Characterization of Pentlandite-Bearing Quartz Veins of Kambalda, Western Australia

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CHARACTERIZATION OF PENTLANDITE-BEARING QUARTZ VEINS OF KAMBALDA, WESTERN AUSTRALIA

By Isaac Simon

A thesis submitted to the Faculty and Board of Trustees of the Colorado School of Mines in partial fulfillment of the requirements for the degree of Master of Science (Geology).

Golden, Colorado

Date ______

Signed: ______Isaac Simon

Signed: ______Dr. Katharina Pfaff Thesis Advisor

Golden, Colorado

Date ______

Signed: ______Dr. M. Stephen Enders Professor and Department Head Department of Geology and Geological Engineering

ABSTRACT The Kambalda Dome of the Yilgarn Craton is located in Western Australia and has been historically mined for Ni after the discovery of magmatic sulfide orebodies in 1966. The Ni orebodies consist of komatiite-associated magmatic sulfide deposits which have experienced deformation, metamorphism and plutonism. A complex geologic history at the Kambalda Dome has resulted in the formation of various types of cross-cutting quartz veins that have not yet been fully described and classified. During recent mining activities at the eastern flank of the Kambalda Dome, three types of quartz veins have been observed: barren quartz veins, auriferous quartz veins and pentlandite-bearing quartz veins. Timing and formation conditions of these quartz veins are unknown and pentlandite-bearing quartz veins are of particular interest as pentlandite usually occurs in magmatic or volcanic settings, such as layered mafic-ultramafic intrusions, mantle xenoliths and komatiite-associated magmatic sulfide deposits. Characterization of these quartz veins and understanding of timing and formation conditions have been determined through detailed optical petrography, scanning electron microscopy techniques, cathodoluminescence imaging and electron microprobe analysis. Microanalytical work indicates that pentlandite occurs in four distinct mineralization styles at Kambalda. Mineralization Style I represents the historically mined komatiite-hosted deposits formed through magmatic processes. Mineralization Style II represents carbonate-rich Ni- arsenide bearing quartz veins that are hydrothermal in nature and bear great similarity to profuse hydrothermal Ni occurrences globally. Mineralization Style III represents quartz veins with pentlandite intergrown with pyrrhotite and gangue minerals such as quartz, feldspar and biotite. Pentlandite, co-genetic pyrrhotite and gangue minerals formed by precipitation from metamorphic fluids. Mineralization Style IV represents fractured or brecciated quartz veins with pentlandite, pyrrhotite, pyrite and chalcopyrite mechanically remobilized and localized in late fractures. This study focused on understanding the formation of hydrothermal pentlandite- bearing quartz veins which, through aforementioned techniques and thermodynamic modelling, have been determined to have formed between 2660 ± 4 Ma and ~ 2,625 Ma, from metamorphic fluids at near-neutral pH and reduced conditions. Age constraints are delimited through hostrock U-Pb zircon ages and deformation event ages available in literature.

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The results of this study are principally important in building on knowledge of crustal fluids that are capable of mobilizing and consequently precipitating Ni minerals such as pentlandite. This study is also relevant to exploration efforts considering such data can be utilized to effectively characterize Ni deposits based on mineralogical and geochemical characteristics of rocks observed.

TABLE OF CONTENTS ABSTRACT ...... iii LIST OF FIGURES ...... vi LIST OF TABLES ...... ix ACKNOWLEDGEMENTS ...... x CHAPTER 1: INTRODUCTION ...... 1 1.1 Behavior of Ni in hydrothermal fluids ...... 1 1.2 Mineral Deposits near Kambalda ...... 4 CHAPTER 2: CHARACTERIZATION OF PENTLANDITE-BEARING QUARTZ VEINS IN KAMBALDA, WESTERN AUSTRALIA ...... 20 2.1 Introduction ...... 20 2.2 Geological Setting ...... 24 2.3 Materials and Methods ...... 33 2.4 Results ...... 41 2.4.1 Petrography ...... 41 2.4.2 CL Imaging ...... 61 2.4.3 Mineral Chemistry ...... 64 2.4.4 Fluid Inclusion Studies ...... 82 2.4.5 Thermodynamic Considerations ...... 83 2.5 Discussion ...... 84 2.5.1 Styles of Pentlandite Mineralization ...... 85 2.5.2 Mineral Chemistry ...... 90 2.5.3 Timing and Conditions of Mineralization Style III Formation ...... 96 2.5.4 Thermodynamic Considerations ...... 104 2.5.5 Formation of Pentlandite-bearing Quartz Veins in Context of Nearby Deposits ...... 105 2.6 Summary and Conclusions ...... 110 CHAPER 3: IMPLICATIONS AND OUTLOOK ...... 114 3.1 Implications and Project Significance ...... 114 3.2 Outlook and future work ...... 115 REFERENCES ...... 118 APPENDIX A: SUPPLEMENTAL ELECTRONIC FILES ...... 128 APPENDIX B: EVOLUTION OF THE YILGARN CRATON THROUGHOUT TIME ...... 129 v

LIST OF FIGURES

Figure 1.1 Map of Australia with major Australian ore deposits highlighted after Phillips (2017b)………………………………………………………………………...…..5 Figure 1.2 Regional geologic map of the Kambalda Nickel Field after Gresham and Loftus- Hills (1981) ……...………..……………………………………………...7 Figure 2.1 Terranes of the Yilgarn Craton after Goscombe et al. (2009)………………..….25 Figure 2.2 Regional geologic map of the Kambalda Nickel Field after Gresham and Loftus-Hills (1981)………………………………………………………….….. 28 Figure 2.3 Geologic map of the Kambalda Dome after Stone et al. (2005) and Staude et al. (2017b)…………………...…………………………………………………...29 Figure 2.4 Volcanic stratigraphy of the Kambalda Sequence after Gresham and Loftus-Hills (1981)…………………………………………………….…..…….31 Figure 2.5 Geologic map of the Kambalda Dome with sample localities and pentlandite- bearing quartz vein images after Stone et al. (2005) and Staude et al. (2017b)... 33 Figure 2.6 Mine wall photographs of sample localities for samples S-137-6 and S-137-7....35 Figure 2.7 Irregular quartz in quartz veins viewed optically…………………………...…...42 Figure 2.8 Pentlandite-bearing quartz vein textures viewed optically and in BSE…..……...43 Figure 2.9 Automated mineralogy scan of sample IS13-7-2 from Long North……………..44 Figure 2.10 Textures in sample IS13-7-2 from Long North viewed optically and in BSE..….45 Figure 2.11 Textures in sample IS13-7-2 from Long North viewed optically and in BSE…...46 Figure 2.12 Automated mineralogy scan of sample IS13-7-2Q from Long North…………...47 Figure 2.13 Textures in sample IS13-7-2Q from Long North viewed in BSE and high resolution automated mineralogy………………………………………………...48 Figure 2.14 Automated mineralogy scan of sample S-460-2 from McLeay………………….48 Figure 2.15 Textures in sample S-460-2 from McLeay viewed optically and in BSE……….49 Figure 2.16 Textures in sample S-460-2 from McLeay viewed optically…………………….50 Figure 2.17 Automated mineralogy scan of sample S-Moran-2 from Moran……….………..51 Figure 2.18 Textures in sample S-Moran-2 from Moran viewed optically and in BSE…….. 52 Figure 2.19 Automated mineralogy scan of sample S-LSU162-1 from Moran…………...….53 Figure 2.20 Textures in sample S-LSU162-1 from Moran viewed optically and in BSE…....53 Figure 2.21 Automated mineralogy scan of sample S-LSU373A-1 from 500m north of Moran………………………………………………………..…………………...54 Figure 2.22 Textures in sample S-LSU373A-1 from 500m north of Moran viewed in BSE...55

Figure 2.23 Textures in sample S-LSU373A-1 from 500m north of Moran viewed in BSE...56 Figure 2.24 Texture in sample S-LSU373A-1 from 500m north of Moran viewed in BSE….56 Figure 2.25 Automated mineralogy scan of sample S-137-7 from Long North……………...57 Figure 2.26 Textures in sample S-137-7 from Long North viewed in BSE………….……….58 Figure 2.27 Textures in sample S-137-7 from Long North viewed in BSE…………………..58 Figure 2.28 Automated mineralogy scan of sample S-137-6 from Long North……………...59 Figure 2.29 Textures in sample S-137-6 from Long North viewed in BSE…………………..59 Figure 2.30 Texture in sample S-137-1 from Long North viewed optically…………….……60 Figure 2.31 Texture in sample S-137-1 from Long North viewed optically………………….60 Figure 2.32 Automated mineralogy scan of sample IS-1569 from Long North……………...61 Figure 2.33 CL images of plagioclase, potassium feldspar, calcite and apatite in sample IS13-7-2 from Long North……………………………………….…………...….63 Figure 2.34 CL image of apatite in sample S-460-2 from McLeay…………………...……...64 Figure 2.35 CL images of quartz in quartz veins …………………………………...………..64 Figure 2.36 Iron and Ni content in pentlandite……………………………………….………67 Figure 2.37 Cobalt content and Fe/Ni ratio of pentlandite……………………………………67 Figure 2.38 Iron and Ni content in pyrrhotite…………………………...……………………69 Figure 2.39 Iron and Co content in pyrite………………………………………………….…71 Figure 2.40 Micro-XRF element map of sample IS13-7-2 from Long North and sample S-460-2 from McLeay……………………………………………………………72 Figure 2.41 Micro-XRF element map of samples S-Moran-2 and S-LSU162-1 from Moran………………………………………………………...... 73 Figure 2.42 Iron and Cu content in chalcopyrite……………………………………………...75 Figure 2.43 Magnesium and Fe content in biotite…………………………………………….78 Figure 2.44 Nickel and F content in biotite…………………………………………………...78 Figure 2.45 Potassium and Na content in potassium feldspar…………………….…...……..80 Figure 2.46 Strontium and Ba content in potassium feldspar………………………...………80 Figure 2.47 Calcium and Na content in albite………………………………………………...82 Figure 2.48 Fluid inclusions in quartz of quartz veins………………………………………..83 Figure 2.49 Oxygen fugacity vs pH diagram of Mineralization Style III pentlandite mineral assemblage………...……………………………………….……………84 Figure 2.50 Drill hole collared at Moran depicting sample location for sample S-LSU162-1.……………………..………………………………………………88 Figure 2.51 Sodium and Ca content in CL albite from Long North Mineralization Style III...91

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Figure 2.52 Generations of potassium feldspar……………………………………………….92 Figure 2.53 Groups of chalcopyrite…………………………………………………………...92 Figure 2.54 Geochemical characteristics of pentlandite across different Mineralization Styles……………………………………………………………………………..94 Figure 2.55 Geochemical characteristics of pyrrhotite across different Mineralization Styles……………………………………………………………………………..95 Figure 2.56 Mineralogy of orogenic gold deposits after Goldfarb (Goldfarb, pers. comm.)..104 Figure 2.57 Oxygen fugacity vs pH diagram for predominant Au and Ni speciation………110

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LIST OF TABLES

Table 2.1 Depositional conditions for some gold deposits in the Kambalda-Kalgoorlie area after Ho et al. (1992)………………………………………………………..22 Table 2.2 Deformational history of the Kambalda Dome after Stone et al. (2005)………...32 Table 2.3 Samples collected with details on sample type, location and analytical technique performed on each sample………………………………………….…34 Table 2.4 Set up conditions for EPMA Session 1 (sulfides)………………….…………….38 Table 2.5 Set up conditions for EPMA Session 2 (biotite)…………………………………39 Table 2.6 Set up conditions for EPMA Session 3 (feldspar)………………….…………....40 Table 2.7 Deformation events in the Eastern Goldfields Superterrane………………..……99 Table 2.8 Metamorphic and metasomatic assemblages associated with gold deposits in the Yilgarn Craton………………………………………………………………103 Table 2.9 Summary of Mineralization style characteristics……………...………………..113

ACKNOWLEDGEMENTS I wish to extend my sincerest gratitude toward Dr. Katharina Pfaff for being an astounding advisor. Her knowledge, guidance and feedback was an essential component to my project and her advice was always thoughtful and beneficial to my overall growth as a geologist. Being her graduate student and learning from her has been an absolute pleasure. Dr. Sebastian Staude is also thanked for serving as a mentor for me while working with him in Tübingen. I am grateful for him for sharing his Kambalda sample suite with me and for allowing me to do research with him. His knowledge of Kambalda, great research skills and overall friendliness made working with him in Germany an absolute pleasure. Dr. Thomas Monecke was instrumental to my research as he was actively involved with my project by providing me with essential input based on his abundant knowledge on economic geology. Additionally, I am grateful toward him and Jim Reynolds as I was able to work with them doing CL imaging and fluid inclusion petrography, respectively. Dr. Thomas Monecke and Jim Reynolds served as crucial resources to my understanding of fluid inclusions and are deeply thanked for the time they dedicated to my project by observing my samples with me. I’d like to thank Dr. Benjamin Walter for teaching me how to use the fluid inclusion petrographic microscope and heating- freezing stage in Tübingen. I wish to thank Dr. Alex Gysi for his involvement in the thermodynamic modelling of my project through GEMS. His vast knowledge of thermodynamics and of the useful software made it possible to model this system in GEMSelektor. Jae Erickson is thanked for his excellent sample preparation and his willingness to answer any questions on sample preparation methods. Colorado School of Mines graduate students Alyssa Smith, Kelsey Livingston, David Hernandez Uribe, Zach Palmer, Jeffery McKeon, Christopher Van Hoozen, Kairan Liu and Duncan McIntire are thanked for providing me with intriguing geologic discussions and for making my experience at Mines memorable. Tübingen PhD students Manuel Scharrer, Tatjana Epp, and Christian Dietzel as well as graduate student Johannes Hecker are thanked for their peer support, for their input and scientific discussions as well as for their overall influence on my happiness for the duration of my Master’s program. I’d like to the thank Professor Dr. Gregor Markl for allowing me to temporarily join his research group and for his feedback on my mineralogical and mineral chemistry results. Dr. Thomas Wenzel is thanked for his help with

x electron probe microanalyses. Dr. Johannes Giebel is thanked for providing his biotite EPMA calculation spreadsheet which was used as a model for my biotite chemical data. Manuel Scharrer is especially thanked for teaching me how to use Geochemists Workbench and for working with me on thermodynamic modelling. His involvement in my project is greatly appreciated and I am grateful for his help and patience while teaching me. I wish to thank my family and close friends for being an active support system during my program. Natasha Galvez and Dr. George Dunne are thanked for their wonderful hospitality and their active involvement in my well-being. My life mentor, Sharon Kinard, is thanked for helping me build the confidence I have in academics and professionalism. I wish to extend my deepest gratitude toward my family, Juan M. Simon, Maria Simon, Juan E. Simon and Kenia Lopez, for their endless support and for always believing in my abilities. To my family - may my accomplishments, such as the completion of this project, reflect your influence on me as encouraging role models in my life. Lastly, I would like to thank the Society of Economic Geologists for providing me with a Graduate Student Fellowship. I also wish to thank the Mineralogical Association of Canada for funding my travel to Tübingen through a Student Travel Grant.

CHAPTER ONE

INTRODUCTION The majority of the world’s nickel is sourced from magmatic Ni sulfide and Ni laterite deposits (Kuck, 2006; Liu et al., 2018). However, there is an increasing interest in processes that are capable of mobilizing and forming hydrothermal Ni deposits since the discovery of the Avebury nickel deposit of Tasmania containing 29 Mt of 0.9% Ni (Keays and Jowitt, 2013; Capistrant et al., 2015). The behavior of Ni in hydrothermal fluids reported in literature is summarized below. An area where abundant Ni deposits occur is the Archean Yilgarn Craton of Western Australia. The Yilgarn Craton is well-known for its abundant economic ore deposits. There are abundant nickel deposits in the Kalgoorlie Terrane of the Craton, which have been mined since the discovery of the magmatic Ni deposits at Kambalda (Hronsky and Schodde, 2006). Furthermore, most of the gold mined in the Yilgarn Craton comes from the Kalgoorlie Terrane with the Kalgoorlie goldfield being the leading gold producer in Australia (Phillips et al., 2017a). In addition to introducing the behavior of Ni in hydrothermal fluids, the various well-known deposits in the Kalgoorlie area are also discussed in this section.

1.1 Behavior of Ni in hydrothermal fluids The element Ni has traditionally been considered to be largely immobile (Le Vaillant et al., 2014) as it is known to be one of the least soluble elements, among the first row transition metals, in most common geological fluids (Liu et al., 2012; Le Vaillant et al., 2016). The assumption that Ni, along with PGE’s, is not easily remobilized by hydrothermal fluids has led to the implementation of using these elements as alteration-stable trace elements in determining petrogenesis of rocks associated with ore deposits (Le Vaillant et al., 2016). However, in the past two decades, there have been an increasing amount of reports which note that many magmatic Ni deposits have been extensively hydrothermally altered, a process during which a number of metals (including Ni) have been remobilized (Tian et al., 2012 and references therein). In addition, there is also published research on hydrothermal Ni occurrences that formed through remobilization of local Ni-rich rocks such as komatiite and peridotite (Tian et al., 2012; Gonzalez-Alvarez et al., 2013a). Evidence of hydrothermal mobilization of Ni was reported from

1 modern hydrothermal settings such as the Rainbow hydrothermal field of the Mid-Atlantic Ridge (Marques et al., 2006). Lastly, Ni-enriched black shales (Jowitt and Keays, 2011), unconformity uranium deposits (e.g. Athabasca Basin; Jefferson et al., 2007), Five Element Vein Deposits (Kissin, 1992; Markl et al., 2016) and sediment-hosted Ni-rich deposits of the Central African Copperbelt (e.g. Enterprise deposit, Capistrant et al., 2015; Menda Central Prospect, Ball et al., 2016) are interpreted to have formed through the precipitation of Ni from an aqueous fluid. These examples, at which Ni-mobility in crustal fluids has occurred, are indicative that at certain conditions, fluids are capable of Ni transport (Le Vaillant, 2014). A variety of geologic processes have been proposed to be capable of forming hydrothermal Ni occurrences globally. Appendix A lists several representative global hydrothermal Ni occurrences and their characteristics. A common feature across these hydrothermal Ni occurrences is the presence of Ni-As-mineral phases (e.g. Coolac ultramafic belt, Ashley, 1973; Beni Bousera, Leblanc, 1986; Pevkos area, Thalhammer et al., 1986; Eastern Metals deposit, Auclair et al., 1993; northwest Nelson area, Grapes and Challis, 1999; and Spessart ore district, Wagner et al., 2008; Bou Azzer deposit, Ahmet et al., 2009; Avebury deposit, Keays and Jowitt, 2013; Doriri Creek prospect, Gonzalez-Alvarez et al., 2013b; Main Urals Fault deposits, Melekestseva et al., 2013; Niutitang Formation, Xu et al., 2013; Miitel hydrothermal Ni halo, Le Vaillant et al., 2015). However, the fluids that have been proposed to be responsible for mobilizing Ni vary greatly across each geologic setting. Fluid types that have been proposed to be capable of Ni transport include basinal brines, magmatic hydrothermal fluids, oceanic fluids and metamorphic fluids. Gonzalez-Alvarez et al. (2013a) suggested that the chemistry of basinal brines could have the appropriate blend of chemical features and available ligands to mobilize Ni at relatively low temperatures. At the Wittichen mining area, highly saline basinal fluids in equilibrium with redbeds mobilized Ni and other metals during diagenesis (Staude et al., 2011). Precipitation of ore minerals occurred through fluid mixing processes as the basinal fluid mixed with ascending brines released from mid-crustal levels (Staude et al., 2011). At the Sudbury Igneous Complex, it is believed that the hydrothermal mobilization of Ni and PGE from the local primary massive sulfides occurred through the interaction of late magmatic and hydrothermal fluids (Le Vaillant, 2014 and references therein). At the Avebury deposit in Tasmania, a late magmatic-hydrothermal fluid is speculated to be responsible for Ni mobilization

(Keays and Jowitt, 2013). At this deposit, Ni-minerals and Ni-As-minerals are interpreted to have formed through the scavenging of metals in mafic and ultramafic rocks in the vicinity, by a magmatic hydrothermal fluid exsolved by an underlying granitic intrusion (Keays and Jowitt, 2013). Peculiarly, pentlandite, a Fe-Ni-mineral, has been reported in some porphyry copper deposits which form from magmatic hydrothermal fluids (e.g. Tertiary Maronia deposit, Melfos et al., 2002; Talitsa Porphyry deposit, Azovskova and Grabezhev, 2006; Sungun Porphyry deposit, Hezarkhani, 2006). Alternatively, at the Talvivaara deposit of Finland, seawater in an anoxic environment is proposed to mobilize Ni from a local Ni-rich rock and consequently form pentlandite in black schists (Loukola-Ruskeeniemi and Lahtinen, 2013). Oceanic water has also been effective in Ni- transport in ancient and modern volcanogenic massive sulfide (VMS) deposit settings such as the Main Urals Fault Deposits (Melekestseva et al., 2013) and the Rainbow hydrothermal field (Marques et al., 2006), respectively. In addition to basinal, magmatic, oceanic and meteoric fluids being proposed as fluid types capable of Ni transport, metamorphic fluids are also suggested as capable of mobilizing Ni. At the Thompson Nickel Belt in Canada, Ni enrichment is attributed to precipitation from Ni-rich fluids that were generated by the metamorphism of the ore and the ore-hosting rocks in the area (Bleeker, 1990; Burnham et al., 2003). At the Sarah’s Find and Miitel deposits of Western Australia, it is suggested the As-rich fluids circulated through the system and mobilized the Ni of the primary komatiite-associated magmatic Ni-sulfide deposits in the vicinity (Le Vaillant, 2014). These As-rich fluids are interpreted to be associated with a regional orogenic gold event (Le Vaillant, 2014). The mobilization of Ni by a metamorphic fluid is in accordance with the occurrence of minor to trace amounts of Ni- and Co-phases, such as pentlandite, at various orogenic gold deposits (Goldfarb, pers. comm.). Examples of pentlandite-formation in such deposits includes the lode gold deposits hosted in the Otago and Alpine schists of South Island (New Zealand; Pitcairn et al., 2006), the schists and phyllites at Kundarkocha in Jharkhand (India; Sahoo et al., 2009) and the Kuhmo greenstone belt (Finland; Novoselov et al., 2013). Studies on the behavior of nickel in hydrothermal fluids, such as aqueous speciation and hydrodynamic properties of such species, are extremely limited. There are very few studies at which Ni transport has been quantitatively modeled in hydrothermal fluids (Le Vaillant, 2014). Experiments involving the addition of HCl to aqueous fluids have resulted in increases in Ni

3 solubility as HCl concentration increases (Lin and Popp, 1984; Fahlquist and Popp, 1989; Liu et al., 2012; Scholten et al., 2018). These experiments are often performed with Pt as well, likely because Ni and Pt have similar behavior in hydrothermal environments as they are transported in the divalent oxidation states as aqua- and/or chlorocomplexes (Tian et al., 2012; Liu et al., 2012; Barnes and Liu et al., 2012; Scholten et al., 2018). These experiments have led scientists to believe that Cl- is a favorable ligand to which Ni would complex with in hydrothermal environments and consequently dissolve and become mobile. However, it is significant to note that bisulfide (HS-) is commonly considered a favorable ligand for Pt and PGE (Barnes and Liu et al., 2012; Scholten et al., 2018) principally because Ni displays chemical behavior that is similar to Pt and PGE in hydrothermal fluids. Results from experiments performed by Scholten et al. (2018) indicate that Ni solubility has a stronger dependence on HCl molality rather than on Cl- activity. The dominant aqueous species formed in Ni dissolution experiments with HCl are: + - NiCl , NiCl2(aq) and NiCl3 (Lin and Popp, 1984; Fahlquist and Popp, 1989; Liu et al., 2012).

1.2 Mineral Deposits near Kambalda The Archean Yilgarn Craton of Western Australia is a highly mineralized area that consists of granite, gneiss and greenstone belts, with deep weathering taking place in the Paleocene and Eocene, resulting in the formation of a regolith blanket (Vearncombe and Elias, 2017; Wyche and Wyche, 2017). A complete summary of the geologic events that occurred at the Yilgarn Craton is provided in Appendix B. This craton is a major producer of bauxite, gold, lithium, nickel and tantalum, and has production of copper, iron ores, silver and zinc (Vearncombe and Elias, 2017). As a component of the Yilgarn Craton, the Kalgoorlie region is an area rich in ore deposits of various commodities (Fig. 1.1). Mines of the Kalgoorlie Region (Fig. 1.1) that are renowned for their significant concentrations of ore include the Kalgoorlie Super Pit gold-mining operation and the copious Ni orebodies of Kambalda. This section highlights the various types of significant deposits of in the Kalgoorlie Region (Fig. 1.1) and details general features of each deposit.

Figure 1.1. Map of Australia depicting, in pink circles, the numerous major known deposits throughout the country. The Yilgarn Craton is highlighted with a pale yellow outline. Figure modified after Phillips (2017b).

Magmatic Ni-Cu-(PGE) Deposits Magmatic Ni-Cu-(PGE) deposits that are associated with komatiitic rocks in Archeaen volcanic belts, are considered to be the most important types of magmatic ore in volcanic rock (Lesher and Barnes, 2008). In 2008, these deposit types served as the source for about 20% of the world’s total identified Ni resources in deposits, while containing copious concentrations of Co and Cu (Lesher and Barnes, 2008). In this subsection, the magmatic sulfide deposits of Kambalda, and those in close proximity to Kambalda, are introduced. The Kambalda Nickel Field (Gresham and Loftus-Hills, 1981) contains abundant komatiite-associated magmatic sulfide ore deposits (Fig. 1.2; Gresham and Loftus-Hills, 1981). The hypogene ore mineralogy of these magmatic orebodies is pyrrhotite, pentlandite, pyrite and

5 minor chalcopyrite with locally significant amounts of millerite (Gresham and Loftus-Hills, 1981). The formation of ore at the Kambalda Nickel Field is attributed to Archean komatiitic volcanic flows and subsequent gravity separation of immiscible sulfide liquid (Gresham and Loftus-Hils, 1981 and references therein). Such volcanic activity, in combination with sedimentary deposition, is responsible for the development of the main stratigraphic units of the Kambalda Nickel Field which include: (1) a primary ore-hosting sequence of mafic-ultramafic volcanic rocks; (2) an overlying sedimentary layer of felsic volcanic rocks, graywackes and argillites; and (3) an uncomformably overlain polymictic conglomerate called the Merougil Beds, among other sedimentary layers like the Lake Dam and Triangle Island Beds (Gresham and Loftus-Hills, 1981; Connelly, 2012). The layers of volcanic, volcaniclastic and sedimentary rocks are overlain by a secondary sequence of mafic-ultramafic volcanic rocks called the Bluebush Sequence (Gresham and Loftus-Hills, 1981). Multiple phases of deformation have resulted in the deformation of massive sulfide orebodies leading to the remobilization of these sulfides to tens of meters from their original position (Gresham and Loftus-Hills, 1981). Additionally, metamorphism has resulted in the recrystallization of ores (Gresham and Loftus- Hills, 1981). The magmatic orebodies of the Kambalda Nickel Field are dominantly concentrated at the Kambalda Dome, St. Ives and the Tramways Dome (Fig. 1.2), largely forming ‘contact ore’ as they form at the base of the lower stratigraphic ultramafic flow and generally occupying elongate troughs in the footwall basalt (Gresham and Loftus-Hills, 1981). The Kambalda Dome is the field area of this study and is host to the most amount of ore shoots, in the Kambalda Nickel Field, which includes the Juan, Otter, Durkin, Gibb, Long and Victor shoots (Gresham and Loftus-Hills, 1981), among others. The orebodies at these shoots are dominantly contact ore with the Juan shoot being associated with sedimentary layers. The ore shoots at the northwest flank of the Kambalda Dome are present as variable ore types and include the Ken, McMahon, Gellatly, Wroth, Gordon and Loreto shoots. The shoots occur in the hanging wall and/or are associated with sedimentary layers (Gresham and Loftus-Hills, 1981). St. Ives is over 10 km south of the Kambalda Dome (Fig. 1.2). St. Ives has been operating as a goldfield since 1980 (Oxenburgh et al., 2017), however the area also contains magmatic Ni deposits hosted in komatiitic flows, including the Jan and Foster orebody shoots (Gresham and Loftus-Hills, 1981). The Foster shoot is located on the southwest flank of the St. Ives structure with most of the ore localized in a well-defined trough and the highest grades being 10-12% Ni

(Gresham and Loftus-Hills, 1981). The Jan shoot is located at the southern end of the St. Ives structure and the ore dominantly occurs as a series of ore horizons in the hanging-wall (Gresham and Loftus-Hills, 1981). The highest ore grades range between 14-20% Ni (Gresham and Loftus- Hills, 1981).

Figure 1.2. Map of the regional geology of the Kambalda Nickel Field located in Western Australia adapted after Gresham and Loftus-Hills (1981).

The Tramways Dome, located about 15 km south of the Kambalda Dome (Fig. 1.2), is an overturned fold structure that hosts magmatic sulfide deposits which are being mined for their

7 nickel content (Hicks and Parkinson, 2017). There are two significant ore shoots in the Tramways Dome called the Lanfranchi and Edwin shoots (Gresham and Loftus-Hills, 1981). The two shoots are located on the southeast flank of the dome structure, forming ‘contact ore’ with varying amounts of massive and matrix sulfides (Gresham and Loftus-Hills, 1981). The typical high grade ore at these shoots contain between 10-15% Ni at Edwin and 10-13% Ni at Lanfranchi (Gresham and Loftus-Hills, 1981). As of 2017, exploration of the Dome has resulted in identification of ten nickel sulfide deposits since its discovery in 1969 (Hicks and Parkinson, 2017). The Archean Widgiemooltha Dome is located about 50 km southwest of Kambalda (Fig. 1.1). The Ni deposits of the Widgiemooltha Dome are located along contacts between basalt and overlying komatiite flows that are considered to be correlative with the footwall basalts and overlying komatiite at Kambalda (Hatfield et al., 2017). The deposits are usually small orebody shoots consisting of disseminated to massive sulfides with ore grades averaging 1-3% Ni (McQueen, 1981). The area has undergone amphibolite facies regional metamorphism and ore mineralogy across the various deposits at the Widgiemooltha Dome dominantly consists of pyrrhotite and pentlandite (McQueen, 1981). The Black Swan Ni sulfide deposits are less than 100 km north of Kambalda (Fig. 1.1) and are hosted within the Black Swan Ultramafic Complex. This ultramafic complex lies on a steeply east-facing limb of a regional anticline having experienced lower greenschist facies metamorphism (Barnes et al., 2017). The Black Swan deposits consist of multiple orebodies including the Cygnet and Black Swan disseminated orebodies and the high-grade Silver Swan pure massive-sulfide ore shoot (Barnes et al., 2017). Various small subsidiary massive ore shoots are also present in the area, like the Gosling and Black duck ore shoots (Barnes et al., 2017). These orebodies were formed from an extensively contaminated invasive komatiite flow and intrusion complex. The external sulfur source, necessary for forming sulfides, is not known but there is pyrite reported in dacitic tuffs and tuff breccias that are intercalated with the youngest komatiite flows. (Barnes et al., 2017).

Laterite Ni Deposits About 60-70% of the global nickel supply is sourced from Laterite Ni ore deposits, which form from the intensive deep weathering of serpentinite under humid tropical conditions (Butt

8 and Cluzel, 2013). There are three general subtypes of this deposit type: oxides, hydrous Mg silicates and clay silicates (Butt and Cluzel, 2013). The Murrin Murrin and Bulong Ni laterite deposits are the best known Australian examples of the clay silicate subtype with Ni being hosted in smectite (Wells and Butt, 2017). The Murrin Murrin deposit is in the NE Yilgarn Region and is over 200 km away from Kambalda (Fig. 1.1), however the Bulong ultramafic complex is only over 80 km northeast of Kambalda. The Bulong ultramafic complex hosts Ni laterite deposits extending over 20 km in the N-S direction (Wells and Butt, 2017). The complex is composed of Archean metamorphosed high-Mg basalts, ultramafic rocks, and a sequence of calc-alkaline, volcanic and metasedimentary rocks (Wells and Butt, 2017 and references therein). The larger ultramafic sills are composed of serpentinized, olivine cumulates and peridotites capped by thin, altered pyroxenitic and noritic units (Wells and Butt, 2017). After the emplacement of these ultramafic rocks, the area was subjected to several periods of deformation producing regional, tight-to-open folding (Wells and Butt, 2017). At Bulong, the Ni laterite resource in June 2003 was 150 Mt, with a mineable reserve of 6.9 Mt at 1.57% Ni (Wells and Butt, 2017 and references therein).

Pegmatite Deposits Pegmatite deposits are important resources for various commodities including Sn, Nb, Ta, Li, Rb, Cs, Be, W, Au and rare earth elements (Dittrich et al., 2019). Such deposits are additionally known for high quality, colorful gemstones and mineral specimens (Dittrich et al., 2019). The most widely accepted genetic model for pegmatites is the formation from residual melts where incompatible components, fluxes, volatiles and rare elements concentrate after the crystallization of granitic plutons (Dittrich et al., 2019 and references therein). The Yilgarn craton, along with the Pilbara craton, is known to host many pegmatites such as the world-class economic Greenbushes and Wodgina deposits (Dittrich et al., 2019). There are at least 125 locations within the Yilgarn Craton that have reported pegmatite deposits (Dittrich et al., 2019 and references therein). Significant pegmatite deposits in the Kalgoorlie Terrane of the Yilgarn Craton’s Eastern Goldfields Superterrane, particularly of lithium-cesium-tantalum (LCT) type, include the Sinclair Cesium Project and the Londonderry, Mount Deans, and Mount Marion pegmatite fields (Dittrich et al., 2019). However, the pegmatite deposits that are in closest proximity to Kambalda are the Mount Marion and Londonderry pegmatite fields.

