Beyond the Confines of the Ore Deposit: Mapping Low Temperature Hydrothermal Alteration Above, Within, and Beneath Carlin‐type Gold Deposits by Ayesha Doris Ahmed B.Sc., The University of British Columbia, 2008 A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in The Faculty of Graduate Studies (Geological Sciences) THE UNIVERSITY OF BRITISH COLUMBIA (Vancouver) October 2010 ©Ayesha Doris Ahmed 2010 ii ABSTRACT Multiple analytical techniques were employed to investigate distal patterns in low temperature hydrothermal fluid flow into and out of Carlin‐type gold deposits in two study areas: the Leeville deposit and the Shoshone range including the Pipeline, Gold Acres and Elder Creek deposits. Previous studies indicate that gold is hosted in lower Paleozoic carbonate rocks overlain by thick sequences of similarly aged siliciclastic rocks. Patterns in δ18O depletion (<20‰VSMOW), and Au, As, Sb, Hg, Tl, and Te concentrations in lower Paleozoic carbonate rock identified three disconnected lateral fluid pathways into the Pipeline deposit: a main conduit providing gold‐bearing fluid to the main ore body, the Abyss fault located ~300m below the main ore zone, and the RMT located underneath the Abyss fault. Following gold precipitation in the Pipeline deposit, gold‐depleted fluids were likely exhausted laterally, at least initially, along the same structures as those that allowed fluid to enter the deposit. Upon intersecting the RMT fault, fluid either exploited the fault to reach surface, or transgressed overlying siliciclastic rocks via small scale faults and fractures. δ18O and δD values of H2O in equilibrium clay minerals, and the concentration and crystallinity of illite outlined multiple zones of hydrothermal alteration in surface rocks from both the Shoshone Range and Leeville study areas, however no genetic link was established to Carlin‐type gold mineralization at depth. Similarities in trace element geochemistry, ore assemblage, and alteration assemblages however, suggest that the Elder Creek deposit may represent low temperature (200°C) gold mineralization resulting from the exhaust of Carlin‐ type ore forming fluid. The region above the surface projection of the Leeville deposit exhibits multiple zones of hydrothermal fluid upflow resulting in pervasive illitization of surface siliciclastic rocks. The Pipeline/ Gold Acres also contain abundant crystalline illite. The presence of highly crystalline illite highlights zones of focused fluid upflow, typically along faults and other secondary permeability structures such as breccias. iii PREFACE Dr. K.H. Hickey identified and designed the research program. The research, sampling, and data accumulation were performed by A. Ahmed in consultation with K.H. Hickey, S.L. and Barker. Chapters 3 and 4 of the thesis are intended for publication in a scientific journal under the same title as each chapter respectively. A. Ahmed will be first author, K.A. Hickey will be second author, and S.L. Barker will be third author. Appendix A of this thesis is currently in press in the Geological Society of Nevada 2010 symposium proceedings, due to be published in November, 2010. iv TABLE OF CONTENTS ABSTRACT ...................................................................................................................................................... ii PREFACE ........................................................................................................................................................ iii TABLE OF CONTENTS .................................................................................................................................... iv LIST OF TABLES ........................................................................................................................................... viii LIST OF FIGURES ............................................................................................................................................. i LIST OF TERMS/ ABBREVIATIONS ................................................................................................................. iii ACKNOWLEDGEMENTS ................................................................................................................................ iv DEDICATION .................................................................................................................................................. v CHAPTER 1 – PROJECT OVERVIEW ............................................................................................................... 1 1.1 RATIONALE FOR STUDY ...................................................................................................................... 1 1.2 OVERVIEW OF CARLIN TYPE GOLD SYSTEMS ..................................................................................... 2 1.2.1 Carlin‐type gold deposits ............................................................................................................. 2 1.2.2 Tectonic framework .................................................................................................................... 4 1.3 THESIS OBJECTIVE ............................................................................................................................... 7 1.4 THESIS ORGANIZATION....................................................................................................................... 7 1.5 SUMMARY .......................................................................................................................................... 9 1.5 REFERENCES...................................................................................................................................... 10 CHAPTER 2‐ CLAY MINERALS AND LOW TEMPERATURE PROCESSES ........................................................ 14 2.1 INTRODUCTION ................................................................................................................................ 14 2.2 CLAY EQUILIBRIA............................................................................................................................... 14 2.3 ILLITE THERMOMETRY ...................................................................................................................... 17 2.4 ILLITE POLYTYPISM ........................................................................................................................... 18 2.5 ILLITE CRYSTALLINITY ........................................................................................................................ 18 2.6 CLAY TEXTURES AND CLAY MORPHOLOGY ...................................................................................... 20 2.7 CHEMICAL COMPOSITION OF ILLITE ................................................................................................. 20 2.8 SUMMARY ........................................................................................................................................ 24 2.9 REFERENCES...................................................................................................................................... 25 CHAPTER 3 – BEYOND THE CONFINES OF THE ORE BODY: SURFACE MAPPING OF LOW TEMPERATURE HYDROTHERMAL FLUID ABOVE MAJOR ORE BODIES USING CLAY ALTERATION ....................................... 28 3.1 INTRODUCTION ................................................................................................................................ 28 v 3.2 GEOLOGICAL SETTING OF CARLIN‐GOLD DEPOSITS ......................................................................... 32 3.3 CHARACTERISTICS OF CARLIN GOLD DEPOSITS ................................................................................ 36 3.3.1 Clay alteration in Carlin‐type systems ....................................................................................... 37 3.4 THERMAL SIGNATURE OF HYDROTHERMAL FLUID FLOW ............................................................... 37 3.5 ANALYTICAL TECHNIQUES ................................................................................................................ 42 3.5.1 Near and Short Wave Infrared Analysis (Terraspec©) .............................................................. 42 3.5.2 X‐ray diffraction ......................................................................................................................... 43 3.5.3 Scanning Electron Microscope (SEM) ........................................................................................ 44 3.5.4 Electron Microprobe Analysis ................................................................................................... 44 3.5.5 Stable Isotope Analysis .............................................................................................................. 45 3.6 SAMPLES ........................................................................................................................................... 45 3.7 RESULTS ............................................................................................................................................ 46 3.7.1 Morphology and textural relationships of clays ........................................................................ 46 3.7.2 Spatial distribution of clays ‐Leeville ......................................................................................... 50 3.7.3 Distribution of clays ‐ Shoshone Range Field Area .................................................................... 56 3.7.4 Comparing the use of XRD vs. Terraspec in the identification of clay minerals ....................... 60 3.7.5. Calculated temperature of illite formation .............................................................................. 63 3.7.6. Oxygen and Hydrogen Stable Isotope data .............................................................................. 64 3.8 INTERPRETATIONS/ DISCUSSION ...................................................................................................... 68 3.8.1 Clay mineral zonation patterns around hydrothermal fluid conduits ....................................... 68 3.8.2 Illite crystallinity halos ............................................................................................................... 70 3.8.3 Crystal morphology ................................................................................................................... 71 3.8.4 The Origin of Clay Minerals ....................................................................................................... 71 3.8.5 Challenges in using the K + ࡲࢋ െ ࡹࢍ thermometer ................................................................ 72 3.8.6 Clay morphology, crystallinity , and composition as a function of reaction progress .............. 73 3.9 IMPLICATIONS .................................................................................................................................. 75 3.9.1 The ability of analytical tools to identify alteration related to Carlin‐Au mineralization.......... 75 3.9.2 Hydrothermal flow on a regional scale ..................................................................................... 77 3.9.3 The pathway of exhausted fluids in Carlin‐type systems .......................................................... 78 3.10 CONCLUSIONS ................................................................................................................................ 78 3.11 REFERENCES ................................................................................................................................... 80 CHAPTER 4 – SHEDDING LIGHT ON THE ABYSS: LATERAL FLUID FLOW UNDERNEATH AND INTO CARLIN‐ TYPE GOLD DEPOSITS ................................................................................................................................. 89 vi 4.1 INTRODUCTION ................................................................................................................................ 89 4.2 GEOLOGICAL SETTING OF THE PIPELINE DEPOSIT ............................................................................ 90 4.2.1 Carlin‐type Au‐Mineralization ................................................................................................... 92 4.2.2 Previous isotope studies of the Pipeline Deposit ...................................................................... 93 4.3 FAULT RELATED FLUID FLOW IN CARLIN SYSTEMS AND THE ABYSS FAULT ..................................... 93 4.4 METHODS and ANALYTICAL TECHNIQUES ........................................................................................ 97 4.4.1 Background 18O, 13C, Au and trace element values for marine carbonate rocks .................. 99 4.5 RESULTS .......................................................................................................................................... 101 4.5.1 18O and 13C values of rocks .................................................................................................. 101 4.5.2 18O and 13C values of veins ................................................................................................... 110 4.5.3 Gold and trace element concentrations near the Abyss fault ................................................ 114 4.6 DISCUSSION .................................................................................................................................... 117 4.6.1 Origin of the 18O depletion in veins ....................................................................................... 120 4.7 IMPLICATIONS ................................................................................................................................ 122 4.8 CONCLUSION .................................................................................................................................. 124 4.9 REFERENCES.................................................................................................................................... 125 CHAPTER 5 –CONCLUSIONS...................................................................................................................... 130 5.1 FLUID FLOW INTO CARLIN‐TYPE GOLD DEPOSITS .......................................................................... 130 5.2 FLUID PATHWAYS OUT OF CARLIN‐TYPE GOLD DEPOSITS ............................................................. 130 5.3 ESTABLISHING VECTORS TOWARD CARLIN‐TYPE ORE ................................................................... 133 5.4 RECOMMENDATIONS FOR FUTURE WORK .................................................................................... 133 5.5 REFERENCES.................................................................................................................................... 135 APPENDIX A: THE ELDER CREEK DEPOSIT: An Upper Plate Expression of an Auriferous Carlin‐type hydrothermal system ............................................................................................................................... 136 INTRODUCTION .................................................................................................................................... 136 REGIONAL GEOLOGICAL SETTING......................................................................................................... 136 GEOLOGICAL SETTING OF THE ELDER CREEK MINE .......................................................................... 140 Clay mineralogy ................................................................................................................................ 144 Lithogeochemistry ............................................................................................................................ 147 METHODOLOGY .................................................................................................................................... 147 RESULTS ................................................................................................................................................ 149 Clay mineralogy ................................................................................................................................ 149 Lithogeochemistry ............................................................................................................................ 155 vii INTERPRETATIONS ................................................................................................................................ 157 Physiochemical nature of fluid ......................................................................................................... 157 Fluid flow pathways .......................................................................................................................... 157 DISCUSSION .......................................................................................................................................... 161 Deposit classification ........................................................................................................................ 163 CONCLUSIONS ...................................................................................................................................... 164 REFERENCES ......................................................................................................................................... 165 APPENDIX B– APATITE FISSION TRACK THERMOCHRONOLOGY DATA .................................................... 168 APPENDIX C – SAMPLE INFORMATION .................................................................................................... 