Orogenic Gold Formation and Tectonic Evolution of the Grass Valley Gold District and Temporal Correlations of Gold Deposits in California

Home , Mother Lode, Platform (geology), Sonoma orogeny

OROGENIC GOLD FORMATION AND TECTONIC EVOLUTION OF THE GRASS VALLEY GOLD DISTRICT AND TEMPORAL CORRELATIONS OF GOLD DEPOSITS IN CALIFORNIA

by Ryan D. Taylor

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

Golden, Colorado

Date ______

Signed: ______Ryan D. Taylor

Signed: ______Dr. Thomas Monecke Thesis Advisor

Golden, Colorado

Date ______

Signed: ______Dr. Paul Santi Professor and Head Department of Geology and Geological Engineering

ABSTRACT

With a total past production of 13 Moz of lode gold, the Grass Valley gold district of the Sierra Nevada foothills province is the historically most productive lode gold source in California. Despite its economic importance, an understanding of the broad processes controlling the gold formation is lacking. Two distinct vein sets are present in Grass Valley: a north-trending set (N-S veins) hosted by the Grass Valley granodiorite and an east-trending set (E-W veins) hosted within mafic-ultramafic rocks. Questions of how these relate to each other and if they are products of the same event or different events remain to be answered. Some of the previously published data are conflicting, and the timing of gold formation for the district seems inconsistent with previous interpretations of orogenic gold formation in the Cordillera of California, particularly when viewed relative to the much better studied Mother Lode belt in the southern Sierra Nevada. A geochemical and geochronological characterization of the ore- hosting granodiorite is also lacking. The present study represents the first detailed modern study on the Grass Valley gold district. The research included a detailed microanalytical and geochronological study of the ore- hosting granodiorite and the orogenic quartz veins. It is shown the ore-hosting Grass Valley granodiorite was emplaced at 159.9 ± 2.2 Ma (U-Pb zircon) at temperatures of nearly 800 ° C and at paleodepths of approximately 3 km. It rapidly cooled to below 300 ° C between 162-160 Ma (40Ar/39Ar hornblende and biotite). After crystallization, the intrusion underwent brittle fracturing concurrent with N-S vein formation. The hydrothermal fluids interacted with the granodiorite and formed monazite and xenotime as alteration products, permitting U-Pb geochronology. An age of 162 ± 5 Ma for vein formation was determined for xenotime. This age is indistinguishable from the intrusive age, but must have occurred after the pluton was cool enough to undergo brittle fracturing. The hydrothermal monazite and xenotime have markedly different geochemical characteristics than magmatic phases. Magmatic monazite from the Grass Valley granodiorite has Th concentrations up to 11.6 wt.%, whereas the hydrothermal monazite has maximum Th concentrations of 0.2 wt.%. The REE profiles are also significantly different, including a strong negative Eu anomaly for the magmatic phases and no Eu anomaly for the hydrothermal phases. Therefore, despite this age overlap between magmatism and hydrothermal activity, they are not genetically related. This implies that the vein-hosted phases are not

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xenocrysts and also did not form from an evolving magmatic-hydrothermal system, but are instead formed by orogenic fluids. A second hydrothermal event formed the E-W veins at ~152 Ma, isolated in time from any regional magmatism. In addition to the geochronological research on the ore-hosting granodiorite and the veins, a detailed paragenetic investigation was performed on the orogenic veins as they are remarkably undeformed. In contrast to typical orogenic gold deposits displaying textures indicating brittle- ductile deformation and recrystallization, those of Grass Valley only display minor brittle fracturing of quartz and pyrite. Optical microscopy and optical cathodoluminescence imaging revealed the presence of multiple generations of quartz characterized by different luminescence responses and concentrations of secondary fluid inclusion trails. Pyrite crystallized following quartz precipitation. Gold precipitates relatively late in the paragenetic sequence entirely independent of quartz and is found within fractures in quartz, in fractures and voids within pyrite, and intergrown with galena and mica. The time of gold mineralization is recorded in pyrite by a chemically distinct growth zone containing arsenic, and nickel and cobalt zones with pyrites found in the E-W veins hosted in mafic-ultramafic rocks. The formation of quartz due to adiabatic decompression indicates the importance of pressure fluctuations in vein formation. The correlation of elements derived from the fluid (Ag, As, and Au) and those from the host rock (Co, Ni, and Pb) indicate the importance of fluid reactions with the local host rock during mineralization. Developing a regional scale view of gold mineralization in the Cordillera of California can help shape the understanding of how gold deposit formation relates to various stages in the late Mesozoic tectonic evolution of California. To constrain the timing of gold mineralization in the other major gold province of California, white mica was separated from samples from eight deposits. Four of these exhibited evidence for excess argon interpreted to result from intense deformation and (or) the presence of mineral inclusions. The other four samples had argon isotope age spectra that provided plateau ages of ~160-140 Ma. The maximum age corresponds to a major plate reorganization and initial gold mineralization in the Sierra Nevada. The minimum age corresponds to the initiation of the lateral offset of the Klamath Mountains westward of the Sierra Nevada and the active arc. This marks the termination of both hydrothermal activity and magmatism in the Klamath Mountains. However, orogenic gold formation within the Sierra Nevada foothills continued as it was still located on the active arc.

TABLE OF CONTENTS ABSTRACT ...... iii LIST OF FIGURES ...... viii LIST OF TABLES ...... x LIST OF ABBREVIATIONS ...... xi ACKNOWLEDGEMENTS ...... xv CHAPTER 1 INTRODUCTION ...... 1 1.1 Orogenic Gold Deposits ...... 1 1.2 Grass Valley Gold District, California ...... 4 1.3 Previous Research ...... 4 1.4 Thesis Objectives ...... 5 1.5 Thesis Organization ...... 6 1.6 References ...... 7 CHAPTER 2 APPLICATION OF U-TH-PB PHOSPHATE GEOCHRONOLOGY TO YOUNG OROGENIC GOLD DEPOSITS: NEW AGE CONSTRAINTS ON THE FORMATION OF THE GRASS VALLEY GOLD DISTRICT, SIERRA NEVADA FOOTHILLS PROVINCE, CALIFORNIA ...... 11 2.1 Tectonic Setting of Grass Valley ...... 12 2.2 Geology of the Grass Valley District ...... 15 2.3 Geology of the Grass Valley Deposits ...... 18 2.4 Previous Geochronological Research ...... 20 2.5 Materials and Methods ...... 22 2.5.1 Sampling and Petrographic Investigations ...... 22 2.5.2 Whole-Rock Geochemistry ...... 22 2.5.3 Electron Microprobe Analysis of Monazite and Xenotime ...... 23 2.5.4 U-Pb Zircon Geochronology ...... 23 2.5.5 U-Pb Geochronology of Monazite and Xenotime ...... 24 2.5.6 40Ar/39Ar Geochronology ...... 24 2.5.7 Hornblende Geobarometry and Plagioclase-Amphibole Geothermometry ...... 25 2.6 Results ...... 25 2.6.1 Petrography and Geochemistry of the Grass Valley Granodiorite ...... 25

2.6.2 Age of the Grass Valley Granodiorite ...... 26 2.6.3 Geochemistry of Vein-Hosted Monazite and Xenotime ...... 31 2.6.4 Age of Vein Formation...... 31 2.6.5 Geothermometry and Geobarometry ...... 40 2.7 Discussion ...... 43 2.7.1 Use of Hydrothermal Phosphate Minerals for Determining Absolute Age of Phanerozoic Ores ...... 43 2.7.2 Timing of Gold Deposit Formation ...... 45 2.7.3 Relationship Between Magmatism and Gold Mineralization ...... 46 2.7.4 Relationship Between Exhumation and Gold Mineralization ...... 48 2.7.5 Middle-Late Jurassic Tectonic Regimes Controlling Gold Formation ...... 49 2.8 Conclusions ...... 50 2.9 References ...... 52 CHAPTER 3 PARAGENETIC EVOLUTION AND FORMATION MECHANISMS OF OROGENIC GOLD DEPOSITS DETERMINED BY MICROANALYTICAL GEOCHEMISTRY AND PETROGRAPHY: NEW PERSPECTIVES FROM THE GRASS VALLEY DISTRICT, CALIFORNIA ...... 59 3.1 Regional Geology ...... 60 3.2 Geology of the Grass Valley Gold District ...... 63 3.3 Materials and Methods ...... 66 3.4 Results ...... 68 3.4.1 Vein Textures ...... 68 3.4.2 Vein Quartz ...... 69 3.4.3 Pyrite Textures and Chemistry ...... 72 3.4.4 Mica and Chlorite ...... 84 3.4.5 Other Sulfide Minerals ...... 84 3.4.6 Gold Textures and Chemistry...... 85 3.4.7 Carbonate Minerals ...... 85 3.4.8 Sulfur Isotopes ...... 87 3.5 Discussion ...... 88 3.5.1 Deformation of Vein Minerals ...... 88 3.5.2 Paragenesis ...... 90 3.5.3 Timing of Gold Deposition ...... 94

3.5.4 Differences Between the Two Vein Sets ...... 96 3.5.5 Evolution of the Systems ...... 97 3.6 Conclusions ...... 100 3.7 References ...... 101 CHAPTER 4 40AR/39AR GEOCHRONOLOGY OF HYDROTHERMAL ACTIVITY RELATED TO GOLD MINERALIZATION IN THE KLAMATH MOUNTAINS, CALIFORNIA ...... 106 4.1 Background ...... 106 4.1.1 Tectonic Setting ...... 107 4.1.2 Klamath Mountains Geology ...... 109 4.1.3 Gold Deposits ...... 110 4.2 Study Sites ...... 111 4.2.1 McKeen Deposit, Callahan District ...... 111 4.2.2 Hickey Deposit, Liberty District ...... 112 4.2.3 Quartz Hill Deposit, Scott Bar District ...... 113 4.2.4 Schroeder Deposit, Yreka-Fort Jones District ...... 113 4.2.5 McKinley Deposit, Humbug District ...... 113 4.2.6 Washington Deposit, French Gulch-Deadwood District ...... 114 4.2.7 Yankee John Deposit, Redding District ...... 114 4.2.8 Walker Deposit, Old Diggings District ...... 115 4.3 Materials and Methods ...... 115 4.4 Results ...... 117 4.5 Discussion ...... 129 4.5.1 Timing of Deposit Formation ...... 129 4.5.2 Timing of Magmatism and Mineralization ...... 132 4.5.3 Tectonic and Metallogenic Relationships ...... 134 4.6 Conclusions ...... 138 4.7 References ...... 140 CHAPTER 5 CONCLUSIONS ...... 145 APPENDIX A SUPPLEMENTAL ELECTRONIC FILES ...... 149

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LIST OF FIGURES Figure 1.1 Map of California ...... 2 Figure 2.1 Geologic map of northern California ...... 13 Figure 2.2 Geologic map of the Grass Valley gold district ...... 16 Figure 2.3 Photographs of typical samples from the Grass Valley district ...... 19 Figure 2.4 BSE images of hydrothermal monazite and xenotime ...... 21 Figure 2.5 Chondrite-normalized REE plots for the Grass Valley granodiorite ...... 28 Figure 2.6 Cathodoluminescence images of Grass Valley granodiorite zircon crystals ...... 28 Figure 2.7 U-Pb isotope data for Grass Valley granodiorite zircon crystals ...... 30 Figure 2.8 40Ar/39Ar age spectra and Ca/K plots for hornblende and biotite ...... 33 Figure 2.9 Trace element characteristics of monazite and xenotime crystals ...... 35 Figure 2.10 U-Pb isotope data for hydrothermal xenotime ...... 39 Figure 3.1 Photomicrographs depicting typical deformation and recrystallization of orogenic deposits ...... 60 Figure 3.2 Geologic map of northern California ...... 62 Figure 3.3 Geologic map of the Grass Valley gold district ...... 64 Figure 3.4 Photomicrographs of quartz veins in Grass Valley ...... 70 Figure 3.5 Back-scattered electron images of pyrite crystals with electron microprobe geochemical data ...... 73 Figure 3.6 Back-scattered electron images of pyrite crystals with LA-ICP-MS geochemical data in different zones ...... 75 Figure 3.7 Selected qualitative EDS elemental maps ...... 77 Figure 3.8 Select geochemical plots for pyrite analyses ...... 82 Figure 3.9 Photomicrographs of gold textural relationships ...... 86 Figure 3.10 Sulfur isotopic composition of pyrite crystals ...... 88 Figure 3.11 Paragenetic sequence ...... 91 Figure 3.12 Photomicrographs of mineral textural relationships...... 93 Figure 3.13 Ternary diagram of arsenic-bearing pyrite crystals ...... 98 Figure 4.1 Geologic map of northern California ...... 108 Figure 4.2 Geologic map of the Klamath Mountains ...... 110 Figure 4.3 Photomicrographs of white mica samples ...... 117 Figure 4.4 Mica compositions on a ternary diagram ...... 124

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Figure 4.5 Age spectra and plateau results of samples with no excess argon ...... 125 Figure 4.6 Disrupted age spectra of samples ...... 126 Figure 4.7 Timeline of events related to gold mineralization in California ...... 136 Figure 4.8 Time slices of northern California and southern Oregon ...... 137

LIST OF TABLE Table 1.1 Summary of regional geology in the Sierra Nevada and Klamath Mountains ...... 3 Table 2.1 Whole rock geochemistry of the Grass Valley granodiorite ...... 27 Table 2.2 SHRIMP U-Th-Pb data for magmatic zircon ...... 29 Table 2.3 Summary of argon isotope analyses of the Grass Valley granodiorite ...... 32 Table 2.4 Electron microprobe geochemical data of hydrothermal monazite ...... 34 Table 2.5 Electron microprobe geochemical data for hydrothermal xenotime ...... 36 Table 2.6 Electron microprobe geochemical data for magmatic monazite ...... 37 Table 2.7 SHRIMP U-Th-Pb data for hydrothermal xenotime...... 38 Table 2.8 SHRIMP U-Th-Pb data for hydrothermal monazite ...... 41 Table 2.9 Electron microprobe geochemical data of magmatic plagioclase ...... 41 Table 2.10 Electron microprobe geochemical data of magmatic amphiboles ...... 42 Table 3.1 The three most productive lode gold districts in California ...... 65 Table 3.2 Laser ablation ICP-MS geochemical data for minor and trace elements in pyrite from Grass Valley ...... 78 Table 3.3 Electron microprobe geochemical data for major and minor elements in pyrite from Grass Valley ...... 80 Table 3.4 Electron microprobe values of gold and silver within gold grains of Grass Valley ....87 Table 4.1 Sample information...... 116 Table 4.2 Electron microprobe geochemical data of hydrothermal white mica ...... 118 Table 4.3 Summary of argon isotope data ...... 120 Table 4.4 40Ar/39Ar and U-Pb ages of orogenic gold deposits in California ...... 135

LIST OF ABBREVIATIONS General Abbreviations: BSE back-scattered electron ca. circa CL cathodoluminescence EDS energy dispersive spectroscopy e.g. exempli gratia EMPA electron microprobe (analysis) ENE east northeast E-W east-west ICP-AES inductively coupled plasma atomic emission spectrometry ICP-MS inductively coupled plasma mass spectrometry i.e. id est LA-ICP-MS laser ablation-inductively coupled plasma-mass spectrometry MC-ICP-MS multi-collector inductively coupled plasma-mass spectrometer MSWD mean square weighted deviation NNW north northwest N-S north-south SEM scanning electron microscope VCDT Vienna Cañon Diablo Troilite WSW west southwest XRF x-ray diffraction

Units: at. % atomic percent cm centimeter g/t grams per tonne kbar kilobar keV kilovolts km kilometer m meter

Ma million years ago mm millimeter Moz million ounces Myr millions of years nA nanoamps oz ounce ppb parts per billion ppm parts per million vol. % volume percent wt. % weight percent °C degrees Celsius δ34S sulfur-34 isotope fraction % per cent ‰ per mil µA/mm2 microamps per cubic millimeter µm micrometer

Chemical Abbreviations: Ag silver Al aluminum Ar argon As arsenic Au gold Bi bismuth Ca calcium

CaF2 calcium fluoride Cd cadmium Ce cerium Co cobalt

CO2 carbon dioxide Cr chromium

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Cu copper Dy dysprosium Er erbium Eu europium Fe iron Gd gadolinium Hg mercury

H2O water

HNO3 nitric acid HREE heavy rare-earth element Ho holmium K potassium

K2O potassium oxide La lanthanum LREE light rare-earth element Lu lutetium Mg magnesium Mn manganese Mo molybdenum MREE middle rare-earth element Nd neodymium Ni nickel Os osmium P phosphorous Pb lead Pr praseodymium Re rhenium REE rare-earth element S sulfur Sb antimony Sc scandium

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Se selenium Si silicon

SiO2 silica Sm samarium Sn tin Sr strontium Tb terbium Te tellurium Th thorium Ti titanium

TiO2 titanium dioxide Tl thallium Tm thulium U uranium Y yttrium Yb ytterbium Zn zinc

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ACKNOWLEDGEMENTS First and foremost, I would like to thank Dr. Richard Goldfarb. Without him, I would have never gotten started in my career with the USGS or have been introduced to orogenic gold deposits. Rich not only ignited my enthusiasm for understanding gold deposits, but also introduced me to innumerable experts and projects regarding other aspects of geology that has opened many doors to help further my career. I appreciate his patience and the amount of time he has spent providing advice and mentorship to help me. It is an honor to have worked with Rich and to learn from him over these years. Dr. Thomas Monecke has been a patient, helpful, and enlightening advisor. He was always available to discuss my project whenever I had questions and helped keep me focused on the tasks at hand. I would also like to thank the other committee members, Drs. Nigel Kelly and Brian Gorman. Nigel provided thoughtful reviews and was available and willing to give advice, especially regarding phosphate minerals and geochronological methods. Brian generously agreed to take time to chair my committee. This project was funded through the Mineral Resources Program of the USGS as a part of various projects. Without help from many of my colleagues, this project would not have succeeded. Erin Marsh joined me in the field for every visit I made to California as part of this study. She was also a great resource for discussing various aspects of orogenic gold formation. Without Heather Lowers, Mike Cosca, Alan Koenig, and Cayce Gulbransen to help with the analytical work, this project would have never left the planning stages. Garth Graham, Karen Kelley, Ed duBray, and Dave Leach were always willing to listen to my ideas and provide feedback on the project. Jim Reynolds provided critical insights into the science of fluid inclusions. Lastly, I would like to thank all of my family and friends. You have all kept me calm and sane throughout this process.

xv CHAPTER 1 INTRODUCTION

This chapter outlines the objectives of this study on the orogenic gold deposits of the Grass Valley district within the Sierra Nevada foothills, California, provides a brief project background, describes previous research, and details the organization of the thesis. 1.1 Orogenic Gold Deposits Orogenic gold deposits have formed episodically during the Archean through the Phanerozoic in relation to accretion of juvenile crust to continental margins (e.g., Goldfarb et al., 2001; Groves et al., 2005). They are typically hosted within rocks that have experienced metamorphism at greenschist facies conditions, and are commonly interpreted to form during the retrograde phase of metamorphism in these terranes. Granitic intrusions related to orogenesis may also host these lode gold systems. Gold-laden fluids are interpreted to be generated at depth during orogenesis as rocks pass through the greenschist-amphibolite facies transition. Hydrous and carbonate minerals break down to release H2O and CO2 and sulfide minerals break down to release Au and other metals (Pitcairn et al., 2006; Phillips and Powell, 2010). These fluids are then concentrated and transported upwards within large, regional-scale fault zones during seismic events associated with earthquakes (Sibson, 1990). The fluids precipitate gangue and ore minerals within secondary and tertiary faults that are related to the large first-order faults (Goldfarb et al., 2005). The gold deposits of California are classified as orogenic gold deposits and placer deposits derived from their erosion (e.g., Goldfarb et al., 2008). They are found within the metamorphic rocks and granitic intrusions of the Sierra Nevada foothills and the Klamath Mountains (Fig. 1.1). These ranges are composed of Devonian through Late Jurassic autochthonous and allochthonous terranes, mostly of oceanic arc affinity, that were progressively accreted to the North American continental margin during complex transpressional and transtensional events (Table 1.1; Dickinson, 2008; Ernst et al., 2008). The accreted terranes now occur as NNW- trending parallel belts of rocks that are separated by fault zones marking the terrane-bounding sutures. Syn- and post-accretionary plutons intruded these terranes within two periods: ~170- 140 Ma and ~120-80 Ma.

Figure 1.1: Map of California showing the locations of the Sierra Nevada foothills, the Klamath Mountains, and the distribution of gold districts. Map modified from Clark (1970).

Over 115 Moz of lode and placer gold has been produced from the deposits in California (Craig and Rimstidt, 1998). Approximately 35 Moz have come from lode gold deposits of the Sierra Nevada foothills. Placer gold derived from the erosion of these lode gold sources has also been historically important and represented a significant resource that initially led to the California gold rush. Substantial lode resources still exist in the gold camps of California. 1.2 Grass Valley Gold District, California The Grass Valley gold district represents the largest historic lode gold producer in the western cordillera of North America. More than 13 Moz of lode gold has been produced from Grass Valley at an average grade of 16.3 grams/tonne (Clark, 1984; Payne, 2000). The largest resource within Grass Valley was the Empire deposit and the second-most productive deposit was the Idaho-Maryland deposit. The gold-bearing quartz veins of Grass Valley fill second-, third-, and fourth-order faults to the east of the terrane-suturing Wolf Creek fault zone. The first-order Wolf Creek structure likely acted as the fluid conduit for hydrothermal fluid flow from a deeper source region. Two distinct vein groups exist within the study area: steeply dipping east-trending veins (E-W veins) in the north (Idaho-Maryland deposit) and north-trending veins with gentler dips averaging 35° (N-S veins) in the southern part of the district (Empire deposit). The E-W veins are found within third- and fourth-order faults connected at depth to the second-order dextral Weimer fault (Payne, 2000). The host rocks are metamorphosed mafic and ultramafic rocks of the Jura- Triassic arc belt that were accreted onto the margin of California in the Mesozoic (Snow and Scherer, 2006). The N-S veins to the south are hosted within both metamorphic rocks of the Jura-Triassic arc belt and within the Grass Valley granodiorite (also known as the Grass Valley pluton of Böhlke and Kistler, 1986, and the La Barr Meadows pluton of Ash, 2001). The surface manifestation of these N-S veins mirrors the geometry of the Grass Valley granodiorite margin, which is elongated in a north-south direction. The majority of the N-S veins dips toward the intrusion and pass from the granodiorite into the metamorphic country rocks with no disturbance to the vein shape, trend, or dip. 1.3 Previous Research Despite the historic economic significance of the Grass Valley district, little research has focused on the genesis of the gold-bearing veins, the ore-hosting Grass Valley granodiorite, and their relation to the tectonic evolution of the Sierra Nevada foothills and the Klamath Mountains.

Prior to this research, the Grass Valley granodiorite has been dated by three different methods, which all provided different results ranging from 110±5 Ma (He-whole rock: Urry and Johnston, 1936) to 126.7±3 Ma (K-Ar hornblende: Böhlke and Kistler, 1986) and ~159 Ma (U-Pb zircon: Irwin and Wooden, 2001). The data for the ~159 Ma age have recently been reinterpreted using updated software and now suggest a crystallization age of 164±2.3 Ma (Wooden, personal commun. 2009). The large variation in these ages makes it extremely difficult to interpret the tectonic history of this region and the relationship between igneous activity and gold mineralization. Previous workers believed that the two distinct orientations of the vein groups were simply the result of differences in host rock competency (e.g., Johnston, 1940). A second possible interpretation is that two distinct gold-forming events occurred within Grass Valley that contributed to it being the largest gold district in California. The E-W veins have previously been dated by three different methods (two-point Rb-Sr isochron quartz; K-Ar mica; 40Ar/39Ar mariposite), but only one date is considered reliable due to the method (152.2±1.2 Ma 40Ar/39Ar mariposite; Snow et al., 2008). The age of ~152 Ma for the E-W veins is significantly older than the ages of gold mineralization elsewhere in the Sierra Nevada foothills province, which cluster in a 20 Myr window between ~135 Ma and ~115 Ma (Marsh et al., 2008). The younger gold events are interpreted to be related to a shift from sinistral to dextral movement along the reactivated thrust fault systems, which represent the terrane-bounding sutures, at approximately 125 Ma (Marsh et al., 2008, and references therein). Beyond the seminal paper written by Johnston (1940), very little additional work has contributed to our understanding of the orogenic deposits in Grass Valley. A comprehensive study involving geochronology, geochemistry, and microanalytical work is necessary to understand the formation and evolution of Grass Valley, in particular its relationship to magmatism, metamorphism, and accretionary tectonics of California. 1.4 Thesis Objectives This thesis aims to constrain major magmatic, tectonic, and hydrothermal events within Grass Valley by dating the emplacement and cooling age of the ore-hosting granodiorite and the formation of the veins. These ages along with geothermobarometric calculations for the host rocks can then be utilized to build a baseline for subsequent tectonic events. Comparing the undeformed veins will also allow an interpretation of the paragenetic sequence and mechanisms

of vein formation. Lastly, additional geochronologic investigations of gold deposits in the Klamath Mountains will build a regional-scale framework for orogenic gold formation in California. 1.5 Thesis Organization Following this initial introductory chapter, this thesis is composed of four additional chapters. The three main chapters (Chapters 2-4) cover in-depth studies of these gold deposits. The final chapter (Chapter 5) concludes and summarizes the work accomplished for this thesis. Chapter 2 is a manuscript already published in Economic Geology (Taylor et al., 2015b). Chapter 3 will be submitted to Economic Geology. Chapter 4 will be submitted to Mineralium Deposita or Geological Society of America Bulletin. Numerous abstracts have been presented describing various aspects of this work and are given in the Appendix (Taylor et al., 2010a, b, 2011, 2012, 2013, 2014, 2015a). Chapter 2 provides geochronological and geochemical constraints on ore formation, host rock emplacement, and geothermobarometry of the intrusive event. The emplacement and cooling history of the Grass Valley granodiorite was examined. Zircon U-Pb ages provide the timing of pluton emplacement and 40Ar/39Ar ages of pristine hornblende and biotite bracket the cooling history of the Grass Valley granodiorite. Pressure and temperature calculations of the Grass Valley granodiorite utilizing the Ti-in-zircon geothermometer (Ferry and Watson, 2007), the plagioclase-amphibole geothermometer (Holland and Blundy, 1994), and the Al-in- hornblende geobarometer (Anderson and Smith, 1995) form a baseline for the pressure and temperature of emplacement, and therefore constraints on unroofing events related to orogenic gold formation. Hydrothermal xenotime and monazite were found as precipitates related to gold mineralization. Analyses from fifteen xenotime crystals were used to calculate a U-Pb age for vein formation. The geochemistry of the monazite and xenotime crystals were also characterized to prove their hydrothermal origin and that they are not xenocrysts or the products of magmatic- hydrothermal activity related to the emplacement of the Grass Valley granodiorite. Chapter 3 focuses on the description of the largely undeformed veins of Grass Valley and uses microanalytical and petrographic investigations to compare and contrast the vein sets. A paragenetic sequence is reconstructed. This information is used to formulate a sequence of events that is responsible for orogenic gold vein formation. Cathodoluminescence observations indicate multiple generations of quartz. Back-scattered electron images of pyrite reveal a

prolonged growth period that includes multiple growth bands and alteration of the pyrite interiors. Electron microprobe and laser ablation ICP-MS analyses of the pyrite crystals show relative enrichments and depletions in certain trace elements, which are interpreted to be related to gold forming stages. It is shown that gold precipitation is genetically related to galena, but not with quartz. The chemistry of the minerals related to gold formation indicates the importance of fluid reaction with the host rock during this event. Variations in sulfur isotope values reflect how the source of the fluid and gold has evolved over the history of hydrothermal activity in Grass Valley. The most important factors in the formation of the veins include pressure fluctuations related to fault movement and the interaction between the hydrothermal fluid and the immediate host rocks. Chapter 4 takes a larger-scale view of gold mineralization in California and its relationship with the tectonic activity that shaped the region during the Mesozoic. Ore and alteration samples were collected from numerous orogenic gold deposits throughout the Klamath Mountains. White mica was separated from the veins and the altered vein selvage. The mineral separates were analyzed by electron microprobe and determined that they represent near end member muscovite with minor celadonite and (or) biotite substitution. New 40Ar/39Ar analyses of white mica from samples representative of eight deposits show the disturbed nature of the argon spectra for some of the deposits, likely as a result of post-depositional deformation. Mica from samples of four deposits are undisturbed with no evidence of excess argon and provide plateau ages for the isotopic data. The ages bracket a ~20 Myr period of orogenic gold formation within the Klamath Mountains. The ages correlate with tectonic events and gold mineralization within the Sierra Nevada foothills, with certain deposits having identical ages of formation to that of Grass Valley. The ages determined also refute the idea that the gold ores of the Klamath Mountains are genetically related to magmatic events. 1.6 References

Anderson, J.L., and Smith, D.R., 1995, The effects of temperature and fO2 on the Al-in- hornblende barometer: American Mineralogist, v. 80, p. 549–559. Ash, C.H., 2001, Relationship between ophiolites and gold-quartz veins in the North American Cordillera: British Columbia Geological Survey, Geological Survey Branch Bulletin 108, 140 p. Böhlke, J.K., and Kistler, R.W., 1986, Rb-Sr, K-Ar, and stable isotope evidence for the ages and sources of fluid components of gold-bearing quartz veins in the northern Sierra Nevada Foothills metamorphic belt, California: ECONOMIC GEOLOGY, v. 81, p. 296–322.