The Mt. Marion lithium pegmatite deposit is located about 25 km northwest of Kambalda (Fig. 1.1). The deposit is hosted in Archean high Mg-basalt and forms along a 10 km-long strike that is subparallel to the Depot Granodiorite intrusion (Smith and Ross, 2017). These deposits typically occur as a series of gently dipping subparallel sills that are 2-30 m thick, although at some locations are as thick as ~225 m (Smith and Ross, 2017). The mineralogy of the deposit consists of spodumene, quartz, albite-oligoclase, potassium feldspar and muscovite with the following accessory minerals: garnet, apatite, columbite-tantalum, beryl, cassiterite, magnetite, tourmaline, and holmquistite (Smith and Ross, 2017). The pegmatite deposits of Mt. Marion are classified as albite-spodumene type of the rare element class of granitic pegmatites discussed in Cerny and Ercit (2005) which form through differentiation from peraluminous S-, I- or mixed S- and I-type granitic plutons (Cerny and Ercit, 2005; Smith and Ross, 2017). It is believed that the pegmatites derived from the local Depot Granodiorite which is age dated at 2689 Ma (Smith and Ross, 2017 and references therein). The Londonderry pegmatite field is located ~55 km northwest of Kambalda. This lithium-cesium-tantalum pegmatite field covers an area of about 5 km2 and outcrops along folds in close proximity to a NNW-SSW trending fault (Dittrich et al., 2019). The host rocks for the pegmatites are dominantly metamorphosed komatiite flows. The Londonderry pegmatite field is composed of at least four known pegmatite sheets called Londonderry Feldspar Quarry, Lepidolite Hill, Tantalum Hill and Bon Ami. Up to until 2017, this pegmatite field was the only locality in Australia at which pollucite (a Cs-bearing zeolite mineral) was reported (Dittrich et al., 2019). Other minerals that formed in this pegmatite field include cassiterite, petalite, columbite-tantalite, beryl, lepidolite, albite, microcline, quartz, zinnwaldite and garnet (Dittrich et al., 2019).

Gold Deposits The Yilgarn Craton (Fig. 1.1) is one of the world’s premier gold-producing regions as it has increased its gold production by tenfold since 1979 (Phillips et al., 2019). The gold deposits of the eastern Yilgarn Craton are structurally controlled and 3D maps show that multi-phase granite-cored domes are present at varying depths beneath all the giant gold deposits (Blewett, 2010a). A quantitative spatial analysis of data was performed to determine parameters that are relevant to the mineral exploration of gold in the Yilgarn Craton (Bierlein et al., 2006). This

10 study revealed that endowment can be correlated with empirical exploration criteria for orogenic gold deposits such as intersections of anticlines and major faults as well as localization of gold in smaller high-order faults in proximity to major first-order faults (Bierlein et al., 2006). However, Blewett (2010a) states that the metamorphic evolution of the eastern Yilgarn Craton does not support a late-stage metamorphic fluid hypothesis for the formation of gold in the area, but may have been responsible for the formation of gold in only one of the multiple gold events (Blewett, 2010a). In this subsection, the gold deposits in and close to Kambalda are introduced. Brief general information is provided for each deposit, except for the longer detailed excerpts of the Kalgoorlie and St. Ives goldfields. These two goldfields will be discussed in further detail as they are the two largest gold-endowed areas in proximity to the Kambalda Dome studied here. The Carosue Dam goldfield is located ~125 km north-northeast from Kambalda (Fig. 1.1). The main gold deposits of Carosue Dam are Karari, Whirling Dervish, Luvironza, Monty’s Damn and Twin Peaks with several smaller gold deposits (Witt and Mills, 2017). The deposits of Carosue Dam occur in the hanging wall of a regional northwest-striking, northeast dipping fault called the Keith Kilkenny fault (Witt and Mills, 2017). The deposits, with the possible exception of the Twin Peaks deposit, are hosted in a succession of metasedimentary rocks of the Carosue Basin, which have been intruded by alkalic dikes and plugs (Witt and Mills, 2017). The Karari, Whirling Dervish, Luvironza and Monty’s Dam deposits are believed to have formed from a high-temperature oxidized magmatic-hydrothermal fluid (Witt and Mills, 2017). These four deposits are intimately associated with subvolcanic alkalic intrusions and have formed in association with potassic alteration and low-grade copper mineralization (Witt and Mills, 2017 and references therein). However, the Twin Peaks deposit holds characteristics that are more typical of orogenic gold deposits as the mineralization is structurally controlled, was formed from a reduced fluid and is associated with hydrothermal alteration typical of orogenic gold deposits (Witt and Mills, 2017). The Kanowna Belle deposit is about 65 km north-northwest of Kambalda (Fig. 1.1) and is a world class deposit located in the Archean Scotia-Kanowna Dome, along with several smaller gold deposits (Davis et al., 2010). This gold deposit is spatially associated with the steeply south- southeast dipping Fitzroy Structural Zone which separates the ore into hanging wall and footwall lithostructural domains (Davis et al., 2010; Bull et al., 2017). The structural zone bisects a sequence of coarse-grained sedimentary and felsic volcanoclastic rocks (Bull et al., 2017). This

11 structure is also noted to have controlled the localization of felsic and intermediate composition intrusions such as the highly mineralized Kanowna Bell Porphyry (Bull et al., 2017). All of these rock types host gold ore, however the majority of the ore is hosted in the Kanowna Bell Porphyry (Bull et al., 2017). This intrusion hosts the highest grade and the most intense alteration in the area, depicting its significance in serving as a major plumbing system for auriferous fluids (Davis et al., 2010). Examples of gold mineralization styles include carbonate vein stockwork, quartz-carbonate veins, silica or carbonate rich breccia, sulfide-quartz-carbonate stringers and arrays of sheeted veins (Bull et al., 2017). The hydrothermal footprint associated with the deposits includes W, Mo, Bi, Te, Sb, As and S (Rogers et al., 2004; Bull et al., 2017). Additionally, the mineralized areas have alkali halos with an enrichment of K, Na, Rb and Ba, being characterized by abundant albite and potassium feldspar with a white mica overprint. The formation of the main gold ore forming process is believed to have been contemporaneous with sinistral reactivation of the Fitzroy Shear Zone, transpiring as a result of a switch in far-field stress axes in the region (Bull et al., 2017). The Kundana goldfield is located about 50 km northwest of Kambalda (Fig. 1.1). The goldfield deposits occur within a 15 km extent encompassing deposits localized in close proximity to the Zuleika Shear Zone Complex which is a large northwest-trending crustal-scale structure that spans a distance of over 150 km (Cooke et al., 2017). Excluding supergene mineralization, there are four types of ore styles at Kundanda: Strzelecki style, K2 style, Pode style and Brittle vein array (Cooke et al., 2017). Strzelecki style mineralization occurs as laminated crack-seal ribbon veins with average grades of 55 g/t gold with a mineral assemblage of quartz, carbonate, pyrite, gold with minor sphalerite, chalcopyrite and galena and locally scheelite (Cooke et al., 2017). K2 style mineralization occurs as laminated crack-seal ribbon veins with post-formation brecciation, typically grading at 25 g/t gold, with a mineral assemblage of quartz, carbonate, chlorite, pyrite, pyrrhotite, arsenopyrite, gold, silver with minor sphalerite and galena and locally scheelite (Cooke et al., 2017). Pode style mineralization are massive quartz breccias with slivers of chlorite-biotite altered hostrock (Cooke et al., 2017). The mineralization type typically grades at 3 g/t with local intersection of greater than 50 g/t and the mineral assemblage is quartz, carbonate, chlorite, pyrite and biotite commonly present in the alteration halo (Cooke et al., 2017). Lastly, brittle vein arrays occur as sheeted to stockwork-style vein arrays which typically grade at about 2 g/t. They consist of quartz, carbonate, pyrite and

12 gold (Cooke et al., 2017). Calcic, sodic and potassic alteration are all observed at Kundana (Cooke et al., 2017). At least four temporally distinct ore forming events are responsible for the formation of the Kundana goldfield and each area associated with certain deformation events (Cooke et al., 2017). The Coolgardie goldfield is approximately 55 km northwest of Kambalda (Fig. 1). The host rocks to the ore at this goldfield are dominantly mafic-ultramafic volcanic and intrusive rocks that have been overlain by sedimentary rocks (Goodz and Money, 2017). These rocks have been metamorphosed at amphibolite facies conditions and have been intruded by I-type granites (Goodz and Money, 2017). The deformation can be described as low strain with weak to moderate foliation (Goodz and Money, 2017). There are about 60 deposits that compose the Coolgardie goldfield and most of them occur within a 5-km arc formed around the Caloolie granite (Knight et al., 2000; Goodz and Money, 2017). The ore dominantly occurs in veins, breccia zones, stockwork veining and shear zones which are hosted in variable lithologies. However, there is a localization of ore adjacent to or within major tectonic zones such as shear zones and fold hinges. Gold is broadly synchronous with peak metamorphism, forming during a transition from ductile to brittle-ductile tectonics (Knight et al., 2000; Goodz and Money, 2017). It has also been suggested that the ore formation coincides with main granite emplacement and regional deformation (Goodz and Money, 2017 and references therein). There are five main mineralization styles in the goldfield: Type 1, Type 2, Type 3, Type 4 and Type 5 (Goodz and Money, 2017). The Type 1 deposits are characterized by being boudinaged and are associated with extensive albite, silica and sulfide alteration (Goodz and Money, 2017). The ore minerals associated with gold are arsenopyrite and pyrrhotite with minor pyrite, galena, sphalerite and chalcopyrite. The Type 2 deposits are characterized by being hosted in fault-bounded quartz stockwork veining and quartz vein arrays in igneous rocks like gabbro, dolerite, basalt or diorite (Goodz and Money, 2017). The ore is hosted in the veins and the vein alteration halos, being associated with albite, carbonate, silica and sulfides (Goodz and Money, 2017). Ore minerals associated with gold include arsenopyrite and pyrrhotite with minor pyrite, galena, sphalerite and chalcopyrite (Goodz and Money, 2017 and references therein). The Type 3 deposits are characterized by being hosted by laminated quartz veins associated with sedimentary horizons at mafic rock contacts (Goodz and Money, 2017). The alteration and the sulfides associated with this type are similar to that of Types 1 and 2 (Goodz and Money, 2017). The Type 4 deposits are

13 located in ductile shear zones and hosted in ultramafic rocks (Goodz and Money, 2017). This type lacks sulfides and quartz, being associated with biotite, chlorite, talc, carbonate and tourmaline in the shear zones (Goodz and Money, 2017). The final type, Type 5, is the least common type and is associated with quartz veins structurally controlled within and near a tonalite intrusion (Goodz and Money, 2017). The ore minerals associated with this mineralization are pyrite and pyrrhotite with minor sphalerite, galena, chalcopyrite and molybdenite (Goodz and Money, 2017). Diopside and grossular sometimes occur near the quartz veins and biotite is a common mineral associated with this mineralization type (Goodz and Money, 2017). The Bullabulling goldfield is located about 75 km west-northwest from Kambalda (Fig. 1.1). The gold deposits at Bullabulling are localized around granitic intrusions (Partington et al., 2017). There is an 11 km-long N-S striking trend that hosts a majority of the gold, being a continuously mineralized zone (Partington et al., 2017). There is laterite that overlays the primary gold mineralization and are often enriched to ore grades (Patington et al., 2017). About 20 m below the gold-enriched laterite is supergene gold deposits being separated by material that lacks significant gold content (Patington et al., 2017). The hypogene ore occurs parallel to N- and NW-trending lithological contacts that wrap around the margin of the underlying Bali Monzogranite (Partington et al., 2017). The gold-bearing ore minerals at the deposit are pyrite with minor chalcopyrite and pyrrhotite (Partington et al., 2017). The gold mineralization formed late in the metamorphic and deformation history and broadly predates the most recent phase of granite and pegmatite intrusive activity (Partington et al., 2017). The ore lenses that host gold are up to 20 m-thick and are characterized by disseminated sulfides with calc-silicate prograde alteration, involving the formation of hornblende, diopside, biotite, albite, carbonate and quartz (Patington et al., 2017). Recent seismic data has revealed that there is a clear spatial relationship between gold mineralization and folds (Patington et al., 2017). It is interpreted that the overprinting of various structures during various deformation events provided structural heterogeneity and permeability, allowing the formation of the broad, low-grade disseminated gold deposits at this goldfield (Patington et al., 2017). The Mt. Monger goldfield is located about 35 km east-northeast of Kambalda (Fig. 1.1). This goldfield can be divided into the following four areas: the central Daisy Milano, the Randalls banded iron formation (BIF), Salt Creek and Wombola (Goodz et al., 2017). The

14 central Daisy Milano area has produced about half of the gold at Mt. Mounger and the Randalls BIF and Salt Creek have produced significant amounts (Goodz et al., 2017). Contrarily, the Wombola area has produced a relatively small amount (8%) of gold at Mt. Mounger (Goodz et al. 2017). The deposits at Daisy Milano are hosted in a series of andesite, epiclastic rocks and conglomerate which are unconformably overlain by komatiite, serving as the hanging wall (Goodz et al., 2017). Two dominant faults occur at this location and appear to be significant in gold endowment as grades increase in proximity to the two faults (Goodz et al., 2017). The gold mineralization here is concentrated in quartz veins containing abundant free gold, pyrite, sphalerite and galena with multiple episodes of dilation and precipitation (Goodz et al., 2017). These veins typically contain alteration halos consisting of carbonate and white mica, including fuchsite (Goodz et al. 2017). The deposits at Randalls are hosted in BIF rocks that display chevron-folding (Goodz et al., 2017). The mineralogy of the ore deposit is dominantly pyrrhotite with some pyrite and arsenopyrite (Goodz et al., 2017). The main chemical mechanism that controls the precipitation of the gold here is sulfidation which is the reason for which BIF has served as an ideal host for gold mineralization (Goodz et al., 2017). The deposits have formed along a NW-trending corridor which extends up to 250 m vertically and extends to at least 1.7 km in length (Goodz et al., 2017). The gold mineralization is in higher grades in the shoots that plunge to the south (Goodz et al., 2017). The ore is localized in the vein alteration halos of tensional quartz veins (Goodz et al., 2017). The gold is associated with a amphibole-biotite- chlorite-carbonate-pyrrhotite-arsenopyrite mineral assemblage forming as a result of alteration of the BIF (Goodz et al. 2017). The last significant gold producing area at Mt. Monger is Salt Creek. The deposits here are hosted in a medium-grained quartz dolerite (Goodz et al., 2017). Similarly, to the Randalls BIF, the deposits here are concentrated along a NW-trending corridor (Goodz et al., 2017). This corridor extends at least 1.2 km in length and extends over 120-170 m vertically (Goodz et al., 2017). The mineralization here occurs in vein selvedges and alteration halos of quartz veins. The gold is associated with a silica-albite-pyrrhotite mineral assemblage formed from the alteration of dolerite (Goodz et al., 2017). Arrays of brittle-ductile and biotite shear zones served as fluid pathways, allowing mineralized fluids to enter the system and produce alteration and veining (Goodz et al., 2017). The Kalgoorlie goldfield is located about 55 km north-northwest of Kambalda (Fig. 1.1) and is the leading gold producer in Australia, with a total production of 2200 t of gold since 1893

(Phillips et al., 2017a). The goldfield is the locality for two world-class deposits: Mt. Charlotte in the north-west and Golden Mile in the south-east (Mueller, 2017). The goldfield also includes the Mt. Percy deposit and several minor producers (Vielreicher et al., 2010). Through studies on the Golden Mile and Mt. Charlotte deposits, the deposits of the Kalgoorlie Gold Field are believed to have formed from the deposition from a juvenile, potassic, low salinity, CO2 (±CH4) bearing aqueous fluid forming around 2.64 Ga (Vielreicher et al., 2016). The Golden Mile of the Kalgoorlie goldfield deposit has produced the most mineralization at 1708 t of gold up until 2015 (Mueller, 2017). The main host rock for this deposit is the Archean Golden Mile Dolerite (Phillips et al., 2017a). The Golden Mile Dolerite is a stratabound unit that is about 600 m thick (Philips et al., 2017a). The dolerite was originally composed dominantly of pyroxenes, plagioclase, quartz and oxide minerals (Phillips et al., 2017a). However regional metamorphism has changed the mineralogy of the dolerite to actinolite, chlorite, albite, quartz and oxides (Phillips et al., 2017a). Because of the wide range of alteration assemblages observed in the Golden Mile Dolerite, an effective classification of alteration is considered to be based on the Fe-bearing minerals that formed during the addition of

CO2, S and Au to regional metamorphic mineral assemblages (Phillips et al., 2017a). The variable alteration assemblages forming reflects the differing degrees of alteration and the significant variations in the original magmatic compositions of the dolerite (Phillips et al., 2017a). However, a general alteration scheme at a large (meter-to-kilometer scale) involves: (1) an inner pyrite alteration zone occurring proximal to the mineralization, (2) a carbonate alteration zone, (3) a chlorite alteration zone, and (4) an unaltered zone of actinolite-bearing rocks (Phillips et al., 2017). The mineralization occurs as a variable combination of vein infill and alteration wallrock in brittle-ductile shear zones (Phillips et al., 2017a). The shear zones have focused strain because of their mica-content and has allowed them to focus fluid flow producing more white mica and producing gold (Phillips et al., 2017a). The mineralization is typically composed of pyrite, minor base metal sulfides, sometimes tellurides and local arsenopyrite and pyrrhotite (Phillips et al., 2017a). The Mt. Charlotte deposit of the Kalgoorlie goldfield has also produced significant amounts of gold at 146 t up until 2015 (Mueller, 2017). The 2655 ± 14 Ma gold orebodies are dominantly hosted in the Golden Mile Dolerite, just like the Golden Mile deposits (Mueller, 2017). At Mt. Charlotte, gold is mined from stockworks of coarse-grained sheeted veins that are

16 spatially associated with dextral, reverse brittle faults (Vielreicher et al., 2016). Mineralization occurs as dilational quartz-, carbonate-, albite- and scheelite-bearing veins with gold being associated with pyrite, pyrrhotite, tellurides, ankerite, sericite and quartz (Vielreicher et al., 2016). Gold is typically observed at the contact between vein and hostrock or in pyrite alteration halos (Vielreicher et al., 2016). The alteration zones surrounding the veins are associated enrichments in K, Rb, Cs, Li, Ba, Ca, Sr, Mg, Ni, V, Cr, W, Te and Au (Mueller, 2015; Vielreicher et al., 2016). The St. Ives goldfield is located in the Kambalda Nickel Field, about 10 km south of the Kambalda Dome (Figs. 1.1 and 1.2). This goldfield is composed of over 60 gold deposits and has been actively mined since 1980, producing over 400 t of gold (Oxenburgh et al., 2017). The stratigraphy of this area includes the mafic-ultramafic rocks of the Kambalda Sequence, felsic to intermediate volcanic and epiclastic sedimentary rocks of the Kalgoorlie Group and rocks of the Merougil Formation that formed from late basin filling (Oxenburgh et al., 2017). The sequences are intruded by felsic to dioritic porphyritic intrusions (Oxenburgh et al., 2017) and the goldfield is bounded by the regionally-extensive Boulder-Lefroy fault, on the east (Cox and Ruming, 2004). The gold mineralization is hosted in rocks of all stratigraphic units at St. Ives (Oxenburgh et al., 2017) and is localized along networks of mixed brittle-ductile shear zones (Cox and Ruming, 2004). Most of the gold deposits within this goldfield are hosted in subsidiary structures of the Boulder-Lefroy fault zone as they are favorable dilational sites for gold deposition (Cox and Ruming, 2004; Neumayr et al., 2008). The deposits formed in low displacement faults by the localization of auriferous fluid flow driven by major slip events along the Boulder-Lefroy fault zone (Cox and Ruming, 2004). A sinistral transpression event is responsible for the main gold mineralization in many areas of the Yilgarn Craton including St. Ives, Kalgoorlie, Wallaby and Sunrise Dam (Goscombe et al., 2009). The deformation event is characterized by a stress switch to N-S contraction resulting in sinistral transpressive reactivation of existing NNW- and N- trending faults (Goscombe et al., 2009). Economic grades of gold dominantly occur in the alteration halos adjacent to faults, shear zones and associated quartz-carbonate vein arrays (Cox and Ruming, 2004). There are at least four distinguishable alteration types, as described in Neumayr et al. (2008): (1) early, porphyry- related epidote-magnetite-carbonate alteration, (2) early, regional shear-zone related carbonate alteration, (3) syn-gold, reduced pyrrhotite-biotite-amphibole and syn-gold, oxidized plagioclase-

17 carbonate-pyrite-magnetite-hematite-biotite-chlorite alteration, and (4) late quartz veins. The deposits in the southwest area of the goldfield contain arsenopyrite as a component of their syn- gold reduced mineral assemblage (Neumayr et al., 2008). The magnetite and magnetite-pyrite alteration likely derived from, or is associated with felsic to intermediate porphyry intrusions (Neumayer et al., 2008). The following five mine areas host the most significant concentrations of gold, hosting more than 2 Moz of gold: Greater Revenge, Invincible, Argo, Junction and Victory (Oxenburgh et al., 2017). Characteristics of the Argo and Invincible areas are discussed below to introduce general characteristics of the deposits at St. Ives. At Invincible, gold mineralization is hosted in fine-grained sedimentary rocks, dominantly mudstone (Oxenburgh et al., 2017). The ore occurs as shear-hosted laminated to brecciated quartz veins associated with pyrite, free gold and albite alteration (Oxenburgh et al., 2017). There are also gentle-dipping extension veins that extend into andesite which host free gold and hematite alteration (Oxenburgh et al., 2017). Carbonate, dolomite, ankerite, tourmaline, actinolite, biotite, chlorite, white mica and pyrrhotite alteration minerals are also present (Oxenburgh et al., 2017 and references therein). The Argo mining area deposits occur in km-scale fault splays off the Boulder-Lefroy fault (Oxenburgh et al., 2017). At the Argo deposits, of the Argo mining area, the ore occurs in 2-5 m wide zones of quartz-carbonate fault fill veins, extension veins and hydraulic breccia (Oxenburgh et al., 2017). A proximal alteration halo is associated with these veins which is dominantly composed of albite, carbonate, biotite, pyrrhotite and pyrite. A distal alteration halo is also present and consists of chlorite and biotite alteration (Oxenburgh et al., 2017). At the Athena and Hamlet deposits of the Argo mining area, the ore is hosted in quartz-albite shear parallel veins, extension veins and breccia veins hosted in moderately east-dipping shear zones within basalt (Oxenburgh et al., 2017 and references therein). The proximal alteration halo is composed of a albite-biotite-actinolite-pyrite-pyrrhotite mineral assemblage with pyrrhotite becoming the dominant sulfide species further down depth (Oxenburgh et al., 2017). The distal alteration halo is characterized by biotite-chlorite-carbonate alteration (Oxenburgh et al., 2017). The primary control on the localization of ore is the intersection of gold-hosted shear zones (Oxenburgh et al., 2017).

Quartz Veins in the Kambalda Dome In addition to the magmatic sulfide orebodies of the Kambalda Dome, there also reports of mineralized quartz veins at this location. Auriferous quartz veins in the Kambalda Dome have been reported at the Lunnon shoot (Cowden et al., 1986), Fisher shoot (Cowden et al., 1986), Beta-Hunt mine (Neall and Phillips, 1987; Oxenburgh et al., 2017) and along the eastern flank of the Dome (auriferous quartz veins of this study; Staude, pers. comm.). To a lesser degree, quartz veins associated with PGE minerals have also been studied at Kambalda. A study by Hudson (1986) reports a quartz-carbonate vein in the Lunnon shoot that is mineralized with PGE- tellurides. Euhedral grains of michenerite-testibiopalladite occur as inclusions in altaite in the quartz-carbonate vein, which cuts primary magmatic ore deposits of the Lunnon shoot (Hudson, 1986). Additional minerals that are associated with this vein include hessite, volynskite, rucklidgeite, melonite, galena, chalcopyrite and gold (Hudson, 1986). An additional study reports small grains of sudburyite in nickeline in a quartz-carbonate vein at the Ken shoot of the Kambalda Dome (Hudson and Donaldson, 1984; Hudson, 1986). The most recent and perhaps significant auriferous quartz veins encountered at the Kambalda Dome include those recently uncovered in 2018 at the Beta-Hunt mine. The Beta- Hunt area is located on the southwestern portion of the Kambalda Dome. Beta-Hunt is host to both komatiite-associated magmatic sulfide deposits as well as auriferous quartz veins. In 1979, the Hunt gold deposit was discovered beneath the historically mined magmatic sulfide nickel orebody (Oxenburgh et al., 2017). However, most recently, in 2018, Royal Nickel Corp announced its discovery of a bonanza-grade gold quartz vein which included a 95-kilogram specimen hosting 2,400 ounces of gold (Resource World, 2018).

CHAPTER TWO

CHARACTERIZATION OF PENTLANDITE-BEARING QUARTZ VEINS IN KAMBALDA, WESTERN AUSTRALIA 2.1 Introduction Gold-bearing quartz veins near the Kambalda Dome have been extensively studied, leading to copious data and documentation of auriferous quartz veins in the vicinity. Studies concerning Ni minerals in quartz veins in the Kambalda Dome have contrarily produced far fewer publications, however Ni mobilization via hydrothermal processes has been documented globally (e.g. United States of America, Candela et al., 1989; New Zealand, Grapes and Challis, 1999; Canada, Molnar et al., 2001; China, Ripley et al., 2005; Brazil, Almeida et al., 2007; Finland, Peltonen et al., 2008; Germany, Staude et al., 2011; Tasmania, Keays and Jowitt, 2013; Russia; Melekestseva et al., 2013; Papua New Guinea, Gonzalez-Alvarez et al., 2013b; Zambia, Capistrant et al., 2015). Evidence for hydrothermal mobilization of Ni is reported in the form of hydrothermal alteration of magmatic sulfide deposits (e.g. North Range, Sudbury Igneous Complex, Molnar et al., 2001; Jinchuan, China, Ripley et al., 2005; Fortaleza de Minas, Brazil, Almeida et al., 2007) and ultramafic rocks (e.g. Avebury, Tasmania, Keays and Jowitt, 2013). There are also a variety of Ni-rich or Ni-bearing deposits that have formed through precipitation of Ni from an aqueous fluid including Ni-enriched black shales (Jowitt and Keays, 2011), unconformity uranium deposits (e.g. Athabasca Basin; Jefferson et al., 2007), Five Element Vein Deposits (Kissin, 1992; Markl et al., 2016), sediment-hosted Ni-rich deposits of the Central African Copperbelt (e.g. Enterprise deposit, Capistrant et al., 2015; Menda Central Prospect, Ball et al., 2016), some porphyry copper deposits (e.g. Tertiary Maronia deposit, Melfos et al., 2002; Talitsa Porphyry deposit, Azovskova and Grabezhev, 2006; Sungun Porphyry deposit, Hezarkhani, 2006), some volcanogenic massive sulfide deposits (e.g. Main Urals Fault deposits; Melekestseva et al., 2013) and many orogenic gold deposits (e.g. Otago and Alpine schists, Pitcairn et al., 2006; Jharkhand, Sahoo et al., 2009; Kuhmo greenstone belt, Novoselov et al., 2013). A common feature across these hydrothermal Ni occurrences is the presence of Ni-As-mineral phases (e.g. Coolac ultramafic belt, Ashley, 1973; Pevkos area, Thalhammer et al., 1986; Beni Bousera, Leblanc,

1986; Eastern Metals deposit, Auclair et al., 1993; northwest Nelson area, Grapes and Challis, 1999; Spessart ore district, Wagner et al., 2008; Bou Azzer, Ahmet et al., 2009; Avebury deposit, Keays and Jowitt, 2013; Doriri Creek prospect, Gonzalez-Alvarez et al., 2013b; Main Urals Fault, Melekestseva et al., 2013; Niutitang Formation, Xu et al., 2013; and the Miitel hydrothermal Ni halo, Le Vaillant et al., 2015). Studies of hydrothermal mobilization of Ni have led to hypotheses regarding fluid conditions that may mobilize Ni. Le Vaillant (2014) attributed Ni mobilization to arsenic-rich metamorphic fluids and Hanley and Bray (2009) proposed the transport of nickel is possible through sodium- and halogen-rich fluids at 390-550⁰C. Studies have indicated that chloride is a likely ligand for nickel transport, as Ni is suggested to be transported in the divalent oxidation state as aqua and/or chlorocomplexes in hydrothermal environments (Tian et al., 2012; Liu et al., 2012; Scholten et al., 2018). Ultimately, a wide range of fluid types with a wide variety of temperatures have been proposed to be capable of Ni transport which includes basinal brines, magmatic hydrothermal fluids, oceanic fluids and metamorphic fluids. Near the Kambalda Dome, at the Widgiemooltha Dome, Ni-rich hydrothermal halos surrounding the Miitel magmatic sulfide deposit was studied by Le Vaillant et al. (2015). Indications for nickel mobilization are observed along vein alteration halos of quartz and carbonate veins, containing gersdorffite, minor nickeline, and rare millerite and pentlandite (Le Vaillant et al., 2015). A similar study was performed on the N-NE side of Kambalda Dome at the Otter-Juan and Durkin komatiite Ni deposits by Le Vaillant et al. (2016). However, the results in this study indicated that these komatiite Ni deposits lack any hydrothermal halos similar to that of the Miitel deposit (Le Vaillant et al., 2016). North of the Kambalda Dome, near Kalgoorlie, quartz-carbonate-nickel sulfide-arsenide ± Au veins with up to 2.0 % Ni and 2.3 % As, are believed to have a late metamorphic-hydrothermal origin, being related to carbonate metasomatism following the peak of regional metamorphism (Marston et al., 1981). Hydrothermal activity at the Kambalda Dome has been reported in the form of auriferous quartz veins, PGE-rich quartz veins and quartz-sulfide veins. Auriferous quartz veins in the Kambalda Dome have been reported at the Lunnon shoot (Cowden et al., 1986), Fisher shoot (Cowden et al., 1986), Beta-Hunt mine (Neall and Phillips, 1987; Oxenburgh et al., 2017) and along the eastern flank of the Dome structure (Staude, pers. comm.). Two large world-class goldfields are also located close to the Kambalda Dome, namely the Kalgoorlie goldfield (about

55 km north-northwest of the Kambalda Dome) and the St. Ives goldfield (about 10 km south of the Kambalda Dome). The Boulder Lefroy-Golden Mile fault system links the two major goldfields with a length of over 100 km (Mueller et al., 2016). This fault system is the most intensely mineralized structure in the Yilgarn Craton yielding more than 2150 t Au, as of 2015 (Mueller et al., 2016). The fault zone is comparable to other gold-rich fault zones in other orogenic belts (e.g. Mother Lode in California and the Porcupine-Destor and the Larder Lake- Cadillac faults of the Abitibi greenstone belt) which are interpreted to be fluid conduits for ascending, predominantly non-magmatic, fluids to sites of gold deposition (Mueller et al., 2016 and references therein). The depositional conditions for well-studied gold deposits in the Kambalda-Kalgoorlie area are listed in Table 2.1. More broadly, gold deposits in the entirety of the Yilgarn Craton are suggested by Ho et al. (1992) to have deposited from a common type of ore fluid that has the following characteristics at the depositional site: an H2O-CO2 ± CH4 (XCO2 = 0.2 ± 0.05) bearing fluid with low- to moderate-salinities (generally <6 wt.% NaCl equivalent), and P-T conditions of 200-300°C and 1-3 kbar. Most of these auriferous fluids, described by Ho et al. (1992), are relatively reduced, although more oxidized fluids are reported at the Golden Mile deposit of the Kalgoorlie Terrane (Ho et al., 1992).