173 APPENDIX D – X‐RAY DIFFRACTION PATTERNS ........................................................................................ 181 APPENDIX E – MICROPROBE DATA ........................................................................................................... 187 APPENDIX F ‐ ISOTOPIC VARIATION BETWEEN LITHOLOGICAL FORMATIONS ........................................ 197 viii LIST OF TABLES Table 2.1 Calibration chart for FWHM values……………………………………………………………………………… 19 Table 2.2 Regression line equations and correlation factors for K+[Fe‐Mg] thermometer…………. 22 Table 3.1 XRD data for the Leeville study area……………………………………………………………………………. 53 Table 3.2 Calibrated FWHM values for the Leeville study area……………………………………………………. 55 Table 3.3 XRD data for the Shoshone Range study area……………………………………………………………… 57 Table 3.4 Calibrated FWHM values for the Shoshone Range study area……………………………………… 59 Table 3.5 XRD vs. Terraspec accuracy…………………………………………………………………………………………. 61 Table 3.6 Averaged illite formation temperatures………………………………………………………………………. 64 Table 3.7 Stable isotope data……………………………………………………………………………………………………… 65 Table 4.1 Trace element geochemistry analytical procedures and detection limits…………………….. 99 Table 4.2 Stable isotope data and geochemistry for background samples from Lone Mtn…………. 100 Table 4.3 δ18O, δ13C, and trace element geochemistry data………………………………………………………… 102 Table A1 Geochemical data comparison between Elder Creek and Carlin‐type deposit…………….. 162 Table B1 Thermochronology data of the Northern Carlin trend…………………………………………………. 168 Table B2 Thermochronology data of the Shoshone Range…………………………………………………………. 170 Table C1 Sample information…………………………………………………………………………………………………….. 173 Table E1 Illite microprobe compositional data…………………………………………………………………………… 187 LIST OF FIGURES Fig. 1.1 Location of Carlin‐type gold deposits in the western United States……………………………….. 3 Fig. 1.2 Major deformation events affecting the western margin of North America…………………… 5 Fig. 2.1 Clay stability phase diagrams………………………………………………………………………………….......... 15 Fig. 2.2 End member morphologies of illite…………………………………………………………………………………. 21 Fig. 2.3 The derivation of the K +[Fe‐Mg] thermometer……………………………………………………………… 23 Fig. 3.1 The distal extent and degree of alteration around ore deposits……………………………………… 29 Fig. 3.2 Regional geology map of north‐eastern Nevada…………………………………………………………….. 32 Fig. 3.3 Geological map and cross section of the Leeville study area…………………………………………… 33 Fig. 3.4 Simplified tectono‐stratigraphic column of the RMT system………………………………………….. 34 Fig. 3.5 Potential fluid evolution pathways…………………………………………………………………………………. 38 Fig. 3.6 Apatite fission track ages across the Northern Carlin trend……………………………………………. 40 Fig. 3.7 Apatite fission track ages across the Shoshone Range……………………………………………………. 41 Fig. 3.8 Sample types………………………………………………………………………………………………………………….. 47 Fig. 3.9 Sample location map from the Leeville study area………………………………………………………….. 48 Fig. 3.10 Scanning electron microphotographs showing the morphology of clay minerals……………. 49 Fig. 3.11 Scanning electron microphotographs showing mineral‐mineral relationships…………........ 51 Fig. 3.12 XRD, Terraspec, FWHM, and microprobe data from across the Leeville area…………………… 52 Fig. 3.13 XRD, Terraspec, and FWHM data from across the Shoshone Range .………………………………. 58 Fig. 3.14 Causes for discrepancies between XRD and Terraspec data……………………………………………. 62 Fig. 3.15 Graph of average potassium content of illite and average oxide totals……………………........ 66 Fig. 3.16 Stable isotope data from this study along with data from other Carlin‐type deposits…….. 67 Fig. 3.17 Zones of hydrothermal alteration identified in the Leeville study area……………………………. 69 Fig. 3.18 Trends in illite compositional data from previous studies……………………………………………….. 74 Fig. 4.1 Regional map of northeastern Nevada with inset of the Pipeline deposit……………………….. 91 Fig. 4.2 Cross section A‐A1 through the Abyss fault………………………………………………………………………. 95 Fig. 4.3 Photographs of core above, at and below the Abyss fault……………………………………………….. 96 Fig. 4.4 Drill hole and cross‐section location map………………………………………………………………………… 98 Fig. 4.5 δ18O and δ13C data…………………………………………………………………….……………………………………. 109 Fig. 4.6 Cross sections through the Pipeline deposit show δ18O depletion zones……………………….. 111 Fig. 4.7 δ18O values relative to the Abyss fault……………………………………………………………………………. 112 ii Fig. 4.8 δ13C values relative to the Abyss fault…………………………………………………………………………….. 113 Fig. 4.9 Trace element geochemistry concentrations relative to the Abyss fault………………………… 115 Fig. 4.10 Potential fluid pathways into the Pipeline ore deposit………………………………………………….. 121 Fig. A1 Regional Geology map of northeastern Nevada……………………………………………………………. 138 Fig. A2 Simplified stratigraphic column of lower Palaeozoic rocks…………………………………………….. 139 Fig. A3 Geology of the Elder Creek deposit……………………………………………………………………………….. 141 Fig. A4 Photograph of sample 456…………………………………………………………………………………………….. 142 Fig. A5 Field photographs of strongly argillized zones………………………………………………………………. 143 Fig. A6 Fluid evolution pathways………………………………………………………………………………………………. 145 Fig. A7 Phyllosilicate phase diagrams………………………………………………………………………………………… 146 Fig. A8 Sample location map…………………………………………………………………………………………………….. 148 Fig. A9 Two XRD standards………………………………………………………………………………………………………. 150 Fig. A10 X‐ray diffraction patterns…………………………………………………………………………………………….. 152 Fig. A11 Hydrothermal illite textures…………………………………………………………………………………………. 153 Fig. A12 Illite crystallinity vs. peak position graph……………………………………………………………………… 154 Fig. A13 Gold ordered trace element diagrams………………………………………………………………………….. 156 Fig. A14 Geochemical halos……………………………………………………………………………………………………….. 158 Fig. A15 Stability field clay phase diagram………………………………………………………………………………….. 160 Fig. F1 Stable isotope value variation with lithology………………………………………………………………… 197 iii LIST OF TERMS/ ABBREVIATIONS ax: activity CIS: crystallinity index standard (a standardized Kubler Index) FWHM: Full width at half the maximum value (at given 2Θ location of an x‐ray diffraction pattern) I‐S: illite‐smectite interlayered clay Lower plate: Lower Paleozoic shelf and slop carbonate rocks that form the footwall to the Roberts Mountains Thrust fault and are the typically hosts for the Carlin‐type gold mineralization PASW: Predictive Analytics Software (an IBM product) PIMA: Portable Infrared Mineral Analyzer RC: Reverse circulation (drill hole) RMT : Roberts Mountains Thrust Fault VPDB: Versus Peedee Belminite (international standard for carbon isotopes) VSMOW: Versus Standard Mean Ocean Water (international standard for oxygen isotopes) Upper Plate: Lower Paleozoic shallow marine and basinal siliciclastic sediments that form the hangingwall to the Roberts Mountains Thrust fault, and can form thick sequences of cover on top of Carlin‐type gold systems. δ = ratio of the heavy isotope (13C or 18O) vs. the light isotope (12C or 16O) ‰ : per mil (stable isotope measurement relative to an international standard) iv ACKNOWLEDGEMENTS This MSc. thesis is part of a larger project coordinated by Dr. Kenneth Hickey, associate professor, University of British Columbia, Mineral Deposit Research Unit entitled ‘Thermal and geochemical footprints of low‐temperature sedimentary rock‐hosted hydrothermal Au‐systems: Identifying far‐field vectors toward ore’. My sincere thanks to those companies and organizations that provided financial support for this project: Barrick Gold Corporation, Newmont Mining Corporation, and Teck Limited with matching funds provided by a Collaborative Research and Development grant from the Natural Sciences and Engineering Research Council. A special thanks to Kevin Creel, Bob Leonardson, Nancy Richter, and Joe Becker for providing support at the Cortez mine. Thank you to the Society of Economic Geologists for providing financial support through scholarships and conference funding. My sincere thanks to Dr. Kenneth Hickey for his continuing support and mentorship. I will fondly remember luring Ken into the office for some quality thesis review time with the help of a either a chocolate bar or biscuit, or both. As busy as he may have been, he always made the time to meet and provide constructive criticism; a truly brilliant man. I thank my committee members Dr. Greg Dipple, Dr. Dominique Weis, and Dr. Shaun Barker for their interest and guidance. I appreciated their feedback during committee meetings and their commitment to keeping my work within the scope of a Master’s project. I am especially grateful for the constant support of Dr. Barker who will be happy never to have to edit anything else I write. I could not have completed any SEM, XRD, or microprobe work without the help of Mati Raudsepp, Edith Czech, Jenny Lai, and Elisabetta Pani. For my work with isotopes, I thank Janet Gabites for sample preparation and data reduction. I am indebted to my fellow lab‐mates Moira Cruickshanks, Will Lepore, and Jeremy Vaughan who provided countless hours of discussion, both thesis related and otherwise. Thanks to good friends Jean Francois Blanchette Guertin, Jaime Poblete, Santiago Vaca, Bram Van Straaten, Esther Bordet, and Tatiana Alva for much needed evening distractions. To Bram and the rest of the ‘orphanage’, thank you for welcoming me into my second home. A shout out to my best friends Devon and Christa for providing a link to a world unrelated to geology. Last and far from least, I will be forever indebted to Shawn Hood whose patience and support was and is never ending. And of course to my family. To both my parents for giving me a desire to never stop learning and for inspiring me to pursue higher education. To my mama for always understanding that in times of despair, it is comforting to hear that you’re right, even when you’re not, and for feeling and experiencing my troubles as much or more than I did. To my Abbu for always providing the extra push, motivation and support to keep me on the right track. And to my special Xantha, who ever since we were young has been a role model for her sense of adventure, and her lack of fear for new experiences and new challenges. v DEDICATION To Mutti, Abbi, and Gogo (+/- Snooks) To Shawny 1 CHAPTER 1 – PROJECT OVERVIEW 1.1 RATIONALE FOR STUDY Exploration for new economic mineral resources in mature mineral terranes has become increasingly difficult and deposits currently exposed at surface have either been discovered or categorized as sub‐economic (Kelley et al., 2006). Recent research has highlighted the importance of discovering new exploration methods and improving on existing exploration methods that look beyond the obvious limits of mineralization to the distal expression of mineralizing systems (Adams and Putnam, 1992; Arehart and Donelick, 2006; Kelley et al., 2006). This task is complicated in the context of low temperature hydrothermal systems (<300C) where conventional alteration or mineral mapping is difficult to employ. Low temperature systems such as active geothermal systems (Simmons and Browne, 2000; Yang et al., 2001), and Carlin‐type gold deposits (Cline et al., 2005) have a tendency to exhibit more subtle alteration halos around ore deposits than their high temperature counterparts such as porphyry deposits (Rose, 1970). This is primarily a function of differences in both thermal and chemical gradients between mineralizing fluid and host rock (Reed, 1982, 1997). Low temperature hydrothermal systems intruding the shallow crust tend to lack large thermal and chemical gradients between hot rock (~50C at 2km given a geothermal gradient of 25C/km) and low temperature (~200C) rock or surface water buffered fluid. Furthermore, the alteration minerals produced from low temperature fluid‐rock interactions are generally fine grained phyllosilicate minerals which are difficult to both identify and analyze, and may not represent equilibrium assemblages (Essene and Peacor, 1995). Identifying subtle expressions of hydrothermal alteration around low temperature mineral deposits may provide robust vectors toward mineralization at depth. These expressions provide insights into to the nature and extent of alteration in low temperature systems that help to further existing models for mineralizing processes. This MSc. thesis is part of a larger project coordinated by Dr. Kenneth Hickey, University of British Columbia, Earth and Ocean Sciences Department, Mineral Deposit Research Unit entitled ‘Thermal and geochemical footprints of low‐temperature sedimentary rock‐hosted hydrothermal Au‐ systems: Identifying far‐field vectors toward ore’. The project is sponsored by Barrick Gold Corporation, Newmont Mining Corporation, and Teck Limited with matching funds provided by a Collaborative Research and Development grant from the Natural Sciences and Engineering Research Council. The project combines a well‐constrained geological understanding of the paleogeographic, tectonic and magmatic environment of gold deposition using a range of thermochronometers, and lithogeochemical, 2 isotopic and mineralogical tracers to delineate the location and scale of low temperature hydrothermal fluid circulation that resulted in Carlin‐type gold deposition, as well as identifying where these deposits are manifested under cover. 1.2 OVERVIEW OF CARLIN TYPE GOLD SYSTEMS 1.2.1 Carlin‐type gold deposits The world‐class Carlin‐type gold deposits of northeastern Nevada provide 9% of the world’s gold production (Cline et al., 2005; Price et al., 2008). The majority of deposits occur along three structural lineaments shown in Figure 1.1: the Carlin trend, the Battle Mountain Eureka trend, and the Getchell trend. The majority of gold is ‘invisible’, hosted within arsenian‐rich pyrite, although the oxidized portion of deposits can contain free gold (Barker et al., 2009; Bettles, 2002). Gold is disseminated and occurs predominantly within lower Paleozoic silty carbonate shelf and slope rocks (Cline et al., 2005). Gold precipitated via sulfidation where a low temperature, slightly acidic fluid reacted with ferrous‐carbonate host rocks (Cline and Hofstra, 2000; Lubben, 2004). Many Carlin‐type deposits are covered by thick sequences of Paleozoic outer shelf and basin siliciclastic rocks which are largely devoid of mineralization (Cline et al., 2005). The namesake for Carlin‐type gold deposits, the ‘Carlin deposit’ is located along the Northern Carlin trend, north‐eastern Nevada. Subsequent to the discovery of the Carlin gold depoist in the 1960’s, a number of other deposits of similar style were discovered in the area (Cline and Hofstra, 2000). Since this type of deposit had not been documented anywhere else, a new classification of hydrothermal ore‐ deposits was created and termed ‘Carlin‐type’. Carlin‐type gold deposits are typically restricted to the shallow crust (1‐4km below surface). The source of fluid and metals is strongly debated. Fluid inclusion and mineral thermometry data indicate that mineralizing fluid forming Carlin deposits in northeastern Nevada was: low temperature (180‐240C), slightly acidic (pH ~4), low‐salinity (~2–3 wt% NaCl equivalent), aqueous fluids that contained CO2 (<4 mol %), and CH4 (<0.4 mol %), and sufficient H2S (10– 1–10–2 m) to transport Au and other bisulfide‐complexed metals (Cline and Hofstra, 2000; Lubben, 2004). No fluid inclusion, mineralogical, or textural evidence exists to indicate fluid boiling or immiscibility. Alteration associated with Carlin‐type mineralization is subtle and includes: pre‐ore 3 Fig.1.1. (Modified from Hofstra et al., 1999) Dots show locations of Carlin‐type gold deposits in the western United States. Deposits occur principally along three structural lineaments: the Battle Mountain – Eureka trend, the Carlin trend, and the Getchell trend. The alignment of Carlin‐type deposits reflects major basement fault fabrics which were established during Neoproterozoic rifting (Roberts, 1966; Tosdal et al., 2000). 4 decalcification, argillization, sulfidation and local silicification. Clay minerals observed in the argillic alteration assemblage include illite, dickite, and kaolinite (Ilchik, 1990; Clode et al., 1997; Folger et al., 1998; Hofstra et al., 1999; Cline and Hofstra, 2000; Cail and Cline, 2001). Thermochronology based estimates of bedrock exhumation suggest that the Carlin‐type Au‐deposits likely formed over a paleodepth range of <1‐4 km (Cline et al., 2005). DH2O values measured on hypogene kaolinite, illite and fluid inclusions from a wide‐range of Carlin deposits suggest evidence of meteoric water with very low DH2O values of < ‐110 ‰ (Hofstra, 1999). Geochemical data and mineral stability also suggest a highly exchanged meteoric source fluid (Ilchik and Barton, 1997). Stable isotope data from fluid inclusions and clay minerals of the the Getchell, Twin Creeks, and Deep Star deposits, however, indicate the additional presence of deeply sourced metamorphic or magmatic fluid (Hofstra et al., 1999; Heitt et al., 2003; Cline et al., 2005). A spatial association between some Carlin‐type deposits and Eocene intrusive centers has also led to the interpretation that Carlin‐type deposits are distal portions of magmatic hydrothermal systems (Radtke et al., 1980; Arehart et al., 1993). Gold is typically hosted within Siluro‐Devonian miogeoclinal carbonate rocks that form the footwall to the Roberts Mountains Thrust (RMT) (the “lower plate”). The structurally overlying “upper plate” is dominated by Ordovician‐Mississippian siliciclastic eugeoclinal rocks (Roberts, 1966). The upper plate is not known to host any major Carlin‐type Au deposits although mineralization does occur locally immediately above the RMT in several of the lower plate‐hosted deposits. The lack of Carlin‐type deposits in the upper plate is thought to reflect the less reactive nature of the siliciclastic rocks that dominate the Roberts Mountains allochthon (Cline et al., 2005). 