Clark, W.B., 1970, Gold districts of California: California Division of Mines and Geology, Bulletin 193, 186 p. Clark, W.B., 1984, Gold mines of Grass Valley, Nevada County, California: California Geology, v. 37, p. 43–53. Craig, J.R., and Rimstidt, J.D., 1998, Gold production history of the United States: Ore Geology Reviews, v. 13, p. 407–464. Dickinson, W.R., 2008, Accretionary Mesozoic-Cenozoic expansion of the Cordilleran continental margin in California and adjacent Oregon: Geosphere, v. 4, p. 329–353. Ernst, W.G., Snow, C.A., and Scherer, H.H., 2008, Contrasting early and late Mesozoic petrotectonic evolution of northern California: Geological Society of America Bulletin, v. 120, p. 179–194. Ferry, J.M., and Watson, E.B., 2007, New thermodynamic models and revised calibrations for the Ti-in-zircon and Zr-in-rutile thermometers: Contributions to Mineralogy and Petrology, v. 154, p. 429–437. Goldfarb, R.J., Groves, D.I., and Gardoll, S., 2001, Orogenic gold and geologic time: A global synthesis: Ore Geology Reviews, v. 18, p. 1–75. Goldfarb, R.J., Baker, T., Dubé, B., Groves, D.I., Hart, C.J.R., and Gosselin, P., 2005, Distribution, character, and genesis of gold deposits in metamorphic terranes: ECONOMIC GEOLOGY 100TH ANNIVERSARY VOLUME, p. 407–450. Goldfarb, R.J., Hart, C.J.R., and Marsh, E.E., 2008, Orogenic gold and evolution of the Cordilleran orogen, in Spencer, J.E., and Titley, S.R., eds., Ores and orogenesis: Circum- Pacific tectonics, geologic evolution, and ore deposits: Arizona Geological Society Digest 22, p. 311–323. Groves, D.I., Condie, K.C., Goldfarb, R.J., Hronsky, J.M.A., and Veilreicher, R.M., 2005, 100th Anniversary Special Paper: Secular changes in global tectonic processes and their influence on the temporal distribution of gold-bearing mineral deposits: ECONOMIC GEOLOGY, v. 100, p. 203–224. Hacker, B.R., 1993, Evolution of the northern Sierra Nevada metamorphic belt: Petrological, structural, and Ar/Ar constraints: Geological Society of America Bulletin, v. 105, p. 637– 656. Holland, T., and Blundy, J., 1994, Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry: Contributions to Mineralogy and Petrology, v. 116, p. 433–447. Irwin, W.P., 2003, Correlation of the Klamath Mountains and the Sierra Nevada: U.S. Geological Survey Open-File Report 02-490. Irwin, W.P., and Wooden, J.L., 2001, Map showing plutons and accreted terranes of the Sierra Nevada, California, with a tabulation of U/Pb isotopic ages: U.S. Geological Survey, Open-File Report 01-229, 1 map. Johnston, W.D., Jr., 1940, The gold quartz veins of Grass Valley, California: U.S. Geological Survey Professional Paper 194, 101 p.

Marsh, E.E., Goldfarb, R.J., Kunk, M.J., Groves, D.I., Bierlein, F.P., and Creaser, R.A., 2008, New constraints on the timing of gold formation in the Sierra Foothills province, central California, in Spencer, J.E., and Titley, S.R., eds., Ores and orogenesis: Circum-Pacific tectonics, geologic evolution, and ore deposits: Arizona Geological Society Digest 22, p. 369–388. Payne, M., 2000, Geology of the Grass Valley mining district, Nevada County, California, in Shaddrick, D.R., ed., Platinum group elements, high grade gold and history: The Sierra Nevada 2000: Geological Society of Nevada Special Publication 32, p. 125–136. Phillips, G.N., and Powell, R., 2010, Formation of gold deposits: a metamorphic devolatilization model: Journal of Metamorphic Geology, v. 28, p. 689–718. Pitcairn, I.K., Teagle, D.A.H., Craw, D., Olivo, G.R., Kerrich, R., and Brewer, T.S., 2006, Sources of metals and fluids in orogenic gold deposits—Insights from the Otago and Alpine schists, New Zealand: ECONOMIC GEOLOGY, v. 101, p. 1525–1546. Saleeby, J.B, and Harper, G.D., 1993, Tectonic relations between the Galice Formation and the Condrey Mountain Schist, in Dunne, G.C., and McDougall, K.A., (eds.), Mesozoic paleogeography of the western United States—II: Pacific Section, Society for Sedimentary Geology, Book 71, p. 197–225. Schweickert, R.A., Armstron, R.L., and Harakal, J.E., 1980, Lawsonite blueschist in the northern Sierra Nevada, California: Geology, v. 8, p. 27–31. Sibson, R.H., 1990, Conditions for fault-valve behavior, in Knipe, R.J., and Rutter, E.H., eds., Deformation Mechanisms, Rheology and Tectonics: Geological Society Special Publication No. 54, p. 15–28. Snow, C.A., and Scherer, H., 2006, Terranes of the Western Sierra Nevada Foothills metamorphic belt, California: A critical review: International Geology Review, v. 48, p. 46–62. Snow, C.A., Bird, D.K., Metcalf, J., and McWilliams, M., 2008, Chronology of gold mineralization in the Sierra Nevada Foothills from 40Ar/39Ar dating of mariposite: International Geology Review, v. 50, p. 503–518. Taylor, R.D., Lee, J.P., Marsh, E.E., and Goldfarb, R.J., 2010a, Geochronologic constraints on the Grass Valley granodiorite, California, and relation to lode gold formation [abs.]: Geological Society of Nevada 2010 Symposium, Reno, Nevada, 2010, Abstract Volume, p. 85. Taylor, R.D., Marsh, E.E., Goldfarb, R.J., and Lee, J.P., 2010b, 40Ar/39Ar geochronology of the Grass Valley granodiorite, California, and relation to lode gold formation [ext. abs.], in Monecke, T., ed., The Challenge of Finding New Mineral Resources: Global Metallogeny, Innovative Exploration, and New Discoveries, Society of Economic Geology 2010 Keystone Conference, Keystone, Colorado, Extended Abstract and Poster, A-48. Taylor, R.D., Goldfarb, R.J., and Lee, J.P., 2011, Age constraints on the emplacement of the Grass Valley granodiorite, California, and relation to lode gold formation [abs.]: Geological Society of America Abstracts with Programs, Minneapolis, Minnesota, v. 43, no. 5, p. 470.

Taylor, R.D., Fletcher, I.R., Goldfarb, R.J., and Monecke, T., 2012, U-Th-Pb SHRIMP evidence for multiple gold-forming events in the Grass Valley gold district, California, USA [abs.]: Society of Economic Geologist 2012 Conference, Lima, Peru: Integrated Exploration and Ore Deposits, Poster 102. Taylor, R.D., Marsh, E.E., and Goldfarb, R.J., 2013, Mesozoic tectonics and geochronology of orogenic gold metallogeny in California [abs.]: Society of Economic Geology 2013 Conference, Whistler, British Columbia.: Geoscience for Discovery, Poster P2.32. Taylor, R.D., Marsh, E.E., and Goldfarb, R.J., 2014, Tectonic implications of 50 million years of gold-forming events in the Sierra foothills and Klamath Mountains, California [abs.]: Geological Society of America Abstracts with Programs, Vancouver, British Columbia, v. 46, no. 6, p. 250 Taylor, R.D., Goldfarb, R.J., and Monecke, T., 2015a, Geochronology and geochemical constraints on formation of the Grass Valley gold district, Sierra Nevada foothills province, California, USA [ext. abs.]: Mineral resources in a sustainable world, SGA 2015 Proceedings, Nancy, France, v. 1, p. 217–220. Taylor, R.D., Goldfarb, R.J., Monecke, T., Fletcher, I.R., Cosca, M.A., and Kelly, N.M., 2015b, Application of U-Th-Pb phosphate geochronology to young orogenic gold deposits: New age constraints on the formation of the Grass Valley gold district, Sierra Nevada Foothills province, California: ECONOMIC GEOLOGY, v. 110. p. 1313–1337. Urry, W.D., and Johnston, W.D., Jr., 1936, Age of the Sierra Nevada granodiorite: Geological Society of America Proceedings for 1935, p. 114. Wright, J.E., and Fahan, M.R., 1988, An expanded view of Jurassic orogenesis in the western United States Cordillera: Middle Jurassic (pre-Nevadan) regional metamorphism and thrust faulting within an active arc environment, Klamath Mountains, California: Geological Society of America Bulletin, v. 100, p. 859–876.

CHAPTER 2 APPLICATION OF U-TH-PB PHOSPHATE GEOCHRONOLOGY TO YOUNG OROGENIC GOLD DEPOSITS: NEW AGE CONSTRAINTS ON THE FORMATION OF THE GRASS VALLEY GOLD DISTRICT, SIERRA NEVADA FOOTHILLS PROVINCE, CALIFORNIA

The Grass Valley gold district within the northern Sierra Nevada foothills province represents the largest historic lode gold producer in the North American Cordillera. Gold- bearing quartz veins were discovered in the district in 1850 at Gold Hill, just two years after the first discovery of placer gold in Wolf Creek. From 1851 until the closure of the mines in 1956, more than 13 million ounces (Moz) of lode gold were produced from Grass Valley at an average grade of 16.3 grams/tonne (g/t; Clark, 1984; Payne, 2000). The oldest mine and most productive of these, the Empire deposit, produced 5.8 Moz of lode gold, and the second-most productive mine, the Idaho-Maryland deposit, produced 2.4 Moz (Pease, 2009). Ten other deposits in the Grass Valley district each yielded between 0.1 and 0.4 Moz and another 24 mines yielded between 10,000 and 100,000 oz of gold (Payne, 2000). A recent report on the Idaho-Maryland deposit defined an additional 470,000 oz of measured plus indicated gold resources and more than 1 Moz of inferred gold resources (Pease, 2009). Significant resources at the Empire deposit are also likely to still exist. For comparison, all of the lode gold deposits in the 190-km-long Mother Lode belt, farther south in the Sierra Nevada foothills province, are estimated to have cumulatively produced 12-13 Moz of lode gold (Landefeld, 1988; Payne, 2000). Despite its historic economic significance and substantial resource potential, little modern- day work has been published on the gold-bearing veins of the Grass Valley district and their relation to the local geology and regional tectonic evolution. The most comprehensive descriptions of the deposits, hosted within granodiorite and Jurassic-Triassic greenschist to amphibolite facies metamorphic rocks, were provided by Lindgren (1896) and Johnston (1940). Since these seminal papers on the geology of the ores, much of the published data regarding temporal relationships are inconsistent. The timing of gold formation in the district seems incompatible with previous interpretations of orogenic gold formation in the cordillera of California, particularly when viewed relative to the much better studied Mother Lode belt in the southern part of the Sierra Nevada foothills province (e.g., Marsh et al., 2008; Snow et al., 2008).

Discrepancies between the published ages of host rocks and auriferous quartz veins, as described by Böhlke and Kistler (1986), have prevented a comprehensive interpretation of the evolution of the Grass Valley district. The presence of two distinct vein sets, with different geometries and different host rocks, also poses the question as to whether they are products of multiple events, or formed as a result of differing host rock properties during a single prolonged hydrothermal event. Determining the age(s) of mineralization is critical in order to unravel the regional metallogeny at the district and orogen scale, which contributes to identification of the most favorable geologic environments for undiscovered deposits. To document the timing of gold formation and its relation to magmatism and the Mesozoic tectonic history of the Grass Valley gold district, detailed geochronological and geochemical studies were conducted on gold-bearing quartz veins and their granodiorite host rocks. Geochronologic investigations of the ore-hosting Grass Valley granodiorite by the U-Pb zircon and 40Ar/39Ar hornblende and biotite methods were employed to determine emplacement age and subsequent cooling history. Whole-rock geochemistry and mineral compositions were used to derive crystallization temperature, depth of emplacement, and the geotectonic environment of formation for the granodiorite. These parameters were used to establish a baseline for the tectonic evolution of the region during the Mesozoic and to compare the granodiorite with similarly aged, but gold-barren plutons elsewhere in the Sierra Nevada. The age of the north-south trending gold-bearing vein set was precisely measured using U-Pb methods on rare-earth element (REE)-bearing hydrothermal phosphate minerals, allowing comparison to the previously dated E-W veins. This study thus defines the timing of major magmatic and hydrothermal events in the Grass Valley district and uses these results to evaluate their geologic and tectonic setting. The data are synthesized to provide a more accurate view of the Mesozoic tectonic evolution of Grass Valley and its historically significant gold deposits. 2.1 Tectonic Setting of Grass Valley Grass Valley is located in the western cordillera of North America (Fig. 2.1), which underwent multiple orogenic events during the Mesozoic and Cenozoic, although its evolution extends back into the Paleozoic. Passive margin sedimentation dominated the development of the western margin of the North American craton from the Neoproterozoic through the Middle to Late Devonian (Dickinson, 2000, 2004, 2008). The middle Paleozoic Antler orogeny involved thrusting of these deep-water sedimentary rocks eastward onto the passive margin. The

Figure 2.1: Geologic map of northern California including the Sierra Nevada Range (modified from Irwin, 2003, and Ernst et al., 2008). In addition to the Sierra Nevada batholith, plutons that intruded across major faults are shown. BMF-Bear Mountains fault; CSFT-Calaveras–Shoo Fly Thrust; DF-Downieville fault; MF-Melones fault; RBF-Rich Bar fault; SF-Sonora fault; SPF- Spencerville fault; TT-Taylorsville fault; WCF-Wolf Creek fault.

subsequent Permian-Triassic Sonoma orogeny led to a major westward shift of the continental margin into what is now California and additional east-directed thrusting of allochthons. Accretion of allochthonous, near-shore Devonian-Permian island arcs during the Antler and Sonoma orogenies formed the Northern Sierra, Central Metamorphic, and Eastern Klamath terranes in the farthest inboard part of the Klamath-Sierra Nevada arc (Dickinson, 2000, 2004, 2008). Subsequent arc magmatism and terrane accretion during the Mesozoic dominated the later evolution of the active continental margin.

Devonian through Late Jurassic autochthonous and allochthonous terranes of the Sierra Nevada foothills were accreted to the western margin of the North Sierra terrane beginning in the Middle Triassic and were sutured to North America during complex transpressional and transtensional processes (Dickinson, 2008; Ernst et al., 2008). Terrane accretion is thought to have been complete by the Late Jurassic (Sharp, 1988), or significantly (see below) by ca. 160 Ma. Because of the geologic complexities that may mask contacts and the shingling of similar but smaller terranes, many of the rocks of the Sierra Nevada foothills have been defined as tectonostratigraphic belts (Ernst et al., 2008), rather than as distinct terranes. The Calaveras complex, accreted to the North Sierra terrane, is a mélange belt dominated by Permian and Triassic chert, argillite, and greenstone. The Melones fault zone separates the complex from the Jura-Triassic arc belt (Snow and Scherer, 2006), which includes a basement of Paleozoic mélange and serpentinite overlain by Mesozoic arc volcanic rocks. These units have been interpreted as oceanic arcs and related subduction complexes (Dickinson, 2008). The Jura- Triassic arc belt is bounded along its western edge by the Bear Mountains fault in the south and Wolf Creek fault in the north. To the west of these faults, the Upper Jurassic accretionary sequence is dominated by fine-grained oceanic sedimentary rocks. The arc belt and accretionary complex are also referred to as the central and western belts, respectively (Day and Bickford, 2004). During the Late Jurassic through Early Cretaceous, oblique convergence along the continental margin of California led to widespread folding, thrusting, and sinistral slip (Glazner, 1991; Umhoefer, 2003). Subsequently, sinistral movement along the terrane-bounding faults of the Sierra Nevada foothills switched to dextral motion at approximately 125 Ma as a result of major plate reorganization in the Pacific basin (Goldfarb et al., 2008). The tectonic reversal may be important for the gold mineralization in the Mother Lode of the southern Sierra Nevada foothills between 135 and 115 Ma (Goldfarb et al., 2008; Marsh et al., 2008). Plutonism within the Sierra Nevada is concentrated in two episodes, mainly between ~170- 140 Ma and ~120-80 Ma, separated by a ~20 million year magmatic lull (Glazner, 1991; Irwin and Wooden, 2001). The older episode formed a semi-continuous volcanic-plutonic arc that spanned the formerly contiguous Sierra Nevada and Klamath Mountains (Ernst, 2013). The younger episode, with peak magmatic activity between 100 and 85 Ma, was a major contributor

to formation of the Sierra Nevada batholith, but there was no associated magmatism in the Klamath Mountains. This tectonic grain of central California is manifest today as NNW-trending parallel belts of terranes that progressively young to the west. Within the Sierra Nevada foothills, these belts are truncated to the east and south by the Sierra Nevada batholith. They are bordered on the west and north by the sedimentary rocks of the Great Valley Group and other Cenozoic sedimentary rocks. 2.2 Geology of the Grass Valley District The Grass Valley gold district is located in the Jura-Triassic arc belt (Fig. 2.2), and consists of Late Triassic–Early Jurassic (ca. 200 Ma) submarine metasedimentary and metavolcanic arc rocks overlying late Paleozoic (ca. 300 Ma) ophiolitic basement rocks (Snow and Scherer, 2006). The belt is interpreted to have formed as an offshore oceanic terrane assembled on older mafic and ultramafic basement rocks (Ernst et al., 2008). The Lake Combie complex, which hosts the Grass Valley district, is probably a ca. 200 Ma mafic arc that is part of the terrane (Edelman et al., 1989; Fagan et al., 2001). These rocks were variably metamorphosed from lower greenschist to amphibolite facies within the Sierra Nevada foothills during and after their accretion to the continental margin between ~200 and 160 Ma, the approximate age of the volcanic arc rocks overlying the Lake Combie complex (Bickford and Day, 1988; Saleeby et al., 1989) and the age of the unmetamorphosed granitic intrusion (this study) emplaced into the deformed and metamorphosed rocks. The 40Ar/39Ar ages of relict amphibole from a lowermost greenschist facies metavolcanic unit suggest a second period of arc volcanism in the Jura-Triassic arc belt at ca. 170 Ma (Fagan et al., 2001), but this second period of volcanism has not been recognized in other parts of the Jura-Triassic arc belt. Fagan et al. (2001) interpret these amphibole crystals to be of volcanic origin, based upon their texture and chemical composition, which suggests rapid cooling during the early portion of their thermal history and subsequent metamorphism that did not reach temperatures high enough to reset the argon systematics. If this age does correspond to a period of volcanism and is not a metamorphic cooling age or if the rocks were faulted into position, then the period of peak regional metamorphism of the Jura-Triassic arc belt is constrained between ~170 and 160 Ma in order to account for the metamorphism of the 170 Ma volcanic rocks. Initiation of subduction and the

Figure 2.2: Geologic map of the Grass Valley gold district (modified from Johnston, 1940, and Saucedo and Wagner, 1992) showing the locations of the two most productive mines, sample sites, distribution of major ore veins, and distribution of various bedrock units. WFZ-Weimar fault zone; GVF-Grass Valley fault; MV-Maryland vein; SHV-Spring Hill vein.

attendant thermal input from magmatism at ~170 Ma supports that peak metamorphism occurred between ~170 and 160 Ma. The Spring Hill tectonic mélange, consisting of an assemblage of tectonic blocks in a serpentinized ultramafic matrix (Payne, 2000), is hosted within the Lake Combie complex. These tectonic blocks consist of metavolcanic rocks, ultramafic rocks, gabbro, and minor metasedimentary rocks that range from 0.1 m to nearly 2.5 by 1 km in size. The matrix is interpreted as an upper mantle harzburgite that intruded under high pressure, low temperature conditions as a cool but ductile mass that entrained fragments from the enclosing Lake Combie complex (Payne, 2000; Pease, 2009). The steeply east-dipping Wolf Creek fault zone (Day et al., 1985) is only a few km west of many of the gold veins, and separates the Lake Combie complex from the Smartville complex of the Upper Jurassic accretionary sequence to the west. The subparallel, second order Weimar and Grass Valley faults, which bound an exposed block of serpentinite and ultramafic-mafic rocks that hosts the Idaho-Maryland deposit, are a few km to the east of the Wolf Creek fault zone. The northern strand of the Melones fault zone is located about 25 km east of Grass Valley. To the north, the 140 Ma Bald Rock pluton cross cuts the Wolf Creek fault zone (Fig. 2.1; Irwin and Wooden, 2001), thus providing a minimum age for displacement along the regional fault system associated with the Grass Valley district. The Bear Mountains fault zone, which is the southern continuation of the Wolf Creek fault zone, may have been active from ~160 to 123 Ma, as indicated by numerous geochronologic determinations including those for crosscutting plutons, their correlative volcanic rocks, deformation, and metamorphism (Miller and Paterson, 1991). Tuminas (1983) determined that the regional Wolf Creek fault zone and the local Weimar fault in Grass Valley were both active in the Late Jurassic, approximately between 160 and 145 Ma as indicated by crosscutting relationships of plutons. In contrast, Day and Bickford (2004) interpreted geochronological data for plutonic rocks to imply displacement along the Wolf Creek fault zone has been less than 1-2 km since approximately 160 Ma and that the Jura-Triassic arc belt and the Upper Jurassic accretionary sequence have been juxtaposed since at least the end of the Middle Jurassic. Reverse movement on the Bear Mountains and Wolf Creek fault zones was likely limited to several km and the extent of strike-slip movement is uncertain due to the lack of known offset markers (Miller and Paterson, 1991; Albino, 1992).