Table 2.1 Depositional conditions for three well-studied gold deposits from the Kambalda- Kalgoorlie area from Ho et al. (1992) and references therein. The symbol N.R. represents that the data are not reported in Ho et al. (1992). Parameter Victory, St. Hunt, Mt. Charlotte, Ives Kambalda Kalgoorlie Temperature (°C) 350 to 430 350 to 400 300 to 350 Pressure (kbar) 1.4 to 2.0 0.8 to 1.8 1.5 to 2.3 XCO2 0.1 to 0.2 0.15 to 0.25 0.25 to 0.3 Bulk salinity (wt. % NaCl 5 to 8 < 3 < 3 equiv.) pH N.R. 6.9 5.8 to 6.1 log fO2 N.R. -29.7 -28.8 to -33.4 Fluid/Rock ratio N.R. 10-100 5000-10000 Gold deposition mechanism sulfidation sulfidation sulfidation

To a lesser degree, quartz veins associated with PGE minerals have also been studied at the Kambalda Dome. A study by Hudson (1986) reports a quartz-carbonate vein in the Lunnon shoot that is mineralized with PGE-tellurides. Euhedral grains of michenerite-testibiopalladite

22 occur as inclusions in altaite in the quartz-carbonate vein, which cuts primary magmatic ore deposits of the Lunnon shoot (Hudson, 1986). Additional minerals that are associated with this vein include hessite, volynskite, rucklidgeite, melonite, galena, chalcopyrite and gold (Hudson, 1986). An additional study reports small grains of sudburyite in nickeline in a quartz-carbonate vein at the Ken shoot of the Kambalda Dome (Hudson and Donaldson, 1984; Hudson, 1986). In and near the Kambalda Dome, hydrothermal mobilization of Ni occurred as a result of hydrothermal alteration of magmatic sulfide deposits (Lesher and Barnes, 2008 and references therein). At the Kambalda Dome, minor hydrothermal veins occur as quartz-sulfide ± carbonate veins which formed during metamorphism by mobilization of Fe, Ni, Cu, Co, Cr, Zn, Pd and Au by hydrothermal fluids (Lesher and Barnes et al., 2008 and references therein). Quartz veins bearing great resemblance to this description are observed during mining, in the form of pentlandite-bearing quartz veins (Staude, pers. comm.). In addition to these pentlandite-bearing quartz veins, barren and auriferous quartz veins have also been observed during mining activity along the eastern flank of the Kambalda Dome (Staude, pers. comm.). The three vein types have been identified in proximity to the magmatic sulfide orebodies (Staude, pers. comm.). Pentlandite-bearing quartz veins are of particular interest because pentlandite is intergrown with pyrrhotite in veins, which is a mineral assemblage that is generally considered to be magmatic or volcanic in origin (e.g. Lesher and Barnes, 2008; Haldar and Tisljar, 2014; Pracejus, 2015). Ultimately, this vein type is not commonly reported in literature and therefore detailed characterization of such a vein type is lacking. Auriferous quartz veins have been extensively studied in the Yilgarn Craton and are attributed to orogenic ore-forming processes (e.g. Groves et al., 1998; Witt and Vanderhor, 1998; Bierlein et al., 2006); however, pentlandite- bearing quartz veins in the area are not well-studied and consequently have not been characterized. Unfortunately, the complex geologic history at the Kambalda Dome poses uncertainty regarding the tectonic control, fluid source and physicochemical processes that are responsible for development of these quartz veins. The aim of this study is to investigate the pentlandite-bearing quartz veins by focusing on the mineralogical, geochemical and textural characterization of this vein type. Microanalytical work was used to constrain the timing of vein formation and to better understand the processes responsible for Ni ore formation.

2.2 Geological Setting The Kambalda Dome is a component of the Kambalda Nickel Field which is located in the Archean Yilgarn Craton of Western Australia. The abundance of economic mineral deposits that has been identified in the craton has led to abundant geologic research of the area, including the reviews of the recent summary on Western Australian ore deposits edited by Phillips (2017b). The following section introduces the geologic setting of the Kambalda Dome and the samples that were collected along the eastern flank of the Dome.

Yilgarn Craton The Archean Yilgarn Craton of Western Australia is a highly mineralized area predominantly composed of gneisses, granitoids and greenstone sequences with an exposed area of about 657,000 km2 (Fig. 2.1; Trendall, 1990; Wyche and Wyche, 2017). The granitoids and granitic gneisses are the dominant rock types in the craton, constituting at least 70% of the preserved crust (Wyche and Wyche, 2017). The greenstone belts are unevenly distributed throughout the Yilgarn Craton and are most abundant in the eastern part of the region (Wyche and Wyche, 2017). The evolution of the Yilgarn Craton stabilized prior to 2400 Ma and has since undergone intense weathering (Wyche and Wyche, 2017 and references therein). The Yilgarn Craton is divided into seven terranes (Fig. 2.1) based on geochemistry, age, as well as sedimentary and magmatic characteristics (Wyche and Wyche, 2017). The western part of the craton is constituted by the Narryer Terrane, the Youanmi Terrane and the South West Terrane (Fig. 2.1). The eastern part of the craton is composed of terranes which constitute the Eastern Goldfields Superterrane (Fig. 2.1). The Eastern Goldfields Superterrane is constituted by the following terranes that are separated by east-dipping, large-scale, crust-penetrating shear zones: the Kurnalpi Terrane, the Burtville Terrane and the Kalgoorlie Terrane (Wyche and Wyche, 2017 and references therein).

Eastern Goldfields Superterrane, Yilgarn Craton Characteristics of the Eastern Goldfields Superterrane have been studied to understand the abundant economically occurring deposits in the area, especially Archean lode gold deposits which are largely considered to be associated with orogenic ore forming processes (e.g. Groves et al., 1998; Witt and Vanderhor, 1998; Bierlein et al., 2006). However, the Eastern Goldfields

Superterrane has been recognized to also be a large source of nickel, particularly in the Kambalda Nickel Field of the Kalgoorlie Terrane (Fig. 2.2; Gresham and Loftus-Hills, 1981; Stone et al., 2005; Lesher and Barnes, 2008; Barnes and Fiorentini, 2012).

Figure 2.1. Subdivision of the Archean Yilgarn Craton in southwest Western Australia. Figure after Goscombe et al. (2009).

The Kambalda Nickel Field, Eastern Goldfields Superterrane, Yilgarn Craton The Kambalda Nickel Field contains abundant komatiite-associated magmatic sulfide ore deposits which includes the Ni-rich orebodies at the Kambalda Dome, St. Ives and the Tramways Dome (Fig. 2.2; Gresham and Loftus-Hills, 1981). In addition to magmatic sulfides, St. Ives contains abundant gold deposits which are actively mined, producing over 400 t of gold since 1980 (Oxenburgh et al., 2017). The magmatic sulfide orebodies at the Kambalda Nickel Field contain varying amounts of pyrrhotite, pentlandite and pyrite with lesser amounts of chalcopyrite and locally occurring millerite (Gresham and Loftus-Hills, 1981). Development of ore bodies are

25 attributed to Archean komatiitic volcanic flows and subsequent gravity separation of sulfide-rich liquid (Gresham and Loftus-Hills, 1981 and references therein). Such volcanic activity, in combination with sedimentary deposition, is responsible for the development of the main stratigraphic units of the Kambalda Nickel Field which include: (1) a primary magmatic ore- hosting sequence of mafic-ultramafic volcanic rocks (the Kambalda Sequence); (2) an overlying sedimentary layer of felsic volcanics, graywackes and argillites (Black Flag Group); (3) and an uncomformably overlain polymictic conglomerate called the Merougil Beds, among other sedimentary units like the Lake Dam and Triangle Island Beds (Gresham and Loftus-Hills, 1981; Miller et al., 2010; Mueller et al., 2016). The volcanic, volcaniclastic and sedimentary rocks are overlain by a sequence of mafic-ultramafic volcanic rocks called the Bluebush Sequence (Gresham and Loftus-Hills, 1981). The volcanic and sedimentary units of the Kambalda Nickel Field have suffered many geological events in the Archean which include multiple intrusive, metamorphic and deformational events (Gresham and Loftus-Hills, 1981). The rocks of the Kalgoorlie-Kambalda area experienced five phases of orogenic deformation, after 2730-2690 Ma back arc rifting (Mueller et al., 2016). In addition to the multitude of events that have affected the rocks of the Kambalda Nickel Field, the following metasomatic replacement occurrences have altered the mineralogy and chemistry of rocks: serpentinization, carbonatization and potassium metasomatism (Gresham and Loftus-Hills, 1981). The complex geologic history of the Kambalda area has resulted in intrusion emplacement, altered mineral assemblages and convoluted geometries in stratigraphy, including the dome complexes in the Kambalda Dome, St. Ives, Tramways and Democrat areas (Gresham and Loftus-Hills, 1981; Fig. 2.2). The Kambalda Nickel Field is also host to various faults and structures with the most prominent structures being highlighted in Figs. 2.2 and 2.3. On the west, the St. Ives goldfield is truncated by the Speedway fault (Miller et al., 2010), also called the Merougil Fault in Gresham and Loftus-Hills (1981), shown in Fig. 2.2. At the southeastern part of the Kambalda Dome is the Alpha Island Fault (e.g. Stone et al., 2005; Staude et al., 2017b), which is a late-stage NNE- trending fault with dextral offset (Miller et al., 2010) shown in Fig. 2.3. On the eastern side of the Kambalda Nickel Field, bounding the St. Ives Goldfield and the Kambalda Dome, is the NNW-trending regional Boulder-Lefroy Fault (Figs 2.2. and 2.3). Major crustal thickening was accommodated through ductile deformation along the Boulder-Lefroy Fault (Miller et al., 2010).

The Boulder-Lefroy regional fault likely originated during the early development of the Eastern Goldfields Superterrane as it parallels the NNW orientation trend that defines the western margin of the Superterrane (Miller et al., 2010). This structure is a regional first order fault with higher order fault splays such as the second order Playa Fault at St. Ives (Miller et al., 2010). The Boulder-Lefroy Fault is significant because many workers have highlighted the controls which this fault has had on major gold systems through localization of gold deposits along dilational and contractional jogs, having a sinistral-slip component during major gold mineralization (Miller et al., 2010 and references therein). This fault has been compared to the Larder Lake-Cadillac deformation zone in the southwest Abitibi as it bears many similarities to the structure (Weinberg et al., 2005 and references therein). From south to north, this fault hosts gold deposits in the following important gold districts: St. Ives, Hampton-Boulder-Jubilee, Golden Mile and Mt. Charlotte (Kalgoorlie goldfield), and Paddington-Broad Arrow (Weinberg et al., 2005). Ultimately, intense mineralization within and around the shear zone, such as the deposits of Kalgoorlie, were formed from efficient, multiscale focusing of mineralizing fluids into the Boulder-Lefroy structure (Weinberg et al., 2005). At a closer scale, mineralizing fluids were focused into regional-scale anticlinal structures, such as the Kambalda anticline (Fig. 2.3) of the Kambalda Dome (Weinberg et al., 2005). Gold mineralization along the Boulder-Lefroy is believed to have formed late in the tectonic history of the Yilgarn craton between 2.64-2.63 Ga (Weinberg et al., 2005).

The Kambalda Dome, Kambalda Nickel Field, Eastern Goldfields Superterrane, Yilgarn Craton The field area for this study is the Kambalda Dome (Figs. 2.2 & 2.3) of the Kambalda Nickel Field, which is a major tectonic structure developed through polyphase tectonism and metamorphism (Cowden and Roberts, 1990; Stone et al., 2005). The generalized stratigraphic components of the Kambalda Sequence of the Kambalda Nickel Field are observed at the Kambalda Dome and are named as follows: Lunnon basalt, Silver Lake komatiite, Tripod Hill komatiite, Devon Consols basalt, Kapai slate and Paringa basalt (Fig. 2.4). The Lunnon basalt is overlain by up to 10 m of sedimentary units that are either chlorite-rich sedimentary rocks or sulfide-rich chert (Gresham and Loftus-Hills, 1981; Staude et al., 2017b). The intrusive activity of the region is represented at the Kambalda Dome by felsic, intermediate and less-common mafic intrusions (Gresham and Loftus-Hills, 1981; Staude, 2015). The most prominent intrusion

27 is the 2660 ± 4 Ma (U-Pb dating on zircon; Cassidy et al., 1991) coarse-grained, weakly porphyritic trondhjemite positioned at the core of the Kambalda Dome (Gresham and Loftus- Hills, 1981; Fig. 2.3).

Figure 2.2. Map of the regional geology of the Kambalda area located in Western Australia adapted after Gresham and Loftus-Hills (1981).

Figure 2.3. Generalized geologic map of the Kambalda Dome modified after Stone et al. (2005) and Staude et al. (2017b).

After the Archean development of the volcanic, sedimentary and intrusive rocks of the Kambalda Dome, the rocks of this area were deformed into a dome-shaped geometry that is steeply dipping or overturned on the east flank and moderately dipping on the west flank of the Dome (Stone et al., 2005). This dome is a part of the regional Kambalda anticline and is

29 structurally truncated to the east by the regional Boulder-Lefroy Fault Zone (Fig. 2.3; Stone and Archibald, 2004; Stone et al., 2005). As mentioned earlier, this fault zone is regionally auriferous as it hosts abundant world-class gold deposits such as the nearby St. Ives and Kalgoorlie goldfields (Weinberg et al., 2005). Stone et al. (2005) detailed the deformational events that are recorded in the rocks of the Kambalda Dome, and suggested its development was a result of four deformational events (Table 2.2). In addition to these deformational events, the Kambalda Dome also underwent upper greenschist-lower amphibolite facies regional metamorphism (Stone et al., 2005 and references therein), and magmatic intrusive events of various ages evident through cross-cutting features (Staude, pers. comm.). Ocean floor metamorphism (hydration) has altered komatiite to serpentine-dominated assemblages and metamorphism has produced talc and carbonate mineral assemblages (Cowden and Roberts, 1990; Stone et al., 2005). It is believed that the Yilgarn Craton and its components, which includes the Kambalda Dome, have been essentially stable since its Archean deformation which occurred before 2.4 Ga (Trendall, 1990). Magmatic sulfides at the Kambalda Dome are formed from Archean komatiitic volcanic activity. The Lunnon basalt serves as the immediate footwall to the Ni deposits (Fig. 2.4) which occurs as massive sulfides or matrix/net-textured sulfides (Staude et al., 2017a). The Silver Lake komatiite, which overlies the Lunnon basalt, hosts the Ni deposits in channelized areas of komatiitic lava flow (Staude et al., 2017a). The deposit model for these magmatic sulfides invokes the formation of such deposits by the presence of a fertile magma which reaches sulfide saturation (prior to extensive olivine crystallization), forming a sulfide liquid that is chemically enriched in metals and is physically transported and accumulated (Le Vaillant et al., 2018). At Kambalda, Ni deposits accumulated in linear trough-shaped embayments, on the top surface of the footwall basalt, which formed through thermomechanical erosion (Staude et al., 2017a). The primary metal source for these deposits is a magma which has greater fertility with greater degrees of partial melting of the mantle source that generated the magma (Le Vaillant et al., 2018). At Kambalda, the metal source is the komatiitic lava flow that is associated with and hosts the Ni deposits. Lastly, it is imperative for thermomechanical assimilation of an external S source to occur during magmatic emplacement in order for sulfide liquids to form and so that chalcophile elements in the magma can partition into the sulfide droplets (Le Vaillant et al., 2018). At Kambalda, this external S source is the pyrrhotite-bearing, sulfide-rich chert sedimentary unit (Fiorentini et al., 2012) which consists of an average of 15-20% sulfides

(Gresham and Loftus-Hills, 1981). At 1050⁰C, solidification of sulfides was complete and at temperatures below 600⁰C, an Fe-Ni-S monosulfide solid solution segregated into mineral phases including pentlandite and pyrrhotite (Marston et al., 1981 and references therein). In summary, the magmatic sulfides hosted in the Silver Lake komatiite (Fig. 2.4) formed during the crystallization of the komatiitic lava which formed at 2.7 Ga (Pb-Pb dating; Chauvel et al., 1985). After the formation of these deposits, a majority of the Ni deposits at Kambalda were overprinted by metamorphism, younger faults and magmatic dikes which have resulted in multiple ore lenses of single deposits and mechanical remobilization of sulfides into surrounding rocks (Staude et al., 2017a).

Figure 2.4. Volcanic stratigraphy of the Kambalda Sequence, the major stratigraphic components of the Kambalda Dome modified after Gresham and Loftus-Hills (1981).

Table 2.2. Deformational history of the Kambalda Dome, developed after the proposed deformational history in Stone et al. (2005). Event Deformation that occurred

D1 South- to north- directed thrusting and distributed deformation in incompetent lithotypes

D2 Inclined thrust-related folding; peak regional deformation event; south- southwest to north-northeast directed thrust-folding Doming and faulting; related to regional normal and reverse faults in some D3 areas; deformation event coincided with the main felsic trondhjemite intrusion, the regional peak metamorphism and the formation of the Kambalda Dome geometry

D4 Oblique strike-slip faulting; dextral shearing on steep, north to northeast trending faults and fault zones

The Kambalda Dome has been continuously mined since the discovery of these deposits in 1966 (Hronsky and Schodde, 2006; Staude, 2017) serving as an economic source of nickel. A single mining company, Independence Group, produced 3.0 Mt of ore containing about 116 kt of nickel metal with averaging grades of 3.9% Ni from 2002 to 2015 (Staude et al., 2017b). Throughout mining operations of the magmatic komatiite-hosted ore bodies of the Kambalda Dome, igneous textures like pillow-structures and flow-top breccia host late minerals localized along permeable sites of such structures (including quartz, carbonate, epidote, scheelite, anhydrite and sometimes garnet) of the Lunnon basalt (Staude, 2015). However, there are also veins that are clearly related to fracturing, which contain quartz and/or carbonate minerals (Staude, 2015). Some of these quartz veins observed in the Kambalda Dome, contain pentlandite and intergrown biotite (Staude, 2015). Other veins have been observed to contain abundant euhedral pyrite, some of which have been assayed and evidently can contain up to 8 ppm of irregularly distributed Au (Staude, 2015). Unfortunately, the various types of quartz veins, including the pentlandite-bearing and auriferous quartz veins, at the Kambalda Dome have not been classified, therefore there is little known about their mineralogy and origin.

2.3 Materials and Methods A total of 12 samples were collected for this study by Dr. Sebastian Staude. Sample localities are highlighted in Fig. 2.5A. Nine of the samples are quartz veins which are categorized into the following quartz vein type: pentlandite-bearing, auriferous and barren quartz veins. Two samples were collected from a magmatic sulfide orebody and one sample was collected from one of the host rocks of a mineralized vein. Table 2.3 denotes the sample numbers, sample type and the location, in the Kambalda Dome, at which each sample was collected.

Figure 2.5. (A) Generalized geologic map of the Kambalda Dome denoting sample locations with black stars, modified after Stone et al. (2005) and Staude et al. (2017a). Underground mine wall photographs, taken by Dr. Sebastian Staude, of (B) the Long North vein that samples IS13- 7-2 and IS13-7-2Q were collected from and (C) the McLeay vein systems in sample S-460-2.

Pentlandite-bearing quartz veins Six samples of pentlandite-bearing quartz veins were collected. Samples IS13-7-2 and IS13-7-2Q from Long North were collected from the vein shown in Fig. 6A, hosted in the 2660 ± 4 Ma (U-Pb dating on zircon; Cassidy et al., 1991) trondhjemite positioned at the core of the

Kambalda Dome. Sample S-460-2 from McLeay shows two generations of quartz veins, both containing pentlandite. The sample locality is displayed in Fig. 6B, showing two veins systems that are hosted in the 2.7 Ga Lunnon basalt (Pb-Pb dating; Chauvel et al., 1985) unit of the Kambalda Sequence. The remaining three samples were collected from Moran. Vein sample S- Moran-2 was collected as a float sample in an underground stope and is hosted in the Lunnon basalt. Sample S-LSU162-1 was collected from an exploration drill hole and was intersected before the magmatic sulfide orebody was intersected further down depth. This vein is hosted in komatiite and an intermediate-composition dike.

Table 2.3. Sample number and sample type with the location at which each sample was collected from. An 'X' is placed next to each sample if the analytical technique, denoted at the column header, was performed on the sample. Key: CL = Cathodoluminescence microscopy, EPMA = electron probe microanalysis, µ-XRF = micro X-ray fluorescence analysis.

Sample µ- Sample Type Location Petrography CL EPMA Number XRF

Pentlandite-bearing IS13-7-2 Long North X X X X quartz vein Pentlandite-bearing IS13-7-2Q Long North X X X X quartz vein Pentlandite-bearing S-460-2 McLeay X X X X quartz vein Pentlandite-bearing S-Moran-2 Moran X X X X quartz vein Pentlandite-bearing S-LSU162-1 Moran X X X X quartz vein 500 m north of S-LSU373A-1 Auriferous quartz vein X X X X Moran S-137-7 Auriferous quartz vein Long North X X X X S-137-6 Barren quartz vein Long North X X X X S-137-1 Magmatic sulfides Long North X X S-157-1 Magmatic sulfides Long North X IS-1569 Trondhjemite hostrock Long North X

Barren and auriferous quartz veins One barren quartz vein (S-137-6) from Long North was collected for this study and was sampled from a mine wall, hosted in 2.7 Ga Kambalda komatiite (Pb-Pb dating; Chauvel et al., 1985), along a granitic dike (Fig. 2.6A). Two samples of auriferous quartz veins were collected for this study. Sample S-137-7 (Long North) was collected from a mine wall with the vein being hosted in the 2660 ± 4 Ma (U-Pb dating on zircon; Cassidy et al., 1991) trondhjemite, shown in Fig. 2.6B. Sample S-LSU373A-1 was collected from a drill hole that was collared 500 meters north of Moran with the vein being hosted in Lunnon basalt.

Figure 2.6. Underground mine wall photographs, taken by Dr. Sebastian Staude, of (A) the barren quartz vein that was sampled for this study (Long North) and (B) the anastomosing vein system from which sample S-137-7 was collected from (Long North).

Magmatic sulfides and trondhjemite hostrock Two magmatic sulfide samples were collected from the Long North massive and disseminated sulfide orebody. The trondhjemite hostrock sample was collected from near the Long North orebody, near the orebody-trondhjemite contact. When collected, the trondhjemite exposed at the mine wall did not appear to be altered by quartz veins (Staude, pers. comm.). However, it is possible for the trondhjemite to have been altered as there are reports of the Long North magmatic sulfide orebody, near the trondhjemite, being altered to millerite, magnetite and pyrite (Staude, 2015).

Analytical Methods The barren, auriferous and pentlandite-bearing quartz vein samples, denoted in Table 2.3, were prepared as thin sections at the Colorado School of Mines Sample Preparation Lab. These thin sections were used for detailed petrography, scanning electron microscopy techniques, micro X-ray fluorescence analysis, and electron probe microanalyses. The magmatic sulfide samples (samples S-137-1 and S-157-1 from Long North) were provided already prepared as thin sections, from Dr. Sebastian Staude. One epoxy mount of trondhjemite rock chips (sample IS- 1569 from Long North) was prepared and analyzed in automated mineralogy.

Field Emission-Scanning Electron Microscopy The quartz vein samples were analyzed in the electron beam laboratory in the Department of Geology and Geological Engineering at the Colorado School of Mines in Golden, Colorado. The laboratory is equipped with a TESCAN MIRA3 LMH Schottky field emission scanning electron microscope (FE-SEM). The FE-SEM features a TESCAN motorized retractable annular, single-crystal YAG backscattered electron (BSE) detector and a Bruker XFlash 6/30 silicon drift detector for energy dispersive x-ray spectrometry (EDS). BSE imaging and EDS analyses were performed at a 15 keV or 20 keV acceleration voltage, a working distance of 10 mm and a beam intensity of 11.

Optical Cathodoluminescence Microscopy Cathodoluminescence (CL) imaging of quartz vein samples was performed using a HC5- LM hot-stage CL microscope by Lumic Special Microscopes, at the Fluid Inclusion and CL Laboratory of the Department of Geology and Geological Engineering at the Colorado School of Mines. The system was operated at an acceleration voltage of 14 kV and a current density of about 10 µA/mm2 (Neuser, 1995). Images of the optical luminescence were captured with a high sensitivity, double-stage Peltier cooled Kappa DX40C CCD camera.

Automated Mineralogy Quartz vein samples and the trondhjemite sample were analyzed using automated scanning electron microscopy in the electron beam laboratory of the Department of Geology and Geological Engineering at the Colorado School of Mines. The automated mineralogy system that

36 was employed is the TESCAN Integrated Mineral Analyzer (TIMA). The system is equipped with four energy dispersive X-ray (EDX) detectors, one BSE detector, a TESCAN VEGA SEM and the TIMA analytic software suite. Overview scans of each of the vein samples and the trondhjemite hostrock were produced with a beam stepping interval (spacing between acquisition points) of 15 µm, an acceleration voltage of 25 keV, a beam intensity of ~14.00, and a working distance of 15 mm. High resolution scans (with a beam stepping interval of 3 µm) of selected areas in two of the pentlandite-bearing quartz veins (samples S-460-2 and IS13-7-2Q) were performed. The collected EDX and BSE data were used for mineral assignment, matched by the TIMA software, through comparison with a library of mineral definitions.

Electron microprobe analysis Major, minor and trace chemical composition of sulfides, biotite, potassium feldspar and albite were determined through the use of a JEOL Superprobe JXA–8900RL at the University of Tübingen in Germany. The analysis of sulfides, biotite and feldspars were performed in three distinct electron probe microanalysis (EPMA) sessions, each with their particular set of conditions. EPMA Session 1 was used to analyze sulfides: pentlandite, pyrrhotite, chalcopyrite and pyrite; whereas EPMA Session 2 was used to analyze biotite and EPMA Session 3 was used to analyze feldspars. The element and oxide energy lines measured, corresponding detection limits (1σ), set up conditions (beam diameter, acceleration voltage, beam current and x-ray counting times) and the crystals used for each session are noted in Tables 2.4, 2.5 and 2.6. Natural and synthetic standards were used for the calibration of each EPMA session and are also noted in Tables 2.4, 2.5 and 2.6. The matrix correction used for data attained is the Phi-Rho-Z method which was applied through the JEOL software used. During Session 1, energy peak overlaps were addressed. The peak width for S was limited to a range where it no longer overlapped with Co. Additionally, a peak overlap correction of 0.001126 for Fe and As was used.

Fluid inclusion investigations Fluid inclusion petrography was performed at the Fluid Inclusion and CL Laboratory at the Department of Geology and Geological Engineering at the Colorado School of Mines. The

Table 2.4. EPMA Session 1 on sulfides with a focused beam diameter at 25 kV acceleration voltage and 20 nA beam intensity. Elemen Crysta Standard Energ Coun Backgroun Beam Peak Detectio t l y Line t d Count Diamete positio n Limits Time Time (s) r n (mm) (1σ) (s)

As TAP A_GaAs25 La 30 15 Focused 105.14 275 ppm 5 S PETJ M_FeS225 Ka 16 8 Focused 172.04 100 ppm 6 Bi LIF M_Bi25 La 30 15 Focused 79.119 500 ppm Zn LIFH M_Zn25 Ka 30 15 Focused 100.25 100 ppm 7 Se TAP M_PbSe25 La 30 15 Focused 97.717 200 ppm Pb PETJ A_Galena2 Ma 30 15 Focused 169.19 600 ppm 5 2 Cu LIF M_Cu25 Kb 30 15 Focused 96.560 700 ppm Co LIFH M_Co25 Ka 30 15 Focused 124.74 100 ppm 6 Ag PETJ M_Ag25 La 30 15 Focused 133.05 300 ppm 8 Ni LIF M_Ni25 Ka 30 15 Focused 115.16 125 ppm 7 Fe LIFH M_FeS225 Ka 16 8 Focused 134.92 100 ppm 9 Te PETJ M_PbTe25 La 30 15 Focused 105.35 450 ppm 8 Cd PETJ M_CdS25 La 30 15 Focused 126.72 450 ppm 5 Pd PETJ M_Pd25 La 30 15 Focused 139.86 375 ppm 0

Table 2.5. EPMA Session 2 on biotite with a 2 µm beam diameter at 15 kV acceleration voltage and 20 nA beam intensity. Element/oxi Cryst Standard Energ Cou Backgrou Beam Peak Detecti de al y nt nd Count Diamet positio on Line Time Time (s) er (µm) n Limits (s) (mm) (1σ) F LDE1 Topaz_Utah Ka 30 15 Focuse 86.978 350 d ppm Na2O TAP A_Plagiocla Ka 30 15 10 129.73 150 se 1 ppm K2O PETJ A_Sanidine Ka 16 8 2 119.76 150 4 ppm Cl PETH A_Tugtupit Ka 30 15 5 151.61 100 e 1 ppm MgO TAP A_Diopside Ka 16 8 Focuse 107.77 200 d 5 ppm CaO PETJ A_Diopside Ka 30 15 Focuse 107.48 200 d 8 ppm TiO2 PETH A_SrTiO3 Ka 30 15 Focuse 88.677 200 d ppm Al2O3 TAP A_Plagiocla Ka 16 8 10 90.935 200 se ppm Cr2O3 PETJ M_Cr Ka 30 15 Focuse 73.032 200 d ppm BaO PETH A_Barite La 30 15 5 89.486 200 ppm SiO2 TAP A_Diopside Ka 16 8 Focuse 77.756 300 d ppm MnO PETJ A_Rhodonit Ka 30 15 Focuse 66.905 200 e d ppm FeO LIF A_Hematite Ka 16 8 Focuse 134.59 350 d 0 ppm NiO LIF M_Ni Ka 30 15 Focuse 115.13 300 d 9 ppm

Table 2.6. EPMA Session 3 on feldspar with a 2 µm beam diameter at 15 kV acceleration voltage and 20 nA beam intensity. Oxid Crysta Standard Energ Coun Backgroun Beam Peak Detectio e l y Line t d Count Diameter positio n Limits Time Time (s) (µm) n (mm) (1σ) (s)

Na2O TAP A_Plagioclas Ka 16 8 10 129.76 200 ppm e 7 K2O PETJ A_Sanidine Ka 16 8 2 119.76 150 ppm 4 CaO PETJ A_Diopside Ka 30 15 Focused 107.48 200 ppm 8 TiO2 PETH A_SrTiO3 Ka 30 15 Focused 88.677 200 ppm

Al2O TAP A_Plagioclas Ka 16 8 10 90.927 250 ppm 3 e BaO PETH A_Barite La 30 15 5 89.486 200 ppm

SiO2 TAP A_Plagioclas Ka 16 8 10 77.748 400 ppm e SrO PETH A_SrTiO3 La 30 15 Focused 219.76 550 ppm 5 FeO LIF A_Hematite Ka 30 15 Focused 134.59 250 ppm 0

laboratory is equipped with an Olympus BX51 microscope that was used for fluid inclusion petrographic studies. Fluid inclusion microthermometry was performed on 300 µm thick doubly polished thick sections of quartz veins, using a Leica DMLP microscope equipped with a Linkam TMS-9 cooling-heating stage at the University of Tübingen. The heating-freezing stage was calibrated through measurement of synthetic fluid inclusions, particularly measuring the triple point of CO2

(-56.6°C), the melting point of pure H2O (±0.0°C) and the critical point of pure H2O (+374.1°C). In order to attain proper microthermometric data, a cooling/heating rate of 0.1°C/minute was employed when recording phase change temperatures.

Micro X-ray Fluorescence Micro X-ray fluorescence (µ-XRF) analyses of the quartz vein samples were done using the high performance spectrometer Bruker© M4 Tornado. The measurements were performed in

40 a vacuum through the X-ray excitation by a high brilliance X-ray tube focused with polycapillary X-ray optics. The target material used to produce the X-rays, during excitation, is Rh with a spot size of 25 µm with the polycapillary lense. The XFlash® silicon drift detector is used and has an energy resolution of < 145 eV at 300,000 counts per second.

Thermodynamic Modelling The thermodynamic modelling was performed on Phase 2 and React software programs of the Geochemists Workbench (GWB) software suite. The Log K values of pentlandite at temperatures between 0-300℃ were calculated and stored in the database at 25℃ intervals. The Log K values were calculated based on thermodynamic data of a stoichiometric pentlandite

[(Ni4.5 Fe4.5)S8] experimentally determined by Klein and Bach (2009) and Berezovskii et al. (2001).

2.4 Results Petrographic work used to characterize samples mineralogically and texturally included optical microscopy, FE-SEM work and automated mineralogy. Before any SEM-based techniques were employed, CL imaging was performed to preserve any short-lived CL signatures that quartz may display. Following CL imaging and petrographic work, EPMA was employed to attain mineral chemical data. Micro X-ray fluorescence was used to supplement EPMA data attained.

2.4.1 Petrography The three types of quartz veins, the trondhjemite hostrock and the samples from the magmatic sulfide orebodies have different mineralogical and textural features that are evident through optical microscopy, automated mineralogy and scanning electron microscopy. The mineralogical and textural features are described through results attained from transmitted light optical microscopy, reflected light optical microscopy, FE-SEM analyses and automated mineralogy. There is a shared characteristic across the three quartz vein types which is the textural features of quartz. Quartz in all vein samples analyzed is anhedral and lobate with irregular grain boundaries (Fig. 2.7). The quartz grains depict undulose extinction and are variably sized from

41 micrometer-sized to centimeter-sized grains. In addition to these features, many of the quartz grains in veins are elongated. The irregular quartz grain textures are similar to images of bulging, recrystallization, subgrain rotation recrystallization and grain boundary migration recrystallization, illustrated in Passchier and Trouw (2005).

Figure 2.7. Representative transmitted cross-polarized light photomicrographs of quartz in quartz veins, depicting irregular grain boundaries, lobate shapes, variable grain size and undulose extinction in samples: A) Pentlandite-bearing quartz vein from Long North (IS13-7-2), B) Pentlandite-bearing quartz vein from Long North (IS13-7-2Q), C) Pentlandite-bearing quartz vein from McLeay (S-460-2), D) Pentlandite-bearing quartz vein from Moran (S-Moran-2), E) Pentlandite-bearing quartz vein from Moran (S-LSU162-1), F) Auriferous quartz vein from 500 m north of Moran ( S-LSU373A-1), G) Auriferous quartz vein from Long North (S-137-7) and H) Barren quartz vein from Long North (S-137-6).