1.2.2 Tectonic framework Figure 1.2 presents a timeline of deformation events leading to the formation of Carlin‐type gold deposits in Nevada. The western margin of continental North America records a history of prolonged episodic rifting and basin subsidence from the Neoproterozoic to Paleozoic as the proto‐Pacific oceanic basin opened following the breakup of Rodinia to form the paleo‐continent Laurentia (Powell et al., 1993; Wingate and Giddings, 2000; Lund, 2008). Geochronologic data indicate the presence of two main rifting events beginning in the Cryogenian and accelerated in the Ediacaran to Cambrian (Stewart, 5 Fig. 1.2. (Modified from Hofstra et al., 1999) Timeline showing the major deformation events affecting the western margin of continental North America. PC = pre‐Cambrian C= Cambrian O = Ordovician D= Devonian M = Mississippian IP = Pensylvannian P = Permian Tr = Triassic J= Jurassic K= Cretaceous T = Tertiary. 6 1972; Thompson et al., 1987; Ross, 1991; Colpron et al., 2002). The earlier rifting event corresponds to the deposition of coarse, partly glaciogenic diamictite and mafic volcanic rocks of the Windermere Supergroup on top of an intercontinental rift margin in British Columbia (Ross, 1991; Colpron et al., 2002). Later rifting in the Neoproterozoic is interpreted to indicate continental breakup and establishment of a passive margin along the western margin of Laurentia (Colpron et al., 2002). Neoproterozoic and Early Cambrian clastic rocks across central Nevada are dominated by quartzite with interstratified argillite and phyllite, and were deposited in a westward‐thickening sedimentary wedge over structurally complex, thinned Paleoproterozoic to Neoproterozoic crystalline basement (Poole et al., 1992). During the Early‐ Middle Devonian, basins formed in the shelf and outer shelf regions of the passive margin, probably as a result of reactivation of Neoproterozoic rift structures in the underlying crystalline basement (Stewart, 1972; Morrow and Sandberg, 2008). Following Neoproterozoic rifting and initial sedimentation, miogeoclinal and eugeoclinal sequence sediments were deposited on top of Cambrian clastic rocks, along the western margin of continental North America (Burchfiel and Davis, 1972). Interbedded carbonate and shale, and silty carbonate were deposited along the shelf and slope, with basinal and shallow marine siliciclastic sediments deposited further west (Burchfiel and Davis, 1972; Morrow and Sandberg, 2008). Subsequent to deposition, the Paleozoic passive margin sequence was subject to multiple episodes of contractional deformation occurring from middle Mississppian to Early Permian (Cashman et al., 2008). During the Late Devonian – Early Mississippian Antler orogeny, lower Paleozoic basin and slope rocks (locally termed ‘upper plate’) were thrust eastward over coeval shelf‐margin and outer‐shelf rocks (lower plate) forming the RMT system (Johnson and Pendergast, 1981; Poole et al., 1992). Lower plate carbonate rocks are the typical host for gold mineralization in Carlin‐type gold deposits. Upper plate rocks were eventually overlain by Mississippian to Permian shelf deposits; the ‘overlap sequence’ described by Roberts et al., (1958). The Late Mississippian to Middle Pennsylvanian Humboldt orogeny resulted in uplift and later subsidence of the overlap sequences. Rocks of the Golconda allochthon were thrust eastward over upper plate and overlap sequence rocks during the Late Permian and Early Triassic Sonoma orogeny (Miller, 1984). Starting in the Late Triassic and continuing through to the Tertiary, the western margin of North America became the site of semi‐continuous east‐directed subduction (Speed et al., 1988) leading to the development of the Cordilleran orogenic belt. Further west, subduction was marked by episodes of thin‐skinned contractional deformation with accompanying extensional faulting and magmatism. 7 Two periods of extension affected Northeastern Nevada during the Cenozoic. Extension commenced in the late Paleogene with the onset of regional magmatism (Wernicke et al., 1987; Christiansen et al., 1992; Sonder and Jones, 1999). The spatial and temporal overlap of Carlin‐type deposits with the onset of Cenozoic volcanism and extension suggests a fundamental link between these phenomena (Hofstra et al., 1999; Cline et al., 2005). Rb/Sr dating of the syn‐ore mineral galkhaite indicates a mineralization age of 38‐40 Myr (Tretbar et al., 2000; Arehart et al., 2003; Cline et al., 2005). Subsequent to Carlin‐type gold deposition, heterogeneous extension of the Great Basin, accompanied by magmatism continued into the Oligocene and ended with the mid Miocene development of the Basin and Range province and the Northern Nevada Rift; a series of north‐northwest‐striking mafic dikes and high angle normal faults, basaltic volcanic flow units related to the dikes and epithermal, volcanic‐hosted mineral deposits (Wernicke et al., 1987; Zoback et al., 1994). 1.3 THESIS OBJECTIVE The goal of this thesis is to investigate and answer two scientific questions: Question 1: Can we identify fluid flow patterns into and out of low temperature hydrothermal systems by means of thermal and/ or chemical alteration halos, and if so how is fluid flow manifested in host rocks? Question 2: What set of tools and analytical techniques may provide vectors toward regions of hydrothermal flow most likely to have precipitated economic grades of mineralization? These two questions are answered with two complimentary studies: (i) by identifying a thermal or chemical alteration halo in upper plate siliciclastic rocks resulting from interaction with exhausted Carlin‐type hydrothermal fluid (ii) by investigating alteration within and above a major fault underlying a giant Carlin‐type deposit to identify large scale fluid flow pathways. The results contained in this thesis further the understanding of fluid rock interaction in Carlin‐type environments and provide information on the practical application of tools and analytical techniques related to identifying low temperature hydrothermal alteration. 1.4 THESIS ORGANIZATION This thesis is arranged in four chapters. Chapter 1 introduced the concepts outlined in this study by presenting a literature review of Carlin‐type gold systems, and the alteration associated with low temperature hydrothermal systems. Chapter 2 provides background information on low temperature clay thermometers. Chapters 3 and 4 are written as stand‐alone manuscripts to be submitted to an international scientific journal for publication, and adhere to the formatting protocol of Economic Geology. 8 Chapter 3 presents a study that tested for the presence of clay alteration phases above regions of known mineralization. A regional study of the Shoshone Range and a deposit scale study of the Leeville deposit (both in northeastern Nevada) investigated the relationship between argillization at surface in upper plate siliciclastic rocks, and argillization typically associated with Carlin‐type Au mineralization at depth. The identification of a hydrothermal clay mineral assemblage at surface may provide insight into the way in which fluid was exhausted following mineralization of Carlin‐type deposits. Identifying zones of surface exhaust provides a direct link to ore deposits at depth. Detailed results of this study include clay mineral zonation patterns, illite formation temperatures, illite crystallinity, textural and morphological analyses of clays, and O and H stable isotope analyses of clay minerals. Results from this chapter indicate that extensive zones of hydrothermal clay alteration at surface are spatially associated to Carlin‐type gold deposits at depth. The chapter concludes with a critical analysis of the results achieved using two different analytical techniques, x‐ray diffraction and near and short wave spectroscopy, to identify clay minerals. Chapter 4 of this thesis extends the work of Arehart and Donelick (2006) and Rye (1995) who studied the stable isotopic depletion signature above and within the giant Pipeline gold deposit, northeastern Nevada. This chapter focuses on defining and identifying fluid pathways leading into Carli‐ type gold systems by studying the Abyss fault; a large thrust structure that separates gold mineralized lower plate on top of largely unmineralized upper plate below the Pipeline deposit. Results are presented from a C and O stable isotope study on carbonate rocks within the hangingwall of the fault to determine whether the Abyss fault acted as a fluid conduit, or whether it played a passive role during mineralization of the giant deposit. Those results indicate that ore‐forming fluid flow did occur along the Abyss fault however the fault was not the main conduit for fluid forming the Pipeline gold deposit. Furthermore, fluid pathways within the Pipeline deposit were dominantly lateral, not vertical. A case study is presented in Appendix A of this thesis of the Elder Creek deposit; a small, sediment hosted disseminated gold deposit in the upper plate siliciclastic rocks of the Shoshone Range. Information is used from chapters 2, 3, and 4 to discuss the potential for Carlin‐type gold mineralization in allochthonous siliciclastic rocks (upper plate). Characteristics of the Elder Creek deposit are considered relative to known characteristics of Carlin‐type gold deposits including trace‐element geochemistry, clay alteration, temperature, and ore mineral assemblages. The study of the Elder Creek deposit has been accepted by the Geological Society of Nevada for publication in the 2010 GSN Symposium proceedings. 9 Although chapters 3 and 4 focus on different aspects of the same scientific questions outlined in section 1.3, each chapter has been prepared as a stand‐alone paper resulting in some inevitable repetition and overlap between the chapters. 1.5 SUMMARY Exploration for low temperature hydrothermal ore systems can be hindered by a lack of visible alteration. Carlin‐type deposits are examples of low temperature hydrothermal systems exhibiting subtle alteration associated with mineralization. In this thesis, we discuss the presence of both a thermal and chemical alteration below, within, above and outboard from large low temperature hydrothermal systems by studying variations in low temperature clay mineral assemblages in siliciclastic rocks and oxygen and carbon isotope depletion in carbonate rocks. In addition, we investigate the role of largescale faults in the process of mineralization to place constraints on pathways for ore forming fluids. 10 1.5 REFERENCES Adams, S.S., and Putnam, B.R., III, 1992, Application of mineral deposit models in exploration: a case study of sediment‐hosted gold deposits, Great Basin, Western United States: Geological Society, London, Special Publications, v. 63, p. 1‐23. Arehart, G., and Donelick, R., 2006, Thermal and isotopic profiling of the Pipeline hydrothermal system: Application to exploration for Carlin‐type gold deposits: Journal of Geochemical exploration, v. 91, p. 27‐40. Arehart, G., Chakurian, A., Tretbar, D., Christensen, J., McInnes, B., and Donelick, R., 2003, Evaluation of radioisotope dating of Carlin‐type deposits in the Great Basin, western North America, and implications for deposit genesis: Economic Geology, v. 98, p. 235‐248. Arehart, G., Eldridge, C., Chryssoulis, S., and Kesler, S., 1993, Ion microprobe determination of sulfur isotope variations in iron sulfides from the Post/Betze sediment‐hosted disseminated gold deposit, Nevada, USA: Geochimica et Cosmochimica Acta, v. 57, p. 1505‐1519. Barker, S., Hickey, K., Cline, J., Dipple, G., Kilburn, M., Vaughan, J., and Longo, A., 2009, Uncloaking Invisible Gold: Use of NanoSIMS to Evaluate Gold, Trace Elements, and Sulfur Isotopes in Pyrite from Carlin‐Type Gold Deposits: Economic Geology, v. 104, p. 897‐904. Bettles, K. 2002, Exploration and geology, 1962–2002 at the Goldstrike property, Carlin Trend, Nevada: Economic Geology Special Publication 9, p. 275–298. Burchfiel, B., and Davis, G., 1972, Structural framework and evolution of the southern part of the Cordilleran orogen, western United States: American Journal of Science, v. 272, p. 97‐118. Cail, T., and Cline, J., 2001, Alteration associated with gold deposition at the Getchell Carlin‐type gold deposit, north‐central Nevada: Economic Geology, v. 96, p. 1343‐1359. Cashman, P., Trexler, J., Snyder, W., Davydov, V., and Taylor, W., 2008, Late Paleozoic deformation in central and southern Nevada: GSA Field Guides, v. 11, p. 21‐43. Christiansen, R., Yeats, R., Graham, S., Niem, W., Niem, A., and Snavely, P., 1992, Post‐Laramide geology of the US Cordilleran region: The Cordilleran Orogen: Conterminous US G‐3, Geological Society of America, Boulder, Colorado, p. 261–406. Cline, J., and Hofstra, A., 2000, Ore‐fluid evolution at the Getchell Carlin‐type gold deposit, Nevada, USA: European Journal of Mineralogy, v. 12, p. 195‐212. Cline, J., Hofstra, A., Muntean, J., Tosdal, R., and Hickey, K., 2005, Carlin‐type gold deposits in Nevada: Critical geologic characteristics and viable models: Economic Geology, 100th Anniversary Volume, p. 451‐484. Clode, C., Grusing, S., Heitt, D., and Johnston, I., 1997, The relationship of structure, alteration, and stratigraphy to formation of the Deep Star gold deposit: Eureka County, Nevada: Society of Economic Geologists Guidebook Series, v. 28, p. 239–256. 11 Colpron, M., Logan, J., and Mortensen, J., 2002, U‐Pb zircon age constraint for late Neoproterozoic rifting and initiation of the lower Paleozoic passive margin of western Laurentia: Canadian Journal of Earth Sciences, v. 39, p. 133‐143. Essene, E., and Peacor, D., 1995, Clay mineral thermometry‐a critical perspective: Clays and Clay Minerals, v. 43, p. 540‐553. Folger, H., Hofstra, A., Eberl, D., and Snee, L., 1998, Importance of clay characterization to interpretation of 40 Ar/39 Ar dates of illite from Carlin‐type gold deposits: Insights from Jerritt Canyon: Contributions to the Gold Metallogeny of Northern Nevada, ed. Tosdal, RM, USGS Open File Rept, p. 98–338. Heitt, D., Dunbar, W., Thompson, T., and Jackson, R., 2003, Geology and geochemistry of the Deep Star gold deposit: Carlin trend, Nevada: Economic Geology, v. 98, p. 1107‐1135. Hofstra, A., Snee, L., Rye, R., Folger, H., Phinisey, J., Loranger, R., Dahl, A., Naeser, C., Stein, H., and Lewchuk, M., 1999, Age constraints on Jerritt Canyon and other carlin‐type gold deposits in the Western United States; relationship to mid‐Tertiary extension and magmatism: Economic Geology, v. 94, p. 769‐802. Ilchik, R., and Barton, M., 1997, An amagmatic origin of Carlin‐type gold deposits: Economic Geology, v. 92, p. 269‐288. Ilchik, R., 1990, Geology and geochemistry of the Vantage gold deposits: Alligator Ridge‐Bald Mountain mining district, Nevada: Economic Geology, v. 85, p. 50–75. Kelley, D., Kelley, K., Coker, W., Caughlin, B., and Doherty, M., 2006, Beyond the obvious limits of ore deposits: the use of mineralogical, geochemical, and biological features for the remote detection of mineralization: Economic Geology, v. 101, p. 729‐752. Johnson, J., and Pendergast, A., 1981, Timing and mode of emplacement of the Roberts Mountains allochthon: Antler orogeny: Geological Society of America Bulletin, v. 92, p. 648‐658. Lubben, J., 2004, Quartz as clues to paragenesis and fluid properties at the Betze‐Post deposit, northern Carlin trend, Nevada: Unpublished M.Sc. thesis, Las Vegas, University of Nevada, p. 155. Lund, K., 2008, Geometry of the Neoproterozoic and Paleozoic rift margin of western Laurentia: Implications for mineral deposit settings: Geosphere, v. 4, p. 429‐444. Miller, E., Holdsworth, B., Whiteford, W., and Rodgers, D., 1984, Stratigraphy and structure of the Schoonover sequence, northeastern Nevada: Implications for Paleozoic plate‐margin tectonics: Bulletin of the Geological Society of America, v. 95, p. 1063‐1076. Poole, F., Stewart, J., Palmer, A., Sandberg, C., Madrid, C., Ross Jr, R., Hintze, L., Miller, M., and Wrucke, C., 1992, Latest Precambrian to latest Devonian time; development of a continental margin: The Cordilleran Orogen: Conterminous US, p. 9–54. 