A conspicuous geologic unit in the district is the Late Jurassic Grass Valley granodiorite, which has also been referred to as the La Barr Meadows pluton (Tuminas, 1983) and the Grass Valley pluton (Böhlke and Kistler, 1986). The Grass Valley granodiorite is immediately east of the steeply dipping Wolf Creek fault zone and intrudes massive diabase and metavolcanic rocks of the Lake Combie complex. The pluton is elongate in the N-S direction and has an irregular dumbbell shape approximately 8-9 km long and ranging from less than 1 km to more than 3 km wide. The granodiorite is a medium-grained, leucocratic intrusion composed of predominantly plagioclase (~45%), quartz (~20%), potassium feldspar (~15%), hornblende (~15%), and biotite (~3%) with trace apatite, zircon, titanite, and magnetite in the unaltered form (Fig. 2.3). Locally, quartz and potassium feldspar intergrowths produced micrographic textures. The entire pluton was variably altered by post-crystallization hydrothermal activity. Biotite and hornblende vary from pristine to entirely chloritized. All of the feldspar crystals are at least partially altered to white mica, clay, and carbonate minerals; most are almost entirely replaced. Additional secondary minerals include disseminated pyrite and epidote locally along fractures. The pluton has not been affected by metamorphism and there is no macroscopically or microscopically notable late fabric. Rounded to subangular oblate mafic enclaves are scattered throughout the Grass Valley granodiorite and range from a few centimeters to nearly a meter in diameter. These enclaves are fine-grained, melanocratic, and contain a mineral assemblage similar to that of the granodiorite, although the enclaves contain a greater proportion of plagioclase, hornblende, and apatite. Contacts between the enclaves and the surrounding granodiorite are sharp and the enclave perimeters include a chilled margin of finer-grained crystals. The shape and texture of the enclaves indicate liquid-liquid interaction with the surrounding granodiorite. 2.3 Geology of the Grass Valley Deposits Gold-bearing quartz veins fill second-, third-, and fourth-order faults about 2-5 kilometers east of the Wolf Creek fault zone. Two distinct vein groups exist within the Grass Valley district: steeply dipping east-west trending veins (E-W veins) in the northern and generally north-south trending veins with gentler dips averaging 35° (N-S veins) in the southern part of the district (Fig. 2.2). The most important E-W veins are associated with the Idaho-Maryland, Brunswick, and Spring Hill deposits. Significant veins of the N-S vein group include those of the Empire, North Star, and W.Y.O.D. deposits. The N-S veins hosted within and immediately

Figure 2.3: Photographs of typical granodiorite and vein samples from the Grass Valley district. (A) Example of characteristically medium-grained Grass Valley granodiorite. (B) Empire vein exposed in the underground tourist adit within the Empire Mine State Historic Park. Both the vein and the altered wall rock are oxidized, producing the brown and black coloring.

adjacent to the Grass Valley granodiorite accounted for roughly 70% of the historic lode gold production in Grass Valley at an average grade of 19.1 g/t, whereas the E-W veins had a slightly lower average grade of 13.1 g/t (Payne, 2000). The E-W veins are in tightly bunched third- and fourth-order faults that are connected at depth to the second-order dextral Weimar fault zone, and are hosted mostly within the Spring Hill tectonic mélange (Payne, 2000; Pease, 2009). These veins occur between the Grass Valley fault and the Weimar fault zone along lithologic contacts of tectonic blocks in the Mélange unit. The E-W veins are mostly composed of quartz, calcite, and ankerite. Pyrite is the dominant sulfide mineral (1-2%). Hydrothermal xenotime and monazite are absent as accessory minerals, but mariposite is present. Locally abundant scheelite and telluride minerals have been historically recovered as sources of W and Au, respectively. The N-S veins are hosted in both greenschist-facies metamorphosed diabase of the Jura- Triassic arc belt and in the Grass Valley granodiorite. The surface orientation of these conjugate N-S veins mirrors the geometry of the Grass Valley granodiorite margin. The majority of the N- S veins dips toward the intrusion and pass from the metamorphic country rocks into the granodiorite with little to no disturbance to the vein shape, trend, or dip (Johnston, 1940). However, some of the veins dip away from the intrusion; they pitch upwards from the metamorphosed diabase into the granodiorite.

The mineralogy of the N-S veins is similar to that of the E-W veins, but host rock chemical differences result in distinct trace hydrothermal phases within the two vein sets. Similar to the E-W veins, pyrite is the most abundant sulfide mineral and generally occurs in higher concentrations (2-3%) than in the E-W veins, with additional arsenopyrite, galena, and chalcopyrite. Galena and electrum form fracture fills within pyrite. Gold occurs as free grains and as inclusions within pyrite. Minor to trace quantities of hydrothermal xenotime and monazite are found within quartz or hydrothermal white mica and are spatially related to sulfide minerals and wall rock slivers (Fig. 2.4). Hydrothermal zircon is a very rare accessory phase. Both vein sets have extraordinary vertical and lateral persistence; individual veins extend for kilometers. Descriptions of the vein sets from numerous historic underground exposures have long emphasized the obvious ductile-brittle nature of the mineralization (e.g., Lindgren, 1896; Howe, 1924; Farmin, 1938, 1941). Elongate wall rock slivers are incorporated into the veins parallel to subparallel to the vein margin. These fault-fill laminated veins are hosted within minor thrust faults. All veins in the district contain some fault gouge (Johnston, 1940; Pease, 2009). 2.4 Previous Geochronological Research Modern research focused on the Grass Valley gold district is limited, and the few published geochronological data are difficult to reconcile with known geological relationships. The ore- hosting Grass Valley granodiorite had been previously dated by three different methods, yielding three different results. The earliest estimated age came from a mafic enclave and the immediately surrounding granodiorite using the “helium method” giving an age of 110±5 Ma (Urry and Johnston, 1936); the applicability of this method for determining absolute ages was challenged soon thereafter (e.g., Johnston, 1940). Böhlke and Kistler (1986) employed the K-Ar geochronological method on hornblende from the granodiorite and reported an age of 126.7±3 Ma, although they noted this result as problematic because it was younger than their estimated age for contained gold-bearing veins (~141 Ma). Another estimate of ~159 Ma was made from U-Pb zircon analysis of four grains (Irwin and Wooden, 2001), an age that has been revised to 164±2.3 Ma using improved data reduction procedures (J. Vazquez, written commun., 2010). Estimates of the time of gold vein formation were derived from an E-W vein in the Brunswick deposit for which a two point Rb-Sr isochron for a micaceous quartz vein sample

Figure 2.4: Backscattered electron images of hydrothermal monazite and xenotime crystals within gold-bearing quartz veins of the Empire mine. (A) Xenotime (xen) entirely surrounded by quartz (qtz). (B) Monazite (mon) with quartz and hydrothermal biotite (bt). (C) Xenotime with small overgrowth of monazite. (D) Monazite and xenotime crystals hosted in quartz.

yielded an age of 140.9±3 Ma and a K-Ar age from unspecified mica gave 143.7±3 Ma (Böhlke and Kistler, 1986). More recently, a 40Ar/39Ar plateau age for mariposite in an E-W vein of the Idaho-Maryland deposit yielded a more robust age of 152.2±1.2 Ma (Snow et al., 2008). Because no age estimates existed for the more abundant and economically more significant N-S veins, Marsh et al. (2008) attempted to date pyrite from the N-S Empire vein by the Re-Os technique. However, the analyzed pyrite only contained 1.07 ppb Re, a concentration too low to produce precise data. Although traces of white mica are also present in these veins, the grains have proven to be too fine for absolute age determination by 40Ar/39Ar methods. Marsh et al. (2008) and Snow et al. (2008) dated other gold deposits within the Sierra Nevada foothills from the Alleghany district in the north through the Mother Lode belt and the Bagby district to the south. These analyses employed 40Ar/39Ar geochronological methods for

hydrothermal mariposite and the majority of determined ages cluster between 135 and 115 Ma. It is important to note that unlike the Grass Valley district, these deposits are associated with the regional Melones fault system to the east and not the Bear Mountains–Wolf Creek fault system. 2.5 Materials and Methods Initially, representative samples were collected from the Grass Valley granodiorite at multiple locations along the length of the pluton both distal and proximal to the ore zones. Mineralized quartz material from the N-S vein set was collected underground at the Empire deposit in the Empire Mine State Park. 2.5.1 Sampling and Petrographic Investigations Thin sections of variably altered granodiorite samples and one mafic enclave were prepared for petrographic analysis and examined using a standard optical microscope. Doubly polished thick sections (100 µm) of vein material were examined by optical microscopy. All sections were studied by scanning electron microscopy using a JEOL JSM-5800LV scanning microscope at the U.S. Geological Survey in Denver, Colorado, to further characterize textural relationships at small scales. In addition, back-scattered electron (BSE) imaging was used to locate and study accessory minerals such as monazite and xenotime. 2.5.2 Whole-Rock Geochemistry The whole-rock composition of variably altered granodiorite samples and a mafic enclave were determined at the U.S. Geological Survey in Denver, Colorado. The concentrations of the major elements were measured by X-ray fluorescence spectrometry (XRF) using a Bruker S8 Tiger spectrometer following preparation of standard glass disks by fusion of sample powder with a lithium metaborate lithium tetraborate flux. The detection limits of the XRF were 0.01% for all major elements. Sample loss on ignition was determined by gravimetry. The concentrations of trace elements were measured by a combination of inductively coupled plasma-atomic emission spectrometry (ICP-AES) and inductively coupled plasma-mass spectrometry (ICP-MS). Powdered samples were decomposed using a sodium peroxide sinter at

450°C and then leached with water and acidifed with HNO3. Prior to sample aspiration, tartaric acid was added. The ICP-AES analyses were conducted on a Perkin Elmer Optima spectrometer and the ICP-MS analyses were performed on a Perkin Elmer Elan spectrometer. The precision and accuracy of the major and trace element determinations by XRF, ICP- AES, and ICP-MS were monitored using the georeference material GSP-2. Repeated analysis

showed that the precision was typically better than 5% for elements occurring at concentrations significantly above their respective detection limits. Close agreement between the analytical data and the reference values suggests that the element determinations were also highly accurate. 2.5.3 Electron Microprobe Analysis of Monazite and Xenotime The composition of monazite and xenotime in the N-S veins, as well as in the Grass Valley granodiorite, was determined using the U.S. Geological Survey (Denver, Colorado) JEOL 8900 Electron Microprobe with five wavelength dispersive analyzers. Operating conditions for the analysis of monazite and xenotime were 20 keV accelerating voltage, a 50 nA current (measured on the Faraday cup), and a focused electron beam. Standards used include MG1 xenotime, BS1 xenotime, 44069 monazite, synthetic REE glasses from the University of Oregon, Taylor metals, and Smithsonian orthophosphates. The following elements were analyzed: Al (123), Si (187), P (226), Ca (156), Y (781), La (364), Ce (339), Pr (1002), Nd (335), Sm (1165), Eu (342), Gd (1057), Tb (1038), Dy (420), Ho (1115), Er (380), Tm (442), Yb (420), Lu (427), Th (1234), U (1158) Pb (1035), As (256), and Sc (192); average 99% confidence detection limits in parentheses are in elemental parts per million and are based upon a total of 107 analyses of samples and standards. 2.5.4 U-Pb Zircon Geochronology Zircon crystals from the Grass Valley granodiorite were physically separated through standard magnetic and density separation techniques before hand-picking and emplacement into an epoxy mount along with the age standard zircon R33 (419 Ma; Black et al., 2004). Prior to analysis, all zircon crystals from Grass Valley were studied by cathodoluminescence (CL) and BSE imaging using the JEOL JSM-5800LV scanning microscope at the U.S. Geological Survey (Denver, Colorado). Geochronological data, trace elements, and Ti concentrations for zircon geothermometry were collected on the SHRIMP RG at the Stanford–U.S. Geological Survey Micro Analysis Center in Palo Alto, California, under standard operating conditions. The compositional standards MAD, SL13, and 91500 were used from an in-house mount. Temperatures obtained from the zircon crystals were determined using the Ti-in-zircon thermometer of Ferry and Watson (2007). Uncertainties associated with age determinations are 1 sigma at 95% confidence and incorporates the uncertainty associated with the analyzed age standard.

2.5.5 U-Pb Geochronology of Monazite and Xenotime Xenotime and monazite of sufficient size (>10 µm in diameter) were drilled out from doubly polished thick sections of vein material into ~2 mm plugs and then cast into 25 mm epoxy mounts along with standards. All mounts were analyzed using the SHRIMP II at the John De Laeter Centre for Isotope Research at Curtin University, Perth, Australia. The SHRIMP analytical and data reduction procedures for xenotime and monazite are discussed in detail by Fletcher et al. (2004, 2010). The primary calibration standards were MG-1 xenotime and z2234 monazite. Trace element concentrations of individually analyzed phosphates grains were considered because matrix effects on the ionization efficiencies of secondary ionic species of U, Th, and Pb can be substantial (Fletcher et al., 2004, 2010). The U and Th abundances used in matrix corrections were derived from the SHRIMP data. The REE abundances used in matrix corrections for xenotime, and Y in monazite are from electron microprobe analyses made adjacent to SHRIMP analytical spots. The mass resolution (1% definition) was >5000 in both ion microprobe analytical sessions. – The small areas of exposed inclusion-free sample required the use of narrow O2 primary ion beams; ~0.2 nA in a spot diameter ≤10 µm for xenotime and ~0.3 nA in a ~12 µm spot for monazite. 2.5.6 40Ar/39Ar Geochronology Hornblende and biotite were analyzed by 40Ar/39Ar methods at the U.S. Geological Survey (Denver, Colorado). Purified mineral grains of unaltered biotite and hornblende were separated from crushed and coarsely milled rock samples that were washed in distilled water and acetone. Individual mineral grains, 1.5 to 2.5 mm in diameter, were hand-picked to ensure purity and were inspected under a microscope to confirm that they were not visibly altered. These, together with grains of the 40Ar/39Ar age standard Fish Canyon Tuff sanidine (applying an age of 28.201 Ma; Kuiper et al., 2008) were irradiated at the U.S. Geological Survey’s TRIGA reactor in

Denver, Colorado. Following irradiation, samples were incrementally heated using a 30W CO2 laser equipped with a homogenizing lens and analyzed using a Mass Analyzer Products 215-50 mass spectrometer. Mass spectrometric analyses were performed by peak hopping using a single electron multiplier operated in analog mode. Correction factors for nucleogenic interferences during

24 irradiation in the TRIGA reactor were determined from irradiated CaF2 and zero-age K-glass. Raw data were corrected for blanks, radioactive decay, and nucleogenic interferences. 2.5.7 Hornblende Geobarometry and Plagioclase-Amphibole Geothermometry Compositions of unaltered hornblende and plagioclase crystals from the Grass Valley granodiorite were determined by electron microprobe analysis, using the facilities described above, for Al-in-hornblende geobarometry and plagioclase-amphibole thermometry. Analysis spots were located at the core and near the rim of hornblende crystals. Three randomly located spots within several crystals were analyzed and averaged to yield plagioclase compositions. Operating conditions included an accelerating voltage of 15 keV and a current of 30 nA (measured on the Faraday cup). 2.6 Results Geochronological data was collected at the U.S. Geological Survey (Denver, Colorado), Curtin University (Perth, Australia), and the Stanford–U.S. Geological Survey Micro Analysis Center (Palo Alto, California). The remainder of the data were collected at the U.S. Geological Survey (Denver, Colorado). 2.6.1 Petrography and Geochemistry of the Grass Valley Granodiorite Data for six samples of the Grass Valley granodiorite provide basic characterization of this rock unit. Sample GVGD-1, collected farthest from gold-bearing quartz veins, is the least altered. Sample GVGD-3 is granodiorite within 10 cm of a mafic enclave within the granodiorite (sample GVGD-3e) and is notably more altered than other granodiorite samples; it also contains several <1 cm mafic enclaves. Samples GVGD-3, -4, and -5 were collected in the northern part of the pluton in close proximity to numerous gold-bearing quartz veins. All whole rock samples are at least partially affected by hydrothermal alteration. Carbonate and sericitic alteration of feldspars and chlorite alteration of biotite and hornblende is common. Much of the granodiorite near the sample sites for GVGD-3 and -3e had significant secondary epidote on fracture surfaces. Sample GVGD-1 contains some unaltered feldspars, whereas feldspars in samples GVGD-4 and GVGD-5 have been nearly completely altered to white mica and carbonate and clay minerals. Sample GVGD-1 has a higher proportion of biotite to hornblende compared to GVGD-4 and GVGD-5. Hornblende is preferentially chloritized in GVGD-1, whereas biotite is preferentially chloritized in samples GVGD-4 and GVGD-5.

Whole-rock geochemical analyses of the Grass Valley granodiorite samples indicate that

these are calc-alkaline igneous rocks (Table 2.1). The SiO2 abundances range from 63.5 to 65.4

wt.%, with sample GVGD-3 being more silica-rich (SiO2=69.7 wt.%). The mafic enclave

sample GVGD-3e is relatively silica-poor (SiO2=61.7 wt.%) and more calcic (CaO=6.01 wt.%) than the granodiorite samples. Chondrite-normalized REE patterns for the Grass Valley granodiorite have moderate negative Eu anomalies (Fig. 2.5). In contrast, the mafic enclave sample has a much flatter overall REE pattern with a more pronounced negative Eu anomaly than the granodiorite samples; the pattern for sample GVGD-3, which hosts the enclave, is similar to that of the enclave but has a smaller Eu anomaly. All analyzed samples have subtly concave-upward middle REE–heavy REE (MREE-HREE) patterns centered on Er. 2.6.2 Age of the Grass Valley Granodiorite A population of 23 zircon crystals from sample GVGD-1 was analyzed to provide a more robust and accurate estimate of ore host rock age (Table 2.2). Cathodoluminescence imaging revealed that most of these zircon crystals have dark euhedral cores and lighter, thin, concentrically zoned rims (Fig. 2.6). All analytical spots were located in the rims to most closely approximate the solidification age of the granodiorite. During the analytical session, 20 analyses of the R33 zircon standard provided a 207Pb- corrected 206Pb/238U weighted average age of 418.9±2.1 Ma (206Pb/238U=0.06716, 419 Ma; Black et al., 2004). Three of the 23 Grass Valley granodiorite zircon spot analyses were excluded because trace element results indicate that the analytical spot impinged on an inclusion within the zircon. The ISOPLOT data reduction (Ludwig, 2012) suggested that ages for the two oldest zircons are statistical outliers; however, no microscopic or geochemical evidence supports their rejection because they appear identical to all other analyzed grains and do not include an inherited component or preserve growth zoning laminations different from that characteristic of other grains. Hence, these data were included in the age determination. Common Pb abundances are negligible. Uranium concentrations are low (57-176 ppm), which suggests radiation damage and concomitant Pb loss are also insignificant. The U-Pb zircon data yield a 207Pb-corrected 206Pb/238U weighted average age for 20 spots of 159.9 ± 2.2 Ma (Fig. 2.7). The 207Pb-corrected ages are preferred for rocks of this age because they are more precise for geologically young zircon samples and discordance of Phanerozoic crystals in individual analyses is typically not detectable within the limits of

Figure 2.5: Chondrite-normalized REE plots for Grass Valley granodiorite sample data. Chondrite values from McDonough and Sun (1995).

Figure 2.6: Cathodoluminescence images of zircon crystals from the Grass Valley granodiorite. White circles represent spot locations for SHRIMP analyses. Zircons lacking a white circle were not analyzed.

Figure 2.7: U-Pb isotope data for zircon from the Grass Valley granodiorite. (A) Tera- Wasserburg Concordia diagram of SHRIMP U-Pb data for 20 zircon analyses. Uncertainty ellipses are 1 sigma. (B) Linearized probability plot for analyzed zircon grains also showing the 207Pb-corrected 206Pb/238U weighted average age. Uncertainty bars are 1 sigma. analytical uncertainty (Ireland and Williams, 2003). The Tera-Wasserburg Concordia plot indicates that Pb loss was insignificant (Fig. 2.7). In addition, it is noteworthy that the corrected 206Pb/238U weighted average ages for 204Pb-corrected, 207Pb-corrected, and 208Pb-corrected ages are statistically indistinguishable.

The ages of other mineral phases in three samples of Grass Valley granodiorite were also determined by 40Ar/39Ar geochronology (Table 2.3, Fig. 2.8). Pristine biotite crystals were separated from sample GVGD-1. Analysis of this biotite yielded a 40Ar/39Ar plateau age of 161.9±1.4 Ma (2 sigma uncertainty) that is interpreted as the time when the southern part of the pluton cooled through the biotite closure temperature (~300°C; Harrison et al., 1985). Unaltered hornblende from samples GVGD-4 and GVGD-5 yielded 40Ar/39Ar plateau ages of 159.7±0.6 and 161.9±0.7 Ma, respectively. These ages are interpreted as the time when the northern part of the pluton cooled through the hornblende closure temperature (~500-600°C; Harrison, 1981). 2.6.3 Geochemistry of Vein-Hosted Monazite and Xenotime Analyzed monazite and xenotime grains are from granodiorite-hosted quartz veins in the Grass Valley district. Monazite from the N-S veins (Table 2.4) has lower concentrations of Th and U than magmatic monazite, between 0.01 and 0.2 wt.% Th (Fig. 2.9a). Vein-hosted monazite has REE patterns with small negative Ce anomalies, but lacks Eu anomalies (Fig. 2.9b). Vein-hosted xenotime also lacks a significant Eu anomaly (Table 2.5, Fig. 2.9c). Magmatic monazites from the Grass Valley granodiorite have a much higher U and Th content than vein-hosted monazites, as much as 11.6 wt.% Th (Table 2.6, Fig. 2.9a). Their REE patterns are gently negatively sloped in the light REE (LREE) to MREE (Pr-Dy) range. These monazite grains are significantly Eu depleted as Eu concentrations are below the electron microprobe detection limit (Fig. 2.9b). No xenotime of magmatic origin was found. 2.6.4 Age of Vein Formation The age of the Empire deposit was determined by analysis of 15 vein-hosted xenotime grains (Table 2.7, Fig. 2.10). Two analyses were disregarded because of very high common Pb contents (>15%). One of these samples has low U, which is of unknown significance, and the other gives a 206Pb/238U age entirely consistent with the main data group, which supports the validity of the common Pb corrections in the other analyses. One analysis yielded an extreme 206 238 (~6σ) outlier. The main group of Pb/ U results (12 analyses from 11 crystals) has MSWD = 1.2, consistent with a single age population and the data are concordant, as well as can be assessed given the poor precision in 207Pb/206Pb and the young age of the samples (Fig. 2.10). The weighted average 206Pb/238U age, with its uncertainty augmented by the uncertainty in the averaged data for the reference standard (1.4%, 95% confidence) and a nominal uncertainty in

Figure 2.8: 40Ar/39Ar age spectra and Ca/K plots for hornblende and biotite from the Grass Valley granodiorite. Height of the age-step rectangles represents 2-sigma analysis uncertainties.

Figure 2.9: Trace element characteristics of phosphate minerals in the Grass Valley gold district. (A) Concentrations of Th versus Th/U for hydrothermal and igneous monazite from Grass Valley. All data obtained by electron microprobe analysis. (B) Chondrite-normalized REE plots for hydrothermal and igneous monazite from Grass Valley. Elemental values below the electron microprobe detection limits are not shown. (C) Chondrite-normalized REE patterns of hydrothermal xenotime from Grass Valley. Chondrite values from McDonough and Sun (1995).

Figure 2.10: U-Pb isotope data for xenotime from the Empire vein in Grass Valley (A) Tera- Wasserburg Concordia diagram of SHRIMP U-Pb data for 12 xenotime analyses. Arrows indicate data omitted from the age determination. Uncertainty ellipses are 1 sigma. (B) Linearized probability plot of the main subset of SHRIMP 206Pb/238U dates from hydrothermal xenotime from Grass Valley also showing the weighted average age. Uncertainty bars are 1 sigma. Excluded data shown in (A) and in Table 2.7 are not displayed. matrix corrections equivalent to one quarter of the average correction (1.1%), is 162 ± 5 Ma (approximately 95% confidence). This age is indistinguishable from that of the host Grass Valley granodiorite and older than any previous estimates for the age of gold deposition in the Sierra Nevada foothills province. Considering all possible analytical uncertainties, the gold

deposit formed no more than approximately 5 million years after crystallization of the granodiorite. Results for monazite from the N-S Empire vein did not provide precise chronologic data. All analyses have low Th and U, with correspondingly low 208Pb and 206Pb. The 204Pb/206Pb values were too high for reliable U–Pb geochronology because of high common Pb levels in the hydrothermal monazite. Four grains were analyzed, and subsequent analyses were discontinued due to high common Pb and low U and Th concentrations. Only a single monazite analysis indicated less than 5% common Pb; this analysis yielded a 208Pb/232Th age within uncertainty of the xenotime data (Table 2.8). Due to the unreliability of the U-Th-Pb data for these monazites, they are not further considered in this investigation. 2.6.5 Geothermometry and Geobarometry Crystallization temperatures for the Grass Valley granodiorite were calculated using the Ti- in-zircon geothermometer (Ferry and Watson, 2007). Titanium concentrations measured concurrently with SHRIMP U-Pb analyses constrain the emplacement temperature of the Grass Valley granodiorite and help establish ore formation within the regional cooling history (Table

2.2). The Grass Valley granodiorite contains titanite and Ti-bearing magnetite, so a TiO2 activity of 0.7 is assumed (e.g., Claiborne et al., 2006). As the granodiorite is quartz bearing, the associated magma was silica saturated, and a SiO2 activity of 1.0 can be assumed. Total Ti was calculated from the SHRIMP-derived concentration of 48Ti, given that 48Ti constitutes 73.72% of total Ti. Using the calibrations of Ferry and Watson (2007), calculated temperatures for individual zircons ranged from 759 to 831°C with an average temperature for 20 crystals of

793±22°C (1 sigma). Using a lower TiO2 activity of 0.5 increases the average temperature by approximately 35°C. Plagioclase crystals in the Grass Valley granodiorite are andesine in composition with an orthoclase endmember component of <4%. The three averaged compositions for each analyzed plagioclase grains are indistinguishable. The anorthite component of the 10 analyzed plagioclase crystals ranged from 37.6 to 51.1%, with an average of 44.4% (Table 2.9). The primary mineral assemblage of the Grass Valley granodiorite includes quartz, potassium feldspar, plagioclase, hornblende, biotite, titanite, and Ti-bearing magnetite. This mineral assemblage provides the complete buffering assemblage required by the Al-in-

hornblende geobarometer. Analyses of Grass Valley amphibole cores and rims indicate that these crystals are calcic amphiboles and, when ferric iron is empirically determined, the majority is composed of magnesiohornblende (Table 2.10). Estimates of pluton emplacement pressures and depths were calculated using the Al-in-hornblende geobarometer of Anderson and Smith (1995), with Al content being normalized to 13 cations (Cosca et al., 1991). Anderson and Smith (1995) recommend that the Al-in-hornblende geobarometer is applicable to amphiboles whose compositions have Fe/(Fe+Mg) <0.65. Given that the Grass Valley amphiboles have Fe/(Fe+Mg) ≤0.5, their compositions are suitable for application of the barometer. Anderson and Smith (1995) further emphasize the importance of a temperature correction for Al-in-hornblende

geobarometry. For this reason, the plagioclase-amphibole temperatures calculated using the thermometer of Holland and Blundy (1994) were included in the pressure calculations because this is the best complement for the Al-in-hornblende geobarometer (Anderson et al., 2008). [6] Analyses of crystals with Al#>0.21 (Al#= Al/AlT) were rejected (e.g., Ridolfi et al., 2010). Five additional hornblende analyses were rejected because calculated pressures are negative and unrealistic. Temperatures derived from the plagioclase-amphibole thermometer range from 787 to 851°C, with an average of 809±15°C, for amphibole cores, and range from 750 to 817°C, with an average of 797±16°C, for the rims. These temperatures are in excellent agreement with temperatures derived from the Ti-in-zircon method. Accounting for the plagioclase-amphibole temperatures, calculated pressures range from 0.4 to 1.8 kbar for the amphibole cores and 0.3 to 1.2 kbar for the amphibole rims; associated uncertainty is ±0.6 kbar. Using the conversion of 0.27 kbar/km for lithostatic pressure in the upper crust, these values provide average solidification depths of 3.4±0.2 and 2.7±0.2 km for the cores and rims, respectively. These relatively shallow depth estimates are consistent with the reported hypabyssal nature of many of the igneous rocks in this part of the northern Sierra and the coeval ca. 160 Ma volcanism immediately to the west of the Wolf Creek fault system (e.g., Day and Bickford, 2004). Importantly, such data also provide a maximum depth estimate for gold formation hosted in the Grass Valley intrusion. 2.7 Discussion The U-Th-Pb method has been successfully used to determine the age of hydrothermal phosphate minerals and therefore, the age of orogenic gold ores in the Grass Valley district of California. Results indicate that gold ore formation is associated with two distinct hydrothermal events at relatively shallow crustal depths. The first event occurred soon after crystallization of the granodiorite that hosts much of the resource for the N-S veins, while the second event (Snow et al., 2008) resulted in the formation of the historically less significant E-W veins. 2.7.1 Use of Hydrothermal Phosphate Minerals for Determining Absolute Age of Phanerozoic Ores Unique chemical compositions distinguish hydrothermal phosphate minerals from those with other origins, such as igneous, metamorphic, or diagenetic (Kositcin et al., 2003; Schandl and Gorton, 2004; Lowers et al., 2008). The chemistry of vein-hosted phosphates in this study

led to their conclusive classification as hydrothermal in origin, which is an important confirmation considering their similar age to that of the host rock. Characterization based on textures, optical microscopy, and scanning electron microscopy suggests that the vein-hosted monazite and xenotime are hydrothermal in origin, but these characteristics do not preclude a xenocrystic origin. All vein-hosted phosphate minerals analyzed in this study are contained in samples that also include minerals clearly of hydrothermal origin, including native gold and sulfide minerals. Xenotime and monazite grains that occur within quartz veins are typically <10 to 20 µm in diameter and inclusion free (Fig. 2.4). These phosphate grains are located within quartz, both along fractures and in unfractured sites within the two-dimensional plane of thin sections. They are consistently proximal to wall rock slivers incorporated in the quartz veins, but are commonly entirely surrounded by quartz. Crystal forms range from euhedral to anhedral. Some grains preserve subtle concentric zoning patterns visible in backscattered electron (BSE) images that are not truncated along crystal edges, which suggests that crystals have not been broken by transportation within the veins, but rather grew in situ. Furthermore, our detailed examination of the Grass Valley granodiorite suggests it contains no xenotime crystals, which precludes vein-hosted xenotime crystals having been derived from the intrusive rock. Rare earth-element abundances for vein-hosted xenotime and monazite from the Grass Valley gold district are distinct relative to those for the igneous monazite from the Grass Valley granodiorite (Tables 2.4 to 2.6, Fig. 2.9b-c). Hydrothermal xenotime has notably lower U and Th concentrations than published results for igneous xenotime (Kositcin et al., 2003) just as hydrothermal monazite from Grass Valley has much lower U and Th concentrations than magmatic monazite from the Grass Valley granodiorite. None of the hydrothermal phosphate minerals are relatively depleted in Eu, whereas all igneous or other phosphate minerals that crystallize in equilibrium with plagioclase have pronounced negative Eu anomalies. Most hydrothermal phosphate minerals do not have negative Eu anomalies because, unlike magmatic plagioclase that preferentially incorporates Eu, non-magmatic hydrothermal systems do not involve plagioclase crystallization. Overall REE abundances are broadly similar among hydrothermal phosphate minerals from a number of different orogenic gold deposits, but local country rock REE contributions can subtly influence these abundances (Kositcin et al., 2003).