Pentlandite-bearing quartz veins All pentlandite-bearing quartz veins observed in thin section share some mineralogical and textural features. The first shared characteristic is the intergrowth of pentlandite with

42 pyrrhotite (Fig. 2.8A). These two sulfide phases are replaced by late pyrite which is intergrown with fine-grained irregular chalcopyrite (Figs 2.8B and C). The fine pyrite-chalcopyrite intergrowth dominantly occurs in the center of the pyrite grains (Fig. 2.8B). The pentlandite and pyrrhotite are also intergrown with minor to trace occurrences of anhedral melonite (NiTe2) grains. Lastly, pyrrhotite regularly displays kinking and twinning (Fig. 2.8D).

Figure 2.8. Representative photomicrographs and backscattered electron (BSE) images of textures observed in pentlandite-bearing quartz veins. A) BSE image of pyrrhotite and pentlandite intergrowth in sample IS13-7-2 (Long North). B) Reflected plane-polarized light image of pyrite intergrown with chalcopyrite in sample IS13-7-2 (Long North). C) Reflected plane-polarized light image of pyrite intergrown with chalcopyrite in sample IS13-7-2Q (Long North). D) Reflected cross-polarized light image of twinning in pyrrhotite of sample IS13-7-2 (Long North).

IS13-7-2 – Pentlandite-bearing quartz vein from Long North Sample IS13-7-2 represents a pentlandite-bearing quartz vein and trondhjemite hostrock (Long North; Fig. 2.9). The pentlandite-bearing quartz vein is dominantly composed of euhedral biotite, anhedral pentlandite, anhedral pyrrhotite, anhedral late pyrite and subhedral to anhedral calcite with minor amounts of albite, quartz and potassium feldspar. Trace amounts of melonite occur intergrown with pentlandite and pyrrhotite. Biotite dominantly occurs as large euhedral

43 blades and is intergrown with pyrrhotite and pentlandite (Fig. 2.10A), apparent in the automated mineralogy scan in Fig. 2.9. Biotite contains inclusions of albite and is commonly kinked. Biotite also occurs as smaller euhedral grains intergrown with anhedral albite, quartz, potassium feldspar, pentlandite and pyrrhotite which is an assemblage that is observed in spatial association with calcite (Figs. 2.10B and C). Anhedral late pyrite is a major component of the vein, as the pyrite replaces pentlandite and pyrrhotite. There are two types of calcite: anhedral crack-filling calcite and coarse subhedral-anhedral calcite grains. The anhedral crack-filling calcite appears to fill cracks that are associated with the replacement of pyrrhotite and pentlandite by late pyrite. Coarse calcite grains occur as subhedral to anhedral grains and are spatially associated with the biotite-albite-pentlandite-pyrrhotite-potassium feldspar assemblage (Fig. 2.10B and 2.10C). Anhedral zones of Mg- and Fe-rich calcite and inclusions of ankerite and strontianite can be observed in coarse calcite (Figs. 2.10D).

Figure 2.9. Automated mineralogy scan of pentlandite-bearing quartz vein sample IS13-7-2 from Long North. The vein and trondhjemite hostrock portions of the sample are labeled.

The trondhjemite hostrock in sample IS13-7-2 (Long North) is dominantly composed of partially sericitized albite, anhedral pyrrhotite and anhedral pentlandite with minor anhedral late pyrite, anhedral quartz, anhedral potassium feldspar and anhedral rounded calcite. Trace amounts of galena, apatite, biotite, melonite, chlorite and muscovite are also present. Euhedral sericitized albite is coarse grained and is overgrown by an albite rim (Fig. 2.11A). Albite composes the majority (~80%) of the hostrock which is evident in the automated mineralogy scan of Fig. 2.9.

Figure 2.10. Photomicrographs and BSE images of textures observed in the pentlandite bearing quartz vein sample IS13-7-2 from Long North. A) BSE image of pentlandite with biotite. B) Transmitted cross-polarized light image of pyrrhotite, biotite and albite in spatial association with calcite. C) Transmitted cross-polarized light image of fine-grained biotite, albite and quartz in spatial association with calcite. D) BSE image of calcite with irregular zones of Mg- and sometimes Fe-rich calcite and with inclusions of strontianite.

Albite is often spatially associated with anhedral irregular potassium feldspar grains which typically rim pentlandite and pyrrhotite. Albite and biotite in the hostrock are commonly intergrown with pyrrhotite or pentlandite (Fig. 2.11B). Potassium feldspar may also occur as euhedral-subhedral grains intergrown with fine-grained biotite, pentlandite and pyrrhotite (Fig. 2.11C). Small amounts of fine-grained biotite occurring in the hostrock are commonly associated with muscovite. Biotite is dominantly chloritized to varying degrees and sometimes pentlandite and pyrrhotite is observed replacing biotite (Fig. 2.11D). Wallrock-hosted calcite is different than vein-hosted calcite as wallrock-hosted calcite dominantly occurs as round grains and occurs as inclusions in the sericitized euhedral albite. Apatite occurs as subhedral, round grains and are spatially associated with pentlandite, pyrrhotite and late pyrite.

Figure 2.11. Photomicrographs and BSE images of textures observed in the hostrock of sample IS13-7-2 from Long North. A) Transmitted plane-polarized light image of euhedral sericitized albite being rimmed by albite. B) BSE image of albite and biotite intergrown with pyrrhotite of the wallrock C) BSE image of potassium feldspar intergrown with biotite and pentlandite. D) BSE image of biotite being replaced by pentlandite.

IS13-7-2Q – Pentlandite-bearing quartz vein from Long North Sample IS13-7-2Q was collected from the same vein as sample IS13-7-2 (Long North pentlandite-bearing quartz vein) and therefore has many mineralogical and textural similarities compared to sample IS13-7-2 from Long North. Sample IS13-7-2Q is also composed of a vein segment and a segment of the trondhjemite hostrock apparent in the automated mineralogy scan in Fig. 2.12. Sample IS13-7-2Q contains a greater variety of trace minerals including tetradymite, galena, zircon, parisite, melonite and apatite in the vein compared to sample IS13-7-2 (Long North pentlandite-bearing quartz vein). Moreover, sample IS13-7-2Q contains quartz intergrown

46 with cm-sized biotite, evident in Fig. 2.12. Similar to IS13-7-2, biotite is kinked and contains inclusions of albite. Finer-grained biotite is also present in the vein in the same albite-quartz- biotite-potassium feldspar-pentlandite-pyrrhotite assemblage that is observed in sample IS13-7- 2. The two types of calcite, noted in IS13-7-2, are also present in this sample: calcite in fractures and coarse subhedral-anhedral calcite. The two calcite types occur in the same manner as described in IS13-7-2, however in this sample one grain of coarse-grained calcite is spatially associated with tetradymite and chalcopyrite (Fig. 2.13A). A final major difference, with sample IS13-7-2, is the occurrence of zoned potassium feldspar grains in the vein (Fig. 2.13B). The mineralogical and textural features of the trondhjemite hostrock of this sample are the same as observed and described for sample IS13-7-2.

Figure 2.12. Automated mineralogy scan of pentlandite-bearing quartz vein sample from Long North (IS13-7-2Q). The vein and trondhjemite hostrock portions of the sample are labeled.

S-460-2 – Pentlandite-bearing quartz vein from McLeay Sample S-460-2 consists of two pentlandite-bearing quartz veins in a basaltic hostrock. The two veins have been assigned names to describe the veins’ respective mineralogical and textural features. The pentlandite-bearing quartz vein on the left side of the McLeay mine wall photograph in Fig. 2.5C is called ‘Upper Vein’. The pentlandite- and scheelite-bearing quartz vein system on the right side of the McLeay mine wall photograph in Fig. 2.5C is called ‘Scheelite Vein’. The Upper Vein and Scheelite Vein are labeled in the automated mineralogy scan in Fig. 2.14.

Figure 2.13. BSE images of textures observed in the vein segment of sample IS13-7-2Q from Long North. A) High-resolution, 3 µm stepping size, automated mineralogy scan of anhedral calcite spatially associated with chalcopyrite and tetradymite. B) BSE image of zoned coarse potassium feldspar with inclusions of zircon, chlorite and muscovite.

Figure 2.14. Automated mineralogy scan of pentlandite-bearing quartz vein sample S-460-2 from McLeay. The two veins and the basaltic hostrock in the sample are labeled.

The Upper Vein is composed of quartz and minor amounts of anhedral pentlandite, subhedral pyrrhotite, and trace late pyrite. Pyrrhotite and pentlandite appear to have formed with the quartz, whereas late pyrite is cross-cutting the quartz and replacing pyrrhotite and pentlandite (Fig. 2.15A). The Scheelite Vein is composed of quartz, anhedral calcite, anhedral scheelite, anhedral pentlandite, anhedral pyrrhotite, anhedral chalcopyrite, subhedral-euhedral pyrite and trace late pyrite. Pentlandite and pyrrhotite are largely cross-cutting quartz but sometimes are

48 intergrown with quartz. (Figs. 2.15B). Chalcopyrite and pyrite are typically intergrown with pyrrhotite and pentlandite. Calcite is commonly spatially associated with scheelite as it either rims scheelite or occurs as intergrowths with the mineral. The basaltic hostrock in sample S-460-2 (McLeay) is dominantly composed of anhedral moderately-sericitized albite, euhedral biotite, euhedral chlorite, subhedral-euhedral coarse- grained apatite, irregular anhedral-euhedral actinolite and anhedral calcite with abundant pentlandite, pyrrhotite, chalcopyrite, subhedral pyrite and trace amounts of late pyrite, rutile, chlorite, potassium feldspar and melonite. In the altered basalt, titanite and rutile are sometimes observed with calcite and scheelite. Apatite is spatially associated with sericitized albite, anhedral calcite and sometimes scheelite. Sericitized albite commonly hosts anhedral irregular lobate inclusions of potassium feldspar. Euhedral chlorite and euhedral actinolite is observed largely replacing and cross-cutting biotite however it is also common for biotite to be intergrown with actinolite and chlorite (Figs. 2.16A, B and C). Rarely, biotite replaces chlorite and calcite is replacing actinolite. The sulfides appear to be concentrated in the altered basalt with pentlandite and pyrrhotite being intergrown with biotite and sometimes quartz (Fig. 2.16D). Chalcopyrite and pyrite are sometimes intergrown with pentlandite, pyrrhotite and quartz. However, pyrite and chalcopyrite are dominantly intergrown with calcite and sometimes scheelite.

Figure 2.15. Photomicrograph and BSE image of textures observed in the pentlandite-bearing quartz vein sample S-460-2 from McLeay. A) BSE image of pentlandite and pyrrhotite with quartz in the Upper Vein. The image depicts late pyrite replacing pentlandite and pyrrhotite while cross-cutting quartz. B) Combined reflected and transmitted plane-polarized light image of pentlandite and pyrrhotite cross-cutting quartz in the Scheelite Vein at the contact with the basaltic hostrock.

Figure 2.16. Photomicrographs and BSE images of textures observed in the basaltic hostrock of the pentlandite-bearing quartz vein sample S-460-2 from McLeay. A) Transmitted plane- polarized light image of biotite dominantly being cross-cut by actinolite; sulfides (opaques) are largely associated with biotite B) Transmitted plane-polarized light image of biotite dominantly being cross-cut by actinolite but also intergrown with some actinolite grains. C) Transmitted plane-polarized light image of biotite being replaced by or intergrown with chlorite. D) Transmitted cross-polarized light image of chalcopyrite and pyrrhotite intergrown with quartz, biotite and actinolite.

S-Moran-2 – Pentlandite-bearing quartz vein from Moran Sample S-Moran-2 is composed of a pentlandite-bearing quartz vein and a segment of the basalt hostrock which is evident in the automated mineralogy scan of Fig. 2.17. The vein segment is entirely composed of quartz. The quartz vein exhibits a prominent late fracture that occurs across the entire sample, splitting the quartz vein in two, highlighted in the automated mineralogy scan in Fig. 2.17. This fracture is infilled by pentlandite, pyrrhotite, late pyrite and melonite (Fig. 2.18A) with trace amounts of scheelite, altaite, calcite and sphalerite. It is noteworthy that quartz is absent in the late fracture infill. Additionally, there are 50

Figure 2.17. Automated mineralogy scan of sample S-Moran-2 from Moran. The vein segment, the basaltic hostrock, and the late fracture splitting the quartz vein are labeled.

late fractures protruding from the basalt which are infilled by calcite and chalcopyrite (Fig. 2.18B). Lastly, the quartz vein hosts wave-like features that are sub-perpendicular to the main large vein-splitting fracture (Fig. 2.18C) and are noticeable below the late fracture and to the left of the basalt in Fig. 2.17. These features are dominantly composed of intensely sericitized albite with inclusions of muscovite and potassium feldspar (Fig. 2.18C). The features may also contain calcite with minor to trace titanite, zircon, rutile, chlorite, irregular anhedral monazite and a La- Ce-Nd-bearing apatite (Fig. 2.18D). The basaltic hostrock is dominantly composed of subhedral-euhedral pyrite, anhedral elongated pyrrhotite, anhedral pentlandite, coarse euhedral biotite and a fine-grained anhedral groundmass of dominantly albite (Fig 2.18E) with minor euhedral titanite, anhedral quartz, and anhedral potassium feldspar. Pyrite often contains fine-grained chalcopyrite that is similar to the late pyrite observed in the other pentlandite-bearing quartz vein samples. Pyrite appears to be stable in a matrix of elongated pyrrhotite and minor pentlandite (2.18F).

Figure 2.18. Photomicrographs and BSE images of textures observed in sample S-Moran-2 from Moran. A) BSE image of the most prominent fracture of the sample being infilled by pentlandite and late pyrite. B) BSE image of the smaller fractures, extending from the altered basalt into the quartz vein, filled with calcite and chalcopyrite. C) Transmitted cross-polarized light image of the wave-like features composed of intensely sericitized albite. D) BSE image of wave-like features composed of sericitized albite and chlorite. E) Transmitted plane-polarized light image of euhedral biotite in fine-grained albite groundmass. F) BSE image of subhedral-euhedral pyrite occurring with elongated pyrrhotite and minor pentlandite.

S-LSU162-1 – Pentlandite-bearing quartz vein from Moran Sample S-LSU162-1 represents a brecciated quartz vein with a sulfide matrix (Fig. 2.19). The clasts are mostly composed of quartz although there is a rock clast occurring in the sample

52 as well (Fig. 2.19). The quartz clasts are composed entirely of quartz and the rock clast is dominantly composed of fine-grained albite, actinolite, Fe-rich actinolite and chlorite with minor clinopyroxene, epidote, biotite, Mn-bearing ilmenite, apatite and molybdenite (Figs. 2.20A). The matrix is composed of pyrrhotite, pentlandite and late pyrite with trace amounts of melonite, ilmenite and Fe-bearing sphalerite (Fig. 2.20B).

Figure 2.19. Automated mineralogy scan of sample S-LSU162-1 (Moran). The quartz and rock clasts are labeled.

Figure 2.20. Photomicrographs and BSE images of textures observed in pentlandite-bearing quartz vein sample S-LSU162-1 from Moran. A) BSE image of actinolite and Mn-bearing ilmenite from the rock clast. B) Reflected plane-polarized light image of matrix pyrrhotite and pentlandite with a quartz clast.

Auriferous quartz veins Both auriferous quartz veins are dominantly composed of quartz and euhedral-subhedral pyrite. In both samples, pyrite is intergrown with subhedral-anhedral galena. Additionally, both samples display galena being replaced by aikinite. Aikinite is broken down into a symplectite of Bi, Ag, Cu, Fe, Pb and S phases. Gold occurs as native Au and as electrum, but electrum is the dominant Au phase. Electrum and native Au occur as inclusions hosted in pyrite and/or galena.

S-LSU373A-1 – Auriferous quartz vein from 500 m north of Moran Sample S-LSU373A-1 from 500 m north of Moran consists of an auriferous quartz vein and a segment of the basalt that the quartz vein is hosted in (labeled in the automated mineralogy scan in Fig. 2.21).

Figure 2.21. Automated mineralogy scan of auriferous quartz vein sample S-LSU373A-1 from 500 m north of Moran. The vein segment and the basalt segment of the scan are labeled.

The quartz vein is largely composed of quartz, coarse subhedral-anhedral galena and coarse subhedral-euhedral pyrite with minor amounts of calcite, albite, molybdenite, chalcopyrite, argentiferous galena, aikinite, and sphalerite with trace native Au, electrum, cervelleite, acanthite or argentite and hessite. Calcite and albite are anhedral and commonly

54 spatially associated with coarse anhedral galena. Galena is usually associated with pyrite which both host precious metal (Ag and Au) minerals or are rimmed by these minerals. Hessite is co- genetic with galena and pyrite, occurring as inclusions or rimming galena (Figs. 2.22A). The pyrite-galena intergrowth is often in assemblage with cervelleite, acanthite or argentite and argentiferous galena (Fig 22B).

Figure 2.22 BSE images of textures observed in the vein segment of the auriferous quartz vein sample S-LSU373A-1 from 500 m north of Moran. A) BSE image of hessite associated with galena. B) BSE image of cervelleite and argentiferous galena at the grain boundary of pyrite and quartz; cervelleite shows a fine rim of acanthite or argentite.

Electrum is the dominant gold phase, however native Au does occur as micron-sized grains. The precious metals and base metals such as sphalerite and molybdenite occur as either inclusions, intergrowths or rims of pyrite and galena (Fig. 2.23A). Chalcopyrite is another base metal mineral present in the vein and is commonly observed being spatially associated with sphalerite, galena and pyrite. Aikinite occurs as either inclusions or occurs as a replacement of galena (Fig. 2.23B). Aikinite is broken down into a symplectite of native Bi, chalcopyrite, galena, bornite and sometimes a Ag-Bi alloy (Fig. 2.23B). The basaltic hostrock of sample S-LSU373A-1 (500 m north of Moran) is composed of coarse subhedral biotite, euhedral fine-grained pyrite, sericitized albite, anhedral irregular potassium feldspar, chlorite, and a rutile-calcite-titanite intergrowth which replaces euhedral- subhedral pyrite (Fig. 2.24) or euhedral biotite. Galena is associated with pyrite and occurs as

55 inclusions in pyrite (Fig. 2.24). Subhedral pyrite is spatially associated with biotite and sericitized albite which is evident in the automated mineralogy scan of Fig. 2.21 and in Fig. 2.24.

Figure 2.23. BSE images of textures observed in the auriferous quartz vein of sample S- LSU373A-1 (500 m north of Moran). A) BSE image of molybdenite rimming pyrite. B) BSE image of aikinite inclusion in galena; aikinite is being replaced by galena, bornite, chalcopyrite and native Bi.

Figure 2.24. Representative BSE image of the basalt hostrock in sample S-LSU373A-1 from 500m north of Moran. The image shows galena inclusions in pyrite and the spatial association between pyrite, biotite and sericitized albite. Pyrite is being replaced by calcite, titanite and rutile.

S-137-7 – Auriferous quartz vein from Long North Sample S-137-7 is composed entirely of the auriferous quartz vein, as illustrated in the automated mineralogy scan in Fig. 2.25.

Figure 2.25. Automated mineralogy scan of auriferous quartz vein sample S-137-7 from Long North.

The auriferous quartz vein is largely composed of quartz and euhedral pyrite intergrown with subhedral galena. The vein also contains calcite, anhedral Fe-oxide (magnetite or hematite), subhedral apatite, subhedral Ba-rich celestite, euhedral biotite, anhedral chalcopyrite, anhedral irregular pyrrhotite, euhedral molybdenite, anhedral sphalerite and anhedral rutile. Rutile is chemically inhomogeneous and shows areas with elevated concentrations of Fe and W. Rutile is being replaced by an intergrowth of titanite and Mg-Fe-bearing calcite, with relict rutile occurring in the intergrowth (Fig. 2.26A). Potassium feldspar is also chemically heterogeneous and shows areas with elevated Ba contents. Ba-rich feldspar is typically associated with minerals such as Ba-rich celestite, molybdenite and muscovite. Galena is commonly being replaced by aikinite which is being replaced by a symplectite of native Bi, chalcopyrite, galena and sometimes bismuthinite (Fig. 2.26B). Pyrite contains inclusions of Fe-oxide, electrum and native Au. Galena contains inclusions of a Bi-telluride phase and electrum (Fig. 2.26B). The minerals that are commonly observed to be associated with galena and pyrite include sphalerite, molybdenite, electrum, native Au, a Bi-telluride phase, muscovite and Ba-rich celestite. Chalcopyrite is typically replacing galena and/or pyrite, while occurring with Mn-Fe-bearing calcite (Fig 2.27A). Chalcopyrite is also observed replacing aikinite, as is tetradymite. Pyrrhotite is also present in trace amounts but is being replaced by galena (Fig. 2.27B).

Figure 2.26. BSE images of the auriferous quartz vein sample from Long North (S-137-7). A) BSE image of rutile and Fe-W-rich rutile being replaced titanite and Mg-Fe-bearing calcite. B) BSE image of electrum and a Bi-Te phase hosted in galena; the image additionally depicts galena and electrum intergrown with molybdenite.

Figure 2.27. BSE images of sample S-137-7 from Long North. A) BSE image of chalcopyrite and Mn-Fe-bearing calcite replacing galena. B) BSE image of pyrrhotite being replaced galena.

Barren quartz vein The barren quartz vein sample S-137-6 from Long North includes the barren quartz vein, the komatiite hostrock, and the contact between the quartz vein and a granitic dike. The three different sections are all labeled in the automated mineralogy scan in Fig. 2.28. The quartz vein is composed of quartz with trace micron-sized calcite, celestite, potassium feldspar, chlorite, muscovite and plagioclase which collectively compose less than 1% of the vein. The komatiite hostrock is dominantly composed of biotite and tremolite with minor amounts of chromite, aluminian chromite, fluorapatite, and pyrite rimmed by millerite (Fig. 2.29A). At the contact between the komatiite and the vein are large euhedral blades of tremolite.

Figure 2.28. Automated mineralogy scan of barren quartz vein sample S-137-6 from Long North. The barren quartz vein, the komatiitic hostrock and the dike-vein contact are labeled.

At the quartz vein-dike contact, zoned euhedral anhydrite, euhedral diopside and euhedral biotite with trace celestite, millerite and pyrite can be found (Fig. 2.29B).

Figure 2.29. BSE images of sample S-137-6 (barren quartz vein from Long North). A) BSE image of the contact between the komatiite and the vein, displaying tremolite and intergrown pyrite with millerite rims. B) BSE image of the quartz vein-dike contact showing zoned euhedral anhydrite, diopside, biotite, and micron-sized millerite and pyrite.

Magmatic sulfides from Long North (samples S-137-1 and S-157-1) Sample S-137-1 represents the massive sulfide orebody at Long North. The sample is dominantly composed of coarse pentlandite and pyrrhotite (Fig. 2.30) with minor pyrite, chalcopyrite, chromite and magnetite. Chromite and magnetite layers are common (Fig. 2.31)

59 and chromite is often rimmed by magnetite. Late pyrite, observed in the pentlandite-bearing quartz veins is also observed replacing pentlandite and pyrrhotite in this sample.

Figure 2.30. Representative reflected plane-polarized light photomicrograph of the magmatic sulfide orebody minerals in sample S-137-1 (Long North magmatic sulfide sample), displaying coarse-grained pentlandite and pyrrhotite.

Figure 2.31. Scan of sample S-137-1 (magmatic sulfides from Long North) and reflected plane- polarized light photomicrograph of one of the magnetite layers in the sample. In pink dotted lines, the magnetite and chromite layers of the sample are highlighted. The reflected light image depicts euhedral-subhedral magnetite grains that occur in these layers.

Trondhjemite hostrock from Long North (sample IS-1569) Sample IS-1569 from Long North consists of an epoxy puck with two polished rock chips from the same trondhjemite intrusion that hosts the pentlandite-bearing quartz veins in samples 60

IS13-7-2 and IS13-7-2Q (Long North) and the auriferous quartz vein in sample S-137-7 (Long North). The rock chips are dominantly composed of euhedral albite, anhedral quartz, anhedral potassium feldspar and euhedral biotite. The albite is intensely sericitized at its core and is zoned. The sample also contains minor amounts of disseminated apatite and sulfides, namely chalcopyrite, pyrite and trace pentlandite. Trace amounts of calcite and actinolite are also present. An overview of the mineralogy can be observed in Fig. 2.32.

Figure 2.32. Automated mineralogy overview scan of trondhjemite sample IS-1569 from Long North.

2.4.2 CL Imaging CL imaging was performed on the quartz vein samples, resulting in the documentation of various CL signatures.

Pentlandite-bearing quartz veins The pentlandite-bearing quartz veins have the largest variety of CL signatures being associated with luminescence of quartz, feldspar, apatite and carbonate. There are eight total distinguishable CL signatures observed in these quartz veins. Apatite occurs as a yellow color and carbonate appears as a bright red-orange color (Figs. 2.33 and 2.34). Quartz grains initially depict a short-lived blue-green color in CL that faded in seconds and turns into a blue-brown color (Fig. 2.35). Potassium feldspar emits a bright neon blue color (Fig. 2.33). The remainder of the signatures are emitted by plagioclase and are restricted to samples IS13-7-2 and IS13-7-2Q (pentlandite-bearing quartz vein samples from Long North). In pentlandite-bearing quartz vein samples IS13-7-2 and IS13-7-2Q from Long North, apatite, carbonate, potassium feldspar and quartz CL signatures are observed. Four different CL signatures are emitted from albite in the wallrock portion of these samples which are depicted by Figs. 2.33A, C and E. The transmitted light photomicrographs that correspond to the CL images are shown in Figs. 2.33B, D and F. One signature emitted by albite, called albite 1, is displayed by a dark red-brown color and is additionally characterized by abundant carbonate (red-orange CL color) and sericite (extremely fine-grained inclusions that are black in CL). The next signature emitted by plagioclase, called albite 2, is displayed by thin dark blue rims of albite that rim albite 1. Albite 3 displays a bright red-brown color that occurs as overgrowths to albite 1 and albite 2. The final signature emitted by plagioclase, called albite 4, is displayed by dark red-black color and occurs as overgrowths to albite 1, albite 2 and albite 3. In sample S-460-2 from McLeay, there are three CL signatures emitted which come from apatite, calcite and quartz. The rest of the pentlandite-bearing samples (S-Moran-2 and S-LSU162-1 from Moran) only have one CL signature, which is emitted by quartz.

Auriferous and Barren quartz veins The auriferous (samples S-LSU373A-1 from 500 m north of Moran and S-137-7 from Long North) and barren (sample S-137-6 from Long North) quartz veins only depict one CL signature emitted which is emitted by quartz. The CL signature emitted by quartz is the same signature that is emitted by quartz in the pentlandite-bearing quartz veins (Fig. 2.35).

Figure 2.33. Photomicrographs in CL and in transmitted plane-polarized light depicting the various CL signatures observed in samples IS13-7-2 and IS13-7-2Q (Long North pentlandite- bearing quartz veins). Images A, C and E are CL images of trondhjemite hostrock in sample IS13-7-2 in Long North. Images B, D and F are transmitted light images corresponding to images A, C and E, respectively. The number 1 corresponds to albite 1. The number 2 corresponds to albite 2. The number 3 corresponds to albite 3. The number 4 corresponds to albite 4.

Figure 2.34. Photomicrographs in transmitted plane-polarized light (A) and in CL (B) depicting the apatite and calcite CL signatures which are observed in sample S-460-2 from McLeay.

Image 2.35. Photomicrographs in CL of the dark blue-brown color observed in all quartz of the quartz vein samples. This is the CL signature that persists after the short-lived CL signature has faded. A) CL signature in quartz in sample S-Moran-2 pentlandite-bearing quartz vein from Moran. B) CL signature in quartz in sample S-137-6 barren quartz vein from Long North.

2.4.3 Mineral Chemistry

The general mineral formula for pentlandite is (Fe,Ni)9S8. The pentlandite geochemical data are converted from element weight percent to element atoms per formula units (a.p.f.u.) normalized to eight S p.f.u. Pentlandite chemical data are reported in Appendix A. Pyrrhotite has a general chemical formula of Fe1-x S. The pyrrhotite geochemical data are converted from element weight percent to a.p.f.u. normalized to one S p.f.u. Pyrrhotite chemical data are reported in Appendix A. Pyrite has a general mineral formula of FeS2. The pyrite geochemical data are converted from element weight percent to a.p.f.u. normalized to three total a.p.f.u. Pyrite chemical data are reported in Appendix A. Chalcopyrite has a general chemical formula of

CuFeS2. The chalcopyrite geochemical data are converted from element weight percent to a.p.f.u. normalized to two S p.f.u. Chalcopyrite data are reported in Appendix A. Mica has the general chemical formula of AM2-3T4O10(OH, F)2. For biotite, the A-site in the general mica chemical formula is typically filled by K however Na, Ba and Ca can also fill the A-site. The M- site is typically filled by Mg and Fe2+ however Al, Cr, Mn and Ni can also fill this site. Silicon and Al typically fill the T-site, however Fe3+ and Ti can also fill this site. The chemical data have been converted to a.p.f.u and normalized to eight ideal total cations. Biotite chemical data are reported in Appendix A. Feldspar has the general chemical formula AT4O8 with the A-site typically being filled by K, Na and/or Ca. The T-site of the formula is typically filled by Si and/or Al. The oxide weight percent values returned by EPMA were converted into a.p.f.u. normalized to eight total O p.f.u. Potassium feldspar and albite chemical data are reported in Appendix A.

Pentlandite Thirteen spots on pentlandite grains were analyzed in samples IS13-7-2 and IS13-7-2Q (Long North pentlandite-bearing quartz veins). Pentlandite in this sample inconsistently incorporates the trace amounts of Se, Cu and Te. Copper is present up to 0.01 Cu p.f.u. Nickel content is higher than Fe with an average of 4.92 Ni p.f.u., in a range of 4.60-5.06 Ni p.f.u. The ratio between Fe and Ni averages at 0.86 (Fig. 2.36), showing that Ni content is dominant over Fe content in the metal cation site of the mineral formula. The Co content in pentlandite of this sample is consistent with a range of 0.02-0.03 Co p.f.u. (Fig. 2.37). Seven total spots on pentlandite grains were analyzed in sample S-460-2 (McLeay pentlandite-bearing quartz vein), however one analysis was not included in Appendix A because the analysis contained an uncharacteristically high amount of sulfur. Pentlandite in this sample inconsistently incorporates trace amounts of Se, Co and Te. Cobalt content is present with up to about 0.03 Co p.f.u. (Fig. 2.37). Compositionally, the pentlandite in this sample contains slightly more Ni than Fe with an average of 4.79 Ni p.f.u. within a range of 4.72-4.91 Ni p.f.u. The ratio between Fe and Ni averages at 0.89 (Fig. 2.36), indicating that Ni content is dominant over Fe but not as dominant compared to pentlandite in samples IS13-7-2 and IS13-7-2Q (Fig. 2.36). Four total spots on pentlandite grains were analyzed in sample S-Moran-2 (Moran pentlandite-bearing quartz vein). Pentlandite in this sample inconsistently incorporates trace

65 amounts of Se, Co and Te. Cobalt content is present up to 0.07 Co p.f.u. (Fig. 2.37). Compositionally, the pentlandite in this sample contains more Ni than Fe with an average of 4.81 Ni p.f.u., within a range 0f 4.76-4.86 Ni p.f.u. The ratio between Fe and Ni averages at 0.87 (Fig. 2.36), indicating that Ni content is dominant over Fe. Four total spots on pentlandite grains were analyzed in sample S-LSU162-1 (Moran pentlandite-bearing quartz vein). Pentlandite in this sample inconsistently incorporates trace amounts of Se, Cu, Co and Te. Copper is present at concentrations of up to 0.01 Cu p.f.u. Cobalt is present at concentrations up to 0.05 Co p.f.u. (Fig. 2.37). Compositionally, the pentlandite in this sample contains more Ni than Fe with an average of 4.75 Ni p.f.u., within a range of 4.68- 4.84 Ni p.f.u. The ratio between Fe and Ni averages at 0.89, indicating that Ni content is dominant over Fe (Fig. 2.36). Two total spots on pentlandite grains were analyzed in sample S-137-1 (Long North magmatic sulfides). Pentlandite in this sample inconsistently incorporates trace amounts of Se, Te and Co. Cobalt is present at concentrations of up to 0.04 Co p.f.u. (Fig. 2.37). Compositionally, the pentlandite in this sample contains more Ni than Fe with an average of 4.97 Ni p.f.u., within a range of 4.91-5.02 Ni p.f.u. (Fig. 2.36). The ratio between Fe and Ni averages at 0.86, indicating that Ni content is dominant over Fe (Fig. 2.37). Eight total spots on pentlandite grains were analyzed in sample S-157-1 (Long North magmatic sulfides). Pentlandite in this sample inconsistently incorporates trace amounts of Se, Cu, Co and Te. Copper is present at concentrations of up to 0.01 Cu p.f.u. and Co is present at concentrations up to 0.03 Co p.f.u. (Fig. 2.37). Compositionally, the pentlandite in this sample contains more Ni than Fe with an average of 4.83 Ni p.f.u. within a range of 4.77-4.86 Ni p.f.u. (Fig. 2.36). The ratio between Fe and Ni averages at ~0.87, indicating that Ni content is dominant over Fe (Figs. 2.36 and 2.37). In Fig. 2.36 it is evident that pentlandite in samples S-137-1, IS13-7-2 and IS13-7-2Q generally have the highest Ni-content. Generally, pentlandite in samples S-157-1, S-460-2, S- LSU162-1 and S-Moran-2 have lower amounts of Ni. Figure 2.37 illustrates that the sample generally with the most Co-content is sample S-Moran-2, which is followed by sample S- LSU162-1. The remaining pentlandite analyses have lower amounts of Co ranging from ~0.01- 0.04 Co p.f.u. Figure 2.37 also depicts pentlandite in samples S-LSU162-1 and S-460-2 generally having the most Fe with maximum Fe/Ni ratios reaching up to ~0.91.