12 Powell, C., Li, Z., McElhinny, M., Meert, J., and Park, J., 1993, Paleomagnetic constraints on timing of the Neoproterozoic breakup of Rodinia and the Cambrian formation of Gondwana: Geology, v. 21, p. 889‐892. Price, J.G. et al., 2008, The Nevada Minerals Industry 2007: Nevada Bureau of Mines and Geology Special Publication MI‐2007. Reed, M., 1982, Calculation of multicomponent chemical equilibria and reaction processes in systems involving minerals, gases and an aqueous phase: Geochimica et Cosmochimica Acta, v. 46, p. 513‐528. —, 1997, Hydrothermal alteration and its relationship to ore fluid composition: Geochemistry of hydrothermal ore deposits, p. 303–365. Roberts, R., 1966, Metallogenic provinces and mineral belts in Nevada: Nevada Bureau of Mines, Rept, v. 13, p. 47‐72. Roberts, R., Hotz, P., Gilluly, J., and Ferguson, H., 1958, Paleozoic rocks of north‐central Nevada: Am: Assoc. Petroleum Geologists Bull, v. 42, p. 2813‐2857. Rose, A., 1970, Zonal relations of wallrock alteration and sulfide distribution at porphyry copper deposits: Economic Geology, v. 65, p. 920‐936. Ross, G., 1991, Tectonic setting of the Windermere Supergroup revisited: Geology, v. 19, p. 1125‐1128. Rye, R., 1995, A model for the formation of carbonate‐hosted disseminated gold deposits based on geologic, fluid inclusion, geochemical, and stable isotope studies of the Carlin and Cortez deposits, Nevada: Nevada: US Geological Survey Bulletin, v. 1646, p. 35–42. Simmons, S., and Browne, P., 2000, Hydrothermal minerals and precious metals in the Broadlands‐ Ohaaki geothermal system: Implications for understanding low‐sulfidation epithermal environments: Economic Geology, v. 95, p. 971‐999. Sonder, L., and Jones, C., 1999, Western United States extension: How the west was widened: Annual Review of Earth and Planetary Sciences, v. 27, p. 417‐462. Speed, R., Elison, M.W., and Heck, F.R., 1988, Phanerozoic tectonic evolution of the Great Basin, in Ernst., W.G., ed., Metamorphism and crustal evolution of the western United States, Volume 7: New Jersey, Prentice Hall. Stewart, J., 1972, Initial deposits in the Cordilleran geosyncline: Evidence of a late Precambrian (< 850 my) continental separation: Geological Society of America Bulletin, v. 83, p. 1345‐1360. Thompson, B., Mercier, E., and Roots, C., 1987, Extension and its influence on Canadian Cordilleran passive‐margin evolution: Geological Society London Special Publications, v. 28, p. 409‐417. Tretbar, D., Arehart, G., and Christensen, J., 2000, Dating gold deposition in a Carlin‐type gold deposit using Rb/Sr methods on the mineral galkhaite: Geology, v. 28, p. 947‐950. Warr, L., 1996, Standardized clay mineral crystallinity data from the very low‐grade metamorphic facies rocks of southern New Zealand: European Journal of Mineralogy, v. 8, p. 115‐127. 13 Wernicke, B., England, P., Sonder, L., and Christiansen, R., 1987, Tectonomagmatic evolution of Cenozoic extension in the North American Cordillera: Geological Society London Special Publications, v. 28, p. 203‐221. Wingate, M., and Giddings, J., 2000, Age and palaeomagnetism of the Mundine Well dyke swarm, Western Australia: implications for an Australia‐Laurentia connection at 755 Ma: Precambrian Research, v. 100, p. 335‐357. Yang, K., Browne, P., Huntington, J., and Walshe, J., 2001, Characterising the hydrothermal alteration of the Broadlands‐Ohaaki geothermal system, New Zealand, using short‐wave infrared spectroscopy: Journal of Volcanology and Geothermal Research, v. 106, p. 53‐65. Zoback, M., McKee, E., Blakely, R., and Thompson, G., 1994, The northern Nevada rift: Regional tectono‐ magmatic relations and middle Miocene stress direction: Geological Society of America Bulletin, v. 106, p. 371‐382. 14 CHAPTER 2‐ CLAY MINERALS AND LOW TEMPERATURE PROCESSES 2.1 INTRODUCTION The purpose of this chapter is to provide background on clay minerals, and their use as indicators of hydrothermal alteration based on different characteristics. The techniques described in this chapter are applied to the surface exploration for Carlin‐type gold systems in Chapter 3. As will be shown in this chapter, clay mineral zonation patterns observed around hydrothermal systems can reflect gradients in temperature, fluid composition (e.g., acidity), and reaction progress. Clay mineral assemblages have been used as proxies for temperature and fluid composition based on observations that show that clay minerals predictably and repeatedly undergo the same sequence of transformations with increasing grade of diagenesis, metamorphism, and changes in temperature and fluid composition (Essene and Peacor, 1995; Simmons and Browne, 2000). Observations relating changes in clay minerals to temperature are based on both empirical calibrations as observed from natural clay‐bearing systems (Rose, 1970; Simmons and Browne, 2000; Battaglia, 2004) and experimental observations (Whitney and Northrop, 1988; Yates and Rosenberg, 1997; Bauer et al., 2000). 2.2 CLAY EQUILIBRIA Figure 2.1 illustrates the different temperatures and fluid compositions at which illite, kaolinite, and smectite are stable resolved from solution equilibration experiments conducted from 100‐250C (Yates and Rosenberg, 1997). As temperature increases, water in the interlayer site of smectite is replaced by potassium. The loss of interstitial water causes the crystal structure of smectite to become more ordered resulting in the reaction smectite illite‐smectite interlayered clay (I‐S) illite muscovite (Lanson and Champion, 1991; Lanson et al., 1998). Well ordered end‐member illite and muscovite contain no expandable phases (H2O, H3O+) (Yates and Rosenberg, 1997). At 200‐250°C pure, end‐ member illite, lacking any expandable phases (no H2O in the interlayers) appears at low aH4SiO4 values and smectite disappears. The smectite field, which occurs at higher aH4SiO4 values than illite is replaced by I‐S. Above 200C, and at high aK+/aH+ ratios K‐feldspar and illite are stable relative to kaolinite. The equilibrium experiments of Yates and Rosenburg (1997), and modeling by others (Varadachari, 2006) highlight the wide ranges in temperature and fluid composition at which certain clay minerals are 15 Fig. 2.1. (after Yates and Rosenberg, 1997): Isothermal, isobaric phase diagrams derived from solution equilibration experiments by Yates and Rosenberg (1997) showing the stability fields of illite, illite‐smectite and smectite in the simple system K2O‐Al2O3‐SiO2‐H2O at 100C, 150C, 200C and 250C. Natural muscovite/ illite, kaolinite, and quartz or amorphous silica samples were equilibrated in a 2M KCL/HCl solution. Phase boundaries are indicated by solid lines. The Carlin‐fluid field (highlighted in yellow) is based on data from previous studies which indicate that mineralizing fluid: (i) is 180‐240C (ii) precipitates quartz, but not amorphous silica as part of the decarbonitization, argillization, silicification and sulfidation alteration sequence(Cline and Hofstra, 2000; Hofstra and Cline, 2000; Lubben, 2004) (iii) does not typically precipitate k‐spar/ adularia. Adularia has been identified only at the Twin Creeks deposit, Nevada (Simon et al., 1999; Stenger et al., 1998). As temperature increases, expandable phases in the illite‐smectite interlayered clay structure decrease and the overall product tends more toward illite. At 200°C pure, end‐member illite appears with no expandable phases; smectite is no longer stable. The smectite field is replaced by illite‐smectite interlayered clay. In the range of temperatures associated with ore deposition in Carlin deposits, smectite is not stable, but illite, illite‐smectite interlayered clay, and kaolinite are stable depending on fluid composition. These equilibrium experiments indicate that illite and kaolinite can form in a hydrothermal environment within accepted ranges for both fluid temperature and silica content. 16 stable, and subsequently the different environments in which they form. Smectite is generally not present in rocks at temperatures above 160‐200C (Reyes, 1990; Essene and Peacor, 1995; Yates and Rosenburg, 1995). Kaolinite can occur in low temperature supergene/ diagenetic environments, but also in higher temperature hydrothermal environments in equilibrium with illite (Yates and Rosenberg, 1997). Illite can form from smectite in the following prograde temperature reaction during early diagenesis or from hydrothermal alteration (Iuoup et al., 1988): smectite + K+ illite + H2O (eqn. 2.2.1) In hydrothermal systems, illite can form via the retrograde alteration of muscovite (Yates and Rosenberg, 1997): muscovite + H+ + silica illite + K+ + H20 (eqn. 2.2.2) Illite can form directly from K‐feldspar under diagenetic conditions (2.2.3a) (Moore and Reynolds, 1997) from the initial weathering of igneous rocks (2.2.3b) (Meunier, 1977), and in the presence of an acidic fluid, either surface derived or hydrothermal (2.2.3c) (Faure, 1998) as follows: K‐feldspar + smectite illite + chlorite + quartz (eqn. 2.2.3a) K‐feldspar (illite) + smectite + kaolinite (eqn. 2.2.3b) K‐feldspar + H2O + H+ illite + K+ + amorphous silica (eqn. 2.2.3c) Equations 2.2.1 and 2.2.2 are favoured for the formation of illite in sedimentary rocks due to the abundance of sedimentary smectite and muscovite in primary sedimentary rock types. In equation 2.3.3b, weathering processes cause K‐feldspar to decay initially to illite, but due to the metastability of illite with respect to a low temperature weathering environment, illite quickly reacts to form smectite + kaolinite (Meunier, 1977). Smectite forms from the low temperature (near surface) alteration of aluminosilicate minerals or from shallow diagenesis in sedimentary basin environments. Inherent problems are associated with studying clay mineral assemblages outside of the laboratory: (i) equilibrium of low temperature clay mineral assemblages may be kinetically inhibited (Essene and Peacor, 1995). There is a risk that low temperature systems never reach equilibrium given that equilibrium is typically attained at elevated temperatures (Essene and Peacor, 1995). Many researchers 17 assume that the reproducibility of a reaction, such as the smectite illite reaction, signifies equilibrium conditions. Essene and Peacor (1995) suggest that reproducibility does not necessarily represent equilibrium and that many reactions in nature can be reproducible without representing equilibrium including: the maturation of hydrocarbons during diagenesis, or the precipitation of magnesian calcite or aragonite from seawater. Changes in clay minerals including grain size and composition, may be a function of kinetic factors such as time and fluid‐rock ratio, not temperature; (ii) certain aluminosilicate poor lithologies are not favourable to argillization i.e. quartzite, chert, limestone etc. Siliciclastic sediments with low aluminosilicate content may exhibit little to no argillic alteration, specifically no illite alteration unless K+ is added to the system; (iii) relationships between minerals are difficult to determine due to the fine grained nature of clay minerals. Determining pre‐existing minerals from which clay minerals grew is made difficult due to a lack of discerning characteristics between some low temperature clays such as illite and muscovite; and (iv) the origin of clay minerals as products of hypogene, supergene, or diagenetic processes is often difficult to determine due to the wide range of temperatures at which most clays are stable. 2.3 ILLITE THERMOMETRY Despite the uncertainty regarding equilibrium clay processes, a number of techniques are available to estimate formation temperatures of clay minerals. Clay thermometry techniques provide consistent information on the relationship between crystallinity, morphology, chemical composition, and formation temperature. As a result, these techniques serve as tools for determining temperature of formation, for differentiating between generations of argillization and for distinguishing between supergene and hypogene clay development. Many techniques involved in the characterization of clay minerals are laborious and time consuming (clay separation of different clay size fraction, TEM, SEM, microprobe) and have not found use in the realm of mineral exploration. The advent of rapid analysis tools such as PIMA (portable infrared mineral analyzer) or Terraspec© to identify clay assemblages in the field, has seen the revival of clay thermometry in the context of exploration. One goal of this paper is to determine whether tools such as the PIMA or Terraspec can provide the accuracy necessary to identify potential ore target zones based on clay mineral assemblages. This is accomplished by comparing Terraspec data to data collected by x‐ray diffraction. 18 2.4 ILLITE POLYTYPISM There are three main illite polytypes: 1Md, 1M, and 2M1 (Meunier and Velde, 2004). The transition from the 1Md to 1M to 2M1 polytype is generally thought to represent a progressive increase in temperature, pressure, reaction time, and/or fluid‐rock ratio (Baronnet, 1980; Lonker et al., 1990; Srondon et al., 2001). Low temperature or early stage diagenetic illite was typically thought to occur as the 1Md polytype while 2M1 illite represented higher temperature hydrothermal environments or late‐ stage diagenesis. Continued study of the transition between polytypes suggests that the thermal stability fields of each polytype are poorly defined and that growth mechanisms and kinetics of polytypic transformations play a key role in polytype determination (Baronnet, 1980). Furthermore, illite polytypes have been observed coexisting in the same environment. In the Broadlands Ohaaki geothermal system, 1M and 2M illite co‐exist on the nanometer scale (Lonker et al., 1990). Owing to the poorly constrained nature of illite polytypism and the questionable relevance of distinguishing between them, polytypes were not identified in this study. 2.5 ILLITE CRYSTALLINITY As temperature increases and the crystal structure of illite becomes more ordered, the crystallite or grain particle size of illite increases (Kubler, 1967; Ji and Browne, 2000). The crystallinity of illite can be measured from x‐ray diffraction patterns by measuring the full width at half the maximum (FWHM) value of the 10Å (001) illite peak measured in 2. Low FWHM values have been correlated to poorly crystallized illite formed at low temperatures, while high FWHM values correspond to crystalline illite formed at higher temperatures (Kubler, 1967; Warr, 1996). The Kubler index uses measured FWHM crystallinity values to infer temperature in diagenetic and low grade metamorphic environments (Kubler, 1967). According to the Kubler Index, the diagenesis/ anchizone boundary corresponds to temperature values of ~200C (Frey et al., 1987) and to an FWHM value of 0.42 2, and the anchizone/epizone boundary corresponds to a temperature of ~300C and to an FWHM value of 0.25 2. Kubler (1967) selected these limits on the basis of certain mineralogical changes. The lower anchizone limit coincides with the upper‐grade limit of the existence of liquid hydrocarbons, the dickite to pyrophyllite transformation, and finally the loss of interlayer water and the conversion of I‐S to pure illite. The upper anchizone limit is associated with the appearance of greenschist facies minerals such as chlorotoid (Kubler, 1967; Kubler and Jaboyedoff, 2000). 19 Since its initial development, the Kubler index has been standardized to ensure precision between measurements made on different diffractometers. The crystallinity illite standard (CIS) scale, used in this study, provides standards used to calibrate individual diffractometers and correct for differences that exist between different machines, and different clay separation techniques (Warr, 1996). The crystallinity results of this study were calibrated to the standardized scale (the crystallinity index standard, CIS) of Warr (1996) by measuring the same six sets of pelitic rock powder standards used by Warr (1996). Once calibrated to the set of standards, the Kubler index can be applied to the standardized experimental FWHM values. Table 2.1 shows the FWHM values of six clay‐separated standards (Warr, 1996) which were used to calibrate the CIS scale for this study. Each standard sample contained illite with a different FWHM value, as determined by Warr (1996). Table 2.1. Calibration chart showing FWHM values for six standards (Warr, 1996) vs. measured FWHM values calculated in this study for the same standards FWHM values FWHM Values Standard Sample ID (Warr, 1996) (2) (This study) (2) SW‐1 0.63 0.28 SW‐2 0.47 0.27 SW‐3 0.46 0.23 SW4‐4 0.38 0.22 SW5 0.36 0.15 SW6‐1 0.25 0.12 A linear regression was calculated using IBM’s statistical software package Predictive Analytics Software (PASW) v.18.0, between the FWHM values measured on the x‐ray diffractometer used for this study and the values determined by Warr (1996) (Eqn. 2.2.4). All statistical calculations in this study were performed using PASW v. 18.0. The regression was then applied to all FWHM values for samples used in this study. ݕ = 1.794ݔ + 0.041 R = 0.89 Where ݔ= FWHM value measured in this study ݕ = standard value from Warr (1996) Eqn. 2.2.4 20 2.6 CLAY TEXTURES AND CLAY MORPHOLOGY Previous studies indicate that two end‐member morphologies of illite exist (Figure 2.2); each representing a different environment of formation (Hancock and Taylor, 1978; Peaver, 1999; Meunier and Velde, 2004; Schleicher et al., 2006). Hexagonal illite is the stable crystal shape of illite and has been empirically observed to form under hydrothermal conditions (Inoue et al., 1988). “Hairy illite” is the metastable crystal shape and is typically observed in the pore‐space of sedimentary rocks likely resulting from unconstrained growth in large pore‐spaces during diagenetic processes (Peaver, 1999; Schleicher et al., 2006). Meunier and Velde (2004) attribute changes in morphology to temperature. As temperature increases, the crystal shape of illite changes from elongated one‐dimensional 'hairy' crystals to more rigid hexagonal laths which increase in width progressively. Bauer et al. (2000) attribute the change between morphologies to reaction progress whereby initial stages of growth result in hairy illite and late stages of growth exhibit hexagonal illite. However, departures from these end‐member morphologies have also been observed. Hancock and Taylor (1978) identified sheeted stacks of illite, typical of hexagonal illite but lacking sharp boundaries. They concluded this was a classic replacement texture. As temperature increased along with depth, a higher temperature mineral (illite) pseudomorphed a pre‐existing lower temperature mineral (kaolinite) no longer stable at those temperatures (Hancock and Taylor, 1978). Hancock (1978) describes four different variations of the hexagonal morphology representing progressive stages of diagenesis and therefore increasing temperature: (i) tangential rims where large flakes appear to be formed by a coalescence of much smaller bladed crystals (ii) radial rims where illite flakes protrude radially from grain margins, typically as overgrowths of tangential rims (iii) illite mesh‐work where interlocking illite crystals extend far out into porespace (iv) dense homogenous illite whereby illite occurs as large, thick, randomly oriented flakes which are often curves and may result from recrystallization of detrital illite. Studies show that absolute temperature cannot be determined using illite morphology but relative temperatures can be established using variations in both texture and morphology. 