The Spring Hill tectonic mélange of serpentinized ultramafic matrix and enclosed metavolcanic blocks hosts the E-W veins and these units have low mobile P and REE abundances. Consequently, as hydrothermal fluids reacted with these ore-host rocks, the opportunity to extract wall rock P and REE was limited and thus hydrothermal phosphates are not present in the younger auriferous veins. In contrast, hydrothermal fluids responsible for veins hosted by calc-alkaline to alkaline igneous rocks have a significant opportunity to extract P and REE from the intrusions because of the elevated background concentrations for these elements. Dating of hydrothermal monazite (Th-Pb) and xenotime (U-Pb) grains has been successfully conducted in geochronologic investigations of a number of Precambrian orogenic gold deposits (e.g., Vielreicher et al., 2003; Salier et al., 2005; Rasmussen et al., 2006; Sarma et al., 2008). These geochronological methods for hydrothermal phosphates have not previously been applied to orogenic gold deposits as young as those in California; very few Phanerozoic ore deposits of any type have been dated using phosphate minerals (e.g., Kempe et al., 2008; Li et al., 2011). Closure temperatures of monazite and xenotime for the U-Th-Pb systems are much higher than the temperature of hydrothermal fluids that form orogenic gold deposits; consequently, geochronology using these systems should provide an unequivocal age for vein formation. Experiments by Cherniak et al. (2004) indicate a closure temperature for Pb in monazite in excess of 900°C. Other studies indicate that U-Pb systematics of monazite are preserved at temperatures up to at least ~700-750°C (Copeland et al., 1988; Parrish, 1990). Cherniak (2006) also demonstrated that diffusion of Pb in xenotime is exceptionally slow and that closure temperatures are similar to those of monazite and zircon. Therefore, hydrothermal monazite and xenotime have significant potential to constrain the timing of gold-forming events. 2.7.2 Timing of Gold Deposit Formation Gold-bearing quartz veins of the Grass Valley gold district were formed in the already rapidly uplifted and partially eroded terranes of the northern Sierra Nevada foothills, as is evidenced by the depth of emplacement of the Grass Valley granodiorite relative to depths and pressures (roughly greater than 7.5 km depth) necessary for the metamorphic facies of the country rocks. An initial gold deposit forming event quickly followed the intrusion, cooling, and faulting of the Grass Valley granodiorite. This event must have occurred no later than 157 Ma,

the youngest permissive xenotime U-Pb age for the N-S veins, and no earlier than the 162 Ma maximum age of granite host solidification. Approximately 5-10 million years later (e.g. Snow et al., 2008), a second major hydrothermal event created another set of gold-bearing veins. The N-S veins at Grass Valley constitute the oldest known orogenic gold veins in California; their formation is coincident with movement along the Bear Mountains–Wolf Creek fault system. Faulting along this regional fault zone and (or) the subsidiary Grass Valley fault and Weimar fault zone must also have been ongoing at 152 Ma when the associated structures served as conduits for hydrothermal fluid flow and the precipitation of the E-W veins between the Weimer and Grass Valley second-order faults. Orientations of veins and faults within specific rock types can be used to indicate prevailing stress regimes during vein formation. The N-S veins are mainly fault-fill veins that dip shallowly to both the east and west. Their strike parallels the elongate Grass Valley granodiorite, and many of them are also approximately parallel to the adjacent Wolf Creek fault zone. These conjugate veins are hosted within a competent rock unit, the Grass Valley granodiorite, and have been interpreted as having formed during E-W to ENE-WSW compression (Hodgson, 1989; Bierlein et al., 2008). Some of these veins preserve evidence of reverse displacement and one of the veins with the largest reverse offset (a few meters) contains mullions and grooves that strike 040°, indicating the direction of movement (Johnston, 1940). The shallow dip of the conjugate fault-fill veins (e.g., Robert and Poulsen, 2001) and the interpretation that these veins formed in a suprahydrostatic fluid pressure regime resulting in self-sealing fault valve behavior (Sibson, 1990) are consistent with a compressive tectonic regime. Nevertheless, given the probable overlap of vein formation with at least the end stage of Middle Jurassic regional uplift (see below), some sort of transpressional component was likely required during the hydrothermal event. It is more difficult to determine the stress regime prevailing during formation of the E-W veins solely from their orientations because these veins preferentially formed along lithologic contacts involving rocks with pronounced differences in rock competency; however, the veins within this set are interpreted to be oblique or extensional in origin (Payne, 2000). 2.7.3 Relationship Between Magmatism and Gold Mineralization Geochronology, major and trace element geochemistry, and isotope data, suggest that magmas responsible for the plutons of the northern Sierra Nevada were derived by melting mafic

and juvenile accreted terranes in the lower crust at depths of >35 km with residual amphibole and garnet in the source region (Cecil et al., 2012). Although Cecil et al. (2012) did not study the Grass Valley granodiorite, they did interpret data for other coeval granodioritic plutons that are located within approximately 10 km of Grass Valley. The major element geochemistry of the Grass Valley granodiorite is similar to that of these and other plutons in the northern Sierra Nevada, except that it is has more Fe, Mg, and a higher Mg# compared to most plutons of similar

SiO2 content (see Cecil et al., 2012, and references therein). Its chondrite-normalized REE patterns are less negatively sloped with a greater spread in values for HREE compared to LREE, but still within the same range of values (Fig. 2.5). The flatter REE profile resulting from slight decreases in LREE and increases in HREE could be the result of chloritization of biotite and amphibole caused by hydrothermal alteration of the granodiorite (Alderton et al., 1980). The Sr and Y contents of the Grass Valley granodiorite are consistent with compositions of normal, intermediate arc magmas (Drummond and Defant, 1990). The geochemical composition of the granodiorite is that of a calc-alkaline, metaluminous volcanic arc granite. Its age and chemistry are commensurate with early stage Sierra Nevada arc magmatism and are similar to those of other arc-related northern Sierra Nevada Jurassic plutons. As such, the Grass Valley granodiorite is not compositionally or genetically unique among plutons of the Sierra Nevada foothills. The approximate overlapping ages between the Grass Valley granodiorite and the older gold event could be taken as evidence for a possible genetic association. However, the chemical signatures of the hydrothermal phosphate minerals in the N-S veins not only preclude them from being xenocrystic, but also from being magmatic-hydrothermal in origin. The composition and geochemical signature of magmatic-hydrothermal fluids and the minerals precipitated from those fluids are mainly controlled by processes operating at the magmatic stage prior to fluid exsolution (e.g., Audétat et al., 2008). Consequently, the REE signature of this magmatic- hydrothermal fluid is inherited from the parental magma from which the fluid exsolved (Reed et al., 2000) and should similarly be recorded in minerals precipitated from the fluid. For example, Pettke et al. (2005) indicated that hydrothermal zircons from the magmatic-hydrothermal Sn-W mineralized Mole Granite, Australia, have nearly identical REE profiles and negative Eu anomalies as late magmatic zircons, and Schaltegger et al. (2005) noted that hydrothermal

monazite and xenotime grains from the same deposit possess large negative Eu anomalies that reflect earlier plagioclase fractionation in the melt. The Grass Valley granodiorite is characterized by a negative Eu anomaly (Fig. 2.5), indicating that the source of the magma had residual plagioclase that acted as a sink for Eu or that early forming magmatic plagioclase was removed from the system. Later during magmatic evolution, the intrusion crystallized more plagioclase that further removed Eu from the remaining melt. Magmatic monazite displays a prominent negative Eu anomaly (Fig. 2.9b) because there was so little Eu remaining in the system that could be incorporated into its crystal structure. However, neither the hydrothermal monazite, nor xenotime display Eu anomalies (Fig. 2.9b-c), as would be expected in a progressively evolving magmatic-hydrothermal system equilibrated with a magma that has had Eu continuously partitioning into plagioclase. This is suggestive that the gold-forming hydrothermal fluids in Grass Valley were not sourced from the Grass Valley magmatic system. 2.7.4 Relationship Between Exhumation and Gold Mineralization Geochronological and geothermobarometric investigations indicate that the Grass Valley granodiorite was emplaced at temperatures of approximately 800°C within approximately 3 km of the surface and rapidly cooled to near the ambient temperatures (below 300°C) of the surrounding country rocks. Structural characteristics of the gold-bearing quartz veins, with incorporated wall rock slivers and breccia fragments reflects ductile to brittle deformation, which indicates that the pluton was solidified prior to vein formation. However, because the hydrothermal xenotime age also overlaps the emplacement and cooling ages of the pluton, hydrothermal activity quickly followed shallow magma emplacement and solidification. The age of volcanogenic debris in the Upper Jurassic accretionary sequence indicates that the rate of arc- related sedimentation was at a maximum between ca. 175 and 160 Ma in the western Sierra Nevada foothills (Ernst, 2011). Consequently, the Grass Valley granodiorite was certainly intruded as the orogen was being unroofed. Crystal sizes of the granodiorite are consistent with emplacement at shallow crustal depths and its elongate nature parallel to the Wolf Creek fault zone suggests emplacement in an anisotropic stress field. It is doubtful that significant exhumation near Grass Valley continued much after ca.160 Ma magmatism and gold deposition. The lack of deformation of the granodiorite indicates any significant compressional to transpressional event had to be waning. Textures and mineral

assemblage characteristics of the N-S and E-W veins in Grass Valley are similar, which suggests formation at approximately the same depths. If so, then the Grass Valley region apparently experienced only minor exhumation between about 160 and 150 Ma, the approximate age of the N-S and E-W vein systems, respectively. In fact, detrital zircon geochronologic data for arc- related sedimentation suggest a very limited final record of unroofing between ca. 158 Ma and the ca. 150 Ma ultimate cessation of deposition of clastic material in the Upper Jurassic accretionary sequence (Ernst, 2011). The subsequent Great Valley sedimentation to the west shows intrusive material being shed from the Jurassic arc at ca. 145-140 Ma (Ernst, 2011). This interval may represent the time when the Grass Valley granodiorite and both sets of auriferous quartz veins were exhumed to the near-surface. Consequently, the first gold episode occurred near the end of a period of rapid exhumation, whereas the second took place during a time of limited exhumation. 2.7.5 Middle-Late Jurassic Tectonic Regimes Controlling Gold Formation The first gold mineralization event, represented by the ca. 160 Ma shallowly dipping, N-S fault-fill veins, correlates with the onset of greater compressive stresses along the North American margin. For the period between 163 and 118 Ma relative velocities between the Farallon (Pacific basin) and North American plates are well constrained (Engebretson et al., 1985). Normal and tangential plate velocities along the continental margin trench indicate that the largest shift in tangential velocity between the two plates occurred between the Middle Jurassic and middle Cretaceous, with a major decrease in the southerly motion of the Farallon plate taking place at ca. 160 Ma (Engebretson et al., 1985), coincident with a major change in the absolute motion of the North American plate (J2 cusp; Beck and Housen, 2003). Any relative increase in compressive stress accompanying oblique plate subduction may have been a potential trigger for fluid overpressuring, hydraulic fracturing, and ductile-brittle vein formation (e.g., Sibson, 2004) along the margins of the slightly older and relatively competent Grass Valley pluton. This change in far-field stress at 160 Ma affected the terrane bounding faults of the Sierra Nevada and coincides with initial movement on the Bear Mountains–Wolf Creek fault zones (Miller and Paterson, 1991). The second gold-vein forming event was coeval with transcurrent activity on the Wolf Creek and Bear Mountains fault zones, and initiation of additional deformation along the Bear Mountains fault zone (Tobisch et al., 1989). Syntectonic emplacement and deformation of the

Guadalupe igneous complex and the Hornitos pluton in the southern part of the Sierra Nevada foothills along the Bear Mountains fault zone took place at ca. 151 Ma and resulted in mylonitically deformed plutons within the fault zone (Vernon et al., 1989). Sinistrally sheared syntectonic dike swarms were also emplaced between 155 and 148 Ma in the Owens Mountain area of the southwestern foothills terrane, probably along an extension of the Bear Mountains and Wolf Creek fault zones (Wolf and Saleeby, 1992). The change to more transcurrent deformation in the Late Jurassic probably reflects a change in the far-field stress regime at this time and may be responsible for steep E-W gold-bearing vein formation. Characteristics of ore-hosting veins in Grass Valley are consistent with their classification as orogenic gold deposits, but these veins are somewhat unusual in having formed at epizonal depths, but with mineral assemblages typical of most mesozonal orogenic gold deposits (e.g., Groves et al., 1998; Goldfarb et al., 2005). This suggests a very high Late Jurassic geothermal gradient within the shallow crust of the northern Sierra foothills, coincident with oblique subduction of paleo-Pacific oceanic lithosphere underneath California from 170 Ma until at least 140 Ma (Ernst et al., 2008; Ernst, 2011). Although the initial N-S veins are broadly coeval with magmatic activity, the geochemistry of the hydrothermal phosphate minerals precludes them from being magmatic-hydrothermal in origin; instead the fluids responsible for vein formation originated from an external source. Local magmatism is absent during development of the younger set of E-W veins, also dismissing their development from a magmatic-hydrothermal source. It is unimaginable that a hydrothermal system could be maintained for more than 5-10 million years by such a volumetrically minor intrusion as the Grass Valley granodiorite to form the E-W veins, particularly as numerical modeling of magmatic-hydrothermal systems suggests that temperatures greater than 200°C are not sustained for more than 800,000 years by a large, single intrusive event (Cathles et al., 1997). Furthermore, the laminated texture of the more ductile veins is more consistent with episodic fluid overpressuring events, which are likely the product of a regional flow event along the deep-crustal Wolf Creek fault system, than with fluid exsolution from a shallow level equigranular granite. 2.8 Conclusions The U-Pb geochronological method was successfully applied to Mesozoic-aged hydrothermal xenotime from the Grass Valley district, historically the most important orogenic lode gold district in the North American Cordillera. Combined textural and geochemical

characterization of vein-hosted monazite and xenotime conclusively established that they are of hydrothermal origin. For example, in contrast to phosphate minerals of igneous or magmatic- hydrothermal origin, the ore-related phosphate minerals have notably low concentrations of U and Th, and distinct REE profiles that lack negative Eu anomalies. Chemical characteristics of igneous or magmatic-hydrothermal phosphate minerals are determined by the chemical evolution of the magmatic system. This would have resulted in a significant Eu anomaly if these hydrothermal phosphate minerals were the product of a magmatic system, as is noted for the Grass Valley magmatic monazite. Hydrothermal activity in Grass Valley commenced at ca. 160 Ma in the already rapidly exhumed greenschist-facies rocks of the Jura-Triassic arc belt and immediately following intrusion, cooling, and faulting of the ore-hosting Grass Valley granodiorite; formation of the N- S veins is coeval with initial movement on the regional Wolf Creek fault zone in a compressional tectonic regime. A second hydrothermal event, 5-10 million years younger, formed the historically less important E-W veins temporally independent of magmatism and developed in a more extensional to transcurrent tectonic regime. The geochemistry of the ore-hosting Grass Valley granodiorite is consistent with it being a product of arc magmatism. The location of the intrusion, its geochemistry, and its competency contrast with the surrounding country rocks made its margin an appropriate structural trap for gold mineralization. Geochemical analysis of zircon, hornblende, and plagioclase from the Grass Valley granodiorite indicate that it was emplaced at elevated temperatures (~800°C) within approximately 3 km of the paleosurface, where it rapidly cooled to below 300°C. Overlapping emplacement, cooling, and hydrothermal ages indicate significant thermal activity in the Grass Valley area during the Late Jurassic, coincident with the end of a period of extensive regional exhumation. The 5-10 million years younger second set of veins formed at a paleodepth that was not much shallower than the first vein set, but this gold event was temporally and spatially unrelated to magmatism or exhumation. Post-mineralization exhumation of both vein sets is likely coeval with Great Valley sedimentation in the Early Cretaceous. The older set of orogenic gold veins in Grass Valley were formed during a compressive regime as relative plate motions between the Farallon and North American plates changed.

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CHAPTER 3 PARAGENETIC EVOLUTION AND FORMATION MECHANISMS OF OROGENIC GOLD DEPOSITS DETERMINED BY MICROANALYTICAL GEOCHEMISTRY AND PETROGRAPHY: NEW PERSPECTIVES FROM THE GRASS VALLEY DISTRICT, CALIFORNIA

The formation of orogenic gold deposits is related to the dehydration of crustal rocks in orogenic systems undergoing prograde metamorphism (e.g., Goldfarb et al., 1991; Phillips and Powell, 2010). The deposits are commonly located near the base of the continental seismogenic regime and the brittle-ductile transition at mesozonal depths of 6-12 km, but can be found to have formed at depths ranging from a couple km to over 15 km (e.g., Groves et al., 1998). Mineral deposition within the fault-hosted deposits is thought to take place as a result of extreme pressure fluctuations associated with major seismic events (Sibson et al., 1988; Cox et al., 2001). The textures of gold-bearing veins in orogenic gold deposits typically record complex processes of deformation caused by repeated seismogenic fault failure during and after mineralization (Fig. 3.1). Quartz, which is the main gangue mineral in these deposits, is particularly susceptible to grain-scale deformation processes as this mineral is mechanically weakened under hydrothermal conditions (Griggs and Blacic, 1965; Luan and Paterson, 1992; Post et al., 1996). Characteristic quartz textures that can be observed in gold-bearing quartz- carbonate veins include patchy or sweeping undulose extinction, deformation lamellae, mechanical Dauphiné twinning, bulging recrystallization, subgrain rotation recrystallization, grain boundary migration, and recrystallization resulting in granoblastic polygonal fabrics (e.g., Graupner et al., 2000). Due to recrystallization of quartz and other vein minerals, paragenetic relationships in lode gold veins are generally not well understood. The lack of primary textures also makes the chemical processes of gold precipitation difficult to ascertain. The present paper documents the occurrence of primary textures in quartz-carbonate veins from the Grass Valley district in California. Gold mineralization in this district took place during the Middle to Late Jurassic (Snow et al., 2008; Taylor et al., 2015) at crustal depth of less than 3 km (Taylor et al., 2015). Based on a combination of optical petrography, optical cathodoluminescence microscopy, scanning electron microscopy, and electron microprobe analysis, the paragenetic relationships for minerals within the two main vein sets in the district

Figure 3.1: Photomicrographs depicting typical deformation and recrystallization textures of orogenic gold vein minerals from deposits in California. (A) Cross-polarized photomicrograph of grain boundary migration and bulging, representing recrystallization of quartz crystal boundaries in a thick section from the Eureka deposit, Mother Lode belt. Trails of secondary fluid inclusions can also be seen to cut across the crystal boundary. (B) Cross-polarized photomicrograph of grain boundary migration and bulging, representing recrystallization of quartz crystal boundaries in a thick section from the Oxford deposit, Downieville district. (C) Cross-polarized light photomicrograph of deformed calcite with bent cleavage from the Washington deposit, French Gulch-Deadwood district, Klamath Mountains. (D) Reflected light photomicrograph of deformed galena with bent cleavage from the Washington deposit, French Gulch-Deadwood district, Klamath Mountains. (E) Back-scattered electron image of pyrite from the Harvard deposit, Mother Lode belt with no chemical zoning. (F) Cross-polarized photomicrograph of brecciated quartz vein in a thick section from the Sixteen to One deposit, Alleghany district. (G) Cross-polarized light photomicrograph of deformation lamellae in quartz from the Washington deposit, French Gulch-Deadwood district, Klamath Mountains. (H) Cross- polarized photomicrograph of undulose extinction in quartz from the Schroeder deposit, Yreka- Fort Jones district, Klamath Mountains.

are established. The obtained paragenetic sequence was used to reconstruct the processes resulting in vein formation and gold precipitation. Through a combination of microanalytical and isotopic studies on pyrite, additional important constraints on the chemical processes of gold deposition are derived. 3.1 Regional Geology The western cordillera in California consists of multiple accretionary terranes incorporated sequentially into the North American continental margin during the Mesozoic. Several NNW-

trending parallel belts of terranes that progressively young from east to west are recognized (Fig. 3.2). The western margin of the North American craton was the site of passive margin sedimentation from the Neoproterozoic through the Middle to Late Devonian, until the middle Paleozoic Antler orogeny thrusted deep-water sedimentary rocks eastward onto the passive margin (Dickinson, 2000, 2004, 2008). Subsequently, the Permian-Triassic Sonoma orogeny was coupled with initial east-directed thrusting of near-shore oceanic arcs and a major westward shift of the continental margin into what is now California. These two orogenies emplaced the Northern Sierra, Central Metamorphic, and Eastern Klamath terranes as the farthest inboard segments of the Klamath Mountains-Sierra Nevada by accretion of near-shore Devonian- Permian island arcs to the margin of North America (Dickinson, 2000, 2004, 2008). Continuing through final terrane amalgamation, likely by ca. 160 Ma (Sharp, 1988; Taylor et al., 2015), Devonian through Late Jurassic autochthonous and allochthonous terranes were accreted to the western margin of North America by complex transpressional and transtensional processes (Dickinson, 2008; Ernst et al., 2008). The host rocks of the orogenic gold veins were metamorphosed from lower greenschist through amphibolite facies during and subsequent to terrane amalgamation. Plutonism within the Klamath Mountains and the Sierra Nevada is focused within two distinct episodes, between approximately 170-140 Ma and 120-80 Ma (Glazner, 1991; Irwin and Wooden, 2001). Magmatic activity between the two periods was rare. Plutons of the older period are found within both the Klamath and the Sierra Nevada ranges, whereas plutons from the younger period are exclusively located in the Sierra Nevada as the Sierra Nevada batholith that intruded and truncated the eastern margin of the accreted terranes of the Sierra Nevada foothills. Emplacement of the Sierra Nevada batholith occurred mostly during peak activity in the younger episode, between 100-85 Ma (Ducea, 2001). Oblique convergence along the cordillera led to widespread folding, thrusting, and sinistral slip along the terrane-bounding faults during the Late Jurassic through Early Cretaceous (Glazner, 1991; Umhoefer, 2003). Major plate reorganization in the Pacific Basin at approximately 125 Ma resulted in a switch from sinistral to dextral movement along these regional faults (Goldfarb et al., 2008).

Figure 3.2: Geologic map of northern California including the Sierra Nevada and the Klamath Mountains (modified from Irwin, 2003, and Ernst et al., 2008). Locations of deposits mentioned in this manuscript included for reference.