Fe vs Ni content 4.40 4.30 4.20

4.10 Fe p.f.u Fe 4.00 3.90 4.50 4.60 4.70 4.80 4.90 5.00 5.10 Ni p.f.u

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU162-1 S-137-1 S-157-1

Figure 2.36. Iron and Ni content in atoms per formula unit in pentlandite of samples analyzed in this study. Circles = pentlandite-bearing quartz veins, triangles = primary magmatic ore; yellow = Moran, purple = Long North, gray = McLeay.

Fe/Ni Ratio vs Co content 0.92 0.91 0.9 0.89 0.88 0.87 0.86 Fe/Ni Fe/Ni Ratio 0.85 0.84 0.83 0.82 0.00 0.01 0.02 0.03 0.04 0.05 0.06 0.07 Co content

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU162-1 S-137-1 S-157-1

Figure 2.37. Co content in pentlandite in atoms per formula unit of samples analyzed in this study plotted versus Fe/Ni ratio. Circles = pentlandite-bearing quartz veins, triangles = primary magmatic ore; yellow = Moran, purple = Long North, gray = McLeay.

Pyrrhotite Ten spots in pyrrhotite grains were analyzed in samples IS13-7-2 and IS13-7-2Q (Long North pentlandite-bearing quartz veins). Pyrrhotite in this sample contains Ni at concentrations of up to 0.01 Ni p.f.u. (Fig. 2.38). Pyrrhotite is compositionally iron deficient, relative to troilite, as it contains an average of 0.88 Fe p.f.u. within a range of 0.86-0.89 Fe p.f.u. (Fig. 2.38). Seven spots on pyrrhotite grains were analyzed in sample S-460-2 (McLeay pentlandite- bearing quartz vein). Trace amounts of Ni are present at up to 0.01 Ni p.f.u. Pyrrhotite is compositionally iron deficient, relative to troilite, as it contains an average of 0.87 Fe p.f.u. within a range of 0.86-0.89 Fe p.f.u. (Fig. 2.38). Four spots on pyrrhotite grains were analyzed in sample S-Moran-2 (Moran pentlandite- bearing quartz vein). Trace Ni is present with concentrations up to 0.01 Ni p.f.u. Pyrrhotite is compositionally iron deficient, relative to troilite, as it contains an average of 0.87 Fe p.f.u., within a range of 0.85-0.87 Fe p.f.u. (Fig. 2.38). Three spots on pyrrhotite grains were analyzed in sample S-LSU162-1 (Moran pentlandite-bearing quartz vein). Nickel is present at concentrations of up to 0.01 Ni p.f.u. (Fig. 2.38). Pyrrhotite is compositionally iron deficient, relative to troilite, as it contains an average of 0.86 Fe p.f.u. within a range of 0.86-0.87 Fe p.f.u. (Fig. 2.38). Two spots on pyrrhotite grains were analyzed in sample S-137-7 (Long North auriferous quartz vein). Unlike the pyrrhotite in the pentlandite-bearing quartz veins and the pyrrhotite in the magmatic sulfide samples, pyrrhotite in this sample does not contain any Ni (Fig. 2.38). Additionally, pyrrhotite is compositionally less iron deficient, when compared to the other pyrrhotite analyses in the pentlandite-bearing quartz veins and the magmatic sulfide samples. Pyrrhotite contains Fe at concentrations of 0.97 Fe p.f.u. (Fig 2.38). Three spots on pyrrhotite grains were analyzed in sample S-137-1 (Long North magmatic sulfides). Trace Ni is present with concentrations up to 0.01 Ni p.f.u. (Fig. 2.38). Pyrrhotite is compositionally iron deficient, relative to troilite, as it contains an average of 0.91 Fe p.f.u., within a range of 0.90-0.92 Fe p.f.u. (Fig. 2.38). Six spots on pyrrhotite grains were analyzed in sample S-157-1 (Long North magmatic sulfides). Trace amounts of Ni are present with concentrations up to 0.01 Ni p.f.u. (Fig. 2.38). Pyrrhotite is compositionally iron deficient, relative to troilite, as it contains an average of 0.88 Fe p.f.u., within a range of 0.87-0.89 Fe p.f.u. (Fig. 2.38).

Fe vs Ni content 1.00 0.95 0.90

Fe p.f.u. Fe 0.85 0.80 0.00 0.01 0.02 Ni p.f.u.

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU162-1 S-137-7 S-137-1 S-157-1

Figure 2.38. Table of Fe content versus Ni content in atoms per formula unit in pyrrhotite of all samples analyzed for this study. Circles = pentlandite-bearing quartz veins, triangles = primary magmatic ore, dash = auriferous quartz vein; yellow = Moran, purple = Long North, gray = McLeay.

In Figure 2.38, it is evident that pyrrhotite in sample S-137-7 contains the least amount of Ni, with no Ni detected in either of the analyses for this sample. The rest of the data contains up to 0.01 Ni p.f.u. Iron content is highest in sample S-137-7, with pyrrhotite consistently containing 0.97 Fe p.f.u. However, Fe content in pyrrhotite in the rest of analyses is highly variable within the same sample. The rest of the analyses yield Fe concentrations ranging between ~0.85-0.93 Fe p.f.u.

Pyrite Eight total spots on pyrite grains were analyzed in samples IS13-7-2 and IS13-7-2Q (Long North pentlandite-bearing quartz veins). Pyrite in this sample contains trace amounts of Co with concentrations of up to 0.03 Co p.f.u. (Fig. 2.39) and Ni up to 0.01 Ni p.f.u. Relative cobalt enrichment is also visible in the micro-XRF scan in Fig. 2.40. Compositionally, pyrite contains an average of 0.99 Fe p.f.u. within a range of 0.97-1.00 Fe p.f.u. (Fig. 2.39). Five spots on pyrite grains were analyzed in sample S-460-2 (McLeay pentlandite- bearing quartz vein). Pyrite in this sample contains trace amounts of Co at up to 0.06 Co p.f.u. (Fig. 2.39) and trace amounts of Ni of up to 0.01 Ni p.f.u. Relative Cobalt enrichment is also

69 visible in the micro-XRF scan in Fig. 2.40. Compositionally, the pyrite contains an average of 0.97 Fe p.f.u., within a range of 0.94-0.98 Fe p.f.u. (Fig. 2.39). Five spots on pyrite grains were analyzed in sample S-Moran-2 (Moran pentlandite- bearing quartz vein). Pyrite in this sample contains inconsistent trace amounts of Co of up to 0.06 Co p.f.u. (Fig. 2.39) and Ni of up to 0.01 Ni p.f.u. Relative Cobalt enrichment is also visible in the micro-XRF scan in Fig. 2.41. Compositionally, the pyrite contains an average of 0.96 Fe p.f.u., within a range of 0.93-0.99 Fe p.f.u. Two spots on pyrite grains were analyzed in sample S-LSU162-1 (Moran pentlandite- bearing quartz vein). Cobalt and Ni content are essentially absent in pyrite of this sample. Relative cobalt enrichment is also visible in the micro-XRF scan in Fig. 2.41. Compositionally, the pyrite contains an average of 1.00 Fe p.f.u. with one analysis containing ~0.99 Fe p.f.u. and the other containing ~1.00 Fe p.f.u. (Figs. 2.39). Nine spots on pyrite grains were analyzed in sample S-LSU373A-1 (500 m north of Moran pentlandite-bearing quartz vein). Cobalt and Ni content is essentially absent in pyrite of this sample. Compositionally, the pyrite contains an average of 1.01 Fe p.f.u., within a range of 1.00-1.02 Fe p.f.u. (Fig. 2.39). Two spots on pyrite grains were analyzed in sample S-137-7 (Long North auriferous quartz vein). Cobalt and Ni content is essentially absent in pyrite of this sample. Compositionally, the pyrite contains an average of 1.00 Fe p.f.u., within a range of 0.99-1.01 Fe p.f.u. (Fig. 2.39). Two spots on pyrite grains were analyzed in sample S-137-6 (Long North barren quartz vein). Pyrite in this sample contains inconsistent traces amounts of Co at up to 0.02 Co p.f.u. (Fig. 2.39) and high Ni content at up to 0.05 Ni p.f.u. Compositionally, the pyrite consistently contains 0.96 Fe p.f.u. which appears to be compensated by the abundant Ni content. Two spots on pyrite grains were analyzed in sample S-137-1 (Long North magmatic sulfides). Pyrite in this sample contains high amounts of Co at up to 0.08 Co p.f.u. (Fig. 2.39) and variable Ni content of up to 0.03 Ni p.f.u. Compositionally, the pyrite is variable in Fe content with concentrations of up to 0.98 Fe p.f.u. (Fig. 2.39). Two spots on pyrite grains were analyzed in sample S-157-1 (Long North magmatic sulfides). Pyrite in this sample contains trace amounts of Co at up to 0.02 Co p.f.u. (Fig. 2.39). Compositionally, the pyrite consistently contains 0.99 Fe p.f.u. (Fig. 2.39).

Fe vs Co content 1.04

1.02

1.00

0.98

Fe p.f.u. Fe 0.96

0.94

0.92

0.90 0.00 0.01 0.02 0.03 0.04 0.05 0.06 0.07 0.08 Co p.f.u.

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU162-1 S-LSU373A-1 S-137-7 S-157-1 S-137-1 S-137-6

Figure 2.39. Iron and Co content in atoms per formula unit in pyrite analyzed in this study. Circles = pentlandite-bearing quartz veins, triangles = primary magmatic ore, dash = auriferous quartz veins; Yellow = Moran, purple = Long North, gray = McLeay, Green = 500m north of Moran.

Figure 2.40. Micro-XRF images of Fe, Ni and Co content for pentlandite-bearing quartz veins in Long North and McLeay (top) with automated mineralogy scans (bottom) for mineral identification. The influence of Co (pink color) is exaggerated, relative to Fe and Ni, in order to visually represent the localization of variable concentration of Co in the sample. The images depict Co being concentrated in pyrite. The two images on the left are of pentlandite-bearing quartz vein sample IS13-7-2 from Long North. The two images on the right are of pentlandite- bearing quartz vein sample S-460-2 from McLeay.

Figure 2.41. Micro-XRF images of Fe, Ni and Co content for pentlandite-bearing quartz veins from Moran (top) with automated mineralogy scans (bottom) for mineral identification. The influence of Co (pink color) is exaggerated, relative to Fe and Ni, in order to visually represent the localization of variable concentration of Co in the sample. The images depict Co being concentrated in pyrite. The two images on the left are of pentlandite-bearing quartz vein sample S-Moran-2 from Moran. The two images on the right are of pentlandite-bearing quartz vein sample S-LSU162-1 from Moran.

Figure 2.39 depicts pyrite in samples S-LSU162-1, IS13-7-2Q, S-157-1, S-LSU373A-1 and S-137-1 containing the most Fe and the least amount of Co. Next, are samples IS13-7-2 and S-460-2 which contain slightly higher amounts of Co but lower Fe content. Pyrite in sample S- 137-6 stands out containing low amounts of Fe and Co as well. Lastly, pyrite in samples S- Moran-2 and S-137-1 are distinguishable as they contain the most Co and the least amount of Fe.

Chalcopyrite Seven total spots on chalcopyrite grains were analyzed in samples IS13-7-2 and IS13-7- 2Q (Long North pentlandite-bearing quartz veins). Two of the seven analyses were not included in Appendix A because of either uncharacteristically high S content or uncharacteristically low metal (Cu + Fe) content. Compositionally, chalcopyrite contains a range of 1.01-1.04 Cu p.f.u. and 0.99-1.02 Fe p.f.u. (Fig. 2.42). The average Cu content is 1.02 Cu p.f.u. and average Fe content is 1.01 Fe p.f.u. Seven spots on chalcopyrite grains were analyzed in sample S-460-2 (McLeay pentlandite-bearing quartz vein). Compositionally, the chalcopyrite contains a range of 1.01-1.03 Cu p.f.u. and 0.99-1.01 Fe p.f.u. (Fig. 2.42). The average Cu content is 1.02 Cu p.f.u. and the average Fe content is 1.00 Fe p.f.u. Four spots on chalcopyrite grains were analyzed in sample S-Moran-2 (Moran pentlandite-bearing quartz vein). Compositionally, the chalcopyrite contains a range of 0.96-1.02 Cu p.f.u. and 1.00-1.01 Fe p.f.u. (Fig. 2.42). The average Cu content is 0.99 Cu p.f.u. and the average Fe content is 1.00 Fe p.f.u. Two spots on chalcopyrite grains were analyzed in sample S-LSU162-1 (Moran pentlandite-bearing quartz vein). One analysis contains 0.96 Cu p.f.u. and 0.99 Fe p.f.u. (Fig. 2.42). The other analysis contains 1.00 Cu p.f.u. and 0.99 Fe p.f.u. (Fig. 2.42). Three spots on chalcopyrite grains were analyzed in sample S-LSU373A-1 (500 m north of Moran auriferous quartz vein). Compositionally, the chalcopyrite contains a range of ~1.02- 1.03 Cu p.f.u. with an average of ~1.02 Cu p.f.u. and all analyses consistently contain 1.00 Fe p.f.u. (Fig. 2.42). Three total spots on chalcopyrite grains were analyzed in sample S-137-7 (Long North auriferous quartz vein) however, one of the three analyses were not included in Appendix A

74 because of low sums in total a.p.f.u. Compositionally, the chalcopyrite consistently contains 0.99 Cu p.f.u. and a range of 0.96-0.98 Fe p.f.u. (Fig. 2.42). Four total spots on chalcopyrite were analyzed in sample S-137-1 (Long North magmatic sulfides) for this study, however two of the four spots were not included in Appendix A because of low totals in a.p.f.u. One analysis yields 1.00 Cu p.f.u. and 1.02 Fe p.f.u., meanwhile the other analysis yields 1.00 Cu p.f.u. and 1.01 Fe p.f.u. (Fig. 2.42). Two spots on chalcopyrite were analyzed in sample S-157-1 (Long North magmatic sulfides). One of the analyses yields 1.00 Cu p.f.u. and 1.01 Fe p.f.u., meanwhile the other analysis yields 1.01 Cu p.f.u. and 1.00 Fe p.f.u. (Fig. 2.42).

Fe vs Cu content 1.03 1.02 1.01 1.00

0.99 Fe p.f.u. Fe 0.98 0.97 0.96 0.94 0.96 0.98 1.00 1.02 1.04 1.06 Cu p.f.u.

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU162-1 S-137-7 S-LSU373A-1 S-137-1 S-157-1

Figure 2.42. Iron and Cu content in atoms per formula unit in chalcopyrite in samples analyzed in this study. Circles = pentlandite-bearing quartz veins, triangles = primary magmatic ore, dash = auriferous quartz veins; yellow = Moran, purple = Long North, gray = McLeay, green = 500m north of Moran.

Figure 2.42 illustrates chalcopyrite in sample S-137-7 distinguishably containing the lowest amount of Fe. With the exception of some outliers, the chalcopyrite in all samples is fairly

75 consistent in Cu content containing between ~0.99-1.03 Cu p.f.u. Samples IS13-7-2 and S-460-2 consistently contain more Cu than samples S-157-1 and S-137-1. Copper content in chalcopyrite of samples S-Moran-2 and S-LSU162-1 is extremely variable, but Fe content is consistently within the range of ~0.99-1.01 Fe p.f.u.

Biotite Thirty-one total spots on biotite were analyzed in samples IS13-7-2 and IS13-7-2Q (Long North pentlandite-bearing quartz veins). However, six of the thirty-one analyses were not considered because of low totals. The A-site of the biotite formula is mostly occupied by K but there are trace amounts of Na and Ba. Potassium content varies between ~0.92-0.97 K p.f.u. Ba content is up to 0.01 a.p.f.u. The M-site is largely occupied by Mg and Fe with variable amounts of Ti, Mn, Al and Ni. Magnesium content, with the exception of one outlier, is at an average of 1.50 Mg. p.f.u. (Fig. 2.43). However, the outlier analysis contains 1.03 Mg p.f.u. With the exception of the same outlier analysis, Fe content varies between 1.05-1.18 Fe p.f.u. with an average of 1.12 Fe p.f.u. (Fig. 2.43). The outlier analysis contains 1.47 Fe p.f.u. Nickel content is up to 0.03 Ni p.f.u. with an average of 0.02 Ni p.f.u. (Fig. 2.44). The hydroxyl-site is occupied largely by OH- however there are trace amounts of F and Cl detected in the analyses. Fluorine content is up to 0.16 F p.f.u. with an average of 0.11 F p.f.u. (Fig. 2.44). Chlorine content is up to 0.01 Cl p.f.u. Nine total spots on biotite were analyzed in sample S-460-2 (McLeay pentlandite-bearing quartz vein). However, one of these nine analyses were removed due to uncharacteristically high Ca content. The A-site of the biotite formula is mostly occupied by K but there are trace amounts of Na and Ba. Potassium content varies between ~0.82-0.94 K p.f.u. averaging at 0.90 K p.f.u. Barium content occurs in a range of ~0.00-0.01 Ba p.f.u. with an average of ~0.00 Ba p.f.u. The M-site is largely occupied by Mg and Fe with variable amounts of Ti, Mn, Al and Ni. Magnesium content varies between ~1.73-1.86 Mg p.f.u. and an average of 1.77 Mg. p.f.u. (Fig. 2.43). Iron content varies between 0.79-1.10 Fe p.f.u. with an average of 0.91 Fe p.f.u. (Fig. 2.43). Nickel content varies between ~0.01-0.02 Ni p.f.u. with an average of 0.01 Ni p.f.u. (Fig. 2.44). The hydroxyl-site is occupied largely by OH- with no F detected in the analyses (Fig. 2.44) and essentially not Cl content.

Three spots on biotite were analyzed in sample S-Moran-2 (Moran pentlandite-bearing quartz vein). The A-site of the biotite formula is mostly occupied by K but there are trace amounts of Na, Ca and Ba. Potassium content varies between ~0.83-0.94 K p.f.u. averaging at 0.90 K p.f.u. Barium content occurs in a range of ~0.01-0.03 Ba p.f.u. with an average of 0.02 Ba p.f.u. The M-site is largely occupied by Mg and Fe with variable amounts of Cr, Ti, Mn, Al and Ni. Magnesium content varies between ~1.35-1.36 Mg p.f.u. and an average of 1.36 Mg. p.f.u. (Fig. 2.43). Iron content is at an average of 1.15 Fe p.f.u. (Fig. 2.43). Nickel content varies between ~0.01-0.02 Ni p.f.u. with an average of 0.02 Ni p.f.u. (Fig. 2.44). The hydroxyl-site is occupied largely by OH-. There is no F detected in the analyses (Fig. 2.44) and essentially no Cl content. Three total spots on biotite were analyzed in sample S-LSU373A-1 (500 m north of Moran auriferous quartz vein). The A-site of the biotite formula is mostly occupied by K but there are trace amounts of Na. Potassium content varies between ~0.80-0.95 K p.f.u. averaging at 0.89 K p.f.u. Barium content is essentially absent in this sample. The M-site is largely occupied by Mg and Fe with variable amounts of Ti, Mn, Al and Ni. Magnesium content varies between ~1.70-1.87 Mg p.f.u. with an average of 1.76 Mg. p.f.u. (Fig. 2.43). Iron content varies between 0.92-0.95 Fe p.f.u. with an average of 0.94 Fe p.f.u. (Fig. 2.43). Nickel content is consistently at ~0.1 Ni p.f.u. (Fig. 2.44). The hydroxyl-site is occupied largely by OH- however F content is present up to 0.01 F p.f.u. (Fig. 2.44) and Cl content is essentially absent. Three total spots on biotite were analyzed in sample S-137-6 (Long North auriferous quartz vein). The A-site of the biotite formula is mostly occupied by K but there are trace amounts of Ca and Na. Potassium content varies between ~0.86-0.91 K p.f.u. averaging at 0.89 K p.f.u. Barium content is essentially absent in the sample. The M-site is largely occupied by Mg and Fe with variable amounts of Ti, Mn, Al and Ni. Magnesium content varies between ~2.58- 2.63 Mg p.f.u. and an average of 2.61 Mg p.f.u. which are much higher values than the Mg in the biotite of the rest of the samples (Fig. 2.43). Iron content varies between 0.35-0.37 Fe p.f.u. with an average of 0.35 Fe p.f.u. which are much lower values than the Fe content of the biotite in the rest of the samples (Fig. 2.43). Nickel content is consistently at 0.01 Ni p.f.u. (Fig. 2.44). The hydroxyl-site is occupied largely by OH- however there are trace amounts of F of up to 0.05 F p.f.u. (Fig. 2.44) and Cl content is essentially absent.

Mg vs Fe content 3.00 2.50 2.00 1.50

1.00 Mg Mg p.f.u. 0.50 0.00 0.00 0.50 1.00 1.50 2.00 Fe p.f.u.

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU373A-1 S-137-6

Figure 2.43. Magnesium and Fe content in atoms per formula unit in biotite of all samples analyzed in atoms per formula unit. The black line is the 1:1 Mg-Fe line. Circles = pentlandite- bearing quartz veins, triangles = primary magmatic ore, dash = auriferous quartz veins, diamonds = barren quartz vein; yellow = Moran, purple = Long North, gray = McLeay.

Ni vs F content 0.04 0.03 0.02

Ni p.f.u. Ni 0.01 0.00 0.00 0.05 0.10 0.15 0.20 F p.f.u.

IS13-7-2 IS13-7-2Q S-460-2 S-Moran-2 S-LSU373A-1 S-137-6

Figure 2.44. Nickel and F content in atoms per formula unit in biotite of all samples analyzed. Circles = pentlandite-bearing quartz veins, triangles = primary magmatic ore, dash = auriferous quartz veins, diamonds = barren quartz vein; yellow = Moran, purple = Long North, gray = McLeay.

Figure 2.43 shows that biotite in sample S-137-6 has the highest Mg content and the lowest concentration of Fe. Biotite in samples S-LSU373A-1 and S-460-2 are next, containing relatively moderate amounts of Mg and Fe. Biotite in samples IS13-7-2, IS13-7-2Q and S-

Moran-2 contain the lowest amounts of Mg, generally consistently containing about 1.5 Mg p.f.u. Biotite in these samples also contain the highest amounts of Fe. Figure 2.44 shows that biotite in samples IS13-7-2 and IS13-7-2Q contain the highest amounts of Ni and F. Biotite in sample S-LSU373A-1 contrarily contains the least amount of Ni. Biotite in samples S-Moran-2 and S-460-2 contain variable amounts of F but contain essentially no F. Biotite in sample S-137- 6 contains moderate amounts of Ni and F.

Potassium Feldspar Fourteen spots in potassium feldspar grains were analyzed in samples IS13-7-2 and IS13- 7-2Q (Long North pentlandite-bearing quartz veins). However, four of the fourteen spots were excluded due to uncharacteristically low K content. The A-site in potassium feldspar is dominantly occupied by K with trace amounts of Na, Ba, Fe and Sr. Potassium content is at ~0.84-0.98 a.p.f.u. with an average of 0.90 a.p.f.u. (Fig. 2.45) Sodium content is in a range of ~0.02-0.10 a.p.f.u. with an average of 0.07 a.p.f.u. (Fig. 2.45). Barium content is in a range of ~0.01-0.05 a.p.f.u. with an average of ~0.03 a.p.f.u. (Fig. 2.46). Strontium content is present at concentrations of up to 0.03 Sr p.f.u. (Fig. 2.46). Four spots on potassium feldspar grains were analyzed in sample S-Moran-2 (Moran pentlandite-bearing quartz vein). Potassium content is at ~0.96-1.01 a.p.f.u. with an average of 0.98 a.p.f.u. (Fig. 2.45). Sodium content is in a range of ~0.01-0.02 a.p.f.u. with an average of 0.01 a.p.f.u. (Fig. 2.45). Barium content is present up to 0.02 a.p.f.u. with an average of ~0.02 a.p.f.u. (Fig. 2.46). Strontium is essentially absent (Fig. 2.46). Six spots on potassium feldspar grains were analyzed in sample S-137-7 (Long North auriferous quartz vein). Potassium content is at ~0.84-1.00 a.p.f.u. with an average of 0.94 a.p.f.u. (Fig. 2.45). Sodium content is in a range of ~0.02-0.10 a.p.f.u. with an average of 0.05 a.p.f.u. which is a higher Na value than Na content in potassium feldspar of the rest of the samples analyzed (Fig. 2.45). Barium content is up to 0.04 a.p.f.u. with an average of ~0.02 a.p.f.u. (Fig. 2.46). Strontium content is up to 0.03 a.p.f.u. with an average of 0.01 a.p.f.u. (Fig. 2.46).

K vs Na content 1.05 1.00 0.95

0.90 K p.f.u K 0.85 0.80 0.00 0.02 0.04 0.06 0.08 0.10 0.12 Na p.f.u.

IS13-7-2 IS13-7-2Q S-Moran-2 S-137-7

Figure 2.45. Potassium and Na content in atoms per formula unit in potassium feldspar in all samples analyzed. Circles = pentlandite-bearing quartz veins, dash = auriferous quartz veins; yellow = Moran, purple = Long North.

Sr vs Ba content 0.04 0.03 0.02

Srp.f.u 0.01 0.00 0.00 0.01 0.02 0.03 0.04 0.05 0.06 Ba p.f.u.

IS13-7-2 IS13-7-2Q S-Moran-2 S-137-7

Figure 2.46. Strontium and Ba content in atoms per formula unit in potassium feldspar in all samples analyzed. Circles = pentlandite-bearing quartz veins, dash = auriferous quartz veins; yellow = Moran, purple = Long North.

Figure 2.45 shows that potassium feldspar in sample S-Moran-2 contains the highest concentration of K and the lowest amount of Na. Meanwhile potassium feldspar in the rest of the samples has greatly variable concentrations of Na and K. Figure 2.46 shows that potassium feldspar in sample S-Moran-2 consistently lacks Sr but contains up to ~0.02 Ba p.f.u. Potassium feldspar in samples IS13-7-2 and IS13-7-2Q generally contains the most Ba and Sr. Potassium

80 feldspar in sample S-137-7 contains variable, however contains the highest maximum amount of Sr at ~0.03-0.04 Sr p.f.u.

Albite Thirteen spots on albite plagioclase grains were analyzed in sample IS13-7-2 (Long North pentlandite-bearing quartz vein). The A-site in albite is dominantly occupied by Na with trace amounts of K, Fe, Ca. Sodium content is at ~0.96-1.00 a.p.f.u. with an average of 0.98 a.p.f.u. (Fig. 2.47). Calcium content is present at concentrations of up to 0.02 a.p.f.u. (Fig. 2.47). Four spots on albite plagioclase grains were analyzed in sample S-460-2 (McLeay pentlandite-bearing quartz vein). Sodium content is at ~0.64-0.70 a.p.f.u. with an average of 0.68 a.p.f.u. (Fig. 2.47). Calcium content is in a range of ~0.28-0.34 a.p.f.u. with an average of 0.31 a.p.f.u. which is a much higher value than the Ca content in albite of sample IS13-7-2 (Long North pentlandite-bearing quartz vein; Fig. 2.47). Three spots on albite plagioclase grains were analyzed in sample S-Moran-2 (Moran pentlandite-bearing quartz vein). With the exception of one outlier, Na content is consistently at ~0.72 a.p.f.u. with the outlier containing 0.95 a.p.f.u. (Fig. 2.47). This outlier contains 0.03 Ca p.f.u. while the other two analysis contain a consistent amount of 0.27 Ca p.f.u. (Fig. 2.47). Three spots on albite plagioclase grains were analyzed in sample S-LSU373A-1 (500 m north of Moran auriferous quartz vein). The A-site in the feldspar is dominantly occupied by Na with trace amounts of Ca. Sodium content is at ~0.76-0.85 a.p.f.u. with an average of 0.80 Na p.f.u. (Fig. 2.47). Calcium content is in a range of ~0.13-0.22 a.p.f.u. with an average of 0.18 a.p.f.u. (Fig. 2.47). Three spots on albite plagioclase grains were analyzed in sample S-137-7 (Long North auriferous quartz vein). The A-site in the feldspar is dominantly occupied by Na with trace amounts of K. Calcium content is essentially absent (Fig. 2.47). Sodium content is at 0.96-0.98 a.p.f.u. with an average of 0.96 a.p.f.u. (Fig. 2.47). Figure 2.47 shows that albite in sample S-460-2 contains the highest amount Ca (ranging between ~0.28-0.34 Ca p.f.u.) and the lowest amount of Na (ranging between ~0.60-0.70 Na p.f.u.). Albite in samples S-Moran-2 and S-LSU373A-1 contain moderate amounts of Na and Ca. Albite in samples IS13-7-2 and S-137-7 contain the highest amount of Na and are essentially absent of Ca.

Ca vs Na content 0.40 0.30 0.20

Ca p.f.u. Ca 0.10 0.00 0.60 0.70 0.80 0.90 1.00 1.10 Na p.f.u.

IS13-7-2 S-460-2 S-Moran-2 S-LSU373A-1 S-137-7

Figure 2.47. Calcium and Na content in atoms per formula unit in albite of every sample analyzed in this study. Circles = pentlandite-bearing quartz veins, dash = auriferous quartz veins; yellow = Moran, purple = Long North, gray = McLeay, green = 500m north of Moran.

2.4.4 Fluid Inclusion Studies A fluid inclusion study was performed on quartz at Colorado School of Mines. Fluid inclusions in quartz are wispy-textured (Fig. 2.48A) and there are no consistent primary Fluid Inclusion Assemblages. There are, however, consistent secondary fluid inclusions that form reliable Fluid Inclusion Assemblages. At Tübingen, fluid inclusion measurements were performed to better understand fluids that are being recorded by quartz. Heating-freezing experiments indicated that many of the fluid inclusions were not fit for microthermometry. Fluid inclusions typically had inconsistent ice melting temperatures, as every freezing-heating trial yielded a different ice melting temperature. The fluid inclusions that did yield consistent ice melting temperatures formed CO2-rich secondary Fluid Inclusion Assemblages (Fig. 2.48B).

Fluid inclusions that are rich in CO2 are common in systems that involve metamorphic fluids (e.g. van den Kerkhof et al., 2014; Goldfarb and Groves, 2015). However, these fluid inclusions occurred as secondary Fluid Inclusion Assemblages, which are not associated with primary quartz growth and do not necessarily give any information about the ore forming fluid. Fluid inclusions are not fit for microthermometry because primary Fluid Inclusion Assemblages are not consistent and because the inclusions were likely affected by the deformation responsible for recrystallization of quartz, potentially re-equilibrating the fluid inclusions (see Goldstein and Reynolds, 1994; van den Kerkhof et al., 2014; Goldfarb and

Groves, 2015). Fluid inclusions in albite, intergrown with pentlandite-pyrrhotite, of Long North Mineralization Style III were also observed and were determined as not being fit for microthermometry either as the inclusions do not occur in consistent Fluid Inclusion Assemblages.

Figure 2.48. Transmitted light images of fluid inclusions hosted in quartz. A) Photomicograph of inconsistent fluid inclusions at room temperature. B) Photomicrograph of CO2-rich fluid inclusions at room temperature.

2.4.5 Thermodynamic Considerations In order to better understand conditions for the formation of observed mineral assemblages associated with pentlandite in Mineralization Style III, Phase 2 of the Geochemists Workbench (GWB) modelling software package was used with calculated log K values at temperatures from 0-300C. The following elements were incorporated into the model, based on the mineral assemblage associated with pentlandite (pentlandite-pyrrhotite-biotite-albite-quartz ± potassium feldspar) in Long North Mineralization Style III: Fe, Ni, S, Mg, K, Al, Si and Na. The mineral assemblage associated with pentlandite in McLeay Mineralization Style III is encompassed by the assemblage in Long North Mineralization Style III. The elements F and Cl were included in the model based on Mineralization Style III biotite chemistry. A moderate salinity (~ 5 wt % NaCl equivalent) was assumed for the model to represent most metamorphic fluids and to account for the abundant albite associated with pentlandite Long North Mineralization Style III. The model was run at 300⁰C and thermodynamic data used was from the Thermoddem geochemical database for Geochemist’s workbench. Thermodynamic modeling results are illustrated in Fig. 2.49 which depict the stability of the mineral assemblage at reduced

83 conditions relative to the hematite-magnetite buffer, and at near neutral conditions at 300⁰C (pH = ~5.6). A temperature of 300⁰C was chosen as it is the maximum temperature at which Geochemist’s workbench can be reliably used. This value represents the minimum temperature for the system as metamorphic fluids in Archean orogenic gold deposits have a mode temperature of 325-400⁰C (Goldfarb et al., 2005). Additionally, as noted in Table 2.1, depositional temperatures of auriferous fluids in the Kambalda area fall within a range of 300- 430⁰C (Ho et al., 1992 and references therein).