2.7 CHEMICAL COMPOSITION OF ILLITE Modern geothermal environments that are currently precipitating phyllosilicate minerals (Broadlands Ohaaki, Salton Sea, Los Azufres, and Coso) provide robust empirical observations on the positive relationship between temperature and clay mineralogy (Simmons and Browne, 2000; Battaglia, 21 Fig. 2.2. (Modified from Bauer et al., 2000) SEM photographs of synthesis products of micas. A. After 30 days the first lath‐shaped “hairy” illite crystals can be observed B. After 210 days the XRD pattern indicates only lath‐shaped mica‐illite C. After 28 months perfect hexagonal mica‐illite crystals can be observed. They are all approximately the same size. 22 2004). Changes in the chemical composition of illite have been shown to correlate with changes in temperature (Cathelineau, 1988; Battaglia, 2004). As temperature increases, and the crystal structure of illite becomes more ordered, larger cations like Fe and Mg are removed from the interlayer site and replaced by available K (Battaglia, 2004). Therefore, high temperature illite is more potassic than low temperature illite as observed in a compilation of data from the Los Azufres, Salton Sea, and Coso geothermal fields by Cathelineau (1988). However, numerous problems exist with the data of Cathelineau (1988) and are outlined by Battaglia (2004): (i) the relationship between K and temperature is consistent within one geothermal field but absolute temperatures vary between geothermal fields, and (ii) the linear regression lines for data from each geothermal field are not parallel. The data for Coso and Los Azufres geothermal fields are convergent at a K value near 1 (measured in cations), while the regression line for Salton Sea does not converge. The lack of convergence of Salton Sea data is attributed to a lack of elemental Ca in the illite samples analyzed. Hence, [K] alone is not an accurate or consistent proxy for temperature. Based on this previous work completed by Cathelineau (1988), Battaglia (2004) developed a cation‐interdependent formula ܭ+|ܨ݁ െ ܯ݃| which appears to be a more robust indicator of temperature. Figure 2.3 shows data from five geothermal fields around the world which exhibit a linear trend between increases in cation content (ܭ+|ܨ݁ െ ܯ݃|) and increases in surrounding temperature. Regression lines are shown in Table 2.2. The best fit linear regression line through illite compositional data acquired by Battaglia (2004) for all five geothermal fields yields a correlation factor of R = 0.84 between temperature and ܭ+|ܨ݁ െ ܯ݃| giving equation (2.2.5): ݕ = 267.95ݔ + 31.46 Where ݔ = ܭ+|ܨ݁ െ ܯ݃| Table 2.2. Regression line equations and correlation factors (R) corresponding to data graphed in Figure 6. Regression Line Equation Average Error (%) * R Broadlands Ohaaki ݕ = 113.19ݔ + 182.59 5.71 0.72 Coso ݕ = 196.47ݔ + 64.00 10.96 0.97 Salton Sea ݕ = 247.68ݔ + 56.36 4.04 0.96 Los Azufres ݕ = 337.93ݔ ‐ 36.57 7.96 0.79 Aluto Langano ݕ = 328.85ݔ ‐ 1.17 6.73 0.84 *= Measure of average difference between calculated temperatures based on the regression line from all data, and measured temperatures from geothermal well. (Eqn. 2.2.5) 23 Fig. 2.3. Graph plotting microprobe composition data ܭ+|ܨ݁ െ ܯ݃| (measured as cations) from five geothermal fields against measured temperatures of those samples presented in Battaglia (2004). Linear regression lines were calculated using PASW. A regression line including data from all five geothermal fields is superimposed on individual data groups. 24 An average error of 7% on the final temperature value was calculated based on the difference between temperature values calculated using the regression line equation, and temperature measured directly from the geothermal well. 2.8 SUMMARY Variations in the texture, morphology, crystallinity, and composition of clay minerals with temperature can provide information on the temperature at which the minerals formed. There is, however, the possibility that variations in the characteristics of low temperature clay minerals are not in fact solely a function of temperature but instead represent reaction progress (Essene and Peacor, 1995). Temperature is important in order to constrain the origin of clay minerals as either products of hydrothermal alteration or otherwise. In the following chapter, the analytical techniques described in this chapter are applied to clay minerals in sedimentary rocks that occur stratigraphically above and outboard of Carlin‐type gold systems to determine whether clay alteration at surface is related to Carlin‐ type gold mineralization at depth. 25 2.9 REFERENCES Baronnet, A., 1980, Polytypism in micas: a survey with emphasis on the crystal growth aspect: Current topics in materials science, v. 5, p. 447‐548. Battaglia, S., 2004, Variations in the chemical composition of illite from five geothermal fields: a possible geothermometer: Clay Minerals, v. 39, p. 501‐510. Bauer, A., Velde, B., and Gaupp, R., 2000, Experimental constraints on illite crystal morphology: Clay Minerals, v. 35, p. 587‐597. Cathelineau, M., 1988, Cation site occupancy in chlorites and illites as function of temperature: Clay Minerals, v. 23, p. 471‐485. Cline, J., and Hofstra, A., 2000, Ore‐fluid evolution at the Getchell Carlin‐type gold deposit, Nevada, USA: European Journal of Mineralogy, v. 12, p. 195‐212. Essene, E., and Peacor, D., 1995, Clay mineral thermometry‐a critical perspective: Clays and Clay Minerals, v. 43, p. 540‐553. Faure, G., 1998, Principles and applications of geochemistry: a comprehensive textbook for geology students: Upper Saddle Riber, New Jersey, Prentice Hall, 600 p. Frey, M., 1987, Very low‐grade metamorphism of clastic sedimentary rocks: Low temperature metamorphism, p. 9–58. Hancock, N., 1978, Possible causes of Rotliegend sandstone diagenesis in northern West Germany: Journal of the Geological Society, v. 135, p. 35‐40. Hancock, N., and Taylor, A., 1978, Clay mineral diagenesis and oil migration in the Middle Jurassic Brent Sand Formation: Journal of the Geological Society of London, v. 135, p. 69–72. Hofstra, A., and Cline, J., 2000, Characteristics and models for Carlin‐type gold deposits, Chapter 5: in Hagemann, S. G. and Brown, PE, eds., Gold in 2000: Reviews in Economic Geology, v. 13, p. 163‐ 220. Iuoup, A., 1988, Mechanism of illite formation during smectite‐to‐illite conversion in a hydrothermal system: American Mineralogist, v. 73, p. 1325‐1334. Ji, J., and Browne, P., 2000, Relationship between illite crystallinity and temperature in active geothermal systems of New Zealand: Clays and Clay Minerals, v. 48, p. 139‐144. Kubler, B., 1967, La cristallinite de l'illite et les zones tout‐a‐fait superieures du metamorphisme., in Schaer, J.P., ed., Colloque sur les etages tectoniques: Baconniere, Neuchatel, p. 105‐122. Kübler, B., and Jaboyedoff, M., 2000, Illite crystallinity: CONCISE REVIEW PAPER: Comptes Rendus de l'Académie des Sciences‐Series IIA‐Earth and Planetary Science, v. 331, p. 75‐89 Lanson, B., and Champion, D., 1991, The I/S‐to‐illite reaction in the late stage diagenesis: American Journal of Science, v. 291, p. 473‐506. 26 Lanson, B., Velde, B., and Meunier, A., 1998, Late‐stage diagenesis of illitic clay minerals as seen by decomposition of X‐ray diffraction patterns: Contrasted behaviors of sedimentary basins with different burial histories: Clays and Clay Minerals, v. 46, p. 69‐78. Lonker, S., FitzGerald, J., Hedenquist, J., and Walshe, J., 1990, Mineral‐fluid interactions in the Broadlands‐Ohaaki geothermal system, New Zealand: American Journal of Science, v. 290, p. 995‐1068. Lubben, J., 2004, Quartz as clues to paragenesis and fluid properties at the Betze‐Post deposit, northern Carlin trend, Nevada: Unpublished M.Sc. thesis, Las Vegas, Universitty of Nevada, p. 155. Meunier, A., 1977, Les mechanismes de l'alteration des granites et le role des microsystemes: tude des arenes du massif granatique de Parthenay (Deux‐Se vres): PhD. Thesis, 248 p. Meunier, A., and Velde, B., 2004, Illite: origins, evolution, and metamorphism: Poitier, France, Springer Verlag, 286 p. Moore, D., and Reynolds Jr, R., 1997, X‐ray diffraction and the identification and analysis of clay minerals, 378 p, Oxford University Press, New York. Peaver, D., 1999, Illite and hydrocarbon exploration: Proceedings of the National Academy of Sciences, v. 96, p. 3440‐3446. Reyes, A., 1990, Petrology of Philippine geothermal systems and the application of alteration mineralogy to their assessment: Journal of Volcanology and Geothermal Research, v. 43, p. 279‐309. Rose, A., 1970, Zonal relations of wallrock alteration and sulfide distribution at porphyry copper deposits: Economic Geology, v. 65, p. 920‐936. Schleicher, A., Warr, L., Kober, B., Laverret, E., and Clauer, N., 2006, Episodic mineralization of hydrothermal illite in the Soultz‐sous‐Forêts granite (Upper Rhine Graben, France): Contributions to Mineralogy and Petrology, v. 152, p. 349‐364. Simon, G., Kesler, S., and Chryssoulis, S., 1999, Geochemistry and textures of gold‐bearing arsenian pyrite, Twin Creeks, Nevada; implications for deposition of gold in Carlin‐type deposits: Economic Geology, v. 94, p. 405‐422. Simmons, S., and Browne, P., 2000, Hydrothermal minerals and precious metals in the Broadlands‐ Ohaaki geothermal system: Implications for understanding low‐sulfidation epithermal environments: Economic Geology, v. 95, p. 971‐999. Srodon, J., Drits, V., McCarty, D., Hsieh, J., and Eberl, D., 2001, Quantitative X‐ray diffraction analysis of clay‐bearing rocks from random preparations: Clays and Clay Minerals, v. 49, p. 514‐528. Stenger, D., Kesler, S., Peltonen, D., and Tapper, C., 1998, Deposition of gold in carlin‐type deposits; the role of sulfidation and decarbonation at Twin Creeks, Nevada: Economic Geology, v. 93, p. 201‐ 215. Varadachari, C., 2006, Fuzzy phase diagrams of clay minerals: Clays and Clay Minerals, v. 54, p. 616‐625. 27 Warr, L., 1996, Standardized clay mineral crystallinity data from the very low‐grade metamorphic facies rocks of southern New Zealand: European Journal of Mineralogy, v. 8, p. 115‐127. Whitney, G., and Northrop, H., 1988, Experimental investigation of smectite‐to‐illite reaction‐‐dual reaction mechanisms and oxygen‐isotope systematics: American Mineralogist, v. 73, p. 77‐90. Yates, D., and Rosenberg, P., 1997, Formation and stability of endmember illite: II. Solid equilibration experiments at 100 to 250 C and Pv, soln: Geochimica et Cosmochimica Acta, v. 61, p. 3135‐ 3144. 28 CHAPTER 3 – BEYOND THE CONFINES OF THE ORE BODY: SURFACE MAPPING OF LOW TEMPERATURE HYDROTHERMAL FLUID ABOVE MAJOR ORE BODIES USING CLAY ALTERATION 3.1 INTRODUCTION The passage of hydrothermal fluid through the Earth’s crust is invariably accompanied by mineral alteration (Giggenbach, 1981, 1984; Reed, 1997). The degree to which mineral assemblages change is a function of multiple factors: temperature, pressure, time, fluid chemistry, initial rock composition and fluid‐rock ratio (Reed, 1997). Alteration associated with mineral deposits commonly extends outboard from the ore body providing robust vectors toward the core of those deposits (Figure 3.1). In alteration halos around low temperature hydrothermal systems, gradients in temperature and chemical composition between mineralizing hydrothermal fluid and host rock can be small. With increasing distance from the core of the hydrothermal system, thermal and chemical gradients between hydrothermal fluid and country rock decrease until a point where no gradient exists between fluid and rock. The distal expression of low temperature hydrothermal systems may be dominated by mineral assemblages also resulting from near surface processes including weathering, shallow diagenesis, and other overprinting hydrothermal events. Tools that identify subtle variations between low temperature minerals which in turn help distinguish between different generations of minerals are required to assess the extent of hydrothermal alteration, and to use that alteration to vector in toward mineralization at depth (Adams and Putnam, 1992; Kelley et al., 2006). However, the use of low temperature minerals to investigate geological problems has been criticized due to the inability of most low temperature assemblages to reach equilibrium causing nanoscale variations in composition (Essene and Peacor, 1995, 1996). Under disequilibrium conditions, properties such as crystallinity, crystal morphology, or composition might represent reaction progress rather than spatial variation in temperature (Essene and Peacor, 1995). Clay minerals form a large component of low temperature alteration assemblages and have been documented in different settings including: geothermal systems (Simmons and Browne, 2000; Yang et al., 2001), epithermal deposits (White and Hedenquist, 1995), the distal extents of porphyry deposits (Rose, 1970; Tosdal, 2009; Sillitoe, 2010), and diagenetic environments (Meunier and Velde, 2004). Geothermal systems provide current information on low temperature systematics including clay 29 Fig. 3.1. The distal extent and degree of alteration around ore deposits is a direct function of thermal and chemical gradients between country rock and hydrothermal fluid. A. Black dashed lines represent steady state thermal contours resulting in mineral alteration. Porphyry type settings exhibit large thermal and chemical gradients. Initially, hot (>700C) magma intrudes cold (<200C) country rock. Hot magmatic fluid alters country rock immediately around the magma body via conduction. Convection cells of circulating pore fluid are generated at the sides of the intrusion. Heat is carried from the intrusion by advection as fluid travels through fractures forming networks of veins. As magmatic fluid travels upward and decreases in temperature, acids dissociate and the fluid becomes acidic. The lower temperature acidic fluid interacts with country rock to produce argillic alteration. As the pluton cools, cold, near‐neutral surface meteoric water is drawn into the core of the system resulting in the hydrolysis of minerals and propyltic alteration which may overprint other stages of alteration. Visible alteration in a porphyry system can extend kilometers away from mineralization, with distal alteration assemblages containing minerals significantly different to unaltered host rock assemblages i.e. sericite ± illite ± kaolinite B. In low temperature systems such as the Carlin‐type Au systems, there may be little temperature gradient between the incoming fluid (200C) and the shallow crust (~75C), resulting in the subtle mineral alteration assemblage illite ± dickite ± kaolinite at the core of the system. Fluid flow is focused along permeable pathways whereby heat is carried primarily by advection. Chemical gradients in low temperature systems may be large, as shown here whereby acidic fluid causes the decarbonatization of permeable carbonate horizons. Visible alteration however is often subtle and may extend as little as a few metres away from mineralized fluid pathways. Distal to the core of the hydothermal system, there may be no difference between the hydrothermal mineral assemblage and the unaltered host rock mineral assemblage (Modified from Sillitoe, 2010; Tosdal et al., 2009; Reed, 1997). 30 mineral reaction kinetics and timescales of fluid flow. Clay minerals have been used to map out networks of fluid flow pathways in geothermal systems by identifying zonation patterns of clay minerals around upwelling regions of hydrothermal fluid (Simmons and Browne, 2000). Illite, a high temperature clay mineral, occurs within and proximal to fluid conduits. Smectite, a low temperature clay mineral, occurs most distal to fluid pathways with illite‐smectite interlayered clay (I‐S) between (Simmons and Browne, 2000). Clay mineral mapping can also be used to determine relative proximity to mineralization (Figure 3.1) (Rose, 1970; Seedorf et al., 2005; Sillitoe, 2010). In the idealized alteration model of a porphyry deposit (Rose, 1970, Tosdal et al., 2009), the phyllic assemblage consisting of quartz‐sericite (muscovite)‐pyrite occurs most proximal to the magmatic heat source where temperature is highest. Argillic alteration assemblages consisting of illite ± kaolinite indicate progressive acidification of the fluid, decreasing temperature with distance from the core of the system, and possibly interaction with surface derived waters (Tosdal et al., 2009). Another method to trace fluid flow is by the advection of heat. Because heat is transported by an infiltrating fluid at a rate greater than all but the most incompatible of geochemical tracers, transient heating associated with hydrothermal flow is quite likely to be one of the most distally developed expressions of a hydrothermal mineral deposit (Bickle and McKenzie, 1987). Because heat is not only transported by the fluid itself, but also diffuses rapidly outward from fluid flow pathways, the volume of rock affected by the thermal energy of a hydrothermal system will be significantly larger than that recorded by mineralogical alteration and isotopic resetting. The thermal footprint of low temperature systems can be identified using low temperature thermochronology including apatite fission tracks and U‐Th/He in apatite and zircon (Chakurian et al., 2003; Cline et al., 2005; McInnes et al., 2005; Arehart and Donelick, 2006; Kelley et al., 2006; Hickey et al., unpublished data). The Carlin deposits of northern Nevada are one of the world’s major sources of gold (Teal and Jackson, 2002; Price et al., 2007). Carlin‐type gold deposits are an example of low temperature (180‐ 240C) hydrothermal systems that exhibit subtle alteration associated with mineralization, making them a difficult target for exploration. In this study, we investigate the ability of low temperature analytical methods to identify hydrothermal alteration manifested distal to the core of Carlin‐type gold mineralization by characterizing clay mineral assemblages with respect to paleo‐formation temperature, chemistry, crystallinity and morphology. We also provide insight into a combination of analytical techniques that exhibit the highest potential to deliver reliable and robust exploration vectors for low temperature hydrothermal ore deposits in the subsurface. 