Formation of orogenic gold deposits in California can be related to movement along the regional terrane-bounding faults (Goldfarb et al., 2008). Most of the Au-bearing quartz- carbonate vein deposits are located in secondary and tertiary faults associated with the larger first-order terrane-bounding faults. About 35 Moz of lode gold and more than 65 Moz of placer

gold has been produced from the deposits within the accreted terranes of California, with Grass Valley representing the historically most productive lode gold district (Table 3.1). 3.2 Geology of the Grass Valley Gold District The Grass Valley gold district represents the largest historic lode gold producer in the western cordillera of North America (Fig. 3.3). Grass Valley is located in the Jura-Triassic arc belt of the Sierra Nevada foothills, which consists of late Paleozoic (ca. 300 Ma) ophiolitic basement rocks underlying Late Triassic-Early Jurassic (ca. 200 Ma) submarine metasedimentary and metavolcanic arc rocks (Snow and Scherer, 2006). This belt is interpreted to be an allochthonous terrane that formed as an offshore arc assembled on older mafic and ultramafic basement rocks (Ernst et al., 2008). The Lake Combie complex within the Jura- Triassic arc belt hosts the gold district and is part of the ca. 200 Ma mafic arc (Edelman et al., 1989; Fagan et al., 2001). These rocks were variably metamorphosed to lower greenschist and amphibolite facies during and after their accretion to the continental margin between ~200 and 160 Ma (Bickford and Day, 1988; Saleeby et al., 1989). Peak metamorphism likely occurred between ~170 and 160 Ma (Fagan et al., 2001; Taylor et al., 2015). The steeply east-dipping Wolf Creek fault zone (Day et al., 1985) is located only a few kilometers to the west of many of the gold-bearing veins, separating the Jura-Triassic arc belt from the Smartville complex of the Upper Jurassic accretionary sequence to the west. The 162-160 Ma Grass Valley granodiorite (Taylor et al., 2015) is located immediately east of the Wolf Creek fault zone. It intruded into diabase and metavolcanic rocks of the Lake Combie complex. The elongate dumbbell-shaped pluton trends N-S for 8-9 km and ranges from less than 1 km to more than 3 km in width. The medium-grained, leucocratic granodiorite is composed of predominantly plagioclase (~45%), quartz (~20%), potassium feldspar (~15%), hornblende (~15%), and biotite (~3%) with trace apatite, zircon, titanite, monazite, and magnetite. However, the entire pluton has been variably altered by hydrothermal activity. Biotite and hornblende are commonly chloritized, whereas feldspar crystals are altered to clay and carbonate minerals as well as white mica. The granodiorite has not been affected by metamorphism and there is no macroscopically or microscopically late fabric. The Spring Hill tectonic mélange is located to the north of the Grass Valley granodiorite and consists of tectonic blocks of metavolcanic rocks, ultramafic rocks, gabbro, and minor metasedimentary rocks within a serpentinized ultramafic matrix hosted within the Lake Combie

Figure 3.3: Geologic map of the Grass Valley gold district (modified from Johnston, 1940; Saucedo and Wagner, 1992) showing the locations of the two vein sets, mines that were sampled in this study, and distribution of various bedrock units. Abbreviations: GVF = Grass Valley fault, WFZ = Weimar fault zone.

Complex (Payne, 2000). The tectonic blocks range in size from 0.1 m to nearly 2.5 by 1 km in size and represent portions of the surrounding Lake Combie complex that were entrained within the serpentinite matrix as it intruded under high pressure, low temperature conditions as a cool but ductile mass (Payne, 2000; Pease, 2009). Two distinctive vein sets occur within the Grass Valley district, namely north-trending veins with gentler dips averaging 35° (N-S veins) in the southern portion and steeply dipping east-trending veins (E-W veins) in the northern part of the district. The most productive mine in the district was developed on the Empire deposit, which produced 5.8 Moz of lode gold from N- S veins. The second-most productive mine was at the Idaho-Maryland deposit which produced 2.4 Moz of lode gold from E-W veins (Pease, 2009). The N-S veins produced at an average grade of 19.1 g/t, whereas the E-W veins produced at a slightly lower grade of 13.1 g/t (Payne, 2000). The N-S veins are hosted in both diabase of the Lake Combie Complex that has been metamorphosed to greenschist facies conditions and in the Grass Valley granodiorite. The surface expression of these conjugate N-S veins mirrors the geometry of the Grass Valley granodiorite margin. Most of the N-S veins dip toward the intrusion and pass from the metamorphic country rocks into the granodiorite with little to no disturbance to the vein shape or orientation (Johnston, 1940). However, some of the veins dip away from the intrusion such that they pitch upwards from the greenschist facies country rocks into the granodiorite. These conjugate veins have been interpreted to have formed during E-W to ENE-WSW compression (Hodgson, 1989; Bierlein et al., 2008); however, the probable overlap of vein formation with the end stages of Middle Jurassic regional uplift suggests the hydrothermal activity occurred in a transpressional setting (Taylor et al., 2015).

The majority of the mineralogy of the N-S veins is similar to that of the E-W veins. The bulk of the veins are composed of quartz, calcite, and ankerite. Sericite and chlorite are minor vein components, and mariposite is absent. Pyrite is the most abundant sulfide mineral and occurs in slightly higher concentrations (2-3%) than in the E-W veins. Arsenian pyrite, galena, chalcopyrite, and sphalerite are found as additional sulfide phases. Microscopic (~10-15 µm) crystals of monazite and xenotime are spatially associated with sulfide minerals and wall rock slivers within the veins, and the xenotime provided a 206Pb/238U age for vein formation of 162 ± 5 Ma (Taylor et al., 2015). Carbonate-sericite-pyrite alteration of wall rock surrounding the veins is comparatively minor compared to the E-W vein set and gold enrichment is restricted to the veins as free gold. The E-W veins in the northern part of the district formed in densely clustered third- and fourth-order faults that are connected at depth to the second-order dextral Weimar fault zone, and are hosted predominantly within the Spring Hill tectonic mélange (Payne, 2000; Pease, 2009). These veins occur between the Grass Valley fault and the Weimar fault zone along lithologic contacts of various tectonic blocks within the Mélange unit and are interpreted to be oblique or extensional in origin (Payne, 2000). The bulk of the E-W veins are composed of quartz, calcite, and ankerite. Additional silicate gangue minerals include mariposite, sericite, and chlorite. Pyrite is the dominant sulfide phase, accounting for roughly 1-2%. Arsenian pyrite, chalcopyrite, and galena are also found. Sphalerite is rare. Locally, abundant scheelite and telluride (Au-Ag and Ag) minerals have been historically recovered in addition to the free gold. Mariposite mineral separates provided a 40Ar/39Ar plateau age of 152 ± 1.2 Ma (Snow et al., 2008). Significant gold grades exist in the wall rocks immediately adjacent to the quartz-carbonate veins and to a lesser extent in mineralized country rock away from veins. 3.3 Materials and Methods Representative quartz-carbonate vein samples of the N-S vein set were collected underground at the Empire deposit in the Empire Mine State Park. Vein material from the E-W vein set was collected from drill core of the Idaho-Maryland mine. Initially, polished (100 µm) sections of the gold-bearing vein material were obtained and examined by standard optical microscopy using transmitted and reflected light to identify fluid inclusion, textural, and paragenetic relationships.

Following carbon coating, the thick sections were investigated by optical cathodoluminescence (CL) microscopy. A HC5-LM hot-cathode CL microscope by Lumic Special Microscopes, Germany, was used which allowed investigation of the sections during electron bombardment with a modified Olympus BXFM-S optical microscope. The instrument was operated at 14 keV and with a current density of approximately 10µA/mm2 (Neuser, 1995). A high sensitivity, double-stage Peltier cooled Kappa DX40C CCD camera was used to capture the CL images digitally. The CL signal of quartz was recorded using exposure times of 8-10 seconds. To capture the short-lived CL signal, images were captured automatically every 10 seconds following initial exposure. Scanning electron microscopy (SEM) on the thick sections was conducted using a FEI Quanta FEG 450 instrument at the U.S. Geological Survey in Denver, Colorado. Back- scattered electron (BSE) imaging was used to locate and study textural relationships in the veins that were too small for examination by standard optical petrography. In addition, BSE imaging was used to study zoning patterns within the vein minerals. Energy dispersive spectroscopy (EDS) elemental maps were obtained for pyrite grains to correlate BSE images with chemical information. These elemental maps were acquired using an accelerating voltage of 20 keV. The major and minor element chemistry of pyrite and the Au/Ag ratio of gold grains were determined using a JEOL 8900 Electron Microprobe with five wavelength dispersive analyzers at the U.S. Geological Survey in Denver, Colorado. Operating conditions for pyrite and gold analyses were 20 keV accelerating voltage, a 100 nA current (measured on the Faraday cup), and a focused electron beam. The Fe, S, As, Ni, Co, and Zn contents of pyrite and Ag and Au contents of the gold were measured this way. The minor and trace element chemistry of pyrite was determined by laser ablation- inductively coupled plasma-mass spectrometry (LA-ICP-MS) using a Photon Machines Analyte (193-nm excimer) coupled to a PerkinElmer DRC-e ICP-MS at the U.S. Geological Survey in Denver, Colorado. Data were obtained using a spot size of 25 µm, and calibrated using external synthetic sulfide calibration material MASS-1 from the U.S. Geological Survey. An average Fe content of 46% (based on EMPA data) was used as an internal standard element for all concentration calculations using the methods outlined by Longerich et al. (1996). Minor and trace elements that were analyzed for included Ag, Au, Bi, Cd, Co, Cr, Cu, Hg, Mn, Mo, Ni, Pb,

Sb, Se, Sn, Te, Tl, and Zn. The concentrations of Bi, Cr, Hg, Mn, Mo, Se, Sn, Te, and Tl were found to be typically below detection and are, therefore, not reported in the present contribution. Native gold inclusions or void fill exist in many of the pyrites that are analyzed. However, care was taken to ensure that any minerals that are found within the pyrite were not in the vicinity of the laser spot to eliminate any nugget effect. As such, any reported Au contents represent solid solution Au within the chemical structure of the pyrite. In situ sulfur isotope analyses of pyrite were conducted using a Nu Instruments high- resolution multi-collector inductively coupled plasma-mass spectrometer (MC-ICP-MS) at the U.S. Geological Survey in Denver, Colorado, following the methods of Pribil et al. (2015). Sample ablation was conducted via a 193 nm wavelength GS excimer laser ablation system using spot sizes of approximately 40-55 µm in diameter. Additional mineral separate isotopic analyses were conducted at the U.S. Geological Survey stable isotope lab in Denver, Colorado. Individual pyrite crystals were hand-picked. The clean separates were then combusted and analyzed for δ34S according to the methods described by Giesemann et al. (1994) using an Elementar iso-vario cube Elemental Analyzer coupled to a ThermoFinnigan Delta Plus XPTM continuous flow mass spectrometer. Isotope compositions are expressed relative to Vienna Cañon Diablo Troilite (VCDT) with a two sigma uncertainty of ±0.3‰. 3.4 Results Observations are combined with analytical results below. The synthesis of these results and interpretations based upon them will be presented in the discussion section. 3.4.1 Vein Textures Quartz represents the principal gangue mineral in the veins from Grass Valley. The bulk of the quartz in both vein sets is massive and shows little evidence of recrystallization. Zones of massive quartz in thin section are composed of prismatic euhedral quartz crystals intergrown with subhedral quartz crystals. These zones of massive quartz grade into regions of comb quartz, which fills open space. Some of the largest euhedral comb quartz crystals are found as a lining of vugs that were later filled by calcite. Locally, the quartz may be brecciated. The presence of sheared ribbon quartz in zones of brecciation indicates minor amounts of brittle and ductile-brittle fault behavior. Brittle deformation variably fractured quartz and pyrite throughout the veins.

Both vein sets contain abundant wall rock slivers. Hydrothermal fluid interaction with these fragments resulted in a complete replacement by alteration minerals such as white mica and sulfide minerals. Pyrite, white mica, and accessory phosphate minerals commonly occur along fractures within the quartz adjacent to altered wall rock slivers. The abundance of these minerals is distinctly lower away from the wall rock slivers. 3.4.2 Vein Quartz Two major types of quartz can be identified, with an additional two types of overprinting quartz (Fig. 3.4). Quartz 1 in the veins is characterized by myriads of secondary fluid inclusion trails, giving the grains a slightly dark color. The trails of fluid inclusions have a wispy appearance leading to the cloudy form of the quartz crystals. Quartz 1 grains have abundant fluid inclusions and have a short-lived bright blue optical CL that rapidly degrades and turns into a dark red-brown stable CL (Fig. 3.4d, f, h, l, n, p). Quartz 1 contains domains of short-lived yellow CL (Fig. 3.4d, h) that are characterized by abundant secondary fluid inclusions. Quartz 1 is commonly rimmed by a second generation of quartz (quartz 2) that is characterized by a clearer appearance due to a lower abundance of secondary fluid inclusion trails (Fig. 3.4a, c, e). This clear overgrowth can form the outer zones of euhedral quartz crystals that have quartz 1 cores. The quartz 2 lacks a visible CL response, resulting in characteristic black rims around the earlier blue quartz 1 in CL images (Fig. 3.4d, f). Obvious pseudosecondary fluid inclusion trails may be found within this quartz generation. Quartz 2 is commonly found adjacent to many of the larger sulfide grains, but not necessarily surrounding the entire sulfide grain (Fig. 3.4e, f, g, h, i, j). In rare cases, the late quartz 2 showing no CL response is crosscut by zones of yellow CL. Cutting all of the earlier generations of quartz 1 and 2 are streaks or tiny veinlets of inclusion-free and poorly luminescent quartz that are found subparallel to the vein margins (Fig. 3.4n). Individual quartz crystals from the N-S veins more commonly have repetitions and cycles of growth between quartz 1 and 2 types. In contrast, quartz crystals from the E-W veins commonly contain a single generation of quartz 1 that is rimmed by a single generation of quartz 2. In all of these veins, quartz 1 is more prevalent than quartz 2. Most of the fluid inclusions in the Grass Valley quartz average 2-3 µm in diameter, with rare examples exceeding 5 µm. Their small size makes them difficult to characterize, but they

Figure 3.4: Plane polarized light (A, B, C, E, G, I, K, and N) and cathodoluminescence (CL) photomicrographs (D, F, H, J, L, M, and O) of quartz veins in Grass Valley. (A) Euhedral quartz crystal growing into carbonate minerals within an E-W vein. A core of cloudy quartz 1 is surrounded by clear quartz 2. (B) Repetitive cycles of growth between cloudy quartz 1 and clear quartz 2 within a N-S vein. (C) Cloudy core of quartz 1 on the left with a thinner growth of clear quartz 2 on the quartz rim in an E-W vein. A thin sliver of carbonate minerals are shown in the lower right corner. (D) CL image of (C). (E) A chalcopyrite grain located along crystal boundaries and within fractures of both quartz 1 and 2 in an E-W vein. Note the euhedral quartz crystal with a cloudy quartz 1 core and a clear quartz 2 rim just to the left of the chalcopyrite grain. (F) CL image of (E). (G) A chalcopyrite grain surrounded by quartz 1 and quartz 2 in an E-W vein. (H) CL image of (G). (I) Pyrite, gold, quartz 1, and quartz 2 in an E-W vein. (J) CL image of (I). (K) Quartz 1 and 2 in a N-S vein. (L) CL image of (K). (M) Cl image of quartz from a N-S vein that shows both quartz 1 and quartz 2 with streaks or veinlets of late quartz with no CL response. (N) Chalcopyrite, pyrite, galena, and gold surrounded by quartz in an E-W vein. (O) CL image of (N).

71 appear to be single-phase or two-phase with ~5-10 vol. % vapor. No examples of fluid inclusions exhibiting double bubbles were found. 3.4.3 Pyrite Textures and Chemistry Pyrite grains from the veins of Grass Valley are characterized by distinctive chemical zoning patterns. These patterns are visible in BSE images, but can also be recognized in reflected light in sections that have been allowed to tarnish. Although the zoning patterns can be complex, the most common pattern visible in BSE images is characterized by a BSE-dark core that is surrounded by a thin BSE-bright band, and an outer euhedral BSE-dark rim (Figs. 3.5 to 3.7). Similarities and differences in the chemistry exist between the two vein sets (Figs. 3.5 to 3.6). The BSE-dark cores of the pyrite crystals are consistently more fractured than the BSE- bright bands or rims, and commonly contain subtle mottled or patchy domains. Many crystals have fractures in the cores that terminate at the boundary of the BSE-dark cores. Rarely, smaller pyrite crystals may have been amalgamated by overgrowth of later pyrite that has a different chemistry, forming large pyrite crystals. Commonly, the pyrite found in the wall rock slivers is more arsenic rich than pyrite that formed surrounded by quartz. Pyrite crystals from the N-S vein set contain a distinct BSE-dark core that is characterized by trace amounts of Zn (up to 0.85 wt.%, but usually much lower) and Cd (up to ~215 ppm) which were positively correlated and elevated within the darker unaltered cores compared to other zones (Table 3.2, Fig. 3.8a). Zinc concentrations were always greater than the Cd contents. Arsenic concentration was low within the cores, always containing less than 0.09 wt.% (Table 3.3). Solid solution gold values for analyses found entirely within cores did not exceed 3 ppm. Solid solution concentrations of Pb in the pyrite crystals was greatest in the dark core with values between 17.3-111 ppm (Fig. 3.8b, Table 3.2), although the level of solid solution Pb in pyrite did not correlate with the presence of precipitated galena. The mottled and patchy portions of the cores that are interior to the bright growth banding contained elevated levels of As and Au in the N-S vein pyrites relative to other chemical zones (Fig. 3.8b, c). Gold contents range from approximately 10-140 ppm and were found here in the highest concentrations of any of the pyrite zones (Table 3.2). No discrete gold grains were noted, so this elevated gold content likely results from solid solution gold within the mineral structure of the pyrite. Arsenic levels were also elevated and range between 1.4 to 2.6 wt.%.

Figure 3.5: Back-scattered electron images of pyrite crystals displaying growth and alteration zoning surrounding a dark core. Electron microprobe chemistry responsible for BSE zoning for individual spots in elemental weight percent. Complete chemical data is listed in Table 3.3 for each spot on each crystal. Samples with names beginning with N-S are from the N-S vein set and samples with names beginning with E-W are from the E-W vein set. The darker cores are characterized by low As, Ni, and Co concentrations whereas the brighter zones are typified by elevated concentrations of As, Ni, and (or) Co. N.D. = not detected.

Figure 3.5: Continued.

Figure 3.6: Back-scattered electron images of pyrite crystals displaying growth and alteration zoning surrounding a dark core. Select laser ablation ICP-MS chemical data shown for each analytical spot. Values are given in parts per million. The chemical values and additional elemental data is listed in Table 3.2. Samples with names beginning with N-S are from the N-S vein set and samples with names beginning with E-W are from the E-W vein set.

Figure 3.6: Continued.

Figure 3.7: Selected qualitative element maps derived from EDS analysis showing a back scattered electron image, Co, Ni, As, and Au content for pyrite crystals from both vein sets.

Figure 3.8: Select geochemical plots for pyrite analyses. Red symbols (NS) are from the N-S vein set. Purple symbols (EW) are from the E-W vein set. Symbols marked as mc are from the mottled core, bb from the BSE-bright band, core from the BSE-dark core, and dr from the BSE- dark rim. (A) Cadmium in parts per million (Cd ppm) versus zinc in parts per million (Zn ppm). Data from Table 3.2. (B) Gold in parts per million (Au ppm) versus lead in parts per million (Pb ppm). Data from Table 3.2. (C) Arsenic weight percent (As wt.%) versus iron weight percent (Fe wt.%). Data from Table 3.3.

Both Zn and Cd were below detection limits in the altered portions and every growth zone peripheral to this. The BSE-bright bands visible in the N-S vein pyrites contain Au contents (1-38 ppm, mostly >15 ppm) that were more elevated than the dark core, but not as elevated as the altered cores. The analyzed spots do not have any gold inclusions, so this gold content is the product of solid solution gold within the mineral structure. This BSE-bright band also had elevated As concentrations ranging from 1.09 to 2.72 wt.%, which is what imparts the brighter BSE response (Fig. 3.7). Discreet inclusions of galena and gold are associated with this thin bright band, forming inclusions of both minerals along the thin band and within fractures and pits in the pyrite interiors. In one notable sample, a fracture that is interior to the BSE-bright band of the pyrite is filled with gold; this fracture ends abruptly at the light overgrowth and doesn’t continue beyond it toward the rim of the crystal (Fig. 3.5, grain E-W 11). Not all of the N-S vein pyrites have a distinct outer BSE-dark rim. But when present, they contained low levels of gold similar to the dark unaltered cores (<5.3 ppm), but with slightly higher As concentrations (~0.5-1.5 wt.%). These euhedral overgrowths rarely contain inclusions or fracture/void fill of gold or galena but may have growths on the outer rim (Figs. 3.5 and 3.6). Nickel and cobalt concentrations were consistently low within all of the different zones of the N-S vein pyrites (Tables 3.2 and 3.3). Nickel is always below EMPA background levels and Co is always below the detection limit. Laser ablation ICP-MS results also find consistently low levels of both elements, commonly found in concentrations below the detection limit. The chemical zoning patterns of the E-W vein set pyrites are even more pronounced. But like the N-S vein pyrites, Cd and Zn were only found within the BSE-dark cores (Fig. 3.8a). In contrast to the N-S vein pyrites, Cd concentrations were always greater than the Zn concentrations but both no greater than hundreds of ppm level. Arsenic concentrations in the dark cores never exceeded 0.06 wt.% and gold values are always below 1 ppm (Tables 3.2 and 3.3). Nickel concentrations range from below EMPA detection limit up to 1.68 wt.% and cobalt concentrations were mostly below the EMPA detection limit but up to 0.5 wt.% during LA-ICP- MS analyses (Tables 3.2 and 3.3). The mottled and patchy domains within the dark cores are found to a much lesser extent in pyrite from the E-W vein set. But like the N-S vein set, this zone within the pyrite crystals

contain the highest levels of gold, ranging from less than 1 ppm up to 25 ppm (Table 3.2). However, As contents are low. Some of the E-W vein set pyrites have multiple textural generations of BSE-dark and bright bands, but most had a single dominant bright band. The BSE-bright bands were characterized by elevated As (1.87 to 5.53 wt.%), Co (0.21 to 2.6 wt.% EMPA), and Ni (1.6 to 3.31 wt.%). Structurally bound gold contents were very low, with one analysis having 1 ppm and the remainder having below 0.2 ppm. However, native gold and galena inclusions are found spatially associated with these bright bands. The euhedral rims rarely contain these inclusions or fracture/void fill of other minerals. They also have low Au contents (<2.5 ppm), low As (<0.41 wt.%), low Co (<0.34 wt.%), and low Ni (<1.6 wt.%). In both vein sets, Cr, Mn, Mo, Sn, Te, Hg, and Bi were rarely detected; Se and Tl were not detected at all in any analyses. Copper and Sb was detected, but most analyses show less than 10 ppm for each. Elemental EDS maps display the importance of As, Ni, and Co in the zoning patterns of these pyrites that have been emphasized above (Fig. 3.7). Zoning patterns with As were visible in pyrite from both vein sets. However, the bright zones in the As column in Figure 3.7 that are found outside of the pyrite crystals are actually hydrothermal mica and represent interferences. This is because the L peak for As, which was used in the mapping, has an interference with Mg that is found in the mica; magnesium is absent from the pyrite crystals. In contrast, Ni and Co zones are visible in pyrite from the E-W vein set but not the N-S vein set. 3.4.4 Mica and Chlorite Different types of fine-grained mica were found within both vein sets. Sericite and chlorite were found in both vein sets and mariposite was also noted in the E-W vein set. In both vein sets, chlorite forms later and was seen replacing K-rich micaceous phases but can also be found intergrown with each other. Aggregates of mica were more likely to have inclusions of pyrite crystals whereas other sulfides commonly occur on the periphery of these aggregates. 3.4.5 Other Sulfide Minerals Although distinct chemical zoning was evident within pyrite crystals, this feature was not noted within other sulfide minerals. Galena, chalcopyrite, and sphalerite all appear homogenous in BSE images. They were always located within fractures or along grain boundaries of quartz,

84 or associated with aggregates of mica. Although pyrite crystals may be euhedral or subhedral, the remaining sulfide minerals are all anhedral in shape. 3.4.6 Gold Textures and Chemistry In both vein sets, gold particles were found within fractures in quartz (Fig. 3.9a, c), within mica aggregates, within fractures and as inclusions within pyrite (Fig. 3.9g, k, l), and as small grains intergrown with galena and sometimes chalcopyrite (Fig. 3.9b, d, e, f, g). The anhedral gold crystals are porous and pitted (Fig. 3.9i, j), although gold intergrown with galena may be dendritic (Fig. 3.9f). Back-scattered electron imaging did not reveal any compositional zoning within the gold particles, which is consistent with the lack of Au and Ag zoning measured by electron microprobe analysis (Table 3.4). Gold and silver contents of gold particles are remarkably consistent within each vein set regardless of the textural setting of the gold and the immediate host rocks of the gold-bearing veins (Table 3.4). Electron microprobe analysis of eight gold grains from the Empire mine of the N-S vein set had Au contents ranging from 78.97-83.68 wt.%, which corresponds to 65.78-70.30 at.%. Zinc contents ranging from ~125-250 ppm were found in the gold grains (detection limit is 115 ppm). Lead was never detected. Electron microprobe analysis of 12 gold grains from the Idaho-Maryland mine of the E-W vein set also showed remarkably consistent values but more elevated than the N-S vein set, with Au contents ranging from 84.55-87.50 wt.%, which corresponds to 72.95-76.79 at.%. Zinc was also detected, with values ranging from ~150-275 ppm. The Pb concentrations never exceeded the detection limit of approximately 380 ppm. Both Au-Ag- and Ag-tellurides are found within the E-W veins, but are absent from the N- S veins. They are found isolated within fractures in pyrite and quartz, and occur as rims on native gold. The abundance of telluride minerals is low; concentrates from the mine tested at 0.03% Te (Johnston, 1940). 3.4.7 Carbonate Minerals Carbonate minerals fill voids and late fractures. Multiple types of carbonate minerals are present, including ankerite, dolomite, and calcite. Although intergrown, Mg-rich carbonate minerals are paragenetically first and Ca-rich carbonate minerals are last based upon inclusions and cross cutting textures. Carbonate minerals that are vug filling may encapsulate euhedral quartz crystals.

Figure 3.9: Photomicrographs of the textural relationships of gold and associated minerals. (A) Reflected light photomicrograph of gold occupying a fracture within quartz in a N-S vein. Two pyrite crystals are located above and below the gold. (B) Reflected light photomicrograph of gold and chalcopyrite infilling a void between quartz crystals in an E-W vein. (C) Reflected light photomicrograph of gold occupying a fracture within quartz in a N-S vein. Pyrite crystals are located on the right hand side. (D) Reflected light photomicrograph of gold (Au), chalcopyrite (cpy), galena (gn), and Ag-telluride (te) minerals within quartz in an E-W vein. (E) Reflected light photomicrograph of gold blebs in galena within quartz in a N-S vein. (F) Reflected light photomicrograph of dendritic gold within galena in quartz in an E-W vein. Chalcopyrite is attached to the galena in the upper part of the photomicrograph. (G) Reflected light photomicrograph of gold within a fracture in pyrite and as blebs within galena formed within fractures and along the edge of pyrite. (H). Backscattered electron image of gold (Au), galena (gn), Ag-Au-telluride (AuTe), and Ag-telluride (te) within quartz in an E-W vein. (I) Backscattered electron image showing the porous nature of gold hosted within a fracture of quartz in a N-S vein. The dark spots are holes and depressions within the gold. (J) Secondary electron image of same grain shown in (I). (K). Backscattered electron image of gold infilling holes and associated with a fracture in pyrite in an E-W vein. (L). Backscattered electron image of gold infilling a fracture within pyrite from an E-W vein. Note that the fracture containing the gold does not extend to the boundary of the crystal and the fractures termination coincides with an As-enriched growth band. The brittle fracturing of the pyrite core either predated or coincided with formation of the As-enriched growth band. Other bright minerals are galena.