Figure 2.49. Oxygen fugacity versus pH diagram, at 300⁰C and 5 wt% NaCl equiv. with the solid black field highlighting the pentlandite mineral assemblage that is observed in Long North Mineralization Style III. McLeay Mineralization Style III pentlandite mineral assemblage is also encompassed by the solid black field. A) Mineral stability diagram only representing silicates with the black field encompassing annite, albite and quartz. B) Mineral stability diagram only representing the Fe-mineral species with the black field encompassing pyrrhotite, pentlandite and annite.

2.5 Discussion The combination of micro-analytical techniques applied to the samples of this study have allowed for interpretation of the formation of the pentlandite-bearing quartz veins observed in Kambalda.

2.5.1 Styles of Pentlandite Mineralization Pentlandite is observed in five pentlandite-bearing quartz vein samples analyzed in this study (samples IS13-7-2 and IS13-7-2Q from Long North, S-460-2 from McLeay, S-Moran-2 and S-LSU162-1 from Moran), in the magmatic sulfide orebody samples (samples S-137-1 and S-157-1 from Long North) and in carbonate vein samples (cross-cutting the magmatic sulfides of Moran; Staude, pers. comm.). The carbonate vein samples were not analyzed in this study but were observed in hand sample for comparison to pentlandite observed in the rest of the samples. All of these samples have been separated into groups based on textural and mineralogical characteristics associated with the pentlandite occurrence. The groups are Mineralization Style I, Mineralization Style II, Mineralization Style III and Mineralization Style IV. The auriferous quartz veins observed in this study are considered Mineralization Style V and lack pentlandite. The mineralization style groups are solely based on mineralogical and textural features without implications on relative timing of formation.

Mineralization Style I Mineralization Style I is represented by samples from the primary magmatic sulfide orebodies which are dominantly composed of pentlandite and pyrrhotite with minor amounts of pyrite and chalcopyrite and variable amounts of magnetite and chromite. Gresham and Loftus- Hills (1981) describe the characteristics of the various orebodies at the Kambalda Dome and describe the massive ores to be generally composed of 80% or more sulfides mainly pyrrhotite, pentlandite and pyrite with minor chalcopyrite, ferrochromite and magnetite. Staude (2015) studied various magmatic orebodies along the eastern flank of the Kambalda Dome and reported that the main sulfide minerals in these orebodies are pentlandite, pyrrhotite and chalcopyrite. The common non-sulfide minerals that occur in the orebodies include chromite, actinolite and magnetite which overprint the primary ore (Staude, 2015).

Mineralization Style II Mineralization Style II includes the carbonate veins with pentlandite which were studied in hand sample. Carbonate veins cross-cut massive primary magmatic ore (Mineralization Style I) and commonly contain nickeline and gersdorffite (Staude, 2015). Hand sample observations of the carbonate vein samples indicate that the veins are composed of carbonate, quartz, pentlandite,

85 pyrrhotite, chalcopyrite, pyrite, gersdorffite, nickeline and chlorite. The occurrence of Ni- arsenides is common in settings where hydrothermal Ni activity is reported such as the Avebury deposit (Keays and Jowitt, 2013), the Pevkos area (Thalhammer et al., 1986), the Bou Azzer deposit (Ahmet et al., 2009), Beni Bousera (Leblanc, 1986), the Coolac ultramafic belt (Ashley, 1973), the Eastern Metals deposit (Auclair et al., 1993), the northwest Nelson area (Grapes and Challis, 1999), veins at the Fortaleza de Minas deposit (Almeida et al., 2007), the Main Urals Fault deposits (Jefferson et al., 2007), the hydrothermal halo at the Widgiemooltha Dome (Le Vaillant et al., 2015) and Five-element Vein deposits (Kissin, 1992; Markl et al., 2016). These Mineralization Style II Ni-arsenide bearing carbonate veins are interpreted to be hydrothermal in origin based on mineralogical similarities to hydrothermal Ni occurrences globally.

Mineralization Style III Mineralization Style III includes samples IS13-7-2 and IS13-7-2Q (Long North), and S- 460-2 (McLeay). This mineralization style is characterized by the intergrowth of pentlandite with pyrrhotite and gangue minerals such as quartz, feldspar and biotite. Mineralization Style III pentlandite-bearing quartz veins formed through at least two generations of fluid pulses as pentlandite is intergrown with two distinct generations of albite overgrowths apparent in CL (Fig. 2.33). Through mineralogical comparison between the trondhjemite hostrock (Fig. 2.32) and the altered trondhjemite segment in Long North Mineralization Style III samples (Figs. 2.9 and 2.12), it is evident that albitization has occurred as a result of vein formation. This mineralization style is characterized by pentlandite and pyrrhotite being co-genetic with quartz, and other gangue minerals including feldspar and biotite, which are thought to be hydrothermal in origin.

Mineralization Style IV Mineralization Style IV is represented by samples S-Moran-2 and S-LSU162-1 from Moran. Pentlandite and pyrrhotite are localized along late structures that cross-cut quartz veins. In sample S-Moran-2, an early quartz vein is fractured and filled by mostly pentlandite, pyrrhotite and late pyrite. The vein infill resembles the mineralogy of the primary magmatic sulfide orebodies (Mineralization Style I). Pyrrhotite in the basaltic hostrock shows foliation, likely due to regional metamorphic conditions (McQueen, 1987; Marshall and Gilligan, 1987 and

86 references therein). In sample S-LSU162-1, the quartz vein is brecciated and cemented by a matrix of dominantly pentlandite, pyrrhotite, and pyrite with late pyrite. The sulfide matrix additionally carries a rock clast of the intermediate-composition dike that serves as a hostrock to the quartz vein. Similar to the late fracture in sample S-Moran-2, the sulfide cement is analogous to the mineralogy of magmatic sulfide orebodies (Mineralization Style I). The mineralogical resemblance between the primary magmatic sulfide orebody (Mineralization Style I) and Mineralization Style IV is striking and is attributed to sulfides being of the same origin. The localization of sulfides into cracks in early quartz veins could be explained by the effects post-depositional deformation and regional metamorphism. Metamorphic temperatures and pressures that are reported to overprint various magmatic nickel deposits fall within a range of 450-600⁰C and 2-3.5 kbar with low strain rates (McQueen, 1987 and references therein). These conditions are reported to cause sulfides to have very low strength and behave ductilly during deformation (McQueen, 1987). Metamorphism and low strain conditions are reportedly capable of ductile deformation of magmatic sulfides at Kambalda, mechanically remobilizing sulfides along fault planes as well as forming breccia ores and divergent veins of massive sulfides (McQueen, 1987 and references therein; Marshall and Gilligan, 1987). The late fracture, filled by sulfides in sample S-Moran-2, is analogous to divergent veins of massive sulfide. Trace amounts of calcite in the late fracture of S-Moran-2 may be evidence for CO2-rich fluids being present during the remobilization process which could have been capable of assisting mobilization processes through liquid-state transport. The sulfide- cemented breccia of sample S-LSU162-1 resembles breccia ores where clasts are typically basalt, ultramafic rocks, porphyritic rocks and vein quartz (McQueen, 1987 and references therein). However, sulfide-cemented breccia, with Lunnon basalt clasts, have been recently interpreted to form during magmatic sulfide ore formation processes by the percolation of sulfide melts, into the Lunnon basalt during thermomechanical erosion (Staude et al., 2017a). The breccia are proposed to have formed as sulfide melt exploits and infills existing fracture networks of basalt formed by typical syn-volcanic processes (e.g. flow-top autobreccia and hyaloclastites) which could be further enhanced in porosity and permeability by flash boiling of water content in basalt (Staude et al., 2017a). However, processes responsible for existing fractures of basalt, which sulfide melts infill, would not be applicable to quartz veins. Additionally, the sulfide-cementing basalt breccia-forming process, described in Staude et al.

(2017a), occurs at the massive sulfide-footwall basalt contact. The quartz vein from this sample is in contact with matrix and disseminated sulfides, with massive sulfides occurring further downhole (Fig. 2.50). Therefore, the sulfide cement in sample S-LSU162-1 is most probably formed from mechanical remobilization processes, comparable to breccia ores described by McQueen (1987).

Figure 2.50. Labeled (left) and non-labeled (right) images of a drill hole collared at the Moran magmatic sulfide orebody from which sample S-LSU162-1 was collected. The beginning of the drill hole begins at the bottom right of the core box with the end of the drill hole at the top left of the core box. The image shows that the brecciated quartz vein with sulfide cement is in contact with matrix and disseminated sulfides as opposed to massive sulfides which are intersected further downhole.

Lesher and Barnes (2008) note that sulfides have been mobilized only up to tens of meters from the original site of magmatic sulfide ore (Mineralization Style I) formation. This is consistent with observations in this study (Staude, pers. comm.). Samples were collected from veins within tens of meters from the primary magmatic orebody, supporting that mechanical mobilization of the Moran magmatic sulfide ore (Mineralization Style I) is responsible for sulfide localization in Mineralization Style IV veins. Lastly, magmatic sulfide orebodies being mechanically mobilized into early quartz veins has been documented by Staude (2015). However, a lack of direct field evidence, supporting this process to be responsible for Mineralization Style IV vein formation, renders this interpretation ambiguous.

Comparison of Pentlandite-bearing Veins to Barren and Auriferous Quartz Veins The barren quartz vein does not contain any pentlandite but does host trace amounts of millerite (NiS) which is reported to occur as a secondary mineral of the primary magmatic sulfide orebody (Mineralization Style I) near the Kambalda Dome trondhjemite intrusion (Staude, 2015). Pyrite in the barren quartz vein is euhedral-subhedral and intergrown with millerite, whereas pyrite in pentlandite-bearing quartz veins is anhedral, replaces pentlandite- pyrrhotite, and hosts a fine-grained chalcopyrite. With the exception of quartz and pyrite, this vein has no similarities to any of the pentlandite-bearing quartz veins and therefore is not likely similar in origin. The presence of anhydrite and the preferential precipitation of pyrite and millerite as opposed to pentlandite indicates that the vein was formed under different conditions compared to any of the pentlandite-bearing quartz vein types. Auriferous quartz vein samples are grouped and termed Mineralization Style V, characterized by the lack of pentlandite and the presence of electrum and native Au. Mineralization Style V quartz veins bear many mineralogical similarities to the hydrothermal pentlandite-bearing quartz veins (Mineralization Style III). Trace galena in calcite and pyrite occur in Mineralization Style III veins which are abundant minerals in Mineralization Style V quartz veins. Non-sericitized, clear albite and potassium feldspar are abundant in Mineralization Style V and are related to ore minerals including electrum. Similar to Mineralization Style V, Long North samples IS13-7-2 and IS-13-7-2Q (Mineralization Style III) contain non-sericitized albite and potassium feldspar intergrown with ore minerals (e.g. pentlandite and pyrrhotite). Strontium is an abundant element in Mineralization Style V in the form of celestine, whereas it occurs in the form of strontianite in Mineralization Style III veins. Tungsten is abundant in Mineralization Style V in the form of W-Fe-rich rutile, however in Moran sample S-460-2 (Mineralization Style III), W occurs in the form of scheelite. Other minerals that are common in both Mineralization Style V and Mineralization Style III include quartz, biotite, apatite, calcite, chalcopyrite and tetradymite. Similarities in fluid chemistry could explain the formation of similar minerals across both Mineralization Style III and Mineralization Style V quartz veins.

2.5.2 Mineral Chemistry Mineral chemistry data from EPMA was useful in distinguishing generations of minerals and determining substitution mechanisms of elements in the investigated minerals. Silicates Biotite mineral chemistry was attained from Mineralization Style III, Mineralization Style IV, Mineralization Style V and barren quartz vein samples. The data reveal that biotite from vein segment and hostrock segment of samples, is slightly more abundant in Mg, than Fe, in Mineralization Style III (samples IS13-7-2, IS13-7-2Q and S-460-2; Fig. 2.43). Biotite in Mineralization Style IV (sample S-Moran-2) from the hostrock segment of the sample, is also slightly more abundant in Mg, than Fe, but contains more Ti, Mn, Al, Ni and especially Cr, which fill the M-site that Mg and Fe fill (Fig. 2.43). The wallrock segment biotite of the barren quartz vein (sample S-137-6) is markedly different with a much higher Mg content than the rest of biotite analyses (Fig. 2.43). Biotite in the hostrock segment of a Mineralization Style V auriferous quartz vein (sample S-LSU373A-1) are similar in Mg and Fe content to Mineralization Style III biotite that is associated with pentlandite formation (Fig. 2.43) which could indicate that the biotite between these two vein types formed from a similar system. The 1:1 Fe to Mg exchange line in Fig. 2.43 can be used to represent degrees of M-site replacement each analysis has been subjected to by elements that are not Fe and Mg, such as Cr. Mineralization Style III biotite chemical data between McLeay sample S-460-2 and Long North samples IS13-7-2 and IS13-7-2Q have some differences in F and Ni concentration (Fig. 2.44). Biotite in samples IS13-7-2 and IS13-7-2Q (Long North) are relatively rich in F and Ni compared to sample S-460-2 (McLeay). Elevated F concentrations in biotite are considered to indicate the involvement of a magmatic hydrothermal fluid. Fluorine could have potentially been sourced from F-content in the mineralizing fluid which is noted to sometimes be a significant fluid constituent of hydrothermal fluids like greisen systems (e.g. Seedorff et al., 2005). The contrasting F-concentration in Mineralization Style III veins from McLeay and from Long North might indicate the involvement of a magmatic hydrothermal fluid in Mineralization Style III formation at Long North. The Long North orebody is in contact, and has interacted with, the Kambalda Dome trondhjemite intrusion (Staude, pers. comm.). It is possible for a magmatic- hydrothermal fluid to have been exsolved from this intrusion, providing the F-content detected in biotite.

Albite mineral chemistry data highlight differences between albite types showing different CL signatures. Albite geochemical data of albite intergrown with pentlandite in Long North (Mineralization Style III veins) shows elevated Fe contents. In Fig. 2.47 it is evident that Ca increases with decreasing Na content. This relationship represents calcic plagioclase that is albitized through Ca-Na exchange from a likely saline fluid. The analyses with highest Na content are hydrothermal albite overgrowths in Mineralization Style III which are intergrown with pentlandite (Fig. 2.51).

Figure 2.51. Calcium and Na content in various albite CL signatures in sample IS13-7-2, with increasing Na-Ca exchange evident by the 1:1 Na to Ca exchange line. The number 1 corresponds to albite 1, the number 2 corresponds to albite 2, the number 3 corresponds to albite 3 and the number 4 corresponds to albite 4.

Potassium feldspar mineral chemistry data reveal two groups likely representing two different generations. Potassium feldspar generation 1 is relatively K-rich and Ba-Sr-poor. This group is interpreted to be early as one of the analyses is from the core of the zoned potassium feldspar in Fig. 11C. Rim analyses clearly fit generation 2 potassium feldspar chemistry (lower K concentrations and elevated concentrations of Sr and Ba). Figure 2.52 shows potassium feldspar generations 1 and 2 geochemical data relative to potassium feldspar zonation presented in Fig. 11C. Figure 2.52 shows that both generations are observed in both Mineralization Style III veins and Mineralization Style V veins. However potassium feldspar in the hostrock segment of Mineralization Style IV is restricted to generation 1. The overlap of potassium feldspar chemical data between the two vein types (Mineralization Style V and Mineralization Style III) builds on the interpretation that the auriferous quartz veins (Mineralization Style V) and hydrothermal pentlandite-bearing quartz veins (Mineralization Style III) were formed from the same fluid system, but at different times. The lack of Au in Mineralization Style III, and the presence of pentlandite, could be explained by varying distances from the fluid source or varying distances to

91 the sulfide orebody (a major Ni source). Additionally, fluid-rock interactions, which would vary across different hostrocks, could influence the resulting mineralogy between Mineralization Styles III and V.

Figure 2.52. Ternary plot of K-Sr-Ba content in potassium feldspar highlighting the two generations of potassium feldspar.

Chalcopyrite and Pyrite Chalcopyrite mineral chemistry is most variable in Co and Cu content. In the Co-Cu plot in Fig. 2.53, three groups of data are evident and labeled. Group 1 is a relatively Co-rich and Cu- poor group of chalcopyrite data. Group 2 is a relatively Co-poor and Cu-rich group of chalcopyrite data. Lastly, group 3 is relatively Co-rich and Cu-rich. The three groups could represent formation at different times, but there is no textural evidence to support this.

Figure 2.53. Cobalt and Cu content in chalcopyrite plotting as three distinct groups of geochemical data.

Pyrite mineral chemistry largely overlaps and does not depict any particular trends. The Co content is highly variable in pyrite and largely replaces Fe, as can be observed in in Fig. 2.39. However, pyrite in Mineralization Style I, Mineralization Style III and Mineralization Style IV (which replaces pentlandite and pyrrhotite) is abundant in Co which is shown in Figs. 2.39, 2.40 and 2.41. The euhedral pyrite in the barren quartz vein also contains Co, but is only minor compared to the aforementioned groups (Fig. 2.39). The only group of pyrite chemical data that distinguishes itself is the pyrite chemical data related to gold in the Mineralization Style V quartz veins as the grains contain minimal to no Co and Cu. The pyrite in Mineralization Style V are of different origin considering the pyrite from Mineralization Styles III and IV are different texturally and geochemically. Additionally, pyrite in Mineralization Style V quartz veins is related to gold, as it hosts electrum and native gold versus in Mineralization Style III, where pyrite replaces ore minerals namely pyrrhotite and pentlandite.

Fe-Ni-S Systematics Chemical data for pentlandite and co-genetic pyrrhotite provide insight on the variable concentrations of Fe, Ni and S in each mineral. The variable concentrations of these elements in the minerals is believed to be an effective tool to differentiate between different populations related to different ore-forming processes. The pentlandite and pyrrhotite analyzed were grouped and interpreted based on the Mineralization Style groups that were determined through mineralogical and textural characteristics. Unpublished EPMA mineral chemistry data of pentlandite and pyrrhotite in Mineralization Style II was attained and provided by Dr. Sebastian Staude (Staude, pers. comm.). In addition to these data, unpublished mineral chemistry data for pentlandite and pyrrhotite of the McLeay and Moran magmatic sulfide orebodies were also provided by Dr. Sebastian Staude (Staude, pers. comm.). Figures 2.54 and 2.55 depict pyrrhotite and pentlandite geochemical data and the fields that are outlined by each Mineralization Style. Mineralization Styles I and IV greatly overlap, a relationship that is illustrated by the tan field. This relationship is indicative of Styles I and IV pentlandite-pyrrhotite likely being of the same origin with Mineralization Style IV being mechanically remobilized Mineralization Style I sulfides. Mineralization Style II creates its own distinct field, illustrated by the light blue field, which represents veins that formed through hydrothermal processes forming abundant Ni- arsenides.

Figure 2.54. Ternary plots of pentlandite chemical data with moles of S, Ni and Fe plotted. The ternary is initially shown (top) with all data points and the fields that each Mineralization Style forms. The ternary plot is then shown (bottom) with just the fields to observe overlaps in geochemical data. Mineralization Styles I (excluding Long North Style I) and IV plot in the same distinct field illustrated in a tan color. Mineralization Style II forms its own distinct field illustrated in a light blue color. Mineralization Style III overlaps greatly with the tan field and the blue field. Mineralization Style I data from Long North is highlighted in a dotted green field and is distinct as it does not plot with the rest of the data from Mineralization Style I.

Figure 2.55. Ternary plots of pyrrhotite chemical data with moles of S, Ni and Fe plotted. The ternary is initially shown (top) with all data points and the fields that each Mineralization Style forms. The ternary plot is then shown (bottom) with just the fields to observe overlaps in geochemical data. Mineralization Styles I (excluding Long North Style I) and IV plot in the same distinct field illustrated in a tan color. Mineralization Style II forms its own distinct field illustrated in a light blue color. Mineralization Style III overlaps greatly with the tan field and the blue field. Mineralization Style I data from Long North is highlighted in a dotted green field and is distinct as it does not plot with the rest of the data from Mineralization Style IV.

Mineralization Style III data plots as its discrete field, illustrated by the light purple field, plotting right in the middle of the light tan field and the light blue field. This observation is indicative of pentlandite-pyrrhotite having formed at different conditions than the magmatic sulfides of Mineralization Style I and the Ni-arsenide-rich hydrothermal veins of Mineralization Style II. Long North Mineralization Style I, illustrated by the dotted green line, overlaps greatly with Mineralization Style II geochemical data. The geochemical signatures of magmatic sulfides vary across magmatic orebodies and vary within an orebody embayment as a result of variable partitioning of metals in the magmatic ore formation system (Staude, pers. comm.). This characteristic, makes it difficult to determine the reason for which Long North Mineralization Style I has a geochemical signature similar to that of Mineralization Style II.

2.5.3 Timing and Conditions of Mineralization Style III Formation Textural features of vein minerals provided indications for timing and conditions at which Mineralization Style III veins formed. Mineral assemblages were additionally useful in postulating a fluid type responsible for vein formation.

Constraints on Timing and Conditions Irregular grain boundaries and undulose extinction observed in quartz (Fig. 2.7) represent deformation textures of quartz in the veins. These irregular grain boundary quartz grain textures are similar to what is described as grain boundary mobility textures noted in Passchier and Trouw (2005). Quartz grains in Mineralization Style III pentlandite-bearing quartz veins exhibit the three different dynamic recrystallization textures expected with increasing temperature and decreasing flow stress: bulging (Fig. 2.7), subgrain rotation and high temperature grain boundary migration (Figs. 2.7F, 2.7G and 2.7H) recrystallization (Passchier and Trouw, 2005). The highest temperature deformation mechanism observed in quartz is grain boundary migration which is associated with relatively high temperature and relatively low strain rate conditions (Figs 2.7 F, 2.7G and 2.7H; Passchier and Trouw, 2005). Additionally, undulose extinction in quartz is evidence for intracyrstalline deformation as it represents slightly bent dislocations in the crystal lattice (Fig. 2.6G; Passchier and Truow, 2005). Lastly, the homogenous CL signature displayed by quartz grains in all quartz veins (Fig. 2.33) is a typical feature of recrystallization during metamorphism (Monecke, pers. comm.). The deformation and recrystallization conditions,

96 indicated by the predominance of grain boundary migration in quartz grains, corresponds to the regional metamorphism at Kambalda, which Gresham and Loftus-Hills (1981) describes as low strain upper greenschist to lower amphibolite metamorphism. This indicates that the quartz, along with pentlandite in Mineralization Style III, formed before or during regional metamorphism. Deformation textures observed in pyrrhotite can also be used as an indicator for timing and conditions of vein formation. Mechanical kinking and twinning in pyrrhotite happens at temperatures above 275⁰C (Marshall and Gilligan, 1987 and references therein). These deformation textures are commonly observed in pyrrhotite from the magmatic sulfide orebodies, occurring as a result of deformation and recrystallization during tectonism and medium- to high- grade regional metamorphism which occurred after primary magmatic ore formation (McQueen, 1987). Hence, these textures represent the deformation and metamorphism of pyrrhotite in Mineralization Styles III and IV indicating that the two Mineralization Styles formed before or during metamorphism. Biotite provides additional insight on the timing of pentlandite-pyrrhotite formation in Mineralization Style III veins. Kinking in biotite, associated with pentlandite and pyrrhotite, in Mineralization Style III indicates that the veins have experienced deformation. Kinking in biotite occurs in the brittle domain which, for biotite, is below 250⁰C (Passchier and Trouw, 2005 and references therein). This temperature provides insight on a lower temperature deformation event (< 250 ⁰C) that has affected the hydrothermal pentlandite-bearing quartz veins (Mineralization Style III). The textures between biotite, chlorite and actinolite (Fig. 2.16) in McLeay sample S- 460-2 (Mineralization Style III) show that biotite formed before and during the chlorite and actinolite. As chlorite and actinolite are typical mineral assemblage of greenschist facies metamorphosed basalt (Frost and Frost, 2014) it is probable that these minerals formed during metamorphism. The intergrowth between actinolite, chlorite, and biotite in McLeay sample S- 460-2 (Mineralization Style III), likely reflects biotite (along with co-genetic pyrrhotite and pentlandite) forming during greenschist-facies metamorphic conditions. This texture therefore eliminates the possibility that the pentlandite assemblage in Mineralization Style III veins formed before regional metamorphism further constraining the timing of vein formation to during metamorphism.

The Lunnon Basalt, which is host to Mineralization Style III veins from McLeay is dated at 2.7 Ga (Pb-Pb; Chauvel et al., 1985) and the Kambalda Dome trondhjemite is dated at 2660 ± 4 Ma (U-Pb; Cassidy et al., 1991) hosting the Long North Mineralization Style III veins. These relationships indicate that the hydrothermal pentlandite-bearing quartz veins (Mineralization Style III) formed after 2660 ± 4 Ma. Timing of vein formation has been constrained to during metamorphism (by pyrrhotite deformation mechanics, dynamic recrystallization in quartz and biotite intergrown with metamorphic assemblages). A minimum constraint on the timing of vein formation is yielded by the late low-temperature deformation event that biotite in the vein is recording (kinking in biotite). The four deformation events that are recorded at the Kambalda Dome is noted in Table

2.2. Deformation event, D3 coincides with the trondhjemite intrusion and because Mineralization Style III veins formed after the trondhjemite, the biotite in these veins were deformed after this deformation event. The only deformation event, in Table 2.2, that occurred after D3 is deformation event D4. This D4 deformation event is interpreted to be responsible for biotite deformation which is characterized by oblique dextral shearing strike-slip movement along steep, north to northeast trending faults (Table 2.2.; Stone et al., 2005). The characteristics of this deformation event can be correlated with deformation events that are record in rocks of the entire

Eastern Yilgarn Craton. The D5 deformation event noted in Table 2.7 corresponds to the deformation event that is attributed to deformation of biotite in Mineralization Style III (D4 noted in Table 2.2). The Eastern Yilgarn Craton D5 event occurred at about 2,650-2,625 Ma (age constraint linked to structures based on U-Pb dating of granites or maximum depositional ages in basins; Blewett et al., 2010b; McGoldrick et al., 2013), serving as a minimum age constraint for the formation for Mineralization Style III. The minimum and maximum time constraints allow for reasoning that Mineralization Style III veins formed after 2,660 ± 4 Ma (Kambalda Dome trondhjemite age) and before ~ 2,625 Ma (youngest age for D5).

Fluid Conditions The possibility of mobilization and precipitation of Ni by a hydrothermal fluid, at Kambalda, is noted in Lesher and Barnes (2008) as they highlight that all magmatic sulfide ore metals, with the exception of Ir and Cr, are potentially mobile during metamorphism at such magmatic sulfide deposits. It is further noted that Fe, Ni, Cu, Co, Cr, Zn, Pd and Au were

Table 2.7. The major deformational episodes correlated across the Eastern Goldfields Superterrane after Blewett et al. (2010b) and McGoldrick et al. (2013).

Deformation Characterization of Event Event

East-northeast extension that lasted until about 2,670 Ma; responsible for D1 volcanism and sedimentation associated with the Kambalda Sequence and Black Flag Supergroup

A relatively weak compressional event which occurred about 2,665 Ma, producing weak fabrics associated with open folds and minor crustal thickening; event is broadly coincident with the final stages of the voluminous D 2 High-Ca granite magmatism. The event is also suggested to have been linked to basin development, linked to the formation of the Merougil Supergroup, as a result of the final stage of episodic granite emplacement

Extensional event beginning at about 2,665 Ma and generating large D 3 extensional shear zones, exhumed granites and greenstone units Inversion of the 2,715-2,655 Ma successions and therefore steepening of D4a Neoarchean stratigraphy, which occurred during east-northeast contraction at about 2,655-2,650 Ma

Sinistral transpression, generally associated with north-northwest-trending D 4b faults, occurred at about 2,650 Ma

Dextral transpression associated with north-northeast-trending faults at about D 5 2,650-2,625 Ma

East-northeast contraction at about 2,400 Ma that occurred coincident with D 6 emplacement of Proterozoic dykes

mobilized and precipitated into quartz-sulfide veins with variable amounts of carbonate, occurring in the footwall of Kambalda magmatic sulfide orebodies (Lesher and Barnes, 2008 and references therein). It is believed that this phenomenon noted by Lesher and Barnes (2008 and references therein) is depicted by Mineralization Style III veins where Fe and Ni have been mobilized and precipitated into quartz-sulfide (pentlandite and pyrrhotite) veins. To further understand the fluid conditions during Mineralization Style III quartz vein formation, a fluid inclusion study was performed on quartz. However, primary Fluid Inclusion

Assemblages were not present and secondary Fluid Inclusion Assemblages are not associated with primary quartz growth. Ultimately, fluid inclusion studies did not provide any meaningful data regarding the fluid conditions responsible for the formation of pentlandite in Mineralization Style III veins. Due to the lack of appropriate fluid inclusions, potential fluid types are proposed based on the mineralogy of the veins and the general timing of vein formation. The following fluids have been proposed to be capable of Ni transport in global hydrothermal Ni occurrences: (I) magmatic hydrothermal fluids (e.g. Keays and Jowitt, 2013), (II) metamorphic fluids (e.g. Le Vaillant et al., 2015), (III) oceanic fluids (e.g. Loukola-Ruskeeniemi and Lahtinen, 2013; Melekestseva et al., 2013) and (IV) basinal brines (Gonzalez-Alvarez et al., 2013a). (I) Magmatic Hydrothermal Fluids: It is highly likely that magmatic-hydrothermal fluids were present during or before metamorphism as there are various cross-cutting dikes and intrusions in the Kambalda area (Staude, pers. comm.). Many of these intrusive rocks, such as the Kambalda Dome trondhjemite, contain abundant biotite and/or amphibole as well as display porphyritic texture (Connelly, 2012) which are strong indications for the intrusion-forming magmas being hydrous and capable of exsolving magmatic hydrothermal fluids (Chang, pers. comm.). Biotite veinlets described in Seedorff et al. (2005) are typical of porphyry copper deposits and are the most mineralogically similar to the veins of Mineralization Style III. At the northern Bushveld Complex, magmatic hydrothermal fluids mobilized Ni and consequently formed an alteration assemblage of pyrite-millerite-chalcopyrite which replaces the magmatic Ni-Cu-PGE sulfide deposits (Holwell et al., 2017). The hydrothermal assemblage is described as low-temperature (< 200⁰C) and having formed through alteration by an acidic and oxidized fluid, mobilizing metals through chloride complexes (Holwell et al., 2017). A similar mineral assemblage is observed at the Kambalda Dome trondhjemite-magmatic orebody contact occurring as a pyrite-millerite-magnetite assemblage which indicates that similar processes could have occurred at Kambalda. The fluid that may have been responsible for the formation of this assemblage near the Kambalda Dome intrusion may be responsible for mobilizing Ni and producing the Mineralization Style III veins. (II) Metamorphic Fluids: Metamorphic fluids were also likely present during metamorphism at Kambalda, which is evident through the abundant CO2-rich fluid inclusions in quartz and the orogenic gold deposits that have formed nearby (e.g. St. Ives and Kalgoorlie). Bavinton and Scott (1979) believe that carbonatization occurred after peak metamorphism.