31 Fig. 3.2. Regional geology map of northeastern Nevada showing location of significant Carlin‐type mineral deposits including those discussed in this paper. 32 Apatite fission track thermochronology data spanning across northeastern Nevada highlighted potential regions of hydrothermal flow around existing Carlin‐type deposits. Based on AFT data, two study areas (Figure 3.2), were chosen with the intent of (i) identifying zones of hydrothermal clay alteration using currently available clay thermometers and (ii) establishing whether hydrothermal alteration at surface exhibited a genetic link to Carlin‐type ore deposits at depth. The Leeville deposit, (Figure 3.3) is located along the northern Carlin trend, south of the Tuscarora Mountains and is hosted in lower plate rocks immediately beneath the RMT with a few hundred metres of upper plate material preserved above. In this field area, upper plate material was sampled above mineralization to determine whether Carlin‐deposit forming fluid had exhausted through the upper plate to surface. The second field area chosen for this study was the Shoshone Range, host to a number of Carlin‐type gold deposits that form the Battle Mountain Eureka mineral belt. The giant Pipeline deposit, located at the southeastern end of the Shoshone Range was a starting point for sample collection in an area of known Carlin‐style mineralization in lower plate carbonate rocks. From the Pipeline deposit, a northwest trending transect was sampled to determine the aerial extent of alteration in upper plate rocks outboard from a site of known mineralization. 3.2 GEOLOGICAL SETTING OF CARLIN‐GOLD DEPOSITS During the period from Cambrian ‐ Devonian, passive margin miogeoclinal and eugeoclinal sequences were deposited along the rifted margin of western North America (Morrow and Sandberg, 2008). Multiple episodes of compression from Devonian to Cretaceous led to eastward directed thrusting of miogeoclinal rocks on top of eugeoclinal rocks. In northeastern Nevada, there is one main thrust sheet, the Roberts Mountain allochthon (locally termed ‘upper plate’) which has the Roberts Mountain Thrust (RMT) fault at the base. Footwall rocks (‘lower plate’) consist of shelf and slope silty carbonate rocks and are the main host for Carlin‐type gold mineralization. The upper plate lacks any major Carlin‐type Au deposits although mineralization does occur locally immediately above the RMT in several of the lower plate‐hosted deposits. A tectono‐stratigraphic column of lower Paleozoic rocks is presented in Figure 3.4. In most areas, thick sequences of upper plate siliciclastic rocks cover lower plate carbonate rocks, however post‐ Antler orogeny tectonism and erosion have exposed geologic windows into the lower plate as exhibited at Goat Peak (north‐western Shoshone Range), and at the Pipeline deposit locality (south‐western Shoshone Range) (Kelson et al., 2008). Structurally, both the upper and lower plates are imbricately thrusted with tight, upright to westward‐inclined folds caused by the eastward transport of material 33 Fig. 3.3. A. Geological map of the Leeville deposit study area. The Roberts Mountain Thrust fault separates Lower Paleozoic upper plate siliciclastic rocks to the North from Lower Paleozoic lower plate carbonate rocks in the South. B. Cross section A –A1 through the Leeville deposit shows the presence of at least 250m of upper plate siliciclastic rock cover above the main ore zones of the Leeville deposit. A. B. 34 Fig. 3.4. (modified from Bettels et al., 2002 in Emsbo et al., 2003) Simplified tectono‐stratigraphic column showing lower Paleozoic rocks of the RMT system. Variation in formation names between the Battle Mountain – Eureka trend (left) and Carlin trend (right) are shown. During the Mississippian Antler Orogeny, lower Paleozoic siliciclastic sediments (upper plate) were thrust eastward overtop of lower Paleozoic carbonate sediments (lower plate). Carlin‐type gold deposits are typically hosted in lower plate rocks. 35 during orogenic events dating from the Devonian to late Cretaceous (Noble and Finney, 1999). Subsequent deformation includes broad open folding accompanied by oblique‐slip faulting, Mesozoic to Cenozoic low angle faulting, and high angle normal faulting associated with Basin and Range extension (Winterer, 1968; Cluer et al., 1997). The lower plate of the RMT can be subdivided into four main units (Emsbo et al., 2003). The Devonian Horse Canyon Formation/Rodeo Creek Formation is a calcareous siltstone, mudstone and chert with local occurrences of sandstone and mudstone. The Devonian Wenban Formation/ Popovich Formation is composed of laminated calcareous to dolomitic siltstones, micritic limestone, and thick bioclasitc debris flows. The Silurian Roberts Mountain Formation is dominated by laminated dolomitic and calcareous, and calcareous mudstone and calcareous siltstone. The Ordovician Hanson Creek Formation is a sandy dolomite. The overlying upper plate comprises dominantly siliciclastic units with the rare occurrence of thin carbonate lenses. The Ordovician Valmi/ Vinini Formation is composed of several thousand meters of chert, quartz arenite, argillite, slate, and greenstone (Roberts, 1951). The Silurian Elder Formation is a set of interbedded shale, siltstone, chert, and feldspathic and calcareous sandstone. The Devonian Slaven Formation is a mixture of highly contorted and broken, black, nodular chert with some carbonaceous shale partings. Paleozoic rocks are intruded by Mesozoic and Cenozoic intrusive rocks (Ressel and Henry, 2006). Igneous rocks of the Shoshone Range include Eocene‐Pliocene intermediate to felsic intrusions with basalt, andesite and rhyolite flows and tuffs. Eocene intrusions were emplaced along a west‐northwest trend and have been identified both proximal and distal to economic gold mineralization (Gilluly and Gates, 1965; Stager, 1977; Kelson et al., 2005). The Gold Acres stock is a Cretaceous quartz monzonitic pluton located at the southeastern edge of the Shoshone Range proximal to the Pipeline deposit (Mortensen et al., 2000). A thorough compilation of data on the igneous rocks of the Carlin trend by Arehart et al. (2004) and Ressel and Henry (2006) suggest three periods of magmatism in the Northern Carlin trend: Jurassic, Cretaceous, and Eocene. Jurassic intrusions consist of the Goldstrike laccolith and related sills and mostly northwest‐striking lamprophyre and rhyolite dikes ~158 Ma (Ressel and Henry, 2006). The Vivian sill exposed at surface and the Little Boulder Basin stock located at depth in the Leeville area have been interpreted as south‐eastern extensions of the Goldstrike laccolith (Ressel and Henry, 2006). Cretaceous magmatism is characterized by one occurrence ~4km south of the Leeville deposit. Concordant Pb/U ages of two zircon fractions demonstrate intrusion at 112.4 +/‐ 0.6 Ma (Mortensen et al., 2000). 36 3.3 CHARACTERISTICS OF CARLIN GOLD DEPOSITS Carlin gold deposits are restricted to a small geographic area in northeastern Nevada. The Carlin, Battle Mountain‐Eureka, and Getchell Trends describe three regional lineaments along which the majority of Carlin‐deposits are focused, including those described in this study (Figure 3.2). The alignment of Carlin‐type deposits is thought to reflect major crustal faults established during Neoproterozoic rifting (Roberts, 1966; Tosdal et al., 2000). Lower plate Paleozoic carbonate shelf rocks are the favourable hosts for Carlin gold mineralization (e.g., Cortez Hills, Pipeline, Carlin, Goldstrike, Leeville, and Jerritt Canyon). This has been attributed to the reactivity of carbonate sediments with acidic gold‐bearing fluid resulting in mass loss and increased permeability, providing robust pathways for Au‐bearing fluid (Cline et al., 2005). Conversely, upper plate rocks are largely devoid of gold mineralization likely due to a lack of reactivity with these same gold bearing fluids. Upper plate rocks of the RMT are however host to base and precious metal deposits including Miocene Au‐Ag epithermal deposits, Eocene porphyry deposits, and a limited number of Carlin‐type gold deposits i.e. Alligator Ridge (Nutt and Hofstra, 2003), Emigrant (Newmont Mining Company, unpublished data), and Mike (Norby and Orobona, 2002). The Elder Creek deposit, considered in this study, is a small gold deposit located in the Central Shoshone Range. Gold is hosted predominantly in quartzite clast breccia of the Valmi Formation and also within pyrite‐rich argillite and shale of the Elder Formation (Ahmed et al., 2010). Fluid inclusion and mineral thermometry data indicate that mineralizing fluid forming Carlin deposits in north‐eastern Nevada were: low temperature (180‐240C), slightly acidic (pH ~4), low‐ salinity (~2–3 wt% NaCl equivalent) aqueous fluids that contained CO2 (<4 mol %) and CH4 (<0.4 mol %), and sufficient H2S (10–1–10–2 mol/kg) to transport Au and other bisulfide‐complexed metals (Cline and Hofstra, 2000; Hofstra and Cline, 2000; Lubben, 2004). Stable isotope analyses have provided insight into the source of mineralizing fluid in Carlin‐type deposit settings. DH2O values measured on ore‐stage kaolinite and fluid inclusions from a wide‐range of Carlin deposits suggest evidence of meteoric water with very low DH2O values of < ‐110‰ (Hofstra et al., 1999). Similar data from the Getchell, Twin Creeks, and Deep Star deposits indicate the additional presence of deeply sourced metamorphic or magmatic fluid (Hofstra, 1999; Heitt et al., 2003; Cline et al., 2005). Clay samples from Carlin‐deposits appear to form along a mixing line between Eocene‐age meteoric water and magmatic or metamorphic fluid. 37 3.3.1 Clay alteration in Carlin‐type systems The following characteristics are known of alteration in Carlin‐type settings: (i) quartz is precipitated as part of the silicification alteration sequence, but not amorphous silica (Bakkan and Einaudi, 1986; Cline and Hofstra, 2000; Cline, 2001; Ye et al., 2003), and (ii) K‐feldspar has not been observed in Carlin type systems except for the presence of adularia at the Twin Creeks deposit (Stenger et al., 1998; Simon et al., 1999). According to clay equilibria (Figure 2.1), constraints on temperature (180‐240C), fluid chemistry (pH~4), and observed mineralogy indicate that smectite is not stable, but illite, I‐S, and kaolinite are stable depending on variations in fluid chemistry. Illite and kaolinite can form from the same hydrothermal event within a confined range of fluid temperature and silica content. For kaolinite and illite to form from the same fluid, without forming K‐feldspar, log aK+/aH+ values are estimated to be in the range of 2.0‐5.0 (Yates and Rosenberg, 1997). Similar to patterns observed in geothermal systems, clay zonation has been observed in lower plate carbonate rocks around ore shoots of the Getchell deposit, a Carlin‐type gold deposit located in North Central Nevada (Cail and Cline, 2001). Illite was shown to occur in higher volume % with ore, and proximal to ore whereas smectite occurred distal to ore. Both 1M and 2M1 polytypes have been observed spatially associated with mineralization in Carlin type environments (Carlin, Kuehn and Rose, 1992; Betze‐Post, Arehart et al., 1993b; Deep Star, Heitt et al., 2003). The absence of major Carlin‐type gold deposits in the upper plate suggests that ore fluids responsible for mineralization of the lower plate must have followed one of three fluid evolution pathways shown in Figure 3.5. Either exhausted fluid did not transgress the upper plate and all fluid was exhausted laterally along the RMT, or exhausted fluid transgressed the upper plate and reacted with siliciclastic rocks forming an alteration halo above lower plate gold mineralization. Upper plate rocks contain little to no carbonate, precluding the potential for observing decalcification. Silicification of upper plate rocks would be a challenge to identify given the high silica content of upper plate lithologies. The stage of alteration with the most potential to be observed in upper plate rocks is argillization: illite ± dickite ± kaolinite due to the presence of pre‐existing aluminosilicate minerals in some lower Paleozoic sedimentary lithologies. 3.4 THERMAL SIGNATURE OF HYDROTHERMAL FLUID FLOW AFT has conventionally been applied to a number of geological problems (Gallaghar et al., 1998): (i) resolving the thermal history of sedimentary basins; (ii) investigating the provenance of rocks; (iii) 38 Fig. 3.5. The absence of major gold deposits in the upper plate suggests that ore fluids responsible for mineralization of lower plate carbonate rocks must have followed one of three fluid evolution pathways: A. following gold mineralization of the lower plate, fluid was dominantly rock buffered but still contained a small amount of gold. Fluid flowed upward to the RMT, exploiting pre‐ existing fault and fracture networks, precipitating gold along the fluid flow‐path. Fluid was exhausted laterally along the RMT to surface. The upper plate was largely impermeable to exhausted ore fluids and behaved as an aquitard. This scenario supports observed mineralization of rocks along the RMT. In the scenario described, upper plate siliciclastic rocks would exhibit no signs of alteration due to a lack of interaction with the exhausted Carlin‐fluid B. mineralizing fluid was dominantly rock buffered but still contained a small amount of gold. Fluid flowed upward through a network of small‐scale fractures becoming increasingly rock‐buffered along the flow path. Upon reaching the RMT the fluid precipitated a small amount of gold. Fluid was exhausted laterally along the RMT and also through faults and fractures in upper plate siliciclastic rocks. The upper plate exhibits no signs of Carlin‐type alteration because the fluid was completely rock buffered by the time it encountered the upper plate C. largely unbuffered mineralizing fluid was partially exhausted laterally along the RMT and the remainder transgressed upper plate siliciclastic rocks. Sustained fluid‐rock interaction between siliciclastic upper plate rocks and partially rock buffered Carlin‐type fluid produced a thermal and chemical signature in upper plate rocks. A. B. C. 39 studying the deformation history of orogenic belts, and (iv) interpreting the conductive heating history around magmatic intrusions. Recent studies show that fission track analysis can also be used to identify regional and deposit scale patterns of hydrothermal fluid flow (Chakurian et al., 2003; Hickey, 2003; McInnes et al., 2005; Arehart and Donelick, 2006; Hickey et al., 2010). The maximum 180‐240C ore‐ stage fluids responsible for Carlin‐type mineralization have the capacity to reset the AFT thermochronometer system on timescales of 104‐106 years (Hickey et al., 2010). Mapping the extent of thermal resetting may assist in delineating the far‐field extent of Carlin hydrothermal systems at distances beyond any wall rock type alteration. Chakurian et al. (2003), Hickey (2003), and Cline et al. (2005) have used AFT data to examine the regional thermal history of the northern Carlin Trend. Results from thermochronological studies of the Northern Carlin trend (Figure 3.6) indicate that areas of Carlin‐type mineralization are spatially coincident with pooled AFT ages ranging from 50‐25 Ma. All thermochronological data used in this study is included in Appendix B. Areas of young AFT ages lie within large scale regions of Cretaceous or older AFT ages reflecting the pre‐mineralization regional cooling history. The thermal effects of the Carlin‐ hydrothermal system were superimposed on the older regional pattern. Young apatite fission track ages correspond to episodes of transient reheating. Advective heating by circulating hydrothermal fluids appears to be the primary cause of thermal resetting, shown by the heterogeneous nature of fission track ages in the region. This is evidenced by the lack of major reheating in unaltered portions of large Jurassic stocks (i.e., Goldstrike, and Vivian stocks), and total thermal resetting of highly altered, mineralized material in the same stocks. A strongly argillized, mineralized dyke from within the Leeville deposit is partially reset to 44.5 ± 3.0 Ma. New apatite fission track data of samples from the Shoshone Range are presented along with published data from Arehart and Donelick (2006) in Appendix B. Figure 3.7 shows a map of the Shoshone Range on which results from a geophysical magnetics survey has been superimposed along with pooled apatite fission track ages. The central part of Shoshone range exhibits a mixed population of ages ranging from 20.3‐77.4 Ma. A single sample from the western margin of the range exhibits an older age of 108.2 Ma. The older age may represent the western margin of hydrothermal resetting. The southeastern Shoshone Range exhibits dominantly Eocene ages, likely related to the same hydrothermal event responsible for gold mineralization of the Pipeline deposit (Arehart and Donelick, 2006). Younger ages within the Pipeline pit and on the margins of the deposit are likely a result of late Miocene extension and uplift. The northwest Shoshone Range exhibits dominantly young Miocene ages which can be explained two ways: (i) The magnetic high observed in the northwestern Shoshone Range in 40 Fig. 3.6. Compilation of apatite fission track ages across the Northern Carlin Trend (Chakurian et al., 2003; Cline et al., 2005; Hickey et al., 2010). A zone of pervasive annealing preserves evidence for an episode of rapid cooling from > 100C at ~40 Ma., with localized annealing of fission tracks as late as ~20 Ma. The zone of annealing is broadly parallel to the Carlin trend and becomes more heterogeneously distributed and less pervasive farther to the northwest, toward the Post‐Betze‐Screamer and Meikle deposits (Hickey et al., 2003b). Dyke samples from within the Leeville deposit are totally reset to 44.6 and 49.5 Ma. The Vivian stock is reset to 88.4 Ma, an age that likely represents exhumation (Hickey et al., 2003). 