3.4.8 Sulfur Isotopes Two pyrite crystals from the N-S vein set were analyzed by LA-ICP-MS to determine sulfur isotope ratios in situ (Fig. 3.10). Cores of the pyrite crystals that are dark in BSE images yielded δ34S values ranging from 0.60 to 0.97‰ (n=5). Mottled pyrite domains that are BSE- bright within the cores contrast have δ34S values of 1.27-1.50‰ (n=2), whereas the bright growth zones had values of 1.77-2.00‰ (n=3). Two mineral separate samples from the N-S vein set had bulk δ34S values of 0.6-1.4‰, within the range of values of the in situ isotope analyses. Three samples from the E-W vein set had values ranging from 2.2-2.7‰, which is distinctively heavier than the values for the N-S vein set.

Figure 3.10: Sulfur isotopic composition of pyrite from Grass Valley given in δ34S relative to VCDT. (A) Backscattered scanning electron (BSE) image of a pyrite from the N-S vein set with a dark core and bright growth rims and in situ isotopic values. (B) BSE image of a pyrite from the N-S vein set with a dark core and a bright alteration overprint and in situ isotopic values. (C) Diagram showing both in situ and partial grain sulfur isotopic analyses and the age dependence of the isotopic values. Events that occur later are characterized by heavier sulfur isotopic values.

3.5 Discussion The results provided above are integrated to discuss various aspects of the Grass Valley veins. Importantly, these provide evidence for a paragenetic sequence due to the uniquely undeformed nature of the veins. 3.5.1 Deformation of Vein Minerals Previous studies have interpreted that the amount of deformation in both of the vein sets from Grass Valley is less than in other gold districts of California, such as Alleghany (Johnston, 1940). The Grass Valley veins are interpreted to have been emplaced within 3 km of the paleosurface at the time of their formation (Taylor et al., 2015), which is in contrast to greater depths of formation (6-12 km) typical for most orogenic gold deposits (e.g., Groves et al., 1998; Goldfarb et al., 2005). This shallow depth of formation may have aided in preserving the original textures as these veins were not subjected to ductile stress. Although ductile deformation

and recrystallization of the veins is largely absent, fractures in quartz and pyrite in the veins occur, indicative of brittle deformation, and served as areas of nucleation for later gold precipitation. Quartz can be easily deformed and can develop textures such as bulging recrystallization, kink bands, undulose extinction, deformation lamellae, and fractures during deformation in subgreenschist facies conditions below 300 °C (Wu and Groshong, 1991; Nishikawa and Takeshita, 1999; Passchier and Trouw, 2005). However, quartz in the Grass Valley veins rarely displays undulose extinction, subgrain formation, grain boundary migration, or other microstructures indicating recrystallization. Many euhedral quartz crystals also preserve clear banding, interpreted to be oscillatory growth zones. The deep blue CL colors of the Grass Valley quartz 1 are more representative of hydrothermal quartz not influenced by recrystallization, as opposed to brownish CL responses from dynamically recrystallized quartz (e.g., Graupner et al., 2000). In addition, no quartz observed in this study exhibits annealed textures such as 120° dihedral angles representative of recrystallization. Additional brittle and brittle-ductile deformation may be localized in certain areas of the quartz veins, although the bulk of the quartz shows no recrystallization or effects of ductile deformation and is massive. Quartz in some areas may appear granular, interpreted to be the result of cataclasis, which is notably more likely to be found in wider veins. Although small, there are abundant secondary fluid inclusions found within the quartz crystals. This is especially true within portions of quartz 1. In addition, streaks of late overprinting quartz veinlets that form subparallel to the vein walls cut across all generations of quartz 1 and 2 growth. These features, particularly the formation of the abundant secondary fluid inclusions in quartz 1, indicate significant additional fluid flow through these crystals early in their history. However, these delicate textures produced by early reopening events are preserved. Sulfide minerals variably react to strain; of the sulfides present in the Grass Valley ores, galena is the weakest and most malleable, pyrite is the strongest and most brittle, and chalcopyrite and sphalerite are moderate with similar strengths to carbonate rocks (Marshall and Gilligan, 1987). Galena is sensitive to strain by fracturing and plastic deformation, even at experimental temperatures as low as 24 °C (e.g., Salmon et al., 1974). The cleavage planes in galena from Grass Valley are not offset by deformation; curvature or offset of planar features

89 such as cleavage planes in a soft mineral such as galena is evidence of deformation (Craig and Vaughan, 1994). Pyrite will deform brittely up to temperatures of 450 °C and confining pressures of 300 MPa (Marshall and Gilligan, 1987). Pyrite crystals from this study preserve complex chemical zoning patterns that would have been modified by later recrystallization. Pyrite recrystallization would have also likely resulted in the formation of euhedral cubes (Craig et al., 1998), which is not noted. Brittle fracturing is more intense in the interiors than in the overgrowths of the pyrite crystals. The variety of minerals found within the Grass Valley veins is able to record deformation over a wide range of pressures and temperatures during stress. In Grass Valley, the weakest (e.g., galena) through the strongest (e.g., pyrite) minerals record primary growth features not altered by significant strain. All of these microtextures are evidence of the minimal amounts of deformation and recrystallization that these orogenic gold veins underwent, features that are rare amongst orogenic gold deposits. 3.5.2 Paragenesis As discussed above, deformation and recrystallization of the veins is minimal, indicating that the observed textural relationships can be interpreted as primary, allowing the establishment of a paragenetic sequence (Fig. 3.11). Microscopic textural observation of these veins allows an untainted view of how orogenic veins form (Figs. 3.4 to 3.9 and 3.12). The paragenetic sequence described below is considered in fixed space. This sequence may repeat itself as multiple laminations are created within a vein, or as the creation of new veins within a deposit. Quartz is the first vein-filling material to form. As outlined in the results section, two main types of quartz have formed. Most orogenic veins have wispy inclusions and wispy CL patterns as the quartz is always overprinted by later events and streaks of secondary inclusions. Grass Valley is different than this and shows the primary growth zoning of both vein sets through CL and petrographic examination (Fig. 3.4a, b, c, d, e, f). Cloudy quartz 1 with fluid inclusions along growth planes first formed and was overprinted by the introduction of many more secondary fluid inclusions and crosscuts the growth bands in the cloudy primary quartz 1. Following this was precipitation of the clear quartz 2 surrounding and rimming the quartz 1 that formed the cores. Overprinting all of this are streaks of clear quartz, which seem to have preferably formed approximately parallel to the vein walls. After formation of all of the types of quartz, a minor dissolution event of quartz precedes precipitation of pyrite.

Figure 3.11: Paragenetic sequence derived from petrographic observations of thin sections. This is a fixed-space diagram indicating the paragenetic sequence found within samples.

Growth zoning in pyrite suggests a protracted growth history, whereas the lack of growth zoning in gold, galena, and chalcopyrite suggests these minerals precipitated within a narrower window of time with no chemical fluctuations. The zoning patterns observed within the pyrite crystals mark numerous growth episodes in which the chemical equilibration between fluid and wall rock is changing (Figs. 3.5 to 3.7). Most pyrite crystallization forms after the end of quartz precipitation, but before gold, galena, sphalerite, and chalcopyrite formation; minor growth extends to a short time period after gold mineralization has ceased. Pyrite forms euhedral to subhedral crystals, commonly in fractures, wall rock slivers, or along grain boundaries in quartz. Other times, they fill voids or are found intermixed with mica minerals. Almost all of the pyrite crystals formed after cessation of quartz precipitation. Only one example from this study had a small pyrite crystal included within an unfractured quartz crystal within the 2-dimensional view of thin sections. This observation indicates that minor pyrite formation might overlap with quartz deposition, but quartz is generally precipitated alone and prior to other silicate and sulfide phases (Fig. 3.11). As a general observation, euhedral pyrite crystals are more likely to be

located intermixed with mica whereas both subhedral and euhedral pyrite is found surrounded by quartz in fractures and along crystal boundaries. After all of the quartz and the majority of pyrite precipitation, fluids percolated along crystal boundaries and through the fractures of the quartz and early formed pyrite. This fluid responsible for later mica, sulfide, and gold precipitation used these micro fluid pathways to precipitate and locally dissolve minor amounts of quartz where this next generation of minerals was precipitating (Fig. 3.12a, b). As the overpressured fluid flows along these fractures and

grain boundaries, the H2O reacts with the quartz surfaces and causes a weakening of the silica with minor and local dissolution and remobilization of the quartz. Additionally, the solubility of silica increases with increasing pressures (e.g., Akinfiev and Diamond, 2009), such as during the infiltration of the overpressured fluid. Only a small solubility change of quartz is required for the minor amounts of dissolution of quartz and space creation; dissolution and remobilization of quartz is certainly not extensive. New minerals formed in the local areas of silica dissolution. Some small fractures within pyrite crystals are filled by the remobilized quartz (Figs. 3.9g, 3.12f). Hydrothermal mica formed contemporaneously with late stages and subsequent to pyrite growth, but after quartz precipitation. Mica is most commonly noted between and along quartz crystal boundaries, along fractures in quartz, and in altered wall rock slivers. Potassium- rich mica (sericite and mariposite) grows intermixed with chlorite. However, the K-rich mica formed before chlorite, and chlorite can be seen replacing sericite and as new crystals. Gold, galena, chalcopyrite, and sphalerite all appear to form entirely after quartz precipitation has ceased and coincident with final stages of mica formation (Fig. 3.12c). These minerals are chemically homogenous with no chemical or growth zoning visible in BSE images. The narrow compositional variation from analyses of spots within the cores and rims of the gold crystals also indicates their chemical homogeneity (Table 3.4). The base metal sulfides and gold are always anhedral, often occupying amoeba-like spaces between other crystals (Fig. 3.12d, e). As noted with the hydrothermal mica, formation of galena, chalcopyrite, and sphalerite appear to corrode into the preexisting quartz crystals. Gold precipitation appears to largely coincide with formation of galena as gold is found as inclusions within galena, both fill fractures and voids within pyrite, and both occupy locations within pyrite coincident with the bright arsenic banding. Although less common than galena, sphalerite also appears contemporaneous with gold;

Figure 3.12: Photomicrographs of textural evidence for a paragenetic sequence within the veins of Grass Valley. (A) Crossed polar photomicrograph of hydrothermal mica corroding into euhedral quartz crystals from the N-S vein set. Within the quartz crystal, growth zoning patterns of included primary fluid inclusions can be noted. The inner zone of quartz contains abundant primary fluid inclusions whereas the outer quartz contains few fluid inclusions. (B) Crossed polar photomicrograph of hydrothermal mica forming between euhedral quartz crystals and corroding into them from the N-S vein set. (C) Reflected light photomicrograph of a clot of hydrothermal mica rimmed by chalcopyrite that is irregularly fingering into previously formed quartz from the E-W vein set. Pyrite crystals are also found within the mica. (D) Plane polarized light photomicrograph of an anhedral sphalerite crystal corroding into earlier formed quartz from the N-S vein set. (E) Reflected light photomicrograph of an anhedral chalcopyrite grain growing between quartz crystals and corroding into them from the E-W vein set. (F) Reflected light photomicrograph of both gold and sphalerite fracture fill within a pyrite crystal from the N-S vein set.

sphalerite is exceedingly rare in the E-W vein set compared to the N-S vein set. Both gold and sphalerite are noted to occupy the same fractures within pyrite crystals (Fig. 3.12f). Chalcopyrite forms during the later stages or after that of gold, galena, and mica formation. Gold is found as elongate grains within and length parallel to fractures in quartz, within fractures in pyrite, as blebs within galena, and much of the gold and sulfide mineralization also appears to be spatially and temporally related to the alteration of wall rock slivers and the formation of hydrothermal mica. Gold mineralization postdates formation of the hosting quartz as is evident from gold exclusively forming within fractures of the quartz and never enclosed within unfractured crystals. Gold appears to have a closer genetic relationship with galena than to pyrite or quartz, and this positive correlation between gold and galena has been noted in other orogenic gold systems (Bierlein et al., 2004). Telluride minerals are found in the E-W vein set but not in the N-S vein set. These minerals are found within fractures in quartz and pyrite, and found growing on the edges of the later sulfides. Silver-tellurides formed after and grow on top of the Au-Ag-tellurides (Fig. 3.9). The last major vein material to form are carbonate minerals. Carbonate minerals are void filling, also corrode into euhedral quartz crystals, and carbonate veinlets crosscut the quartz veins, all indicating that carbonate minerals formed at the end of these hydrothermal events. Carbonate is also noted to enclose brecciated quartz fragments, likely brecciated during this carbonate forming event. Gold and other sulfides are rarely noted within carbonate, but gold has been noted to be surrounded by carbonate in at least one specimen from the N-S vein set (Johnston, 1940). The carbonate minerals may grow intermixed and contain variable concentrations of Ca, Mg, and Fe. When they are intermixed, the Mg- and Fe-rich carbonates form smaller subhedral crystals within a larger mass of more Ca-rich carbonate. 3.5.3 Timing of Gold Deposition Petrographic observations indicate that gold and quartz did not coprecipitate. Gold is paragenetically late and associated with base metal-sulfide precipitation and arsenic zoning during late pyrite growth. Because of this, studies of primary quartz-hosted fluid inclusions do not yield information regarding the fluid at the time of gold deposition, although the fluid composition should remain fairly constant during the life of the hydrothermal system. Fluid inclusion studies do, however, provide information regarding the properties of the early fluid within the hydrothermal environment during quartz precipitation. The paragenetic sequence

indicates that if primary fluid inclusions can be found within sphalerite, then they may represent the fluid at a time close to the gold precipitation. It is important to note that pyrite content is not identical to gold content. Pyrite cores are darker in BSE images, are fractured, contain pits, and have subtle mottled zoning with BSE- brighter patches containing elevated solid solution gold contents that are interpreted to be the result of chemical alteration by the hydrothermal fluid. Surrounding this is a thin bright band that contains elevated contents of As, Ni, and Co compared to the core. The event that formed this bright overgrowth band likely altered the core of the pyrites and produced the mottled zoning that could not have been produced simply by growth processes; the euhedral growth rims peripheral to the bright band are homogenous in BSE images and do not have the splotchy alteration patterns. Inclusions of galena and gold are associated with this BSE-bright band, forming inclusions of both within the band and within fractures and pits in the pyrite interiors. Overgrowing this is another darker pyrite growth without as many pits and fractures. The BSE- bright growth rim consistently contains elevated As but low gold levels. And even though Ni and Co contents in the E-W pyrites coincide with the alteration and growth zoning, there is no clear correlation with Au concentration within the pyrite crystals. The consistent Au/Ag content of the gold grains hosted within quartz, pyrite, chalcopyrite, and galena suggest that the gold was not remobilized from these crystals (Table 3.4). This was tested to determine if multiple generations could be documented with distinct Au/Ag contents, or if galena hosted gold may be more Ag-rich than pyrite, arsenopyrite, chalcopyrite, quartz, or mica hosted gold within each of the vein sets. It would be expected that the galena-hosted gold would be more Ag-rich if the gold was exsolved from the galena, which is not the case. Groupings of grains with different Au/Ag ratios within a single vein set could also suggest that the gold was introduced during different events or during different physical and chemical environments, which is also not the case. It is clear that the Au/Ag ratio is entirely independent of what the hosting mineral is. Simple mass balance calculations indicate that gold must have been introduced as a native phase and not remobilized in situ from the pyrite crystals; recrystallization did not occur and even if it did, original invisible gold would have needed to be found at unrealistic levels within the pyrite to account for the amount of gold found in these veins if it formed as remobilized invisible gold originally hosted in pyrite. The unaltered cores of the pyrites contain very low or non-detectable levels of gold; it is the altered portions that

95 contain higher levels of Au, indicating that it was precipitated subsequent to crystallization of the early pyrite. Additionally, solid-state remobilization of gold would also remobilize other metals and would blur the chemical zonation patterns that we see, another feature that is absent from these samples. 3.5.4 Differences Between the Two Vein Sets Although the overall mineralogy and textures noted in the two vein sets is very similar, some chemical characteristics of the minerals differentiate the vein sets and show that not all orogenic veins will display the same exact chemical features. These differences may be due to variable host rocks, physical differences in the environment of formation, or slightly different source regions for the fluids and metals. However, these differences would not have been caused by fluid:rock interaction along the flow path as the hydrothermal fluids resulting in both vein sets in Grass Valley traversed the same rocks as they flowed upwards through the Wolf Creek fault zone as a fluid-dominated system. Basic mineralogical differences include arsenopyrite, sphalerite, monazite, and xenotime being found in the N-S veins and absent in the E-W veins. Telluride minerals, scheelite, and mariposite are found in the E-W veins but are rare to absent in the N-S veins. Silver-bearing minerals, such as silver bichromite minerals, pyrargyrite, stephanite, and argentite have been more commonly noted within the N-S veins (Johnston, 1940), although silver-bearing telluride minerals are restricted to the E-W veins. The fineness of the gold between the two vein sets differs, with a larger Ag content within the N-S vein set. However, this difference in Au contents is only a few weight percent. Gold is commonly transported as bisulfide [Au(HS)-2], whereas silver is transported as a Cl complex [AgCl-2] (Morrison et al., 1991). Rare earth elements are also more soluble in fluids with more chloride as the REE’s also complex with chloride (Morrison et al., 1991; Reed et al., 2000). The larger silver content in the gold, the presence of distinct silver minerals, REE phosphate minerals, and the more likely appearance of the base metal mineral sphalerite within the N-S veins suggests that the hydrothermal fluids were more Cl- rich than those for the E-W veins. Pyrite is the most abundant sulfide mineral in both vein sets, but the chemistry and zoning patterns are distinct. Growth zoning and alteration zoning within N-S vein pyrites is solely due to arsenic and arsenic contents within the growth bands do not vary as wildly as those in the E-W pyrites; this has resulted in distinct thin bright bands in a darker pyrite for BSE images of E-W

96 pyrites and more subdued and diffuse bands of varying grey color for BSE images of pyrite from N-S veins (Fig. 3.7). In addition to As, both Ni and Co enrichments influence growth and alteration zoning in the E-W vein pyrites. Local host rock influences likely account for this, as the Ni- and Co-bearing pyrite of the E-W veins are found hosted within ultramafic rocks and the Ni- and Co-poor pyrite of the N-S veins are found within granodiorite. Increased interaction of the fluid with the wall rock and equilibration between the two may aid in diffusion of Ni and Co and the formation of these Ni- and Co-rich growth zones. This increased fluid/wall rock interaction would also lead to additional destabilization of the gold-transporting ligand and a correlation of gold precipitation with these later growth bands. Arsenic occurs in different valence states within pyrite between the two vein sets (Fig. 3.13). Arsenic substitutes for S as As1- in pyrite from the N-S vein set and substitutes for Fe as As2+ in pyrite from the E-W vein set; although, the trends in Figure 3.13 suggest that some of the arsenic can be As1- in pyrite from the E-W vein set. Arsenopyrite [Fe(As,S)] is rarely found in the E-W veins, but arsenopyrite and arsenian pyrite are more likely to be found in the N-S veins (Johnston, 1940; Pease, 2009; this study). The reason for the reduced As in the N-S veins and oxidized As in E-W veins is unknown. 3.5.5 Evolution of the Systems Despite minor to trace mineralogy differences, the two different vein sets share an identical paragenetic sequence in regards to both major gangue and ore. A repetition of the same processes that led to this specific paragenetic sequence of formation of both gangue and ore occurred in the same district during two distinct events 8 myr apart. Repetitions of these events are known to occur in individual laminated veins, in individual deposits as cross cutting veins during a single hydrothermal event (e.g., Muruntau; Graupner et al., 2005), and within districts due to multiple hydrothermal events (e.g., Grass Valley; this study). The mineralogy within the rock package at depth changes and releases metals and volatiles as they undergo the metamorphic transition from greenschist to amphibolite facies. Fluids are produced from the destruction of minerals such as mica and carbonate minerals, releasing H2O and CO2, respectively. The Au, As, CO2, and S, which account for mineralization and alteration, are sourced from the metamorphic rocks at depth undergoing this metamorphic transition (Böhlke and Kistler, 1986; Böhlke, 1989). But different minerals will devolatize at different times during an evolving metamorphic environment as the temperatures and pressures gradually

Figure 3.13: Ternary diagram of arsenic-bearing pyrite crystals. The arsenic speciation for each vein set is different.

change; for example, chlorite has a stability limit of 2-3 GPa of pressure whereas white micas can remain stable at pressures exceeding 3 GPa (e.g., Ernst, 2010). The progression of sulfur isotopic values obtained in this study show a similar evolution of sulfides that devolatize at depth (Fig. 3.10). The sulfur isotopic signature is lightest for the earliest forming pyrite in Grass Valley and progressively gets heavier over time, with the youngest pyrite crystals having the heaviest sulfur isotopes. Within sulfur-bearing phases, the bonds with the lightest sulfur are weaker and require less energy to break. As such, it is reasonable to assume that the first sulfur to be introduced by these hydrothermal systems is the lightest and the latest sulfur would necessarily be heavier and required more energy to be mobilized from the source rock. These metal-laden fluids move upwards from their source region by buoyantly rising within large-scale regional fault zones in concert with the earthquake cycle. Silica and metals are precipitated within smaller lower-order faults that are associated with the larger first-order

faults. Based upon paragenesis and textures, a sequence of events is proposed within a fixed- space framework of the mineralized veins. Fault valve behavior within the faults creates highly permeable fluid pathways and space for mineralization within the smaller faults (Sibson, 1990). This pulsed action allows batches of fluid to be pressed through the system and is linked to the earthquake cycle. Pressure changes accompanying this rapid and large volume expansion will result in temperature changes induced by adiabatic decompression (Ridley and Diamond, 2000) and rapid crystallization of quartz. The degree of equilibrium between the fluid and the wall rock will dictate the mineralogy that is precipitated and the chemistry of the minerals precipitated at that time. Quartz is formed by a silica-laden fluid that is out of equilibrium with the host rocks. Multiple generations of quartz can be created by repetitions of large earthquakes and the creation of space within a fault, leading to the common occurrence of laminated veins during these fluid-dominated events. Once the overpressured fluids in the system exceed the tensile strength of the crystallized quartz veins, they will cause an abrupt fracturing of the quartz and early formed pyrite. Increased pressure of the system by infiltration of the pressurized fluid along these fractures and along grain boundaries will increase the solubility of silica in a fluid, and interaction of an H2O-

CO2 fluid with quartz along grain boundaries and fractures will lead to weakening of silica; these factors combined will result in local dissolution of quartz. On the back end of the fluid infiltration, the pressure will decrease to regain equilibrium and allow Au and sulfides to precipitate depending on the degree of chemical equilibrium of the fluids with the host rock. Variations in equilibrium explain differences in the chemical zoning of pyrite. Lead isotopes within sulfide minerals of Phanerozoic orogenic gold deposits have been shown to correlate with the immediate host rocks and indicate a local derivation of the Pb in contrast to the fluid source for the Au and As (e.g., Goldfarb et al, 1997, 2005; Haeberlin et al., 2003, 2004). The Ni and Co concentrations and zones within pyrite of Grass Valley are also controlled and sourced from the host rocks. The period of gold deposition is marked by galena precipitation and chemical growth zones of As, and growth zones of Ni and Co in the ultramafic rock-hosted E-W vein set pyrites. This correlation of metals derived from the fluid (Au, Ag, and As) with metals derived from the local host rocks (Co, Ni, and Pb) during the Au precipitating event reveal the importance of fluid-rock interaction and equilibration between the two components. This mechanism would account for the paragenetically late appearance of gold, base metal

99 sulfides, and micas, and the minor dissolution of quartz accompanying the precipitation of these late minerals. The timescales necessary to create the observed paragenetic sequence are unknown, but can be repeated through time. Pressure fluctuations are important for creation of the veins and the correlation of mica, sulfides, and gold show the importance of wall rock and fluid interaction in the genesis of orogenic gold veins. 3.6 Conclusions The present study demonstrates that the quartz-carbonate veins of the Grass Valley district in California are fairly unique amongst orogenic gold veins as they have experienced minimal deformation, allowing reconstruction of primary textural and paragenetic relationships. Brittle deformation is restricted to fracturing of the quartz-carbonate veins. However, evidence for ductile deformation is largely absent. Extensive recrystallization of the quartz, resulting in grain boundary migration and granoblastic polygonal fabric, which is typical of quartz contained in orogenic quartz-carbonate veins is notably lacking at Grass Valley. The preservation of primary textures has allowed the development of a paragenetic sequence. Coupling this with microanalytical and geochemical work, the formation of the veins and the timing of gold mineralization can be established. Importantly, quartz did not coprecipitate with gold. Quartz is the first mineral to form and is largely formed independently of any other mineral. Only insignificant amounts of pyrite may form coevally with quartz. The quartz forms through fault-valve processes related to the earthquake cycle that significantly lowers the silica solubility through adiabatic decompression and pressure drops within a fluid-dominated system. Multiple generations of quartz are likely to form via this mechanism. This rapid precipitation minimizes the amount of interaction between the hydrothermal fluid and the wall rock. Subsequent overpressuring events lead to fracturing of the quartz and early forming pyrite, and forms tiny fluid pathways that hydrothermal fluid can percolate through in addition to following grain boundaries. Increased pressures due to infiltration of an overpressured fluid will slowly weaken the surfaces of the quartz to create small voids. Chemical interaction of the fluid with rock fragments during decreasing pressures will crystallize mica, sulfides, and gold along the fractures and grain boundaries. Slivers of wall rock will be heavily altered.