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Carbonatization of mafic and ultramafic rocks could be evidence for abundant CO2-rich fluids circulating during regional metamorphism, after peak metamorphism. Metamorphic fluids are common in areas that have undergone regional greenschist-amphibolite facies metamorphism and are released as a result of dehydration of hydrous minerals, such as chlorite, at metamorphic conditions (Goldfarb and Groves, 2015). Calcite in Mineralization Style III veins could be recording such metamorphic CO2-rich fluids circulating during metamorphism. Most significantly, alteration mineral assemblages in Mineralization Style III veins and vein minerals are similar to those observed in orogenic gold deposits which are formed by metamorphic fluids. Alteration minerals in our veins are particularly similar to those reported in systems related to Au deposits in the Yilgarn Craton (Table 2.8; Witt and Vanderhor, 1998) many of which have been considered to be orogenic deposits. The assemblages described under ‘Mid to upper greenschist 400⁰C’ in Table 2.8 are markedly similar to the secondary minerals that are associated with hydrothermal pentlandite-bearing quartz (Mineralization Style III) veins. Significantly, one of the examples used to define the alteration assemblages of this facies in Table 2.8 includes the auriferous quartz veins of Hunt located in the Kambalda Dome (Witt and Vanderhor, 1998). Additionally, pentlandite is commonly reported as a trace-minor phase in many orogenic gold deposits (e.g. Pitcairn et al., 2006; Novoselov et al., 2013) which indicates that the fluid responsible for forming these deposits is capable of mobilizing and precipitating Ni. In Fig. 2.56 the general mineralogy of orogenic gold deposit giants are displayed while highlighting the various mineral phases that are similarly observed in Mineralization Style III veins. The similarities between orogenic gold deposits and our hydrothermal pentlandite-bearing quartz (Mineralization Style III) veins imply that a metamorphic fluid is highly likely to have been responsible for forming these veins. (III) Oceanic Fluids: Oceanic fluids are capable of mobilizing Ni in settings like volcanogenic massive sulfide (VMS; e.g. Melekestseva et al., 2013) deposits and by chemical precipitation in anoxic/euxinic oceanic basin conditions (e.g. Xu et al., 2013). The VMS deposits of the Main Urals Fault are Ni-rich and there is textural evidence for Ni-minerals having formed with other sulfides as sulfide mounds and chimneys on the oceanic seafloor (Melekestseva et al., 2013) which is not the case for our hydrothermal pentlandite-bearing quartz veins (Mineralization Style III). The Ni-rich black shales of the Niutitang Formation contains Ni minerals in a sedimentary polymetallic ore layer where pentlandite is not a major Ni phases (Xu

101 et al., 2013). These are distinguishable characteristics when compared to Mineralization Style III pentlandite-bearing quartz veins. Therefore, it is unlikely that these processes involving oceanic fluids are responsible for forming our hydrothermal pentlandite-bearing quartz veins (Mineralization Style III). (IV) Basinal Brines: Basinal fluids are considered to have the appropriate chemical characteristics and available ligands to mobilize Ni at relatively low temperatures (Gonzalez- Alvarez et al., 2013a). Additionally, basinal brines play a significant role in the mobilization of Ni at many five-element vein deposits such as the veins at Wittichen (see Kissin, 1992; Staude, et al., 2011; Markl et al., 2016). In Goscombe et al. (2009) it is noted that reduced, acidic, hydrous and relatively highly saline basinal fluids were present in the Eastern Yilgarn Craton. However, hydrothermal pentlandite-bearing quartz vein (Mineralization Style III) mineralogy is not similar to deposits that are formed by basinal fluids such as Mississippi Valley Type (MVT; Leach et al., 2010a) and clastic-dominated lead-zinc (CD Pb-Zn; Leach et al., 2010b) deposits. In five-element vein deposits (i.e., Wittichen; Staude et al., 2011) Ni occurs as Ni-arsenides with abundant U and As minerals such as uraninite and loellingite. Because of these major differences in mineralogy, it is not likely that basinal brines are responsible for the formation of hydrothermal pentlandite-bearing quartz veins (Mineralization Style III). Ultimately, it is proposed that metamorphic fluids are responsible for mobilization of Ni and Fe and consequent precipitation of pentlandite in the Mineralization Style III hydrothermal pentlandite-bearing quartz veins. However, magmatic-hydrothermal fluids cannot be excluded as a viable Ni-transport mechanism. There is a great amount of evidence for metamorphic fluids being the principal transportation mechanism involved in the formation of Mineralization Style III veins, however, magmatic hydrothermal fluids may be involved at Long North. The metamorphic fluid responsible for forming the pentlandite-bearing quartz veins is likely moderately to highly saline, as assumed by the abundance of albite and the variable Cl- content in biotite in Long North Mineralization Style III. Albitization of granitic rocks is a process that typically takes place during reaction of a felsic rock and a saline fluid (Putnis and Austrheim, 2010). Goldfarb et al. (2005) describe metamorphic fluids, responsible for the formation of orogenic gold deposits, as typically of low salinity. However, Fougerouse et al. (2016) note that orogenic fluids, typically containing 3-7 wt. % NaCl equivalent, is capable of mobilizing Ni as chloride complexes.

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Table 2.8. Metamorphic and metasomatic assemblages that are associated with gold mineralization in mafic rocks of the Yilgarn Craton from Witt and Vanderhor (1998). Metasomatic minerals are in upper case lettering, commonly forming >10% of the alteration assemblage. Metasomatic minerals are in low case lettering, normally forming < 10% of the alteration assemblage. Metamorphic Metamorphic Outer alteration zone Inner alteration zone facies assemblage (unaltered) mid to upper hornblende, HORNBLENDE, MICROCLINE, amphibolite plagioclase, PLAGIOCLASE, DIOPSIDE, GARNET, 600⁰C ilmenite BIOTITE, calcite, EPIDOTE, QUARTZ, titanite, pyrrhotite plagioclase, calcite, pyrrhotite, pyrite, arsenopyrite, loellingite Low to mid hornblende, BIOTITE, QUARTZ, BIOTITE, amphibolite plagioclase, HORNBLENDE, PLAGIOCLASE, 500⁰C ilmenite PLAGIOCLASE, calcite, calcite, titanite, garnet, ilmenite, pyrrhotite, pyrite, pyrrhotite, pyrite arsenopyrite Mid to upper actinolite or BIOTITE, CHLORITE, QUARTZ, CALCITE, greenschist 400⁰C hornblende, albite CALCITE, ALBITE OR ALBITE, BIOTITE, or plagioclase, PLAGIOCLASE, titanite, rutile pyrite ilmenite, epidote epidote, pyrrhotite, pyrite Low to mid actinolite, albite, CHLORITE, CALCITE QUARTZ, greenschist 300⁰C ilmenite, epidote OR ANKERITE, MUSCOVITE, ALBITE, epidote, ANKERITE, ilmenite, pyrrhotite SIDERITE, albite, pyrite, pyrrhotite, arsenopyrite

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Figure 2.56. The mineralogy of orogenic gold deposit giants adapted after Goldfarb (Goldfarb, pers. comm.). A purple star is placed next to each mineral phase that is observed in McLeay Mineralization Style III whereas a green star is placed next to each mineral phase that is observed in Long North Mineralization Style III.

The source of the Ni, Fe and S is likely the primary magmatic sulfide orebody, which is composed of mostly pyrrhotite and pentlandite. An additional Ni source, however, includes the serpentinized komatiite that hosts the magmatic sulfide orebody. It is likely that a fluid that is capable of mobilizing Ni will incorporate significant amounts of Ni into the fluid through interaction and potential chemical equilibration with the magmatic sulfide orebody which contains abundant weight % concentrations of Ni.

2.5.4 Thermodynamic Considerations Thermodynamic modeling performed with GWB allowed for understanding conditions of formation for pentlandite and associated minerals in Mineralization Style III veins. According to Fig. 2.49 the observed mineral assemblage (albite, quartz, annite and pyrrhotite) in equilibrium with pentlandite are stable at reduced and near neutral pH conditions. Precipitation of pentlandite in Mineralization Style III is likely associated with fluid-rock interaction. Pentlandite and

104 pyrrhotite are abundant (e.g. Long North samples IS13-7-2 and IS13-7-2Q) or almost entirely localized (e.g. McLeay samples S-460-2) in the altered hostrocks which is an indication that fluid-rock interaction is favorable to ore mineral precipitation. Pentlandite precipitation could have occurred as a result of changes in redox-conditions, particularly reducing the Ni-bearing fluid would allow pentlandite to form. The local serpentinized komatiite could serve as a reducing agent to the fluid and may have assisted in precipitation processes.

2.5.5 Formation of Pentlandite-bearing Quartz Veins in Context of Nearby Deposits Mineralization Style III has similar mineralogical characteristics and similar timing as various hydrothermal deposits in and near the Kambalda Dome, particularly sharing similarities with orogenic gold deposits of the Kalgoorlie Terrane. The gold deposits of the Kalgoorlie Terrane are characterized by association with potassic alteration, post-peak metamorphic timing in mostly greenschist facies rocks and occurrence of gold-rich quartz veins next to barren first- order crustal-scale faults (Vielreicher et al., 2016 and references therein). These characteristics match Mineralization Style III pentlandite-bearing quartz veins (associated with potassic alteration, defined by biotite), formation after peak metamorphism in greenschist facies metamorphic rocks near the Boulder-Lefroy large crustal scale fault. The Boulder-Lefroy fault is closely associated with orogenic gold formation at the world class Kalgoorlie and St. Ives goldfields (Weinberg et al., 2005). Below are the hydrothermal deposits/occurrences that are most similar to Mineralization Style III veins occurring within or in proximity to the Kambalda Dome.

Kalgoorlie Goldfield: Within the Kalgoorlie Terrane, ~55 km north-northwest of Kambalda, is the Kalgoorlie goldfield, the leading gold producer in Australia which consists of two world class deposits: Mt. Charlotte and Golden Mile (Phillips et al., 2017a; Mueller, 2017). The Golden Mile and Mt. Charlotte deposits are believed to have formed during fold-belt compression by metamorphic fluids by many workers (e.g. Groves et al., 1998, 2003; Goldfarb, 2005). The deposits of the

Kalgoorlie goldfield formed from the deposition of potassic, low salinity, CO2 ± (CH4) bearing aqueous fluids (Vielreicher et al., 2016). Ore deposition at Golden Mile and at Mt. Charlotte occurred within 10 m.y. of each other at around 2.64 Ga which was derived from compiling ages

105 determined through dating of monazite, zircon and xenotime (U-Pb) as well as dating gold- related mica (40Ar/39Ar; Vielreicher et al., 2016 and references therein). Timing of gold formation lies within the broad general sense of timing at which Mineralization Style III pentlandite-bearing quartz veins are thought to have formed. The mineralogy of auriferous quartz veins at Mt. Charlotte is strikingly similar to hydrothermal pentlandite-bearing quartz veins (Mineralization Style III) at Kambalda. Auriferous quartz veins at Mt. Charlotte occur as dilational quartz-, carbonate-, albite- and scheelite-bearing quartz veins with gold having formed with pyrrhotite, pyrite, chalcopyrite, tellurides, ankerite, sericite, scheelite and quartz (Vielreicher et al., 2016 and references therein). All of these minerals, except for gold, are observed in Mineralization Style III vein samples. The depositional conditions for the auriferous gold veins at Mt. Charlotte are reported to be 300-350⁰C, 1.5-2.3 kbar, pH 5.8-6.1, and fluid conditions at XCO2 = 0.25-0.3 and < 3 wt % NaCl equiv. (Ho et al., 1992).

The Widgiemooltha Dome: The Widgiemooltha Dome is a magmatic Ni deposit in close proximity to the Kambalda Dome. The Archean Widgiemooltha Dome is located about 50 km southwest of Kambalda and hosts magmatic sulfide deposits along contacts between basalt and overlying komatiite flows that are considered to be correlative with the footwall basalts and overlying komatiite at Kambalda (Hatfield et al., 2017). The origin and mineralogy of the magmatic sulfides of the Widgiemooltha Dome are similar to the primary magmatic sulfides at the Kambalda Dome. Le Vaillant et al. (2015) report a Ni-As-rich hydrothermal halo around the Miitel magmatic sulfide deposit of the Widgiemooltha Dome, which was caused by hydrothermal alteration of the magmatic sulfide ores resulting in the remobilization of Ni-As-Pd-Pt (Le Vaillant et al., 2015). The Ni-mobilizing fluids are As-rich and are proposed to be associated with the regional orogenic gold event that was responsible for the formation orogenic gold deposits such as the St. Ives goldfield (Le Vaillant et al., 2015). The hydrothermal halos consist of small quartz and carbonate veinlets with gersdorffite and minor nickeline (Le Vaillant et al., 2015). Mineralogically, the Miitel hydrothermal halo is similar to the carbonate-quartz vein of Mineralization Style II, containing abundant gersdorffite and nickeline. It is possible that Mineralization Style II formed from similar processes described in Le Vaillant et al. (2015) and thus represents the same hydrothermal Ni mobilization event as the Miitel hydrothermal halo.

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St. Ives St. Ives is the closest world class goldfield to the Kambalda Dome. St. Ives is over 10 km south of the Kambalda Dome and is composed of over 60 gold deposits and also contains primary magmatic Ni deposits hosted in komatiitic flows (Gresham and Loftus-Hills, 1981; Oxenburgh et al., 2017). Most of the auriferous quartz vein within this goldfield are hosted in subsidiary structures of the Boulder-Lefroy fault zone as they are favorable dilational sites for gold deposition (Cox and Ruming, 2004; Neumayr et al., 2008). Many of the largest Au-quartz vein deposits at St. Ives are spatially associated with the Kambalda Anticline, a regional anticline that is also present at the Kambalda Dome (Fig. 2.3; McGoldrick et al., 2013). The deposits formed in low displacement faults by the localization of auriferous fluid flow driven by major slip events along the Boulder-Lefroy fault zone (Cox and Ruming, 2004). A sinistral transpression event, characterized by Goscombe et al. (2009) and McGoldrick et al. (2013) as

D4b (Table 2.7), is noted to be responsible for ore formation in many areas of the Yilgarn Craton including St. Ives, Kalgoorlie, Wallaby and Sunrise Dam. This deformation event occurred at ~2650-2655 Ma (age constraint linked to structures based on U-Pb dating of granites or maximum depositional ages in basins; Blewett et al., 2010b; McGoldrick et al., 2013). Most of the gold ore was deposited during D4b, but gold was also deposited during D3, D4a and D5, characteristics of which are detailed in Table 2.7. Economic grades of gold dominantly occur in the alteration halos adjacent to faults, shear zones and associated quartz-carbonate vein arrays (Cox and Ruming, 2004). There are at least four distinguishable alteration types (Neumayr et al., 2008): (1) early, porphyry-related epidote- magnetite-carbonate alteration, (2) early, regional shear-zone related carbonate alteration, (3) syn-gold, reduced pyrrhotite-biotite-amphibole and syn-gold, oxidized plagioclase-carbonate- pyrite-magnetite-hematite-biotite-chlorite alteration, and (4) late quartz veins. The early carbonate alteration, described in alteration type (2), is potentially represented by early calcite in Long North Mineralization Style III. Alteration type (3), syn-gold reduced, is most similar to what is observed in the veins of Mineralization Style III, excluding the occurrence of amphibole.

Fluid characteristics are described as low to medium salinity, CH4-N2 ± CO2 fluids for syn-gold reduced type 3 alteration.

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Kambalda Dome Various Au-mineralized hydrothermal veins have been reported from the Kambalda Dome (e.g. Cowden et al., 1986; Neall and Phillips, 1987; Oxenburgh et al., 2017 and this study). Mineralization Style V quartz veins are similar to Au-quartz vein deposits at Kalgoorlie and St. Ives. The ore at Golden Mile is described to typically be composed of pyrite, minor base metal sulfides, sometimes tellurides and local arsenopyrite and pyrrhotite (Phillips et al., 2017a). This is extremely similar to Mineralization Style V as the veins are dominantly composed of pyrite intergrown with base metal sulfides, molybdenite, ± tellurides and trace pyrrhotite. Mineralization Style V bears some similarities to the Mt. Charlotte auriferous quartz veins. Both quartz veins contain gold that is associated with quartz and pyrite. Additionally, carbonate and albite are common minerals associated with gold at Mt. Charlotte which are both components of Mineralization Style V. Mineralization Style V quartz veins are most similar to the St. Ives gold quartz veins which is the most proximal world-class gold field to the Kambalda Dome. Mineralization Style V quartz veins of this study is mineralogically similar to St. Ives alteration type (3) as the veins in both localities contain carbonate, pyrite, Fe-oxide and biotite. Due to the astounding mineralogical similarities between Mineralization Style V and the auriferous quartz veins nearby (Kalgoorlie and St. Ives goldfields), it is interpreted that Mineralization Style V auriferous quartz veins formed by orogenic gold forming processes, likely at the same time as the other nearby gold deposits: ~2.65 Ga (St. Ives) and ~2.64 Ga (Kalgoorlie). The local proximity of the Boulder-Lefroy fault and the shared mineralogy and timing between the Mineralization Style III veins and Mineralization Style V auriferous quartz veins at the Kambalda Dome are interpreted to be a result of the two vein types forming from the same metamorphic fluid. The local remobilization of Ni by metamorphic fluids at the Miitel Deposit of the Widgiemooltha Dome augments the argument that metamorphic fluids are capable of mobilizing Ni and could be responsible for forming pentlandite-bearing Mineralization Style III veins.

Accounting for Differences Between Mineralization Style III and Orogenic Gold If the same metamorphic fluids are responsible for forming the orogenic gold veins and pentlandite-bearing Mineralization Style III veins, the contrasting mineralogy in these veins must

108 be explained. One explanation for the lack of As and Au minerals in pentlandite-bearing Mineralization Style III veins is that Au and As were fluid constituents that were simply not precipitated during pentlandite-bearing quartz vein formation (Mineralization Style III). Another potential explanation is that As and Au were precipitated out before the formation of Mineralization Style III veins. Figure 2.57 illustrates the predominant fields for various aqueous Ni- and Au-complexes and stability fields of pentlandite and gold at 350C and 2 kbar (Fougerouse et al., 2016). Solubility contours of Au indicate that highest solubility is reached around neutral pH and oxygen fugacity conditions below the hematite-magnetite buffer. Figure 2.57 illustrates that gold precipitation occurs under reducing and alkaline conditions. It has been proposed that dissolution of sulfide grains may generate a local redox trap that can be effective in gold precipitation from ore fluids (Goldfarb et al., 2005 and references therein). The pentlandite stability field is much larger in contrast to the gold stability field and spans above the hematite- magnetite buffer to more reducing conditions and covers all pH conditions that are more alkaline than pH 3 (Fig. 2.57) with high Ni solubility at reduced acidic conditions and under oxidizing conditions in general. Depending on the conditions of the ore forming metamorphic fluid, it can possibly be able to transport both, Ni and Au, whereas contrasting precipitation mechanism may attribute for Ni- versus Au-bearing quartz veins. The most effective precipitation mechanism for quartz is a decrease in temperature and/or decrease in pressure (Bons, 2000), whereas the most effective precipitation mechanisms for gold seems to be shift in pH concomitant with a drop in oxygen fugacity. Pentlandite precipitation can also be explained by a shift in pH, if the ore forming fluid is rather acidic in nature, or, more likely, a drop in oxygen fugacity. Previous Geochemists Workbench thermodynamic considerations found reducing conditions at a near neutral pH of around 6 to be the most likely precipitation conditions. This is in accordance with Fig. 2.57, which indicates that Au is highly soluble under these conditions, whereas pentlandite and pyrrhotite represent the stable mineral assemblage. While this study was not able to shed much light on the ore forming fluid conditions, ore formation of pentlandite-bearing quartz veins (Mineralization Style III) seem to have formed under reducing conditions within the pyrrhotite stability field at around pH 6, representing the pH of a rock buffered fluid.

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Figure 2.57. pH-log fO2 activity diagrams for the predominant Au (A) and Ni (B) speciation at conditions of 350⁰C and 2kbar. Figure adapted after Fougerouse et al. (2016).

2.6 Summary and Conclusions The samples that were collected were grouped in Mineralization Styles that are observed in the Kambalda Dome, based on mineralogical and textural characteristics. These characteristics are noted in Table 2.9. Pentlandite-bearing quartz veins are a vein type that is not commonly reported in literature, consequently there are few studies that characterize the veins and processes responsible for their formation. Additionally, pentlandite intergrown with pyrrhotite is a mineral assemblage that is typically considered to be magmatic in origin making the formation of such vein types more curious. Ultimately, this study shows that pentlandite-pyrrhotite mineral assemblages are not solely restricted to a magmatic origin and supports newer research that Ni can be mobile in crustal fluids (Gonzalez-Alvarez et al., 2013a; Le Vaillant, 2014; Scholten et al., 2018), contrary to the general belief that the element is immobile in most geologic fluids. The processes and fluids that are capable of mobilizing Ni are a topic of great debate. Despite such debate, most commonly magmatic hydrothermal fluids and/or metamorphic fluids are suggested to be capable of mobilizing Ni and depositing Ni minerals (e.g. Grapes and Challis, 1999; Keays and Jowitt, 2013; Gonzalez-Alvarez et al., 2013b; Le Vaillant et al., 2015; Fougerouse et al., 2016; Holwell et al., 2017). This study provides evidence for metamorphic fluids being responsible for the mobilization of Ni and formation of pentlandite-bearing Mineralization Style III veins. Additionally, some pentlandite-bearing quartz veins

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(Mineralization Style IV) are speculated to have formed from mechanical mobilization of early magmatic sulfide orebodies (Mineralization Style I) and localization into open cracks in early quartz veins. Evidence of Ni mobility by metamorphic fluids is also represented in Mineralization Style II with the distinction that Ni-arsenides and carbonate are the main vein components. The mineralogy of Mineralization Style II is more typical of what is observed in hydrothermal Ni occurrences globally, with Ni-arsenides being prominent Ni-phases (e.g. Coolac ultramafic belt, Ashley, 1973; Pevkos area, Thalhammer et al., 1986; Beni Bousera, Leblanc, 1986; Eastern Metals deposit, Auclair et al., 1993; northwest Nelson area, Grapes and Challis, 1999; Spessart ore district, Wagner et al., 2008; Bou Azzer, Ahmet et al., 2009; Avebury deposit, Keays and Jowitt, 2013; Doriri Creek prospect, Gonzalez-Alvarez et al., 2013b; Main Urals Fault, Melekestseva et al., 2013; Niutitang Formation, Xu et al., 2013; Miitel hydrothermal Ni halo, Le Vaillant et al., 2015). With the exception of the primary magmatic sulfides (Mineralization Style I), characteristics of pentlandite-bearing ore assemblages are rather poorly described. It was the aim of this study to better describe and characterize the different pentlandite assemblages that can be found in the Kambalda Dome. Pentlandite occurs in four distinct styles which formed through different processes. The primary magmatic orebody (Mineralization Style I) is well-documented and it is widely accepted that it is magmatic in origin. Mineralization Style II is represented by gersdorffite-nickeline-pentlandite-pyrrhotite-bearing carbonate veins and is hydrothermal in origin, originating from a metamorphic fluid, analogously to the Ni-As-rich hydrothermal halo quartz-carbonate veins at the Miitel deposit (see Le Vaillant et al., 2015). Mineralization Style III pentlandite occurs in assemblage with pyrrhotite and biotite in quartz veins forming from metamorphic fluids. Pentlandite in Mineralization Style IV only occurs in close proximity to the primary magmatic ore body in late fractures and in breccias of early quartz veins. Mineralization Style IV veins are ambiguous in origin but mechanical remobilization of primary magmatic ore is proposed as a formation mechanism. Mineralization Styles I and II are fairly well studied and occur as such elsewhere, whereas Mineralization Style IV is rather ambiguous in origin. Mineralization Style III pentlandite is co-genetic with quartz and biotite and is thought to have formed from the same metamorphic fluids that formed auriferous quartz veins in adjacent orogenic gold fields. Both Mineralization Style III veins and auriferous quartz veins from Kalgoorlie and St. Ives are

111 intimately associated with the large crustal scale Boulder-Lefroy fault (Weinberg et al., 2005) which coincidently truncates that eastern flank of the Kambalda Dome, being located less than 1 km away from pentlandite-bearing Mineralization Style III quartz vein sample localities. Additionally, timing of Mineralization Style III (forming after 2660 ± 4 Ma and before ~ 2,625 Ma) and auriferous quartz vein formation (main gold mineralization in the Yilgarn Craton occurring at ~2650-2655 Ma) are concurrent, following peak regional metamorphism. Lastly, gold mineralization in Kalgoorlie is accompanied by widespread potassic alteration (Vielreicher et al., 2016) which is exhibited in our hydrothermal pentlandite-bearing quartz (Mineralization Style III) veins as biotite intergrown with pentlandite and pyrrhotite. Mineralization Style III vein formation is interpreted to have formed at temperatures below 510 ± 20⁰C and pressure conditions below 2.5 ± 1kbar, which are peak metamorphic conditions (Gresham and Loftus-Hills, 1981 and references therein). The metamorphic fluid, considered responsible for mobilization of Ni, is moderate to high salinity (~ 3-12 wt. % NaCl equiv.), which encloses general salinities for orogenic fluids (Goldfarb et al., 2005). The fluid was likely composed of H2O with abundant CO2 and H2S as these are typical fluid constituents in orogenic gold forming metamorphic fluids (Goldfarb et al., 2005). At the point of pentlandite precipitation, fluid conditions are thought to have been near-neutral in pH and under reduced conditions relative to the magnetite-hematite buffer. Nickel was likely sourced from the primary magmatic sulfide orebodies. However, the local serpentinized komatiite may also have served as an abundant Ni source. The absence of As and Au in Mineralization Style III, which are typical metamorphic fluid constituents, could either be the result of (1) a lack of precipitation of these elements due to unfavorable P-T-X conditions or (2) Au and As being precipitated while Ni remained in the fluid.

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Table 2.9. Summary of characteristics that define each Mineralization Style.

Mineralization Pentlandite Gangue Hostrock Ore Mineralogy Style Occurrence Mineralogy

pentlandite, intergrown with pyrrhotite, I Komatiite none ore minerals chalcopyrite and pyrite pentlandite, Lunnon basalt associated with Ni- pyrrhotite, carbonate, and magmatic arsenides, sulfides II chalcopyrite, pyrite, quartz and sulfide and gangue gersdorffite and chlorite orebody minerals nickeline quartz, biotite, calcite, Kambalda pentlandite, ankerite, Dome intergrown with pyrrhotite, pyrite, apatite, albite, III trondhjemite gangue minerals chalcopyrite, potassium and Lunnon and pyrrhotite melonite and galena feldspar, basalt scheelite, rutile and titanite Komatiite, pentlandite, intermediate- calcite, cross-cuts quartz pyrrhotite, pyrite, IV composition scheelite and vein chalcopyrite and dike and altaite melonite Lunnon basalt pyrite, galena, aikinite, native Bi, chalcopyrite, bornite, calcite, albite, bismuthinite, Ag-Bi potassium Kambalda alloy, electrum, feldspar, Dome native Au, Bi- absent of biotite, rutile, V trondhjemite telluride phase, pentlandite titanite, Fe- and Lunnon tetradymite, oxide, Ba-rich basalt molybdenite, celestite and sphalerite, muscovite cervelleite, acanthite or argentite, pyrrhotite and hessite

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CHAPTER THREE

IMPLICATIONS AND OUTLOOK

The results of this study are meaningful both scientifically and in applied economic geology. Despite the detailed study performed on the various quartz vein samples at Kambalda, there are still questions that have yet to be answered with the limited underground rock exposures and the limited vein samples available from the mine.

3.1 Implications and Project Significance The transition metal Ni is an important metal in society as it is dominantly used for nickel steel, nonferrous alloy, superalloy and electroplating (Liu et al., 2012). Typical industrial applications of Ni includes utilizing nickel sulfide semiconductors for applications such as solar cells, supercapacitors, rechargeable lithium batteries and photocatalysts (Liu et al., 2018 and references therein). As the demand for such apparatuses increases, a demand of Ni will also increase. Nickel is mined from two main types of deposits: magmatic sulfide deposits and lateritic Ni deposits (Tian et al., 2012 and references therein). About 60% of the Ni is sourced from laterites and about 40% from magmatic sulfide deposits (Liu et al., 2018 and references therein). At present, hydrothermal Ni deposits are subeconomic (Scholten et al., 2018). As technology advances and the demand for Ni rises, Ni demand and prices may drive lower grades of Ni to become economic to mine at modern subeconomic grades found in Hydrothermal Ni deposits. Therefore, a better understanding of the processes and the crustal fluids that are capable of mobilizing Ni and forming hydrothermal Ni occurrences are essential to the development of useful deposit models for this deposit type. The fluid conditions proposed for Mineralization Style III formation is significant as it builds on the limited knowledge available on Ni mobility. The lack of understanding of processes and fluids significant for Ni transport has led to a variety of fluids being proposed for mobilizing the transition metal. In literature Ni mobility has been reported to occur with As as a fluid constituent, precipitating Ni-arsenides. The lack of Ni-arsenides in Mineralization Style III veins is not equivalent with the literature reports of hydrothermal Ni occurrences with abundant

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As. The results of this study allow us to understand that Ni may be transported in crustal fluids without necessarily precipitating as a Ni-arsenide. The identification of Mineralization Style III veins as a product of precipitation from metamorphic fluids has significance to the industry of economic geology. In exploration it is important to distinguish between magmatic sulfide deposits (which are typically highly economic) and hydrothermal Ni deposits (which are commonly subeconomic). The presence of Ni-arsenides is a useful mineralogical indication for distinguishing between the two Ni deposit types, however Mineralization Style III vein formation is proof that Ni may precipitate from metamorphic fluids, forming a pyrrhotite-pentlandite assemblage which could be mistaken as magmatic in origin. Proper mineralogical and textural identification of these veins would therefore be useful to exploration geologists attempting to distinguish between magmatic and hydrothermal Ni deposits. Additionally, proper identification of both Mineralization Style III and Mineralization Style IV veins could be essential to exploration geologists as both vein types occur in proximity to magmatic sulfide orebodies at Kambalda. During field mapping or core logging, identification of these veins could be used as a potential vector toward magmatic sulfide orebodies at Kambalda. Ultimately, the results that metamorphic fluids are capable of mobilizing Ni may be used in the future to identify potentially economic resources of Hydrothermal Ni deposits. The implication of this result is that Hydrothermal Ni deposits could preferentially occur in greenstone belts of Archean cratons where there are abundant ultramafic rocks or magmatic sulfides which can serve as Ni sources. Such an environment would also have experienced regional metamorphism of at least greenschist facies conditions which are ideal for release of metamorphic fluids which could mobilize and potentially concentrate the Ni from a local Ni source.

3.2 Outlook and future work Field work at the underground mines at Kambalda would be useful to identify Mineralization Style IV veins where magmatic sulfides are clearly being mechanically emplaced into early quartz veins. Preferably such evidence would be observed at Moran, where samples S- LSU162-1 and S-Moran-2 were collected. Further field work at Kambalda would be useful to define the extent of hydrothermal pentlandite-bearing quartz veins (Mineralization Style III) from the magmatic sulfide orebodies

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(Mineralization Style I). It would be useful to delineate how far from the orebodies these veins occur if they are to be used as a potential vector for magmatic sulfide orebodies. Additionally, the extent of these veins from the orebody would have indications on how far Ni could be transported in a metamorphic fluid and precipitate pentlandite. This field study would also be useful in evaluating the role that the Boulder-Lefroy fault serves in the formation of these veins. The fluids that formed these veins are believed to have been metamorphic fluids associated with orogenic gold deposits, which are believed to have traveled through the Boulder-Lefroy fault as the main 1st order fault structure fluid conduit. Theoretically, the presence of Mineralization Style III veins would decrease, with increasing distance away from this fault, which is a hypothesis that could be tested through such field work. Sulfur isotope values could also be measured in order to determine approximate temperatures which can be calculated through the knowledge of sulfur isotope fractionation values between co-genetic sulfides. Sulfide-sulfide thermometry between pyrrhotite and pentlandite would therefore be a useful tool for constraining the temperature of formation of Mineralization Style III veins. Laser ablation-inductively coupled plasma mass spectrometry could be useful in detecting trace Ir content in sulfides which has been utilized to discriminate between magmatic and hydrothermal nickel deposits (Keays et al., 1982). Further analytical work could be performed on the barren quartz vein as it records Ni- mobilization in the form of millerite intergrown with pyrite. The two minerals similarly occur in a millerite-magnetite-pyrite alteration of magmatic sulfide orebodies, spatially associated with the Kambalda Dome trondhjemite. It is possible that magmatic hydrothermal fluids, exsolved by the trondhjemite, caused this alteration and formation of the barren quartz vein. However more work should be performed on these millerite occurrences to better understand the Ni-mobility processes involved. Further analytical work should also be performed on Mineralization Style II veins to better characterize the Mineralization Style texturally and mineralogically. Such work would help with confirming the relationship between this Mineralization Style and the Ni-As- rich hydrothermal halo at the Widgiemooltha Dome. Lastly, the thermodynamic data for pentlandite used in the thermodynamic modelling could be improved. The pentlandite data used in both modelling software is not internally consistent with the other minerals of each respective database. Much of pentlandite data available is restricted to lower temperatures and therefore high temperature thermodynamic data,

116 such as heat capacity constants, are necessary for more reliable modelling of pentlandite in a metamorphic setting.

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REFERENCES

Ahmet, A. H., Arai, S., Ikenne, M., 2009. Mineralogy and paragenesis of the Co-Ni arsenide ores of Bou Azzer, Anti-Atlas Morocco. Economic Geology, v. 104, pp. 249-266. Almeida, C. M., Olivo, G. R., Carvalho, S. G., 2007. The Ni-Cu-PGE sulfide ores of the komatiite-hosted Fortaleza de Minas Deposit, Brazil: evidence of hydrothermal remobilization. The Canadian Mineralogist, v. 45, pp. 751-773. Ashley, P. M., 1973. Petrogenesis of sulphide-bearing reaction zones in the Coolac Ultramafic Belt, New South Wales, Australia. Mineralium Deposita, v. 8, pp. 370-378. Auclair, M., Gauthier, M., Trottier, J., Chartrand, F., 1993. Mineralogy, geochemistry and paragenesis of the Eastern Metals serpentinite-associated Ni–Cu–Zn Deposit, Quebec Appalachians. Economic Geology, v. 88, pp. 123-138. Azovskova, O. B., Grabezhev, A. I., 2008. The Talitsa Porphyry Copper-Molybdenum Deposit, the first object of a subalkaline porphyry system in the Central Urals. Doklady Akademii Nauk, v. 418, no. 3, pp. 237-240. Ball, S. M., Hitzman, M., Monecke, T., Pfaff, K., 2016. Hydrothermal nickel sulfide hosted in Neoproterozoic carbonate and evaporitic rocks of the Menda Central Prospect, Democratic Republic of Congo. Unpublished M.S. thesis at Colorado School of Mines, 112 p. Barnes, S. J., Fiorentini, M. L., 2012. Komatiite magmas and sulfide nickel deposits: a comparison of variably endowed Archean Terranes. Economic Geology, v. 107, pp. 755- 780. Barnes, S. J., Hutchison, N., Warriner, S., 2017. Black Swan nickel sulfide deposits in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 123-126. Barnes, S. J., Liu, W., 2012. Pt and Pd mobility in hydrothermal fluids: evidence from komatiites and from thermodynamic modelling. Ore Geology Reviews, v. 44, pp. 49-58. Bavinton, O. A., Scott, I. R., 1975. An orientation-scale assessment of some ultramafic rocks, Kambalda, for factors reflecting mineralisation. Report K2120, 16 p. Berezovskii, G. A., Drebushchak, V. A., Kravchenko, T. A., 2001. Low-temperature heat capacity of pentlandite. American Mineralogist, v. 86, pp. 1312-1313. Bierlein, F. P., Murphy, F. C., Weinberg, R. F., Lees, T., 2006. Distribution of orogenic gold deposits in relation to fault zones and gravity gradients: targeting tools applied to the Eastern Goldfields, Yilgarn Craton, Western Australia. Mineralium Deposita, v. 41, pp. 107-126. Bleeker, W., 1990. Evolution of the Thompson nickel belt and its nickel deposits, Manitoba, Canada. Unpublished PhD thesis, University of New Brunswick, 356 p.