41 Fig. 3.7. Compilation of apatite fission track pooled ages across the Shoshone Range plotted on a reduced to pole magnetics map with a DBM (database management) overlay. The image was provided by Placer Dome Inc. The northwestern Shoshone Range is dominated by younger ages representative of either Miocene magmatism and related Au‐Ag mineralization, or extensive exhumation during the Miocene. The southeastern portion of the Shoshone Range, the location of the Pipeline and Gold Acres deposits is reset to ages between 36‐47Ma, coincident to Eocene Carlin‐type gold mineralization. The rest of the Shoshone range exhibits ages of mixed populations. A group of older AFT ages (108.2, 77.4, and 55.1 Ma) occur at the far western margin of the Shoshone Range. 42 Figure 3.7 is a large buried intrusion of Miocene age. Conductive heating and hydrothermal fluid flow associated with the intrusion thermally reset fission track ages in that region. Dissipation of heat outboard from the intrusion resulted in a mixed population of ages in the central Shoshone Range, or (ii) The Miocene was a period of major uplift and exhumation of the northwestern Shoshone Range. A study of the regional topography of the area shows a N‐S trending fault along the length of the central Shoshone Range which may have caused significant uplift of the upthrown block, resulting in a ‘younging’ of AFT ages across the fault from SE to NW . Caetano Tuff, dated at 32.3 Ma (Naeser and Mckee, 1970), has been identified at the eastern and western margins of the Shoshone Range and suggests that Miocene exhumation would have to have been restricted to only the northwestern region of the Shoshone Range. 3.5 ANALYTICAL TECHNIQUES A number of techniques were employed to measure the variation in clay type, illite crystallinity, illite morphology, and illite composition across the Leeville and Shoshone Range areas. Two different methods were applied to identify clay minerals: the Terraspec analytical spectral device and x‐ray diffraction. These two methods were chosen to compare the accuracy of rapid analysis mineral identification tools such as the Terraspec© and PIMA© to the more conventional and standardized technique of x‐ray diffraction. A review of the techniques used in this study is presented in Chapter 2. 3.5.1 Near and Short Wave Infrared Analysis (Terraspec©) The Terraspec analytical spectral device (ASD) uses near and shortwave infrared technology to measure the vibrational energy between bonds in a mineral. The Terraspec and similar tools such as the PIMA have been used as a rapid analysis technique to indentify alteration mineralogy in a variety of ore deposit settings (Uranium deposits: Zhang et al., 2001; Pb‐Zn‐Ag deposit: Sun et al., 2001, Geothermal systems: Yang et al., 2001). In this study, ~2.5cm3 blocks of each sample were analyzed using the Terraspec© analytical spectral device. Multiple readings were taken of each sample and the nature of the material being analyzed (fracture, matrix, vein) recorded. Smear mounted samples of whole rock and clay separated fractions were analyzed with the Terraspec©, however this type of sample preparation yielded nearly aspectral results, the reasons for which are unknown. The reflectance spectra collected using the 43 Terraspec were interpreted both manually and with the aid of the interpretive software The Spectral Geologist© (Merry et al., 1999). 3.5.2 X‐ray diffraction Clay minerals were identified using x‐ray diffraction in both whole rock samples and clay separates. Clay separation was performed on whole rock samples, according to the methods outlined by Moore and Reynolds (1997) to obtain the <2 m fraction, the accepted maximum particle size for clay minerals based on a spherical volume diameter (Johns et al., 1954). 50‐100g of sample was disaggregated in a Blendtec© blender with 200ml of distilled water for a period of time between 0.5‐1 minute. Higher silica content required longer periods of disaggregation. Following initial disaggregation, the solution was probed with a 500 watt ultrasonic probe for a time between 2 minutes and 10 minutes depending on silica content. The fines from this solution were then decanted into 50ml test tubes. The test tubes were placed in a centrifuge at 2000rpm for 2 minutes. The test tubes were removed and the fines decanted, probed by the ultrasonic device and distributed into new test tubes. The solute remaining at the bottom of the test tube was categorized as the coarse fraction (>25). The second round of centrifugation lasted for 5 minutes. The solute remaining from this run was termed the moderate fraction (5‐25). Following separation of the moderate fraction, the fines were decanted, probed, and centrifuged for 1‐2 hrs or until the water remaining ran clear. The remaining solute was termed the fine fraction (<2‐5). If after 2 hrs the remaining water was still cloudy, the solution was left to settle by gravity. The solute remaining following settling was termed the ultra fine fraction and contained an average grain size of 1. Grain sizes were determined by scanning electron microprobe imaging of powdered clay‐separated grain mounts using the ruler application. Air‐dried smear mounts were prepared for each fraction of all clay separated samples. 1‐2 grams of clay‐separated material was mixed with ethanol by mortar and pestle. The resulting paste was placed onto a glass slide according to the methods of Moore and Reynolds (1997). Smear mounts were analyzed by a Bruker D500 and Bruker D8 diffractometer for angles between 0 and 80° 2 for 18 minutes. Following initial XRD analyses, samples were saturated with glycol to determine the presence of smectite. Samples were placed in a petri dish with a small amount of glycol and set in an oven set at 60°C for 5hrs. The resulting diffraction patterns were then analyzed using the interpretation software EVA©. The location and FWHM of the 001 smectite, illite, and kaolinite peaks were measured. I‐S was distinguished from discrete illite and smectite phases by characteristic differences in their respective diffraction patterns. I‐S was present if one or more of the following criteria were met: a significant 44 decrease in the FWHM value of illite following glycolation due to a loss of interstitial water during heating, joint 001 illite and smectite diffraction peaks with no separation between peaks, and/or asymmetry of the 001 illite and smectite peaks (Meunier and Velde, 2004). X‐ray diffraction peak intensity has been directly correlated to abundance and is a method by which to estimate relative quantities of minerals (Alexander and Klug, 1948; Pierce and Siegel, 1969; Ouhadi and Yong, 2003). Relative abundances of illite, smectite, and kaolinite were calculated using the area measured underneath the 001 diffraction peak of each mineral in the <2 µm clay fraction. The amount of I‐S was not calculated due to presence of overlapping peaks between IS, illite, and smectite. Calculation of relative peak areas for the quantification of minerals has been used by many researchers however this method has been shown to underestimate certain minerals while overestimating others (Ouhadi and Yong, 2003). We do not suggest that these values are perfectly accurate, but provide information on relative abundances of minerals. In a subset of samples, relative abundances calculated using peak areas were compared to modal abundances estimated by thin sections and were determined to be in agreement. In the majority of samples analyzed in this study, one mineral occurs in much greater abundance than other minerals in the same sample. Accordingly, any error associated with quantitative calculation using peak areas is unlikely to affect the category into which each clay assemblage is placed. 3.5.3 Scanning Electron Microscope (SEM) Scanning electron microscopy (SEM) was used to characterize the morphology of illite in this study. Additionally, SEM was used to identify the existence of paragenetic relationships between clay minerals. Polished thin sections were analyzed using backscattered electron imaging and smear mounted samples, and randomly oriented powder mounts were analyzed using secondary electron imaging. A 15kv 10 beam was used. Energy dispersive spectra (EDS) and chemical element maps were analyzed to determine the bulk chemical composition of minerals. 3.5.4 Electron Microprobe Analysis Mineral analyses of illite were obtained by wavelength dispersive x‐ray analysis on a Cameca SX‐ 50 Scanning Electron Microprobe with 4 vertical wavelength‐dispersion x‐ray spectrometers and a fully‐ integrated SAMx energy‐dispersion X‐ray spectrometer. The following parameters were used: an accelerating voltage of 15kV, a beam current of 10nA, and a beam diameter of 10. 20‐50 points were 45 probed from each sample depending on the abundance of illite. Two types of illite textures were probed: tight knit aggregates of illite and long euhedral laths of illite. Tight knit aggregates of illite consisted of dozens of grains of illite woven together into 10‐100 µm wide masses. Multiple probe points were collected from within each aggregate and from the same grains. Only one type of texture occurred in each sample. Some illite could not be probed because the aggregates were too small (less than the spot size of the beam), or the texture of the illite surface was too rough and would cause inconsistencies in the compositional results. Illite compositions were carefully checked for evidence of any contamination from grains of other minerals inadvertently included in the broad microprobe beam. Initial statistical analyses were conducted on all data points. Specific points were later rejected if the data did not satisfy the following criteria: 85.0% < oxide total < 100% determined by a lack of interlayer H2O in pure end‐member illite and oxide totals used in other studies (Gaudette et al., 1966; Hunziker et al., 1986), and K:Al:Si:Mg:Na:Ca similar to the structural formulae outlined by Gaudette et al. (1966) and Meunier and Velde (2004) where K‐values are ~0.8 (cations). 3.5.5 Stable Isotope Analysis Nine samples (~20mg each ) of clay separated (<2) material were analyzed at the Queen’s Facility for Isotope Research at Queen’s University in Kingston, Ontario for oxygen and hydrogen stable isotope analysis using GasBench II, EA, and TC/EA technology and a DELTAplusXP Stable Isotope Ratio Mass Spectrometer. The purpose of these analyses was to help constrain the origin of clay minerals in this study. Along with chemical composition, the oxygen isotope signature of clay minerals can provide information on the origin of water from which the clay formed. Fractionation factors calculated for the distribution of 18O/16O and D/H between kaolinite‐water, illite‐water, and smectite‐water are a function of source fluid temperature; the effects of pressure are less than analytical uncertainty (Taylor, 1974; Gilg and Sheppard, 1996; Hoefs, 2009). Fractionation decreases as temperature increases (Gilg and Sheppard, 1996; Taylor, 1974). As such, supergene clay minerals can be distinguished from hypogene clay minerals by observing a relative enrichment or depletion in 18O and D. 3.6 SAMPLES The samples collected in this study are largely restricted to siliciclastic sedimentary rocks of the lower Paleozoic Roberts Mountain Allochthon (upper plate). A small subset of samples was taken from lower Paleozoic carbonate rocks (lower plate) in and around the Pipeline and Gold Acres pits, and Goat 46 window. The rationale for sample collection in this study was to identify variation in clay species over a broad spatial extent while sampling a diverse range in primary rock types both proximal and distal to hydrothermal centres. Over 450 samples were collected for Terraspec analysis and a subset of 73 samples was selected for x‐ray diffraction analysis. Not all samples collected for Terraspec analysis could be analyzed using x‐ray diffraction due to difficulty in processing quartz‐rich samples for clay separation. As such, sample density is low in the area between the Pipeline deposit and the Elder Creek deposit owing to the low clay content and the presence of abundant quartzite and chert. Samples represent a broad range of primary rock types with varying percentages of total clay content (Figure 3.8). Clay minerals occur in the pore‐space and matrix of sedimentary rocks (S), as thin (<20cm) clay seams (CS) occurring parallel to sedimentary bedding, as alteration products of feldspar and muscovite in intrusive rocks (I), as matrix material in breccia (BX), as a thin veneer along fault surfaces and in fault gouge (FG). Brackets denote the abbreviation used in subsequent tables. Samples collected with low percentages of total clay were mostly confined to aluminosilicate poor, quartzose sedimentary rocks including quartzite, chert, and mudstone. Samples were not taken directly above the Leeville property owing to the presence of mine infrastructure; however two samples were collected from mineralized dykes within the Leeville deposit. A list of all samples collected providing location, and lithology is included in Appendix C. A map of samples locations from the Leeville deposit is shown in Figure 3.9. 3.7 RESULTS 3.7.1 Morphology and textural relationships of clays Scanning electron microphotographs of the clay minerals in this study are shown in Figure 3.10. In all samples imaged of both sedimentary and non‐sedimentary protolith, illite, illite‐smectite and smectite form stacks of ‘pseudo‐hexagonal’ shaped crystals. The terms ‘hexagonal’ and ‘hairy’ were defined in section 2.2. The term ‘pseudo’ is employed here to denote a lack of definite hexagonal grain boundaries (Schleicher et al., 2006; Hancock, 1978; Hancock and Taylor, 1978). Crystal boundaries are 47 Fig. 3.8. Samples were collected from zones of strong argillization and from phyllosilicate‐bearing rocks representing a wide range of lithologies. A. Strong kaolinite and illite altered dyke at the southeast end of the Leeville property B. Magnified photograph of the same argillized intrusion from A. Illite and smectite are replacing k‐feldspar phenocrysts and intrusive matrix material C. Argillized fault gouge (white, left) adjacent to a strongly illite and smectite altered dyke (right) with relatively unaltered chert immediately adjacent to the dyke D. Mudstone clast breccia (top) with illite matrix from the North end of the Leeville property. Illite occurs as a thin veneer coating the bottom of intact mudstone beds (bottom) E. Illite‐bearing siltstone from the Elder Creek deposit, Shoshone Range F. Quartzite clast breccia with illite and quartz matrix from the Elder Creek deposit, Shoshone Range. 48 Fig. 3.9. Sample location map for the Leeville study area. 49 Fig. 3.10. Scanning electron microscope images showing the morphology of clay minerals. A‐B. Randomly oriented powder grain mounts of the <2µ size fraction of illite. A. The crystal shape of illite is pseudo‐hexagonal with rough crystal edges. All the illite analyzed in this study is of the same morphology. B. Illite crystals are not randomly oriented but occur as stacks of pseudo‐hexagonal illite grains. C‐F Back scatter images C. Illite occurs as tightly knit aggregates of illite booklets. D. The tightness of each aggregate appears to be related to the amount of space available for growth of the illite crystals. On the left side of the photograph, illite crystals are not as tightly spaced as on the right hand side. E. Muscovite occurs as long euhedral laths with one good cleavage. F. Kaolinite occurs in a similar texture to illite, in tightly knit aggregates of crystals. The morphology of kaolinite crystals could not be resolved from SEM imaging. 