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The rocks at depth are the ultimate source of fluids, S, and Au and will evolve over time through metamorphism. Different minerals break down at differing pressures and temperatures. The correlation between S isotope values and age indicate that minerals with lighter S isotopes will react and release S into the hydrothermal system first, as minerals with heavier isotopic values are preserved until a later time when they are broken down. Gold precipitation is not related to quartz formation, and gold content does not necessarily relate to pyrite content. Some pyrite formed earlier within the paragenetic sequence with negligible invisible gold content and no gold growths. A rock-buffered event with equilibration between the fluid and the host rock formed a thin pyrite growth zone associated with formation of Au and galena growths and inclusions that also altered the barren early pyrite by introducing arsenic and gold into the altered pyrite cores. Nickel and Co introduction is correlated with the arsenic in the E-W veins but is absent from the N-S veins; this feature is a product of the differing host rocks of the two vein sets. But the correspondence of these Ni, Co, and (or) As rich growth periods with major gold precipitation likely results from accentuated reactions between the hydrothermal fluid and wall rock. Gold remobilization from the pyrite or other sulfide minerals did not occur. 3.7 References Akinfiev, N.N., and Diamond, L.W., 2009, A simple predictive model of quartz solubility in water-salt-CO2 systems at temperatures up to 1000 °C and pressures up to 1000 MPa: Geochimica et Cosmochimica Acta, v. 73, p. 1597–1608. Bickford, M.E., and Day, H.W., 1988, Jurassic ages of arc-ophiolite complexes, northern Sierra Nevada: Implications for duration of the Nevadan “orogeny” [abs.]: Geological Society of America Annual Meeting, Denver, Colorado, 1988, Abstracts with Programs, v. 20, p. A274. Bierlein, F.P., Christie, A.B., and Smith, P.K., 2004, A comparison of orogenic gold mineralisation in central Victoria (AUS), western South Island (NZ) and Nova Scotia (CAN): implications for variations in the endowment of Palaeozoic metamorphic terrains: Ore Geology Reviews, v. 25, p. 125–168. Bierlein, F.P., Northover, H.J., Groves, D.I., Goldfarb, R.J., and Marsh, E.E., 2008, Controls on mineralisation in the Sierra Foothills gold province, central California, USA: A GIS-based reconnaissance prospectivity analysis: Australian Journal of Earth Sciences, v. 55, p. 61– 78. Böhlke, J.K., 1989, Comparison of metasomatic reactions between a common CO2-rich vein fluid and diverse wall rocks; intensive variables, mass transfers, and Au mineralization at Alleghany, California: ECONOMIC GEOLOGY, v. 84, p. 291–327.

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Böhlke, J.K., and Kistler, R.W., 1986, Rb-Sr, K-Ar, and stable isotope evidence for the ages and sources of fluid components of gold-bearing quartz veins in the northern Sierra Nevada Foothills metamorphic belt, California: ECONOMIC GEOLOGY, v. 81, p. 296–322. Cox, S.F., Knackstedt, M.A., and Braun, J., 2001, Principles of structural control on permeability and fluid flow in hydrothermal systems: REVIEWS IN ECONOMIC GEOLOGY, v. 14, p. 1–24. Craig, J.R., and Vaughan, D.J., 1994, Ore Microscopy and Ore Petrography, John Wiley & Sons, New York, NY, 434 p. Craig, J.R., Vokes, F.M., and Solberg, T.N., 1998, Pyrite: physical and chemical textures: Mineralium Deposita, v. 34, p. 82–101. Day, H.W., Moores, E.M., and Tuminas, A.C., 1985, Structure and tectonics of the northern Sierra Nevada: Geological Society of America Bulletin, v. 96, p. 436–450. Dickinson, W.R., 2000, Geodynamic interpretation of Paleozoic tectonic trends oriented oblique to the Mesozoic Klamath-Sierran continental margin in California, in Soreghan, M.J., and Gehrels, G.E., eds., Paleozoic and Triassic paleogeography and tectonics of western Nevada and northern California: Geological Society of America Special Paper 347, p. 209– 245. ——, 2004, Evolution of the North American Cordillera: Annual Review of Earth and Planetary Sciences, v. 32, p. 13–45. ——, 2008, Accretionary Mesozoic-Cenozoic expansion of the Cordilleran continental margin in California and adjacent Oregon: Geosphere, v. 4, p. 329–353. Ducea, M., 2001, The California arc: thick granitic batholiths, eclogitic residues, lithospheric- scale thrusting, and magmatic flare-ups: GSA Today, v. 11, p. 4–10. Edelman, S.H., Day, H.W., and Bickford, M.E., 1989, Implications of U-Pb zircon ages for the tectonic settings of the Smartville and Slate Creek complexes, northern Sierra Nevada, California: Geology, v. 17, p. 1032–1035. Ernst, W.G., 2010, Late Mesozoic subduction-induced hydrothermal gold deposits along the eastern Asian and northern Californian margins: Oceanic versus continental lithospheric underflow: Island Arc, v. 19, p. 213–229. Ernst, W.G., Snow, C.A., and Scherer, H.H., 2008, Contrasting early and late Mesozoic petrotectonic evolution of northern California: Geological Society of America Bulletin, v. 120, p. 179–194. Fagan, T.J., Day, H.W., and Hacker, B.R., 2001, Timing of arc construction and metamorphism in the Slate Creek Complex, northern Sierra Nevada, California: Geological Society of America Bulletin, v. 113, p. 1105–1118. Giesemann, A., Jager, H.J., Norman, A.L., Krouse, H.R., and Brand W.A., 1994, On-line sulfur- isotope determination using an elemental analyzer coupled to a mass spectrometer: Analytical Chemistry, v. 66, p. 2816–2819. Glazner, A.F., 1991, Plutonism, oblique subduction, and continental growth: An example from the Mesozoic of California: Geology, v. 19, p. 784–786.

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Goldfarb, R.J., Snee, L.W., Miller, L.D., and Newberry, R.J., 1991, Rapid dewatering of the crust deduced from ages of mesothermal gold deposits: Nature, v. 354, p. 296–298. Goldfarb, R.J., Miller, L.D., Leach, D.L., and Snee, L.W., 1997, Gold deposits in metamorphic rocks of Alaska: ECONOMIC GEOLOGY MONOGRAPH 9, p. 151–190. Goldfarb, R.J., Baker, T., Dubé, B., Groves, D.I., Hart, C.J.R., and Gosselin, P., 2005, Distribution, character, and genesis of gold deposits in metamorphic terranes: ECONOMIC GEOLOGY 100TH ANNIVERSARY VOLUME, p. 407–450. Goldfarb, R.J., Hart, C.J.R., and Marsh, E.E., 2008, Orogenic gold and evolution of the Cordilleran orogen, in Spencer, J.E., and Titley, S.R., eds., Ores and orogenesis: Circum- Pacific tectonics, geologic evolution, and ore deposits: Arizona Geological Society, Digest 22, p. 311–323. Graupner, T., Götze, J., Kempe, U., and Wolf, D., 2000, CL for characterizing quartz and trapped fluid inclusions in mesothermal quartz veins: Muruntau Au ore deposit, Uzbekistan: Mineralogical Magazine, v. 64, p. 1007–1016. Graupner, T., Kempe, U., Klemd, R., Schuelssler, U., Spooner, E.T.C., Götze, J., and Wolf, D., 2005, Two stage model for the Muruntau (Uzbekistan) high grade ore structures based on characteristics of gold, host quartz and related fluids: Neues Jahrbuch für Mineralogie Abhandlungen, v. 181, p. 67–80. Griggs, D.T., and Blacic, J.D., 1965, Quartz: Anomalous weakness of synthetic crystals: Science, v. 147, p. 292–295. Groves, D.I., Goldfarb, R.J., Gebre-Mariam, M., Hagemann, S.G., and Robert, F., 1998, Orogenic gold deposits: A proposed classification in the context of their crustal distribution and relationship to other gold deposit types: Ore Geology Reviews, v. 13, p. 7–27. Haeberlin, Y., Moritz, R., and Fontbote, L., 2003, Paleozoic orogenic gold deposits in the eastern Central Andes and its foreland, South America: Ore Geology Reviews, v. 22, p. 41–59. Haeberlin, Y., Moritz, R., Fontbote, L., and Cosca, M., 2004, Carboniferous orogenic gold deposits at Pataz, eastern Andean Cordillera, Peru—geological and structural framework, 40 39 paragenesis, alteration, and Ar/ Ar geochronology: ECONOMIC GEOLOGY, v. 99, p. 73– 112. Hodgson, C.J., 1989, The structure of shear-related, vein-type gold deposits: A review: Ore Geology Reviews, v. 4, p. 231–273. Irwin, W.P., 2003, Correlation of the Klamath Mountains and the Sierra Nevada: U.S. Geological Survey Open-File Report 02-490, scale 1:1,000,000, 1 plate. Irwin, W.P., and Wooden, J.L., 2001, Map showing plutons and accreted terranes of the Sierra Nevada, California, with a tabulation of U/Pb isotopic ages: U.S. Geological Survey, Open-File Report 01-229. Johnston, W.D., Jr., 1940, The gold quartz veins of Grass Valley, California: U.S. Geological Survey Professional Paper 194, 101 p.

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Longerich, H.P., Jackson, S.E., and Gunther, D., 1996, Inter-laboratory note. Laser ablation inductively couple plasma mass spectrometric transient signal data acquisition and analyte concentration calculation: Journal of Analytical Atomic Spectrometry, v. 11, p. 899–904. Luan, F.C., and Paterson, M.S., 1992, Preparation and deformation of synthetic aggregates of quartz: Journal of Geophysical Research: Solid Earth, v. 97, p. 301–320. Marshall, B., and Gilligan, L.B., 1987, An introduction to remobilization: Information from orebody geometry and experimental considerations: Ore Geology Reviews, v. 2, p. 87–131. Morrison, G.W., Rose, W.J., and Jaireth, S., 1991, Geological and geochemical controls on the silver content (fineness) of gold in gold-silver deposits: Ore Geology Reviews, v. 6, p. 333–364. Neuser, R.D., 1995, A new high-intensity cathodoluminescence microscope and its application to weakly luminescing minerals: Bochumer Geologische und Geotechnische Arbeiten, v. 44, p. 116–118. Nishikawa, O, and Takeshita, T., 1999, Dynamic analysis and two types of kink bands in quartz veins deformed under subgreenschist conditions: Tectonophysics, v. 301, p. 21–34. Passchier, C.W., and Trouw, R.A.J., 2005, Microtectonics: Springer, Berlin, 366 p. Payne, M., 2000, Geology of the Grass Valley mining district, Nevada County, California, in Shaddrick, D.R., ed., Platinum group elements, high grade gold and history: The Sierra Nevada 2000: Geological Society of Nevada Special Publication 32, p. 125–136. Pease, R.C., 2009, Idaho-Maryland mine project, Grass Valley CA, technical report: Prepared for Emgold Mining Corporation, 110 p. (Also available at www.emgold.com/s/TechnicalReport.asp; accessed 12/2013). Phillips, G.N., and Powell, R., 2010, Formation of gold deposits: a metamorphic devolatilization model: Journal of Metamorphic Petrology, v. 28, p. 689–718. Post, A.D., Tullis, J., and Yund, R.A., 1996, Effects of chemical environment on dislocation creep of quartzite: Journal of Geophysical Research, v. 101, p. 22,143–22,155. Pribil, M.J., Ridley, W.I., and Emsbo, P., 2015, Sulfate and sulfide sulfur isotopes (δ34S and δ33S) measured by solution and laser ablation MC-ICP-MS: An enhanced approach using external correction: Chemical Geology, v. 412, p. 99–106. Reed, M.J., Candela, P.A., and Piccoli, P.M., 2000, The distribution of rare earth elements between monzogranitic melt and the aqueous volatile phase in experimental investigations at 800 °C and 200 MPa: Contributions to Mineralogy and Petrology, v. 140, p. 251–262. Ridley, J.R., and Diamond, L.W., 2000, Fluid chemistry of orogenic lode gold deposits and implications for genetic models: Gold in 2000, p. 141–162. Saleeby, J.B., Shaw, H.F., Niemeyer, S., Moores, E.M., and Edelman, S.H., 1989, U/Pb, Sm/Nd, and Rb/Sr geochronological and isotopic study of northern Sierra Nevada ophiolitic assemblages, California: Contributions to Mineralogy and Petrology, v. 102, p. 205–220. Salmon, B.E., Clark, B.R., and Kelly, W.C., 1974, Sulfide deformation studies: II. Experimental deformation of galena to 2,000 bars and 400 °C: ECONOMIC GEOLOGY, v. 69, p. 1–16.

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Saucedo, G.J., and Wagner, D.L., 1992, Geologic map of the Chico quadrangle: California Department of Conservation, Division of Mines and Geology, Regional Geologic Map Series, Map No. 7A, scale 1:250,000. Sharp, W.E., 1988, Pre-Cretaceous crustal evolution of the Sierra Nevada region, in Ernst, W.G., ed., Metamorphism and crustal evolution of the western United States: Englewood Cliffs, New Jersey, Prentice-Hall, p. 824–864. Sibson, R.H., 1990, Conditions for fault-valve behavior, in Knipe, R.J., and Rutter, E.H., eds., Deformation Mechanisms, Rheology and Tectonics: Geological Society Special Publication No. 54, p. 15–28. Sibson, R.H., Robert, F., and Poulsen, K.H., 1988, High-angle reverse faults, fluid-pressure cycling, and mesothermal gold-quartz deposits: Geology, v. 16, p. 551–555. Snow, C.A., and Scherer, H., 2006, Terranes of the Western Sierra Nevada Foothills metamorphic belt, California: A critical review: International Geology Review, v. 48, p. 46–62. Snow, C.A., Bird, D.K., Metcalf, J., and McWilliams, M., 2008, Chronology of gold mineralization in the Sierra Nevada Foothills from 40Ar/39Ar dating of mariposite: International Geology Review, v. 50, p. 503–518. Taylor, R.D., Goldfarb, R.J., Monecke, T., Fletcher, I.R., Cosca, M.A., and Kelly, N.M., 2015, Application of U-Th-Pb phosphate geochronology to young orogenic gold deposits: New age constraints on the formation of the Grass Valley gold district, Sierra Nevada Foothills province, California: ECONOMIC GEOLOGY, v. 110, p. 1313–1337. Umhoefer, P.J., 2003, A model for the North America Cordillera in the Early Cretaceous: Tectonic escape related to arc collision of the Guerrero terrane and a change in North America plate motion, in Johnson, S.E., Paterson, S.R., Fletcher, J.M., Girty, G.H., Kimbrough, D.L., and Martín-Barajas, A., eds., Tectonic evolution of northwestern México and southwestern USA: Geological Society of America Special Paper 374, p. 117–134. Wu, S., and Groshong, R.H., Jr., 1991, Low temperature deformation of sandstone, southern Appalachian fold-thrust belt: Geological Society of America Bulletin, v. 103, p. 861–875.

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CHAPTER 4 40AR/39AR GEOCHRONOLOGY OF HYDROTHERMAL ACTIVITY RELATED TO GOLD MINERALIZATION IN THE KLAMATH MOUNTAINS, CALIFORNIA

Extensive, structurally-controlled gold mineralization stretches for hundreds of kilometers in the Sierra Nevada foothills and the Klamath Mountain gold provinces of northern California. Over 115 Moz of gold have been produced from mining in California (Craig and Rimstidt, 1998). The Klamath Mountains are the second most productive lode and placer gold province in California, second only to the Sierra Nevada foothills. In the Klamath Mountains, a total of more than 7 Moz of lode plus placer gold is estimated to have been produced (Hotz, 1971a). Most of this was a product of placer mining, but 27% has been produced directly from the gold- bearing quartz veins (Silberman and Danielson, 1993). Numerous studies have characterized the diachronous timing of gold mineralization within the Sierra Nevada foothills (Marsh et al., 2008; Snow et al., 2008; Taylor et al., 2015); however, the timing of gold mineralization within the Klamath Mountains and how this relates to the tectonic and metallogenic evolution of gold deposition within the Sierra Nevada foothills remains poorly constrained. The Morrison-Carlock deposit of the Oro Fino district is the only currently dated deposit in the Klamath Mountains, with a K-Ar age constraint on the timing of mineralization ranging from 145.8 ± 3.0 and 147.7 ± 2.8 Ma (Elder and Cashman, 1992). This study reports new 40Ar/39Ar geochronological data from white mica interpreted to have formed during hydrothermal alteration associated with orogenic gold deposits spread throughout the Klamath Mountains gold province of California. Samples from eight different districts were collected that contain hydrothermal white mica of suitable size for analysis. In all of the deposits studied here, sulfide formation and gold mineralization is associated with formation of white mica. These new ages constrain the timing of gold mineralization in the Klamath Mountains and coupled with ages of hydrothermal activity in the Sierra Nevada foothills, allow an interpretation of the tectonic and metallogenic evolution of California. 4.1 Background The Klamath Mountains are found within northern California and southern Oregon. They are located further west than the Sierra Nevada and closer to the coast. However, their

106 tectonic fabric of approximately north-south trending lithotectonic belts is the same as those found in the Sierra Nevada. 4.1.1 Tectonic Setting Present-day California is a region that has undergone a complex geological evolution since the Paleozoic (Fig. 4.1). Accretion of Devonian-Permian allochthonous near-shore island arcs onto the continental margin of North America formed the Northern Sierra terrane in the Sierra Nevada and the correlative Central Metamorphic and Eastern Klamath terranes in the Klamath Mountains during the late Paleozoic (Dickinson, 2000, 2004, 2008). The Devonian to Late Jurassic allochthonous and autochthonous terranes that comprise the remainder of the Sierra Nevada foothills and the Klamath Mountains were incrementally accreted beginning in the Middle Triassic (Dickinson, 2008). The exact timing of the termination of terrane amalgamation in the Sierra Nevada and the Klamath Mountains remains controversial, but is thought to have occurred by the Late Jurassic (Sharp, 1988; Taylor et al., 2015). Transcurrent movement along the terrane-bounding faults continued well beyond the end of terrane accretion (Tobisch et al., 1989) and is interpreted to represent a key process in the formation of the orogenic gold deposits of California. Two distinct episodes of syn- to post-accretionary plutonism are recognized in the Sierra Nevada (Glazner, 1991; Irwin and Wooden, 2001). One episode occurred mainly between 170- 140 Ma and at the time, formed a semi-continuous volcano-plutonic arc spanning the Klamath Mountains and Sierra Nevada (Ernst, 2013). A 20 Myr magmatic lull occurred before the next magmatic episode from 120-80 Ma, which led to the formation of the Sierra Nevada batholith. Plutonism during this younger episode is absent in the Klamath Mountains. Strike-slip processes along both the northern and southern margins of the Klamath Mountain terranes translated them approximately 200 km to the west of their original Jurassic arc position to their current outboard location (Ernst, 2013). The correspondence of approximately 170-140 Ma plutonism between the two mountain ranges, and the absence of 120-80 Ma plutonism in the Klamath Mountains helps to constrain the translation to the time of the 20 Myr magmatic lull. Ernst (2013) showed that the pattern of sedimentation along the southeastern edge of the Klamath Mountains further constrains the bulk of the translation to 140-136 Ma, which coincides with the age of the youngest dated plutons in the Klamath Mountains (Irwin and Wooden, 2001).

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Figure 4.1: Geologic map of northern California including the Sierra Nevada and the Klamath Mountains. Modified from Irwin (2003) and Ernst et al. (2008).

Gold mineralization within the Sierra Nevada foothills province has been noted to incrementally span from the first plutonic episode through separation with the Klamath Mountains during the magmatic lull to the start of the second magmatic episode, with punctuated episodes of vein formation beginning at 160 Ma and continuing through about 115 Ma (Marsh et

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al., 2008; Snow et al., 2008; Taylor et al., 2015). The Grass Valley district in the Sierra Nevada foothills province had two gold forming episodes prior to the westward translation of the Klamath Mountains and contains the oldest known orogenic gold deposits in California, one vein set having formed at ~160 Ma and another at ~152 Ma (Snow et al., 2008; Taylor et al., 2015). The deposits of the Grass Valley district represent the most productive lode gold district in the entire western cordillera of North America. These are the only dated gold deposits in the Sierra Nevada that are known to have formed prior to the lateral offset of the Klamath Mountains. 4.1.2 Klamath Mountains Geology Nine distinct tectonostratigraphic units are recognized in the Klamath Mountains that correlate with units in the Sierra Nevada foothills (Scherer et al., 2006; Ernst et al., 2008; Ernst, 2013). These were juxtaposed during eight accretionary episodes (Irwin and Wooden, 1999). The nucleus of the Klamath Mountains is the Eastern Klamath terrane (Fig. 4.2), which is composed of multiple subterranes including the Trinity subterrane. During the middle Paleozoic, the Central Metamorphic terrane, composed of metamorphosed oceanic crust and pelagic sediment (Barrow and Metcalf, 2006), was subducted beneath and accreted to the Eastern Klamath terrane. After a hiatus, terrane accretion resumed as the Stuart Fork terrane (also known as the Fort Jones terrane) was accreted during a Permian-Triassic event. The North Fork terrane, the Condrey Mountain terrane, the Eastern Hayfork terrane, the western Hayfork terrane, the Rattlesnake Creek terrane, and the Western Klamath terrane (Fig. 4.2) were incrementally accreted to the westward growing margin during the Jurassic. The metamorphosed terranes of the Klamath Mountains contain pre-accretionary plutons and accretionary plutons associated with the subduction and accretion processes. The accretionary plutons range from ~170 Ma to ~136 Ma in age (Lanphere and Jones, 1978). The majority of these plutons are quartz diorite or granodiorite, but range from gabbro to quartz monzonite in composition (Hotz, 1971b). Lode gold deposits in the Klamath Mountains are located within all of the accreted terranes. These terranes in the Klamath Mountains are thought to correlate with terranes of the Sierra Nevada foothills such that the Northern Sierra terrane correlates with the Eastern Klamath terrane, the Red Ant schist correlates with the Stuart Fork terrane, the Calaveras Complex correlates with the North Fork and Eastern Hayfork terrane, the Feather River terrane correlates with the Condrey Mountain terrane, the Jura-Triassic arc belt correlates with the Rattlesnake

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Figure 4.2: Geologic map of the Klamath Mountains with sample locations. Modified from Irwin (2003) and Ernst et al., 2008. Locations for all of the gold deposits in the Klamath Mountains from Hotz (1971a).

Creek terrane, and the Jurassic accretionary arc sequence correlates with the Western Klamath terrane (Fig. 4.1; Irwin, 2003; Snow and Scherer, 2006). The most important gold-hosting terranes in the Sierra Nevada foothills are the Calaveras Complex and the Jura-Triassic arc belt. The most important gold-hosting terranes in the Klamath Mountains are the North Fork, Central Metamorphic, and Eastern Klamath terranes (Fig. 4.2). 4.1.3 Gold Deposits The highest concentration of significant lode gold deposits of the Klamath Mountains occurs in the southeast near Redding (Fig. 4.2). The northern Klamath Mountain lode gold

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deposits in Oregon are generally smaller. With the exception of some of the larger deposits, lode gold deposits in Oregon appear to be spatially less restricted to known major structures (Fig. 4.2; Hotz, 1971a). The veins of the Klamath Mountains fill fractures or faults that can be centimeters to many meters wide. Most commonly, they are closer to 0.3 to 2.5 m wide. The majority of the veins are steeply dipping, although veins of gentle inclination are observed. Gold in the lode deposits of the Klamath Mountains is typically in the native form (Clark, 1970). The majority of veins are found within the metamorphosed country rock, with lesser amounts hosted within granitic intrusions. Overall, the sulfide contents of the veins are low (<3 to 5%) and are composed predominantly of pyrite with lesser amounts of arsenopyrite, galena, sphalerite, chalcopyrite, pyrrhotite, molybdenite, and telluride minerals (Hotz, 1971a). Elder and Cashman (1992) provided the only published radiometric age estimates for gold mineralization within the Klamath Mountains; a K-Ar age from hydrothermal mica at 145.8 ± 3.0 and 147.7 ± 2.8 Ma from the Morrison-Carlock deposit in the Oro Fino district. 4.2 Study Sites There were attempts to visit nearly 50 mine sites during the course of the field work. Access to the vast majority of these deposits was restricted due to private property, road closures, and ongoing forest fires. Of these, access was available to 15 deposits where samples could be collected. Upon microscopic characterization of the collected samples, it was determined that only eight of the samples contained hydrothermal mica of sufficient size for analysis that were also paragenetically related to sulfide formation. 4.2.1 McKeen Deposit, Callahan District The Callahan district is known for the substantial amounts of gold recovered from gravel benches; however numerous lode deposits are also located here. The McKeen lode gold deposit (also known as the Cummings deposit) of the Callahan district is hosted within the northern portion of the Craggy Peak trondhjemite pluton. The trondhjemite yielded U-Pb zircon crystallization ages ranging from 141 Ma (reported in Barnes et al., 1996) to 138.2 ± 1.3 Ma (Allen and Barnes, 2006) and a K-Ar biotite cooling age of 136 Ma (Lanphere et al., 1968). The Craggy Peak pluton intruded into the Trinity subterrane of the Eastern Klamath terrane, which is mostly comprised of serpentinized peridotite.