118

Blewett, R. S., 2010a. Archean gold mineral systems in the eastern Yilgarn Craton. AusGeo News, issue 100, 8 p. Blewett, R. S., Czarnota, K., Henson, P. A., 2010b. Structural-event framework for the eastern Yilgarn Craton, Western Australia, and its implications for orogenic gold. Precambrian Research, v. 183, pp. 203-229. Bons, P. D. 2000. The formation of veins and their microstructures in Stress, strain and structure, a volume in honour of W D Means, edited by Jessell, M.W. and Urai, J. L., Journal of the Virtual Explorer, vol. 2, Electronic Journal, 47 p. Bull, A., Rogers, J. R., Ross, A., Tripp, G., 2017. Kanowna Belle gold deposit in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 229-234. Burnham, O. M., Halden, N., Layton-Matthews, D., Lesher, C. M., Liwanag, J., Heaman, L., Hulbert, L., Machada, N., Michalak, D., Pacey, M., Peck, D., Potrel, A., Theyer, P., Toope, K. L., Zwanzig, H., 2003. Geology, stratigraphy, petrogenesis, and metallogenesis of the Thompson nickel belt, Manitoba. Report for CAMIRO Project 97E-02, Sudbury, Mineral Exploration Research centre, 410 p. Butt, C. R. M., Cluzel, D., 2013. Nickel laterite ore deposits: weathered serpentinites. Elements, v. 9, no. 2, pp. 123-128. Candela, P. A., Wylie, A. G., Burke, T. M., 1989. Genesis of the ultramafic rock-associated Fe– Cu–Co–Zn–Ni deposit of the Sykesville district, Maryland Piedmont. Economic Geology, v. 84, pp. 663–675. Capistrant, P. L., Hitzman, M. W., Wood, D., Kelly, N. M., Williams, G., Zimba, M., Kuiper, Y., Jack, D., Stein, H., 2015. Geology of the Enterprise Hydrothermal Nickel Deposit, North- Western Province, Zambia. Economic Geology, v. 110, pp. 9-38. Cassidy, K. F., Barley, M. E., Groves, D. I., Perring, C. S., Hallberg, J. A., 1991. An overview of the nature, distribution and inferred tectonic setting of the granitoids in the late-Archean Norseman-Wiluna Belt. Precambrian Research, v. 51, pp. 51-83. Cerny, P., Ercit, T. S., 2005. The classification of granitic pegmatites revisited. The Canadian Mineralogist, v. 43, pp. 2005-2026. Chauvel, C., Bernard, D., Jenner, G. A., 1985. The Sm-Nd age of Kambalda volcanics is 500 Ma too old! Earth and Planetary Science Letters, v. 74, pp. 315-324. Connelly, R., 2012. Long North mapping project. Unpublished Postgraduate Diploma Applied Geology Project, Curtin University, 46 p. Cooke, D. P., Ramsey, T., Miller, J. M., Bath, A. B., 2017. Kundana goldfield in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 199-205. Cowden, A., Donaldson, M. J., Naldrett, A. J., Campbell, I. H., 1986. Platinum-group elements and gold in the komatiite-hosted Fe-Ni-Cu sulfide deposits at Kambalda, Western Australia. Economic Geology, v. 81, pp. 1226-1235.

119

Cowden, A., Roberts, D. E., 1990. Komatiite-hosted nickel sulphide deposits, Kambalda. Australian Institute of Mining and Metallurgy Monography 14, v. 1, pp. 567-582. Cox, S. F., Ruming, K., 2004. The St Ives mesothermal gold system, Western Australia – a case of golden aftershocks? Journal of Structural Geology, v. 26, pp. 1109-1125. Davis, B. K., Blewett, R. S., Squire, R., Champion, D. C., Henson, P. A., 2010. Granite-cored domes and gold mineralization: architectural and geodynamic controls around the Archaean Scotia-Kanowna Dome, Kalgoorlie Terrane, Western Australia. Precambrian Research, v. 183, pp. 316-337. Dittrich, T., Seifert, T., Schulz, B., Hagemann, S., Gerdes, A., Pfänder, J., 2019. Archean rare- metal pegmatites in Zimbabwe and Western Australia - geology and metallogeny of pollucite mineralisations. Springer Briefs in World Mineral Deposits, 134 p. Fahlquist, L. S., Popp, R. K., 1989. The effect of NaCl on bunsenite solubility and Ni- complexing in supercritical aqueous fluids. Geochimica et Cosmochimica Acta, v. 53, pp. 989–995. Fiorentini, M., Beresford, S., Barley, M., Duuring, P., Bekker, A., Rosengren, N., Cas, R., Hronsky, J., 2012. District to camp controls on the genesis of komatiite-hosted nickel sulfide deposits, Agnew-Wiluna Greenstone Belt, Western Australia: insights from the multiple sulfur isotopes. Economic Geology, v. 107, pp. 781-796. Fougerouse, D., Micklethwaite, S., Tomkins, A. G., Mei, Y., Kilburn, M., Guagliardo, P., Fisher, A., Halfpenny, A., Gee, M., Paterson, D., Howard, D. L., 2016. Gold remobilisation and formation of high grade ore shots driven by dissolution-reprecipitation replacement and Ni substitution into auriferous arsenopyrite. Geochimica et Cosmochimica Acta, v. 178, pp. 143-159. Frost, R. B., Frost., C. D., 2014, Essentials of Igneous and Metamorphic Petrology. Cambridge University Press, 303 p. Goldfarb, R. J., Baker, T., Dube, B., Groves, D. I., Hart, C. J. R., Gosselin, P., 2005. Distribution, character, and genesis of gold deposits in metamorphic terranes. Economic Geology, 100th Anniversary Volume, pp. 407-450. Goldfarb, R. J., Groves, D. I., 2015. Orogenic gold: common or evolving fluid and metal sources through time. Lithos, v. 223, pp. 2-26. Goldstein, R. H., Reynolds, T. J., 1994. Systematics of fluid inclusions in diagenetic minerals. SEPM short course 31, 199 p. Gonzalez-Alvarez, I., Pirajno, F., Kerrich, R., 2013a. Hydrothermal nickel deposits: secular variation and diversity. Ore Geology Reviews, v. 52, pp. 1-3. Gonzalez-Alvarez, I., Sweetapple, M., Lindley, I.D., Kirakar, J., 2013b. Hydrothermal Ni: Doriri Creek, Papua New Guinea. Ore Geology Reviews, v. 52, pp. 37-57. Goodz, M. D., Cohalan, L., Thynne, G., Summerhayes, E., Benton, J., Mcahren C., Jones, M., 2017. Mount Monger goldfield including Daisy Milano and Randalls in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 223-228.

120

Goodz, M. D., Money, L. A., 2017. Coolgardie goldfield in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 207- 210. Goscombe, B., Blewett, R. S., Czarnota, K., Groenewald, P. B., Maas, R., 2009. Metamorphic evolution and integrated terrane analysis of the Eastern Yilgarn Craton: rationale, methods, outcomes and interpretation. Geoscience Australia, Record 2009/23, 270 p. Grapes, R. H., Challis, G. A., 1999. Gersdorffite with pentlandite, violarite, pyrrhotite, and pyrite, northwest Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, v. 42, pp. 189-204. Gresham, J.J., Loftus-Hills, G. D., 1981. The geology of the Kambalda nickel field, Western Australia. Economic Geology, v. 76, pp. 1373-1416. Groves D. I., Goldfarb R.J., Robert F., Hart C. J. R., 2003. Gold deposits in metamorphic belts: overview of current understanding, outstanding problems, future research, and exploration significance. Economic Geology, v. 98, pp. 1–29. Groves, D. I., Goldfarb, R. J., Gebre-Mariam, M., Hagemann, S. G., Robert, F., 1998. Orogenic gold deposits: A proposed classification in the context of their crustal distribution and relationship to other gold deposit types. Ore Geology Reviews, v. 13, pp. 7-27. Haldar, S. K., Tisljar, J., 2014. Introduction to mineralogy and petrology, Elsevier Publication, 356 p. Hanley, J. J., Bray, C. J., 2009. The trace metal content of amphibole as proximity indicator for Cu-Ni-PGE mineralization in the footwall of the Sudbury igneous complex, Ontario, Canada. Economic Geology, v. 104, pp. 113-125. Hatfield, R. A., Muccilli, P., Muller, M., 2017. Widgiemooltha nickel deposits in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 117-122. Hezarkhani, A., 2006. Petrology of the intrusive rocks within the Sungun Porphyry Copper Deposit, Azerbaijan, Iran. Journal of Asian Earth Sciences, v. 27, pp. 326-340. Hicks, J. D., Parkinson, R. G., 2017. Deacon nickel sulfide deposit, Tramways Dome, Kambalda in Australian Ore Deposited edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 113-116. Ho, S. E., Groves, D. I., McNaughton, N. J., Mikucki, E. J., 1992. The source of ore fluids and solutes in Archaean lode-gold deposits of Western Australia. Journal of Volcanology and Geothermal Research, v. 50, pp. 173-196. Holwell, D. A., Adeyemi, Z., Ward, L. A., Smith, D. J., Graham, S. D., McDonald, I., Smith, J. W., 2017. Low temperature alteration of magmatic Ni-Cu-PGE sulfides as a source for hydrothermal Ni and PGE ores: a quantitative approach using automated mineralogy. Ore Geology Reviews, v. 91, pp. 718-740. Hronsky, J. M. A., Schodde, R. C., 2006. Nickel exploration history of the Yilgarn Craton: from the nickel boom to today in Nickel Deposits of the Yilgarn Craton: geology, geochemistry, and geophysics applied to exploration edited by Barnes, S. J., Society of Economic Geologists Special Publication 13, pp. 1-11.

121

Hudson, D. R., 1986. Platinum-group minerals from the Kambalda nickel deposits, Western Australia. Economic Geology, v. 81, pp. 1218-1225. Hudson, D. R., Donaldson, M. J., 1984. Mineralogy of platinum group elements in the Kambalda nickel deposits, Western Australia in Sulphide deposits in mafic and ultramafic rocks edited by Buchanan, D. L. and Jones, M. J., London Inst. Mining Metallurgy Publication, pp. 55-61. Jefferson, C. W., Thomas, D. J., Gandhi, S. S., Ramaekers, P., Delaney, G., Brisbin, D., Cutts, C., Quirt, D., Portella, P., Olson, R. A., 2007. Unconformity-associated uranium deposits of the Athabasca Basin, Saskatchewan and Alberta in Mineral deposits of Canada: a synthesis of major deposit-types, district metallogeny, the evolution of geological provinces, and exploration methods edited by Goodfellow, W.D., Geological Association of Canada Mineral Deposits Division Special Publication No. 5, pp. 273-305. Jowitt, S.M., Keays, R. R., 2011. Shale-hosted Ni-(Cu-PGE) mineralization: a global overview. Applied Earth Science, v. 120, pp. 187-197. Keays, R. R., Nickel, E. H., Groves, D. I., McGoldrick, P. J., 1982. Iridium and palladium as discriminants of volcanic-exhalative, hydrothermal, and magmatic nickel sulfide mineralization. Economic Geology, v. 77, pp. 1535-1547. Keays, R. R., Jowitt, S. M., 2013. The Avebury Ni deposit, Tasmania: a case study of an unconventional nickel deposit. Ore Geology Reviews, v. 52, pp. 4-17. Kissin, S. A., 1992. Five-element (Ni-Co-As-Ag-Bi) veins. Geoscience Canada, v. 19, pp. 113- 124. Klein, F., Bach, W., 2009. Fe-Ni-Co-O-S phase relations in peridotite-seawater interactions. Journal of Petrology, v. 50, no. 1, pp. 37-59. Knight, J. T., Ridley, J. R., Groves, D. I., 2000. The Archean amphibolite facies Coolgardie Goldfield, Yilgarn Craton, Western Australia: nature, controls and gold field-scale patterns of hydrothermal wall-rock alteration. Economic Geology, v. 95, pp. 49-53. Kuck, P. H., 2006. Mineral commodity summaries: Nickel. United States Geological Survey Report, 2 p. Le Vaillant, M., 2014. Hydrothermal remobilisation of base metals and platinum group elements around komatiite-hosted nickel sulphide deposits: applications to exploration methods. PhD thesis, The University of Western Australia, 641 p. Le Vaillant, M., Barnes, S. J., Fiorentini, M. L., Miller, J., Campbell-McCuaig, T., Muccilli, P., 2015. A hydrothermal Ni-As-PGE geochemical halo around the Miitel komatiite-hosted nickel sulfide deposit, Yilgarn Craton, Western Australia. Economic Geology, v. 110, pp. 505-530. Le Vaillant, M., Barnes, S. J., Fiorentini, M. L., Santaguida, F., Törmänen, T., 2016. Effects of hydrous alteration on the distribution of base metals and platinum group elements within the Kevitsa magmatic nickel sulphide deposit. Ore Geology Reviews, v. 72, pp. 128-148. Le Vaillant, M., Fiorentini, M. L., Barnes, S. J., 2018. Review of predictive and detective exploration tools for magmatic Ni-Cu-(PGE) deposits, with a focus on komatiite-related systems in Western Australia in Processes and ore deposits of ultramafic-mafic magmas

122

through space and time edited by Mondal S. K. and Griffin W. L., Elsevier Publication, pp. 47-78. Leach, D. L., Bradley, D. C., Huston, D., Pisarevsky, S. A., Taylor, R. D., Gardoll, S. J., 2010b. Sediment-hosted lead-zinc deposits in Earth history. Economic Geology, v. 105, pp. 593- 625. Leach, D. L., Taylor, R. D., Fey, D. L., Diehl, S. F., Saltus, R. W., 2010a. A deposit model for Mississippi Valley Type lead-zinc ores in Mineral Deposit Models for Resource Assessment, Scientific Investigations Report 2010-5070-A. Leblanc, M., 1986. Co–Ni arsenide deposits, with accessory gold, in ultramafic rocks from Morocco. Canadian Journal of Earth Sciences, v. 23, pp. 1592–1602. Lesher, C., Barnes, S., 2008. Komatiite-associated Ni–Cu–PGE deposits in Komatiite edited by Arndt, N., Cambridge University Press Publication, pp. 27-120. Lin, S., Popp, R.K., 1984. Solubility and complexing of Ni in the system NiO-H2O-HCl. Geochimica et Cosmochimica Acta, v. 48, pp. 2713–2722. Liu, S., Li, Y., Ju, Y., Liu, J., Liu, J., Shi, Y., 2018. Equilibrium nickel isotope fractionation in nickel sulfide minerals. Geochimica et Cosmochimica Acta, v. 222, pp. 1-16. Liu W., Migdisov A., Williams-Jones A., 2012. The stability of aqueous nickel (II) chloride complexes in hydrothermal solutions: results of UV-visible spectroscopic experiments. Geochimica et Cosmochimica Acta, v. 94, pp. 276-290. Loukola-Ruskeeniemi, K., Lahtinen, H., 2013. Multiphase evolution in the black-shale-hosted Ni-Cu-Zn-Co deposit at Talvivaara, Finland. Ore Geology Reviews, v. 52, pp. 85-99. Markl, G., Burusch, M., Neumann, U., 2016. Natural fracking and the genesis of five-element veins. Mineralium Deposita, v. 51, pp. 703-712. Marques, A. F. A., Barriga, F., Chavagnac, V., Fouquet, Y., 2006. Mineralogy, geochemistry, and Nd isotope composition of the Rainbow hydrothermal field, Mid-Atlantic Ridge. Mineralium Deposita, v. 41, pp. 52-67. Marston, R. J., Groves, D. I., Hudson, D. R., Ross, J. R., 1981. Nickel sulfide deposits in Western Australia: a review. Economic Geology, v. 76, pp. 1330-1363. Marshall, B., Gilligan, L. B., 1987. An introduction to remobilization: information from orebody geometry and experimental considerations. Ore Geology Reviews, v. 2, pp. 87-131. McGoldrick, K. L., Squire, R. J., Cas, R. A. F., Briggs, M., Tunjic, J., Allen, C. M., Campbell, I. H., Hayman, P. C., 2013. The largest Au deposits in the St Ives Goldfield (Yilgarn Craton, Western Australia) may be located in a major Neoarchean volcano-sedimentary depo-centre. Mineralium Deposita, v. 48, pp. 861-881. McQueen, K. G., 1981. Volcanic-associated nickel deposits from around the Widgiemooltha Dome, Western Australia. Economic Geology, v. 76, pp. 1417-1443. McQueen, K. G., 1987. Deformation and remobilization in some Western Australian nickel ores. Ore Geology Reviews, v. 2, pp. 269-286. Melekestseva, I. Y., Zaykov, V. V., Nimis, P., Tret'yakov, G. A., Tessalina, S. G., 2013. Cu-(Ni- Co-Au)-bearing massive sulfide deposits associated with mafic-ultramafic rocks of the 123

Main Urals Fault, South Urals: geological structures, ore textural and mineralogical features, comparison with modern analogs. Ore Geology Reviews, v. 52, pp. 18-36. Melfos, V., Vavelidis, M., Christofides, G., Seidel, E., 2002. Origin and evolution of the Tertiary Maronia porphyry copper-molybdenum deposit, Thrace, Greece. Mineralium Deposita, v. 37, pp. 648-668. Miller, J., Blewett, R., Tunjic, J., Connors, K., 2010. The role of early formed structures on the development of the world class St Ives Goldfield, Yilgarn, WA. Precambrian Research, v. 183, pp. 292-315. Molnar, F., Watkinson, D. H., Jones, P. C., 2001. Multiple hydrothermal processes in footwall units of the North Range, Sudbury Igneous Complex, Canada, and implications for the genesis of vein-type Cu-Ni-PGE deposits. Economic Geology, v. 96, pp. 1645-1670. Mueller, A. G., 2015. Structure, alteration, and geochemistry of the Charlotte quartz vein stockwork, Mt Charlotte gold mine, Kalgoorlie, Australia: time constraints, down-plunge zonation, and fluid source. Mineralium Deposita, v. 50, 221-224. Mueller, A. G., Hagemann, S. G., McNaughton, N. J., 2016. Neoarchean orogenic, magmatic and hydrothermal events in the Kalgoorlie-Kambalda area, Western Australia: constraints on gold mineralization in the Boulder Lefroy-Golden Mile fault system. Mineralium Deposita, Published Online, 31 p. Mueller, A. G., 2017. Mount Charlotte deposit, Kalgoorlie goldfield in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp.195-198. Neall, F. B., Phillips, G. N., 1987. Fluid-wall rock interaction in an Archean hydrothermal gold deposit: a thermodynamic model for the Hunt Mine, Kambalda. Economic Geology, v. 82, pp. 1679-1694. Neuser, R. D., 1995. A new high-intensity cathodoluminescence microscope and its application to weakly luminescing minerals. Bochumer Geol. Geotech. Arb. 44, pp. 116-118. Neumayr, P., Walshe, J., Hagemann, S., Peterson, K., Roache, A., Frikken, P., Horn, L., Halley, S., 2008. Oxidized and reduced mineral assemblages in greenstone belt rocks of the St. Ives gold camp, Western Australia: vectors to high-grade ore bodies in Archaean gold deposits? Mineralium Deposita, v. 43, pp. 363-371. Novoselov, K., Belogub, E., Ermolina, O., Mikhailov, A., 2013. Gold-bismuth mineralization in the orogenic gold occurrences of the Kuhmo greenstone belt (Finland). 12th Biennial SGA meeting at Uppsala, v. 3, 3 p. Oxenburgh, S. K., Falconer, M., Doutch, D., Edmonds, P., Foley A., Jane, M., 2017. Kambalda- St. Ives Goldfield in Australian Ore Deposits, Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 215-222. Partington, G. A., Davis, T. P., Pilcher, T. W. B., 2017. Bullabulling goldfield in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 211-214. Passchier, C. W., Trouw, R. A. J., 2005. Microtectonics. Springer Science + Business Media Publication, 366 p.

124

Peltonen, P., Kontinen, A., Huhma, H., Kuronen, U., 2008. Outokumpu revisited: new mineral deposit model for the mantle peridotite-associated Cu–Co–Zn–Ni–Ag–Au sulphide deposit. Ore Geology Reviews, v. 33, pp. 559–617. Phillips, G. N., 2017b. Australian Ore Deposits. The Australasian Institute of Mining and Metallurgy: Melbourne, 864p. Phillips, G. N., Hergt, J., Powell, R., 2017a. Kalgoorlie goldfield – petrology, alteration and mineralization in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 185-194. Phillips, G. N., Vearncombe, J. R., Eshuys, E., 2019. Gold production and the importance of exploration success: Yilgarn Craton, Western Australia. Ore Geology Reviews, v. 105, pp. 137-150. Pirajno, F., González-Álvarez, I., 2013. A re-appraisal of the Epoch nickel sulphide deposit, Filabusi Greenstone Belt, Zimbabwe: a hydrothermal nickel system? Ore Geology Reviews, v. 52, pp. 58-65. Pitcairn, I. K., Teagle, D. A. H., Craw, D., Olivo. G. R., Kerrich, R., Brewer T. S., 2006. Sources of metals and fluids in orogenic gold deposits: insights from the Otago and Alpine Schists, New Zealand. Economic Geology, v. 101, pp. 1525-1546. Pracejus, 2015. The ore minerals under the microscope: an optical guide. Second Edition. Elsevier Publication, 1098 p. Putnis, A., Austrheim, H., 2010. Fluid-induced process: metasomatism and metamorphism. Geofluids, v. 10, pp. 254-269. Resource World, 2018. RNC sees massive increase in Beta Hunt gold production: Resource World-Investment Opportunities and News Magazine: http://resourceworld.com/index.php/rnc-sees-massive-increase-in-beta-hunt-gold- production/ (accessed March, 2019). Ripley, E. M., Sarkar, A., Li, C., 2005. Mineralogic and stable isotope studies of hydrothermal alteration at Jinchuan Ni-Cu Deposit, China. Society of Economic Geologists, v. 100, pp. 1349-1361. Rogers, J. R., Joyce, K. A., Masterman, G. J., Halley, S. W., Walshe, J. L., 2004. Kanowna Belle Gold Mine. Geological Survey of Western Australia Report, Record 2004/16, pp. 47-51. Sahoo, P. R., Prasad, J., Prakasam, M., Singh, S., Venkatesh, A. S., 2009. Orogenic gold mineralisation in and around Kundarkocha, East Singhbhum, Jharkhand. Journal of the Indian Academy of Geoscience, v. 52, No. 1, pp. 11-18. Scholten, L., Watenphul, A., Beermann, O., Testemale, D., Ames, D., Schmidt, C., 2018, Nickel and platinum in high-temperature H2O+HCl fluids: Implications for hydrothermal mobilization. Geochimica et Cosmochimica Acta 224, 187-199. Seedorf, E., Dilles, J. H., Proffett Jr., J. M., Einaudi, M. T., Zurcher, L., Stavast, W. J. A., Johnson, D. A., Barton, M. D., 2005. Porphyry deposits: characteristics and origin of hypogene features. Society of Economic Geology, 100th Anniversary Volume, pp. 251- 298.

125

Smith, B., Ross, J., 2017. Mount Marion lithium pegmatite deposit in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 161-162. Staude, S., 2015. Rock description, Long-Victor Mine. Independence Group Project Geologist Report, 48 p. Staude, S., Barnes, S. J., Le Vaillant, M., 2017a. Thermomechanical erosion of ore-hosting embayments beneath komatiite lava channels: textural evidence from Kambalda, Western Australia. Ore Geology Reviews, v. 90, pp. 446-464. Staude, S., Sheppard, S., Parker, P., Paggi, J., 2017b. Long-Victor Fe-Ni-Cu-(PGE) sulfide complex Kambalda, Western Australia in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 107-112. Staude, S., Werner, W., Mordhorst, T., Wemmer, K., Jacob, D. E., Markl, G., 2011. Multi-stage Ag-Bi-Co-Ni-U and Cu-Bi vein mineralization at Wittichen, Schwartzwald, SW Germany: geological setting, ore mineralogy, and fluid evolution. Mineralium Deposita, v. 47, pp. 251-276. Stone, W. E., Archibald, N. J., 2004. Structural controls on nickel sulphide ore shoots in Archean komatiite, Kambalda, WA: the volcanic trough controversy revisited. Journal of Structural Geology, v. 26, pp. 1173-1194. Stone, W. E., Beresford, S. W., Archibald, N. J., 2005. Structural setting and shape analysis of nickel sulfide shoots at the Kambalda Dome, Western Australia: implications for deformation and remobilization. Economic Geology, v. 100, pp 1441-1455. Thalhammer, O., Stumpfl, E. F., Panayiotou, A., 1986. Postmagmatic, hydrothermal origin of sulfide and arsenide mineralizations at Limassol Forest, Cyprus. Mineralium Deposita, v. 21, pp. 95-105. Tian Y., Etschmann B., Liu W., Borg S., Mei Y., Testemale, D., O’Neill, B., Rae, N., Sherman, D. M., Ngothai, Y., Johannessen, B., Glover, C., Brugger, J., 2012. Speciation of nickel (II) chloride complexes in hydrothermal fluids: in-situ XAS study. Chemical Geology, v. 334, pp. 345-363. Trendall, A. F., 1990. Yilgarn Craton in Geology and Mineral Resources of Western Australia, Western Australia Geological Survey Publication, memoir 3, pp. 11-13. van den Kerkhof, A., Kronz, A., Simon, K., 2014. Deciphering fluid inclusions in high-grade rocks. Geoscience Frontiers, v. 5, pp. 683-695. Vearncombe, J. R., Elias, M., 2017. Yilgarn Craton – mineral deposits and metallogeny in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 95-106. Vielreicher, N. M., Groves, D. I., Lawrence, W. S., Fletcher, I. R., McNaughton, N. J., 2010. Broad synchroneity of three gold mineralization styles in the Kalgoorlie Gold Field: SHRIMP, U-Pb, and 40Ar/39Ar geochronological evidence. Economic Geology, v. 105, pp. 187-227. Vielreicher, N. M., Groves, D. I., McNaughton, N. J., 2016. The giant Kalgoorlie Gold Field revisited. Geoscience Frontiers, v. 7, pp. 359-374.

126

Voroshin, S. V., Tyukova, E. E., Newberry, R. J., Layer, P. W., 2014. Orogenic gold and rare metal deposits of the Upper Kolyma District, Northeastern Russia: relation to igneous rocks, timing, and metal assemblages. Ore Geology Reviews, v. 62, pp. 1-24. Wagner, T., Okrusch, M., Weyer, S., Lorenz, J., Lahaye, Y., Taubald, H., Schmitt, R. T., 2008. The role of the Kupferschiefer in the formation of hydrothermal base metal mineralization in the Spessart ore district Germany: insight from detailed sulfur isotope studies. Mineralium Deposita, v. 45, pp. 217-239. Weinberg, R. F., Van Der Borgh, P., Bateman, R. J., Groves, D. I., 2005. Kinematic history of the Boulder-Lefroy Shear Zone system and controls on associated gold mineralization, Yilgarn Craton, Western Australia. Economic Geology, v. 100, pp. 1407-1426. Wells, M. A., Butt, C. R. M., 2017. Murriin Murrin and Bulong smectite-hosted nickel laterite deposits in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 149-152. Witt, W. K., Mills, D., 2017. Carosue gold deposits in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 245- 248. Witt, W. K., Vanderhor, F., 1998. Diversity within a unified model for Archaean gold mineralization in the Yilgarn Craton of Western Australia: an overview of the late- orogenic, structurally-controlled gold deposits. Ore Geology Reviews, v. 13, pp. 29-64. Wyche, N. L., Wyche S., 2017. Yilgarn Craton geology in Australian Ore Deposits edited by Phillips, G. N., The Australasian Institute of Mining and Metallurgy: Melbourne, pp. 89- 94. Xu, L., Lehmann, B., Mao, J., 2013. Seawater contribution to polymetallic Ni-Mo-PGE-Au mineralization in Early Cambrian black shales of South China: Evidence from Mo isotope, PGE, trace element and REE geochemistry. Ore Geology Reviews, v. 52, pp. 66- 84.

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APPENDIX A

SUPPLEMENTAL ELECTRONIC FILES

The supplemental electronic files consist of (1) a table of characteristics of various representative hydrothermal Ni occurrences reported in literature (2) and the data attained through electron probe microanalyses. The table of hydrothermal Ni occurrences is provided to demonstrate the wide variety of environments, processes and fluid types that have been considered to be capable for Ni mobilization via crustal fluids and consequent formation of hydrothermal Ni occurrences. Additionally, this table shows the general similarities across hydrothermal Ni occurrences, globally, including the common occurrence of Ni-arsenide phases. The collective electron probe microanalyses are provided to denote the data that were attained and used for geochemical plots.

Hydrothermal Ni Occurrences File File containing various hydrothermal Ni occurrences and their respective characteristics. HydrothermalNioccurrences.pdf PDF file containing the table of hydrothermal Ni occurrences. Geochemical Data Files File containing the geochemical data attained for this study. EPMAdata.xlsx Excel spreadsheet containing electron probe microanalysis data. The spreadsheet is organized with each Sheet containing mineral chemical data for a particular mineral. The Sheets analytical data are provided for pentlandite, pyrrhotite, pyrite, chalcopyrite, biotite, potassium feldspar and biotite.

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APPENDIX B

EVOLUTION OF THE YILGARN CRATON THROUGHOUT TIME

Evolution of the Yilgarn Craton after Wyche and Wyche et al. (2017) and references therein). Question marks indicate inferences of geologic events responsible for the geologic phenomenon detailed in the comments column. The authors responsible for these inferences are noted in Wyche and Wyche (2017). Age (Ma) Events Comments After 2500 Post-cratonization - Phanerozoic weathering, erosion, sedimentation - Several generations of Proterozoic mafic dikes - Carbonatites in the east: rare earth element mineralization - Proterozoic sedimentation, deformation 2630-2500 Cratonization; minor - Li-Sn-Ta mineralization (Pegmatite) in extension the south-west (ca 2527 Ma) - End of widespread low-Ca granite magmatism (ca 2600 Ma) 2640-2630 Crustal reworking; high - Low-Ca granite magmatism heat flow; transtension - Brittle faults - Gold mineralization 2650-2640 Crustal reworking; high - Low-Ca granites; alkali granites in the heat flow Kurnalpi Terranes - Gold mineralization 2655-2650 Crustal reworking; - Low-Ca granites; alkali granites in the transpression Kurnalpi Terrane - Large-scale shear zones - Gold mineralization 2675-2655 Crustal reworking and - High-Ca and alkali granites extension - Clastic sedimentary basins with minor chemical sedimentary rocks in the Eastern Goldfields Superterrane 2680-2675 End of volcanism - Felsic volcanic rocks, mafic intrusions, high-Ca granite magmatism in the Kalgoorlie and Yamarna Terranes - Felsic volcanic rocks in the South West Terrane

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2690-2680 Youanmi and Burtville - Felsic and mafic magmatism (volcanic Terranes come together? and intrusive) in the Kurnalpi and Waning plume? Kalgoorlie Terranes - Volcanogenic massive sulfide (VMS) mineralization in the Kurnalpi Terranes 2720-2690 Rifting in the eastern - Komatiitic, mafic and felsic volcanism in Yilgarn: mantle plume? the Kalgoorlie and Kurnalpi Terranes - Nickel mineralization in the Kalgoorlie and Kurnalpi Terranes 2735-2720 Last volcanism in the - Mafic and felsic volcanism, clastic Youanmi Terrane sedimentary rocks, high-Ca granites in the Youanmi Terrane 2800-2735 Waning plume? - Mafic to felsic magmatic cycle and Convergence of Narryer associated clastic and chemical Terrane and rifted parts of sedimentary rocks, including banded iron the Youanmi Terrane? formation, in the Youanmi and Burtville Terranes - High-Ca granites - VMS mineralization in the Youanmi Terrane 2825-2800 Rifting in the western and - Layered mafic-ultramafic intrusions: eastern Yilgarn: mantle vanadium and platinum group element plume? mineralization - Volcano-sedimentary successions in the Youanmi Terrane After 2960 Ultramafic magmatism in - Komatiitic volcanism: nickel the Burtville and southern mineralization Youanmi Terranes: mantle plume? 2980-2920 Felsic and mafic - Felsic volcanic centers and associated magmatism in Youanmi granites and Burtville Terranes; - Basalt, clastic and chemical sedimentary ultramafic magmatism in rocks including banded iron-formation

southern Youanmi - VMS mineralization in the Youanmi Terrane Terrane Before 3080 Volcanism and - First preserved Yilgarn greenstones sedimentation - Oldest banded iron formation; detrital zircons 3730-3300 Earliest Yilgarn rocks - Granitic and mafic magmatism

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- Detrital zircons in the Narryer, Youanmi and South West Terranes Before 4000 First crust formation on - Earliest Yilgarn crust preserved as detrital Earth zircons in the Narryer and Youanmi Terranes

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