal mobilization of high field strength elements in alkaline igneous systems: Evidence from the Tamazeght Complex (Morocco): ECONOMIC GEOLOGY, v. 95, p. 559—576. 170

Cite

Citation Scheme:

        

Citations by CSL (citeproc-js)

Usage Statistics

Share

Embed

Customize your widget with the following options, then copy and paste the code below into the HTML of your page to embed this item in your website.
                        
                            <div id="ubcOpenCollectionsWidgetDisplay">
                            <script id="ubcOpenCollectionsWidget"
                            src="{[{embed.src}]}"
                            data-item="{[{embed.item}]}"
                            data-collection="{[{embed.collection}]}"
                            data-metadata="{[{embed.showMetadata}]}"
                            data-width="{[{embed.width}]}"
                            data-media="{[{embed.selectedMedia}]}"
                            async >
                            </script>
                            </div>
                        
                    
IIIF logo Our image viewer uses the IIIF 2.0 standard. To load this item in other compatible viewers, use this url:
https://iiif.library.ubc.ca/presentation/dsp.24.1-0052382/manifest

Comment

Related Items