50 ragged and uneven similar to the pseudomorphing replacement texture described by Hancock and Taylor (1978) and possibly akin to the dissolved edges observed between end‐member growth stages by Bauer et al. (2000). Needle‐shaped or ‘hairy’ illite were not observed in any of the samples in this study. The pseudo‐hexagonal shape of illite may represent the middle stage of a reaction from ‘hairy illite’ to hexagonal illite (Bauer et al., 2000) or single stage illite formation at temperatures on the higher end of illite stability (Meunier and Velde, 2004). In both sedimentary and igneous rocks, muscovite occurred as elongate euhedral laths with visible unidirectional cleavage. Muscovite was differentiated from illite and smectite by shape and size. Furthermore, muscovite, often partially replaced by illite, exhibited higher potassium contents in energy dispersive spectra (EDS) than illite and smectite. For the purposes of this study, muscovitic illite is defined as very crystalline illite replacing euhedral muscovite laths visible in thin section and SEM. The term ‘Illite’ is restricted to samples containing tightly knit aggregates of illite. The fine grained nature of kaolinite inhibited the imaging of its crystal shape. Illite, smectite and kaolinite occur in spatially distinct, tightly knit fibrous aggregates set between quartz, feldspar, and muscovite crystals in both igneous and sedimentary rocks. Aggregates of illite, smectite, and kaolinite were similar to the dense homogeneous texture observed by Hancock (1978) indicative of the latest stage of growth formation in the context of diagenesis. Paragenetic relationships between illite and kaolinite were difficult to determine based on SEM imaging (Figure 3.11). X‐ray element mapping of potassium helped to distinguish between illite and kaolinite in samples containing both minerals. Illite can occur without kaolinite. Kaolinite can also occur without illite. Illite and kaolinite are observed texturally intergrown as fine grained aggregates, and occur simultaneously replacing laths of euhedral muscovite. The spatial occurrence of illite and kaolinite together does not necessarily indicate the synchronous growth of the two minerals. Synchronous growth of illite and kaolinite is however, indicative of specific temperatures and fluid chemistry (Figure 2.1). 3.7.2 Spatial distribution of clays ‐Leeville X‐ray diffraction pattern data for samples from the Leeville area are outlined in Table 3.1, which includes calculated relative abundances of illite, kaolinite, and smectite in the <2 µm clay fraction. Estimates of total undifferentiated clay content in each sample are given in Appendix C. Figure 3.12a 51 Fig. 3.11. Scanning electron microscope (SEM) back‐scatter images with x‐ray element maps for potassium overlain, illustrating the complex textural relationship between illite and kaolinite. The timing of growth for these minerals could not be determined by SEM imaging. Illite, muscovite and k‐feldspar are highlighted in blue, while kaolinite (potassium deficient aluminosilicate) and quartz are grey. A. Intergrown kaolinite and illite partially replacing muscovite lath in strongly argillized intrusive rock. Illite also occurs as fine grained aggregates between quartz grains. B. Fine grained aggregates of kaolinite between long euhedral laths of muscovite in an arkosic sandstone. Energy dispersive spectrometry data indicate a low potassium content of muscovite in this sample as compared to other muscovite samples indicating the potential of a partial replacement by illite C. Intergrown fine grained aggregates of kaolinite and illite in the same lithology as B. Illite appears to be the dominant mineral replacing a broken euhedral lath of muscovite. D. Illite and kaolinite are intergrown as fine grained aggregates in a strongly fractured and oxidized quartzite sample. E. Arkosic sandstone sample containing kaolinite feldspar and muscovite, and lacking illite. F. Patchily hematized clay fault gouge containing quartz and kaolinite, lacking illite. 52 Fig. 3.12. Maps showing the distribution of A. clay minerals identified by XRD diffraction data B. Clay minerals identified by Terraspec data C. Contoured Full Width Half the Maximum values (2), D. Illite formation temperatures calculated from electron microprobe data. 53 Table 3.1. XRD data from the <2 µm clay fraction and whole rock samples from the Leeville field area. n=39 Sample Rock I ■ K S I‐S + Illite Peak (2 Illite Peak* (2 Illite FWHM (2) Illite FWHM *(2) ΔFWHM Illite (2) Smectite Peak (2 Smectite Peak* (2 Smectite FWHM (2) Smectite FWHM* (2) Kaolinite Peak (2 Kaolinite Peak* (2 % I %K % S Type ▲ ◊ 472 I ■ ▲ 10.109 9.968 0.667 0.603 0.064 ‐ ‐ ‐ ‐ 7.146 7.149 57.6 42.4 0.0 473 I ■ ▲ 10.003 9.924 0.228 0.245 0.017 ‐ ‐ ‐ ‐ 7.109 7.104 12.4 87.6 0.0 477a FG ■ ▲ ◊ 10.107 9.975 0.568 0.497 0.071 19.658 18.856 0.978 0.648 7.164 7.169 1.5 92.4 6.1 478 I ■ ▲ ◊ 10.189 9.987 0.584 0.514 0.07 14.98 16.788 1.324 0.913 7.16 7.165 9.8 31.8 58.4 484 I ■ ▲ 10.091 9.971 0.498 0.51 ‐0.012 ‐ ‐ ‐ ‐ 7.154 7.147 81.1 18.9 0.0 486a× S ■ 10.039 0.494 ‐ ‐ ‐ ‐ ‐ ‐ 100 0.0 0.0 487 CS ■ ▲ 10.012 9.989 0.502 0.48 0.022 ‐ ‐ ‐ ‐ 7.202 7.209 41.2 58.8 0.0 492 BX ■ ▲ ◊ 10.005 9.995 0.266 0.265 0.001 ‐ ‐ ‐ ‐ 7.157 7.167 98.1 0.9 1.0 495 S ■ ▲ 10.04 9.998 0.455 0.452 0.003 ‐ ‐ ‐ ‐ 7.142 7.145 74.2 25.8 0.0 496 S ■ ▲ ◊ 9.986 9.972 0.344 0.352 ‐0.008 14.783 16.982 1.322 0.861 7.166 7.164 3.1 32.4 64.5 498× I ■ ▲ 9.968 9.934 0.247 0.255 ‐0.008 ‐ ‐ ‐ ‐ 7.140 7.117 21.2 78.8 0.0 499 I ■ ▲ ◊ + 9.999 10.01 0.398 0.338 0.06 14.299 16.625 1.128 1.101 7.167 7.155 4.3 25.1 70.6 502 BX ■ 9.923 9.919 0.214 0.217 ‐0.07 ‐ ‐ ‐ ‐ ‐ ‐ 100.0 0.0 0.0 506 S ■ 9.97 0.951 0.247 0.213 0.034 ‐ ‐ ‐ ‐ ‐ ‐ 100.0 0.0 0.0 507 CS ■ 10.079 9.944 0.574 0.445 0.129 ‐ ‐ ‐ ‐ ‐ ‐ 100.0 0.0 0.0 508 I ■ ◊ + 10.045 10.013 0.322 0.324 ‐0.002 16.521 16.023 1.077 1.824 ‐ ‐ 10.1 63.9 26.0 509 ¥ I ▲ ◊ ‐ ‐ ‐ ‐ ‐ 14.225 16.334 ‐ ‐ ‐ ‐ 5.6 28.4 66.0 510a FG ■ ▲ ◊ 10.062 10.013 0.389 0.352 0.037 14.376 16.521 1.041 1.089 7.165 7.172 62.6 7.2 30.2 510c CS ■ ▲ ◊ 10.093 10.047 0.394 0.365 0.029 14.792 17.142 2.517 1.151 7.176 56.3 9.9 33.8 511c FG ■ ◊ 9.888 9.935 0.143 0.231 ‐0.108 15.007 16.785 ‐ ‐ ‐ ‐ 9.4 0.0 90.6 513b FG ■ ▲ 10.042 9.989 0.406 0.403 0.003 ‐ ‐ ‐ ‐ 7.157 7.156 92.6 7.5 0.0 513c FG ■ ▲ 10.06 10.026 0.38 0.359 0.021 ‐ ‐ ‐ ‐ 7.211 7.156 77.5 22.5 0.0 516b CS ■ ▲ ◊ 10 9.98 0.283 0.271 0.012 12.486 16.625 1.063 1.479 7.172 7.159 17.0 52.0 31.0 516d FG ■ ▲ ◊ 9.989 9.999 0.238 0.235 0.003 12.427 16.729 1.107 1.348 7.169 7.169 2.2 6.6 91.2 519 S ■ ▲ ◊ 10.113 10.003 0.426 0.458 ‐0.032 14.454 16.842 1.123 0.59 7.159 7.168 1.5 48.9 49.6 521a BX ■ ◊ 10.037 10.01 0.303 0.293 0.01 ‐ ‐ ‐ ‐ ‐ ‐ 99.1 0.9 523 I ▲ ‐ ‐ ‐ ‐ ‐ 13.888 16.749 1.144 0.448 7.152 7.158 0.0 0.6 99.4 524 FG ■ ▲ ◊ + 10.023 9.909 0.246 0.355 ‐0.109 14.245 16.521 0.976 1.464 7.169 7.116 18.0 14.4 67.6 525a‐1 FG ■ 10.079 10.007 0.437 0.421 0.016 ‐ ‐ ‐ ‐ ‐ ‐ 100.0 0.0 0.0 525a‐2 FG ■ ▲ 10.064 9.996 0.419 0.421 ‐0.002 ‐ ‐ ‐ ‐ 7.152 7.155 88.8 11.2 0.0 525b FG ■ 10.056 9.994 0.405 0.392 0.013 ‐ ‐ ‐ ‐ ‐ ‐ 100.0 0.0 0.0 526a I ◊ + ‐ ‐ ‐ ‐ ‐ 14.684 16.637 0.731 0.5 ‐ ‐ 0.0 0.2 99.8 526b I ▲ ◊ ‐ ‐ ‐ ‐ ‐ 14.895 16.685 0.994 0.819 7.21 ‐ 0.0 0.1 99.9 527 FG ■ ▲ ◊ 10.037 9.927 0.616 0.578 0.038 14.613 16.625 1.774 1.345 7.171 7.142 29.9 30.7 39.4 528a CS ■ ▲ ◊ + 10.075 9.916 1.231 0.788 0.443 15.384 17.052 1.439 1.158 7.176 7.176 83.6 4.9 11.5 528c S ■ ▲ ◊ + 10.113 9.998 1.121 0.999 0.122 16.521 17.162 1.91 1.577 7.194 0.724 27.1 36.5 36.5 528d I ■ ◊ 10.087 9.945 0.549 0.532 0.017 14.694 16.625 1.001 0.898 ‐ ‐ 83.0 0.0 17.0 UL‐18 I ■ ▲ ◊ 10.056 10.022 0.143 0.153 ‐0.01 14.228 16.893 0.159 0.329 7.165 7.149 5.4 89.4 5.2 UL‐19 I ■ ▲ ◊ + 10.005 10.002 0.167 0.154 0.013 14.144 16.811 0.193 0.474 7.133 7.133 8.5 49.6 42.0 I = illite, I‐S=illite‐smectite, S=smectite, K=kaolinite × = whole rock XRD analysis * = glycolated ¥ = contains chlorite, no FWHM values calculated for smectite Rock type: I = intrusive, S= sedimentary, CS= clay seam, FG = fault gouge, BX = breccia 54 shows the distribution of clay minerals identified from x‐ray diffraction patterns of the fine fraction (<2m) across the map area. Most samples contained all three of the dominant clay types: illite, kaolinite, and smectite in varying amounts. No obvious correlation became evident between lithology and associated clay mineralogy although smectite does appear most abundant in altered intrusive samples, including those from in and around the Vivian sill: 499, 509, 523, 526, and 527. Dyke samples UL‐18 and UL‐19 from within the Leeville underground deposit containing elevated gold values (0.094ppm and 0.005ppm respectively) contained abundant kaolinite with smectite and illite. Sample UL‐19 contains significantly more smectite and I‐S, and less gold than sample UL‐18. Illite dominated samples were present within close proximity to the surface projection of the Leeville deposit. Sample 525, on the edge of the projected deposit only comprised illite. Samples 510 and 513 contained abundant illite with lesser amounts of kaolinite, and smectite. Sample 528 contained more illite than smectite. Figure 3.12b shows the distribution of clay minerals identified from Terraspec reflectance spectra of whole rock samples. Smectite dominates the northwestern part of the Leeville study area. Some illite occurs above the surface projection of the Leeville deposit. The most prominent trend outlined by Terraspec data is a SW‐NE trending region of kaolinite dominated samples starting at the Leeville deposit. This trend in kaolinite correlates to the illite trend observed in the same samples using x‐ray diffraction. Many samples returned aspectral data characterized by a reflectance spectrum with no identifiable absorption peaks. Comparison of results from the Terraspec data and the XRD indicate that the majority of samples analyzed by the Terraspec produced reflectance spectra indicative of a different clay assemblage than that returned by XRD. Terraspec data returned numerous false positives when in fact illite occurs only in discrete areas. Inconsistency between XRD and Terraspec data are discussed in detail below. FWHM values of illite in samples from the Leeville area, calibrated to the CIS scale are shown in Table 3.2. Results indicate that FWHM values are dominantly above the accepted value of 0.422 for the boundary between diagenesis and low grade anchizonal metamorphism. This boundary coincides with a temperature of ~200C (Kisch, 1990) which is within the temperature range measured from fluid inclusions paragenetically related to Carlin‐type Au‐mineralization (Hofstra and Cline, 2000; Cline et al., 2005). Two samples within close proximity to the surface projection of the Leeville deposit show lower FWHM values indicative of increased crystallinity and higher temperature than surrounding samples. Contouring of FWHM values across the Leeville deposit, shown in Figure 3.12c, an extensive zone of high crystallinity illite; above and outboard of the surface projection of the Leeville deposit. Contoured 55 Table. 3.2. FWHM values of illite from clay separated samples collected from the vicinity of the Leeville deposit calibrated to the CIS‐scale. Samples lacking illite are not included in this table. Sample Rock Experimental Calibrated Inferred ID Type FWHM Value FWHM Value T (°C)* (2) (2) 472 I 0.67 1.24 Diagenetic 477a I 0.57 1.06 Diagenetic 478 FG 0.58 1.08 Diagenetic 484 I 0.5 0.94 Diagenetic 486 S 0.49 0.92 Diagenetic 487 CS 0.5 0.94 Diagenetic 492 BX 0.27 0.53 Diagenetic 495 S 0.46 0.87 Diagenetic 496 S 0.34 0.65 Diagenetic 499 I 0.4 0.76 Diagenetic 502 BX 0.21 0.42 Diagenetic ‐ Anchizonal 506 S 0.25 0.49 Diagenetic 507 CS 0.57 1.06 Diagenetic 508 I 0.32 0.62 Diagenetic 510a FG 0.39 0.74 Diagenetic 510c CS 0.39 0.74 Diagenetic 511c FG 0.14 0.29 Epizonal 513b FG 0.41 0.78 Diagenetic 513c FG 0.38 0.72 Diagenetic 516b CS 0.28 0.54 Diagenetic ‐ Anchizonal 516d FG 0.24 0.47 Diagenetic ‐ Anchizonal 519 S 0.43 0.81 Diagenetic 521a BX 0.3 0.58 Diagenetic ‐ Anchizonal 524 FG 0.25 0.49 Diagenetic ‐ Anchizonal 525a‐1 FG 0.44 0.83 Diagenetic 525a‐2 FG 0.42 0.79 Diagenetic 525b FG 0.41 0.78 Diagenetic 527 FG 0.62 1.15 Diagenetic 528a CS 1.23 2.25 Diagenetic 528c S 1.12 2.05 Diagenetic 528d I 0.55 1.03 Diagenetic UL‐18 I 0.14 0.29 Epizonal UL‐19 I 0.15 0.31 Epizonal n=33 Rock type: I = intrusive, S= sedimentary, CS= clay seam, FG = fault gouge, BX = breccia 56 FWHM values increase away from the Leeville deposit, with small zones of crystalline illite proximal to faults and other permeable pathways. 3.7.3 Distribution of clays ‐ Shoshone Range Field Area XRD pattern data for samples from the Shoshone Range field area are outlined in Table 3.3, Appendix D. The distribution of clays identified using XRD is shown in Figure 3.13a. Abundant quartz content and lack of aluminosilicate minerals in this area restricted the samples analyzed by XRD. There are two main sites of argillization along this transect, both of which are known occurrences of gold; the giant Carlin Au Pipeline deposit and the Elder Creek gold deposit. Samples from the Pipeline and Gold Acres deposits contained illite and some sedimentary muscovite. At the Elder Creek deposit, illite was identified in the matrix of gold‐hosting quartzite clast breccia, in strongly altered siltstone, and in dyke material. To the north of the Elder Creek deposit, unaltered siltstone contained no illite, only smectite and kaolinite. Unaltered siltstone and argillite from west of the Pipeline deposit contained illite‐ smectite, smectite and kaolinite. The assemblage contained by samples 134 and 123 appears to be the background clay content of most upper plate rocks. Calibrated FWHM values from the Shoshone Range study area are presented in Table 3.4, Figure 3.13b and indicate a range of metamorphic grades. Samples from the Gold Acres deposit exhibited the lowest FWHM values indicative of the highest crystallinity illite. Within the Elder Creek deposit FWHM values vary between 0.45‐0.94 2. Illite occurring within the main zone of mineralization at Elder Creek is less crystalline than illite from the Pipeline and Gold Acres deposits. Just outside the main zone of mineralization, one sample contains poorly crystalline illite with an FWHM value of 1.12 2however, ~1km northwest of the main zone of mineralization, sample 230 (quartzite clast breccias) exhibits highly crystalline illite with an FWHM value of 0.45 2. Over 300 samples were collected from the Shoshone Range field area for Terraspec© analysis. The data from these analyses are shown in
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Beyond the confines of the ore deposit : mapping low temperature hydrothermal alteration above, within,… Ahmed, Ayesha Doris 2010
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| Title | Beyond the confines of the ore deposit : mapping low temperature hydrothermal alteration above, within, and beneath Carlin-type Gold deposits |
| Creator |
Ahmed, Ayesha Doris |
| Publisher | University of British Columbia |
| Date Issued | 2010 |
| Description | Multiple analytical techniques were employed to investigate distal patterns in low temperature hydrothermal fluid flow into and out of Carlin-type gold deposits in two study areas: the Leeville deposit and the Shoshone range including the Pipeline, Gold Acres and Elder Creek deposits. Previous studies indicate that gold is hosted in lower Paleozoic carbonate rocks overlain by thick sequences of similarly aged siliciclastic rocks. Patterns in δ¹⁸O depletion (<20‰VSMOW), and Au, As, Sb, Hg, Tl, and Te concentrations in lower Paleozoic carbonate rock identified three disconnected lateral fluid pathways into the Pipeline deposit: a main conduit providing gold-bearing fluid to the main ore body, the Abyss fault located ~300m below the main ore zone, and the RMT located underneath the Abyss fault. Following gold precipitation in the Pipeline deposit, gold-depleted fluids were likely exhausted laterally, at least initially, along the same structures as those that allowed fluid to enter the deposit. Upon intersecting the RMT fault, fluid either exploited the fault to reach surface, or transgressed overlying siliciclastic rocks via small scale faults and fractures. δ¹⁸O and δD values of H₂O in equilibrium clay minerals, and the concentration and crystallinity of illite outlined multiple zones of hydrothermal alteration in surface rocks from both the Shoshone Range and Leeville study areas, however no genetic link was established to Carlin-type gold mineralization at depth. Similarities in trace element geochemistry, ore assemblage, and alteration assemblages however, suggest that the Elder Creek deposit may represent low temperature (200°C) gold mineralization resulting from the exhaust of Carlin-type ore forming fluid. The region above the surface projection of the Leeville deposit exhibits multiple zones of hydrothermal fluid upflow resulting in pervasive illitization of surface siliciclastic rocks. The Pipeline/ Gold Acres also contain abundant crystalline illite. The presence of highly crystalline illite highlights zones of focused fluid upflow, typically along faults and other secondary permeability structures such as breccias. |
| Genre |
Thesis/Dissertation |
| Type |
Text |
| Language | eng |
| Date Available | 2010-10-21 |
| Provider | Vancouver : University of British Columbia Library |
| Rights | Attribution-NonCommercial-NoDerivatives 4.0 International |
| DOI | 10.14288/1.0052809 |
| URI | http://hdl.handle.net/2429/29423 |
| Degree |
Master of Science - MSc |
| Program |
Geological Sciences |
| Affiliation |
Science, Faculty of Earth, Ocean and Atmospheric Sciences, Department of |
| Degree Grantor | University of British Columbia |
| GraduationDate | 2010-11 |
| Campus |
UBCV |
| Scholarly Level | Graduate |
| Rights URI | http://creativecommons.org/licenses/by-nc-nd/4.0/ |
| AggregatedSourceRepository | DSpace |
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