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At least four veins are known from the McKeen deposit, but only one has been developed that averages 0.6 m in width, trending 035°, and dipping 85° southeast (Logan, 1925; Averill, 1931). Nearly 25,000 oz of gold has been produced from the McKeen deposit, making it the second most productive mine in the district behind the Dewey mine, which is also hosted by the Craggy Peak pluton (Hotz, 1971a). Most of the other deposits in the district are located outside of the Craggy Peak pluton within the serpentinite country rocks. The laminated white quartz-carbonate veins primarily consist of deformed quartz displaying evidence for dynamic recrystallization with undulose extinction, deformation lamellae, and bulging recrystallization. Pyrite is common in the quartz veins, mostly occurring along fractures in the quartz veins. Carbonate minerals commonly fill open space between euhedral quartz crystals in the veins. Formation of the quartz-carbonate veins at the McKeen deposit was associated with the widespread alteration of feldspar in the trondhjemite and the formation of secondary white mica, carbonate minerals, and clay minerals. Primary biotite in trondhjemite is replaced by white mica adjacent to the veins, but only partially chloritized away from the veins. Pyrite is widespread in sulfidized wall rock. 4.2.2 Hickey Deposit, Liberty District Approximately 300,000 oz of lode gold is estimated to have been produced from the Liberty district, with over 1,700 oz derived from the Hickey deposit (Ferrero, 1990). It is the second most prolific lode gold producing district in the Klamath Mountains, behind the French Gulch/Deadwood district. The deposits of the Liberty district are mostly situated within a thrust fault shear zone separating the siliceous schist of the overlying Stuart Fork terrane from metavolcanic and metasedimentary rocks of the underlying North Fork terrane to the north and west. The siliceous schist of the Stuart Fork terrane represents the main host rock for the Hickey deposit. However, the country rock has disintegrated to clays that contain fragments of vein quartz mixed in. The average ore grade is 0.16 oz/t (Averill, 1935). Deformation of the quartz-carbonate veins of the Hickey deposit is severe. The quartz displays dynamic recrystallization textures to a significant degree including subgrain formation, lobate intergrowths, grain boundary migration, and deformation lamellae. White mica and sulfide minerals occupy fractures within the quartz vein and along contacts between different

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types/generations of quartz. The sulfide minerals, including pyrite and sphalerite, and white mica are spatially and genetically related. 4.2.3 Quartz Hill Deposit, Scott Bar District The Quartz Hill lode gold deposit is located in the Condrey Mountain terrane immediately adjacent to the boundary with the Eastern Hayfork terrane. The host rock of the deposit is micaceous schist. An extensive system of quartz veins and lenses with rich pockets of gold is found at Quartz Hill (Averill, 1931; Obrien, 1947). The gold-bearing veins have abundant quartz and ankerite. The most highly sulfidized ore comes from host rock inclusions within the veins. Country rock inclusions within the veins are comprised of abundant pyrite, carbonate minerals, muscovite, and quartz, with minor molybdenite. The highest concentration of pyrite and molybdenite is found associated with muscovite. Both Au- and Ag-tellurides have been noted from this deposit (Averill, 1931). 4.2.4 Schroeder Deposit, Yreka-Fort Jones District The Schroeder deposit is hosted within the North Fork terrane, to the south of the 166.9 ± 1.9 Ma (Allen and Barnes, 2006) Vesa Bluffs pluton. The ore-hosting country rocks in this district are predominantly composed of metabasalt. Multiple types of host rock are found at the Schroeder deposit, including various types of intrusive, but altered metabasalt represents the dominant lithology. The main orebody strikes 080° and dips 60-70° to the south with an average grade of 0.43 oz/t (Averill, 1931). The veins display a laminated texture with minor deformation and recrystallization of the quartz. Slickensides on the vein margins indicate movement during or after emplacement. Vein and country rock of altered intrusive was collected, with visible white mica alteration. The white mica was coarsest near the vein margin and was produced by alteration of feldspar. Pyrite and sphalerite dominate the sulfide assemblage. 4.2.5 McKinley Deposit, Humbug District The Humbug district is located in close proximity to the Yreka-Fort Jones district. Like the Schroeder deposit, the McKinley deposit is located within greenstone of the North Fork terrane just to the south of the Vesa Bluffs pluton. Petrographic investigation determined that quartz within the veins formed early in the paragenesis, followed by white mica and sulfide minerals that formed within fractures

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crosscutting the quartz. As with the Schroeder deposit, the deformation and recrystallization of the veins is minor. Later carbonate minerals filled in fractures and open space cavities between euhedral quartz crystals. The white mica separated from these veins is spatially isolated from hydrothermal chlorite aggregates, but is associated with pyrite. The laminated quartz veins from the McKinley deposit contain significant concentrations of galena. 4.2.6 Washington Deposit, French Gulch-Deadwood District The French Gulch-Deadwood district is the richest gold district in the Klamath Mountains, having produced between 800,000 and 1,500,000 oz from both lode and placer deposits. The Washington deposit has produced at least 150,000 oz of lode gold (Jenks and Tregaskis, 2007). The lode deposits are oriented along an east-west trending fracture system within the Eastern Klamath terrane associated with the Spring Creek Thrust between the Bragdon Formation and the Copley Greenstones. However, the Dean vein is found entirely within metavolcanic rocks of the Copley Greenstone. This is unique because most of the veins that cut and are entirely hosted within the Copley Greenstone are thin and of low grade (Hotz, 1971a). Numerous dikes and sills intruded into this highly fractured zone. The district spans about 15 km, with the western portion in Trinity County known as Deadwood and the central and eastern portions in Shasta County known as French Gulch. The Washington deposit is located in the central part of the district near the village of French Gulch. Total historic gold production from the Washington deposit, which consists of at least six veins, is estimated to be 300,000 oz (Shasta Gold Corp, 2014). Samples for the present study were collected from the Dean 3NS vein of the Washington deposit. The quartz vein contains highly fractured, euhedral Fe-bearing clinozoisite (or Fe-poor epidote). Carbonate minerals occur as late fracture and open-space fill within the deformed and recrystallized quartz. White mica is present in the same fractures and open spaces as the carbonate minerals. The white mica is the coarsest of any found during this study, and has abundant mineral inclusions homogenously distributed throughout the mica crystals. Pyrite is the dominant sulfide at the Washington deposit, with lesser amounts of galena, sphalerite, and arsenopyrite. 4.2.7 Yankee John Deposit, Redding District The deposits of the Redding district are hosted within the 400 Ma Mule Mountain quartz diorite-trondhjemite stock (Albers et al., 1981) and the surrounding greenstone and rhyolite of the Eastern Klamath terrane that the stock intruded into.

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The parallel veins of white quartz at Yankee John average 1.5-2 m in width with an average grade of 0.6 oz/t (Averill, 1933). They formed in the Copley greenstone. White mica, chlorite, sulfide minerals, and carbonate minerals occur as fracture fill and open space fill within the quartz vein. White mica, chlorite, and sulfide minerals preferentially occur within fractures, while carbonate minerals fill open space between quartz crystals. White mica and chlorite are commonly intergrown and the pureness of the mineral separates can be determined simply by the color; more chlorite-rich aggregates display a dark green color whereas sericite-rich aggregates have a silky white luster. Minor amounts of chlorite were likely included within the mineral separates for analysis due to the small-scale intergrowth. 4.2.8 Walker Deposit, Old Diggings District The Old Diggings (Buckeye) district produced 200,000 oz of lode gold, representing the third most productive district in the Klamath Mountains behind the French Gulch-Deadwood district and the Shasta copper-zinc-gold belt (Hotz, 1971a). This district is to the north-northeast of the Devonian Mule Mountain stock in the Eastern Klamath terrane. Much of the country rock in the district is Devonian greenstone (Clark, 1970). At least six subparallel veins are known at the Walker deposit, which strike northeast and dip 60° to the northwest (Averill, 1939). Veins can range from approximately 1 to greater than 10 m in width. Within the quartz veins, sulfide minerals and associated white mica occur in masses and as fracture fill within the quartz. The aggregates of fine-grained white mica and sulfide minerals may be completely altered slivers of host rock incorporated into the vein. The ore displays textures characteristic of significant deformation and recrystallization. Quartz is dynamically recrystallized by subgrain rotation. 4.3 Materials and Methods Quartz-carbonate vein samples and the adjacent hydrothermally altered host rock were collected from the gold deposits of the Klamath Mountains described above (Table 4.1). Initial petrographic examination aided in identifying samples with white mica of sufficient size that could be hand-picked (Fig. 4.3). These samples were crushed and sieved to 44-60 and 60-100 mesh-size fractions. Both fractions were placed in an ultrasonic bath to aid in the removal of conjoined mineral phases and then washed in distilled water and acetone to remove clay minerals and other fine particles. Individual mineral grains were then hand-picked under a binocular microscope.

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Figure 4.3: Photomicrographs in cross polarized light of hydrothermal white mica analyzed in this study. (A) Schroeder deposit. (B) McKeen deposit. (C) Yankee John deposit. (D) White mica included within and growing around pyrite crystals at the Quartz Hill deposit. (E) Hickey deposit with highly recrystallized quartz. (F) Walker deposit with highly recrystallized quartz. (G) Fine-grained mineral inclusion-rich white mica from the Washington deposit. Note the abundant mineral inclusions scattered throughout the mica. (H) McKinley deposit.

The composition of the separated white mica was determined using the U.S. Geological Survey (Denver, Colorado) JEOL 8900 Electron Microprobe with five wavelength dispersive analyzers. Operating conditions for the analysis were 15 keV accelerating voltage, a 20 nA current (measured on the Faraday cup), and a focused electron beam. The purified white mica mineral separates were dated by 40Ar/39Ar geochronology in the Western Australia Argon Isotope Facility at Curtin University (Perth, Australia). Samples were analyzed with a MAP215-50 mass spectrometer coupled with a NewWave Nd-YAG dual IR and UV laser. Fish Canyon Tuff sanidine (28.294 ± 0.37 Ma; Renne et al., 2011) was used as an age standard. 4.4 Results Analytical results from this study are listed in Tables 4.2 and 4.3, with graphical representations of mineral chemistry in Figure 4.4 and the age spectra shown in Figures 4.5 and 4.6 with 2σ uncertainties. Figure 4.4 shows that all separated crystals are near the muscovite endmember composition with minor celadonite or biotite substitution. Figure 4.5 shows data

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Figure 4.4: Mica compositions on a ternary Ca-Na-K diagram showing that the white mica crystals used in this study are nearly pure endmember muscovite.

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Figure 4.5: Age spectra with plateau results and inverse isochrons for the Yankee John, McKeen, Schroeder, and Quartz Hill deposits. These deposits all provide plateau ages and no evidence for excess argon. In the inverse isochron plots, green data points represent those that are used and are part of the age plateau, whereas blue data points are rejected.

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Figure 4.6: Age spectra and inverse isochron results for the Washington, McKinley, Hickey, and Walker deposits. The Washington deposit age spectra formed a plateau age but has evidence for excess argon. The McKinley deposit age spectra do not provide a plateau age but provides a reasonable total fusion age. The Hickey and Walker deposits have age spectra that are disrupted and do not form a plateau or inverse isochron. In the inverse isochron plots, green data points represent those that are used and are part of the age plateau, whereas blue data points are rejected.

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that resulted in an age plateau with no evidence of excess argon in the samples. Figure 4.6 shows age spectra that do not result in an age plateau and (or) show evidence for excess argon in the sample. Only the data from the heating steps that contribute to a plateau age are used for the inverse isochron diagrams; all other data points are rejected. As such, no inverse isochron was calculated for the samples that do not provide a plateau age. The age data described below represents the formation timing of white mica alteration products. This is used to constrain the timing of gold mineralization based upon the paragenetic relationship between gold mineralization and mica formation (see Chapter 3). However, it is acknowledged that potential problems arise due to the formation mechanisms of mica alteration. Since they form hydrothermally and the grains likely grow quickly, it is possible that additional phases may become included within the growing muscovite. This is noted for the sample from the Washington deposit. Additionally, it is possible that pre-existing mica from the country rock could potentially be transported into the veins and contaminate the sample. Although, the lack of a large phengitic component in the analyzed samples makes it less likely that they are metamorphic in origin and derived from the local country rocks.

Muscovite from the McKeen deposit contains between 8.18-11.01 wt.% K2O. The mineral separates provide age spectra that quickly climb to a plateau. Only the first 1.03% of cumulative 39Ar released do not form part of the age plateau; the remaining 98.97% form the plateau. This accounts for 12 of the 14 heating steps. A plateau age of 141.37 ± 0.84 Ma with a MSWD of 1.16 is interpreted to be the age of mica precipitation. Additionally, a total fusion age of 140.56 ± 0.96 Ma and an inverse isochron age of 141.64 ± 0.86 Ma were calculated and are both within uncertainty of the plateau age. The quartz veins of the Hickey deposit are the most deformed samples collected for this study. The quartz is highly deformed and recrystallized, displaying subgrain rotation recrystallization, grain boundary migration, deformation lamellae, and lobate intergrowth (Fig.

4.3). Muscovite from the Hickey deposit contains 9.51-9.76 wt.% K2O and slightly elevated

SiO2 concentrations of 48.70-51.64 wt.%. Mineral separates did not provide meaningful results. An age plateau could not be calculated, and the 19 individual steps ranged in age from 232 to 130 Ma in age. An inverse isochron age could not be calculated. A total fusion age of 214.77 ± 1.17 Ma is interpreted to be geologically meaningless due to the disturbed nature of the argon data. No meaningful age could be established for this sample.

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The Quartz Hill deposit contained muscovite with 10.40-11.43 wt.% K2O that developed an age plateau of 144.74 ± 0.74 Ma in the higher temperature steps that accounted for 86.48% of the 39Ar released with an MSWD of 1.55. This accounts for 8 of the 19 heating steps. This plateau age is interpreted as the age of mica precipitation for the sample. A total fusion age of 143.00 ± 0.76 Ma and an inverse isochron age of 144.69 ± 0.88 Ma were calculated. The inverse isochron age is within uncertainty of the plateau age whereas the total fusion age is close but outside of the 2σ uncertainty.

Muscovite from the Schroeder deposit contains 10.15-10.86 wt.% K2O. The mineral separates yielded a plateau age of 159.58 ± 0.58 Ma that is interpreted as the age of mica precipitation. The plateau corresponds to 85.22% cumulative release of 39Ar and a MSWD of 0.78. This represents ten of the 16 heating steps. A total fusion age of 160.07 ± 0.63 Ma and an inverse isochron age of 159.74 ± 1.27 were calculated and are both within uncertainty of the plateau age.

Muscovite separated from the McKinley deposit contains 10.28-10.62 wt.% K2O. The isotopic data and age spectra for muscovite did not provide an age plateau or an inverse isochron age. The ages of individual heating steps ranged from 185 Ma to 130 Ma, although 16 of 18 heating steps provided apparent ages between 162-150 Ma. A single heating step that accounts for 62.4% of the 39Ar release has an age of 155.04 ± 0.30 Ma, but was not part of three contiguous steps that accounted for greater than 50% of the 39Ar release and overlapping in age at the 2σ uncertainty. A total fusion age of 155.54 ± 0.48 is calculated, and likely approximates the age of mica formation. Muscovite from the Washington deposit is the coarsest of any samples collected in this

study and contains abundant mineral inclusions that could not be removed (Fig. 4.3). The K2O content of three spot analyses ranged from 10.16-10.51 wt.%. The mineral separates have an age plateau of 288.92 ± 0.66 Ma, with an MSWD of 1.96. This plateau accounts for 73.43% of the cumulative 39Ar released, which corresponds to eight of the 20 heating steps. An inverse isochron age of 288.24 ± 2.29 Ma was calculated and is within uncertainty of the plateau age, but the total fusion age of 284.60 ± 0.50 is outside of the 2σ uncertainty. The initial 40Ar/36Ar is 449.73 ± 392.63 compared to the expected value of atmospheric argon of 298.56. Due to the large uncertainty in the calculated 40Ar/36Ar intercept it is within error of the atmospheric value,

128 but the calculated value is still substantially higher than expected for a sample that does not contain excess argon.

Sericite from the Yankee John deposit contained 10.02-11.03 wt.% K2O and provided an age plateau of 152.49 ± 1.71 Ma, with an MSWD of 1.22, which is interpreted to be the age of mica formation. With the exception of the first two and the last heating steps, the plateau includes 92.4% of the cumulative 39Ar released during the 15 heating steps. A total fusion age of 151.86 ± 1.70 Ma and an inverse isochron age of 142.20 ± 11.71 Ma are both within uncertainty of the plateau age, although the uncertainty associated with the inverse isochron age is significantly larger. Muscovite from the Walker deposit did not provide meaningful results and was also one of the most deformed samples collected. The K2O content ranges from 9.23-9.76 wt.%. The quartz displays characteristic textures of dynamic recrystallization similar to that of the Hickey deposit (Fig. 4.3). An age plateau could not be calculated and ages of the 15 individual heating steps ranged from 336 to 265 Ma. An inverse isochron age could also not be calculated. A total fusion age of 313.36 ± 0.53 Ma was calculated, but is interpreted to be geologically meaningless given the disturbed nature of the argon data. No meaningful age interpretations can be established for this sample. 4.5 Discussion Multiple interpretations can be drawn from the new geochronological results for gold deposits in the Klamath Mountains. These range from the local-scale to the regional-scale and the deposits relationships to magmatic and tectonic events. 4.5.1 Timing of Deposit Formation Both the Schroeder and McKinley deposits are located within metabasalt to the south of, but in close proximity to the 166.9 ± 1.9 Ma (zircon U-Pb; Allen and Barnes, 2006) Vesa Bluffs pluton. Altered intrusive rocks were noted near the Schroeder deposit and in the waste rock pile, but are volumetrically minor. The mineralization age of the McKinley deposit is interpreted to be similar to the 159.58 ± 0.58 Ma age obtained for the Schroeder deposit; although no plateau was formed for the sample from the McKinley deposit. A total fusion age of ~155.5 Ma and numerous heating steps that include ages of ~160 Ma suggest a similar mineralization age. A hydrothermal event at ca. 160 Ma is significantly younger than the ca. 167 Ma age of the spatially associated Vesa Bluffs pluton. Experimental work by Cathles et al. (1997) indicated

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that a hydrothermal system with temperatures in excess of 200°C cannot be sustained for more than 800,000 years by a large, single intrusive event. Therefore, the hydrothermal event responsible for mineralization south of the Vesa Bluffs pluton and the intrusion itself are not likely genetically related. The age of ~160 Ma is identical to the onset of orogenic gold formation within the Sierra Nevada foothills in the N-S veins of the Grass Valley district (Taylor et al., 2015). The Walker and Yankee John deposits are found close to the lithologic contact of the ~400 Ma Mule Mountain stock and the surrounding greenstone of the Eastern Klamath terrane, but are located wholly within the greenstone host rocks. In addition to greenschist at Yankee John, unaltered and unmineralized porphyritic diorite clasts were found in the waste rock pile. Veins are exclusively found within the schist and show slickensides along the vein margins. Only schist country rock was noted at the Walker deposit. The age of 152.5 Ma for mica formation at the Yankee John deposit coincides with early orogenic gold formation of the E-W veins in Grass Valley and the second known orogenic gold event in the Sierra Nevada foothills province (Snow et al., 2008). The Condrey Mountain schist hosts the Quartz Hill deposit, but the deposit is located in close proximity to the Condrey Mountain thrust fault on its eastern side. The country rock schist protolith is dated at ~170 Ma, with metamorphism beginning at ~160 Ma (Saleeby and Harper, 1993). At approximately 160 Ma, thrusting of the Rattlesnake Creek terrane over the Condrey Mountain schist along the Condrey thrust fault began, with peak metamorphism along the thrust occurring at 157 +3/-2 Ma and the end of amphibolite grade metamorphism and deformation of the upper plate rocks occurring at 152 ± 1 Ma (Saleeby and Harper, 1993). However, other metamorphic ages (Rb-Sr, K-Ar) of the Condrey Mountain schist range from the Late Jurassic into the mid-Cretaceous (Helper et al., 1989; Hacker et al., 1995). The plateau age of 144.7 ± 0.7 Ma for hydrothermal muscovite is found within the broad range of metamorphic cooling ages that previous researchers have assigned for the large Condrey Mountain schist. The new age could be interpreted as a cooling age, and thus, as the minimum age of mineralization. In this scenario, the maximum age would be constrained by initial movement on the Condrey thrust and would bracket the age of mineralization between ~160-145 Ma. However, a mineralization age of 145 Ma is plausible and fits within the range of other mineralization ages constrained in this study. An age of 144.7 ± 07 Ma Ma for mineralization is also within analytical uncertainty of

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the 145.8 ± 3.0 and 147.7 ± 2.8 Ma K-Ar ages for mica formation of the Morrison-Carlock deposit in the Oro Fino district (Elder and Cashman, 1992). Of all of the deposits studied, the McKeen deposit is the only one with an age overlap with a spatially associated intrusive event. The age of mica formation at the McKeen deposit (141.37 ± 0.84 Ma) is indistinguishable from that of the Craggy Peak pluton host rock age (141 to 138.2 ± 1.3 Ma; Barnes et al., 1996; Allen and Barnes, 2006). Considering the crosscutting and brittle nature of the veins, they must have formed subsequent to emplacement and solidification of the pluton. Most of the lode deposits in this district are found within the serpentinite country rocks, although the McKeen deposit itself is located near the lithologic contact. The competency contrast between the two units was likely a preferential zone controlling fluid flow and a structural trap for mineralization to occur in the brittle intrusive rock. The Hickey deposit is located within the Siskiyou fault zone that separates the Stuart Fork and the North Fork terranes. Suturing of these two terranes is suggested to be in the Early Jurassic based upon radiolarian chert ages in the seaward North Fork terrane and volcanic rocks related to this suturing event (Irwin, 2003). The total fusion age for the Hickey deposit of 214.77 ± 1.17 Ma, along with over 90% of the argon released are in the Triassic. These ages are older than the suture zone that hosts the deposit and renders any interpretation of the data as geologically unfeasible. The ~289 Ma plateau age for the Washington deposit is pre-accretion of the Eastern Klamath terrane to the North American margin and the Permian-Triassic Sonoma Orogeny (Saleeby, 1983; Stevens et al., 1990; Wyld, 1991), and younger than the Early Devonian Copley Greenstone that is the host rock (Kinkel et al., 1956; Boucot et al., 1974; Lapierre et al., 1984). No plutonic events within the Eastern Klamath terrane correlate with the age obtained for the Washington deposit (Irwin, 1985, 2003), nor do any accretionary events (Irwin and Wooden, 1999). The early Permian McCloud limestone of the Eastern Klamath terrane suggests that a shallow carbonate platform deposited on top of an island arc was forming during tectonic subsidence at this time (Miller, 1989). This district is located to the northwest of the ~400 Ma Mule Mountain stock (Albers et al., 1981) and to the northeast of the ~136 Ma Shasta Bally batholith (Lanphere and Jones, 1978). At this point, the significance and validity of the 289 Ma age is questionable since a correlation with magmatism, metamorphism, or tectonic activity at this time cannot be established.

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The altered quartz diorite and diorite “birds eye” porphyry dikes of French Gulch have been associated with the gold-bearing quartz veins of the district and have been conjectured to be hypabyssal offshoots of the 136 Ma Shasta Bally batholith (Albers, 1965; Lanphere and Jones, 1978). Others suggested that these porphyry dikes may be older than the Shasta Bally batholith based upon deformation and intrusive contacts (Jenks and Tregaskis, 2007). Silberman and Danielson (1991) presented preliminary K-Ar geochronological data suggesting that some of the “birds eye” porphyry dikes formed around 160 Ma and some quartz porphyry dikes in the area are approximately 135 Ma. If the hydrothermal event forming the Washington deposit is responsible for the alteration of these porphyry dikes, then either the 289 Ma age must be interpreted as being inaccurate and the product of significant excess argon incorporated into the hydrothermal muscovite, or that the preliminary ages for the porphyry dikes are incorrect, or both. The calculated initial 40Ar/36Ar value suggests that abundant excess argon is found within the analyzed hydrothermal muscovite, providing additional evidence for the plateau age being inaccurate. The true mineralization age is likely much younger. The homogenously distributed mineral inclusions found within the coarse hydrothermal mica might contain significant contributions to the total argon that may have shifted the age beyond the formation age of the white mica. Alternatively, the excess argon could also be housed within the crystal structure of the mica. Incorporating excess argon to this extent is plausible and has been displayed in previous studies. Within hydrothermal environments related to metamorphic fluids, excess argon can increase plateau ages by more than 100 Myr (e.g., Cumbest et al., 1994). Even within magmatic- hydrothermal environments, the fluids may add substantial levels of excess argon into hydrothermal phases (e.g., Kendrick et al., 2001a, b). 4.5.2 Timing of Magmatism and Mineralization Provincially, syn- to post-accretionary plutons of the Klamath Mountains intruded mostly between 170-140 Ma, but the number of dated plutons peaked around 170-160 Ma. Numerous dikes that are interpreted to be derived from the larger intrusive bodies (Hotz, 1971a; Silberman and Danielson, 1991) are found in proximity to many gold deposits. The close spatial relationship between many gold deposits and these granitic intrusions has led to models suggesting a genetic relationship between the two (e.g., Clark, 1970; Hotz, 1971a). However,

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this new age data makes a direct genetic connection between magmatism and gold mineralization tenuous. The time period that gold mineralization occurred within the Klamath Mountains overlaps with the first period of post-accretionary magmatism within both the Klamath Mountains and the Sierra Nevada. However, the McKeen deposit is the only deposit from this study with an overlap in age of local magmatism and mica formation, both occurring towards the end of the gold mineralization window in the Klamath Mountains. All of the other sampled deposits may sit adjacent to plutons, but with age discrepancies between magmatism and mineralization. Previous geochronological and geochemical studies of the gold deposits in the Sierra Nevada foothills have shown that they are not genetically related to magmatism and are instead related to faults and tectonic activity (e.g., Marsh et al., 2008; Taylor et al., 2015). The gold deposits in the Klamath Mountains also fit this model of being orogenic gold deposits that are not genetically related to magmatism. Broadly, there is no correlation between magma influx with gold mineralization in California (Fig. 4.7). However, an interesting pattern regarding widespread occurrences of Late Jurassic plutonism and gold mineralization does exist. Grass Valley and the Klamath Mountains which have Late Jurassic through earliest Cretaceous orogenic gold mineralization are also regions that contain Middle to Late Jurassic plutons, both indicative of high regional heat flow. The regions of younger orogenic gold mineralization, such as the Bagby, Confidence, and Coulterville districts, and the Mother Lode Belt, are areas that lack significant Late Jurassic-Cretaceous plutonism (Fig. 4.1). This spatial and broad temporal relationship between early magmatism and gold mineralization should not be overlooked. While the Klamath Mountains and the Sierra Nevada formed a single contiguous arc, it was the northern and central portion of this arc that was magmatically and hydrothermally active. This broad spatial and temporal link between magmatism and orogenic gold mineralization has been noted throughout the world (Goldfarb et al., 2001, 2005). This relationship within the orogeny continued until the Early Cretaceous lateral offset of the Klamath Mountains from the Sierra Nevada and the active arc. This correlation between magmatism and hydrothermal activity is not evident post-separation of the Klamath Mountains and the Sierra Nevada. The magmatism that is generally slightly older than the hydrothermal events would have added extra heat to the crust which would have aided in

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metamorphism and devolatilization as orogenesis was occurring. An offset in ages, with older ages of peak metamorphism of host rocks and younger ages of hydrothermal events, is a common to ubiquitous feature of orogenic gold deposits (Goldfarb et al., 2005). 4.5.3 Tectonic and Metallogenic Relationships Most of the ages from this study correspond to the Late Jurassic through earliest Cretaceous, signifying a major gold event within California at this time. The earliest known gold-forming hydrothermal event in California occurred in Grass Valley at ~160 Ma after final terrane amalgamation in the Sierra Nevada and Klamath Mountains, and within a compressional tectonic regime (Table 4.4, Fig. 4.7; Taylor et al., 2015). This corresponds to the J2 cusp and a major shift in the plate motions of the North American and Farallon plates (Engebretson et al., 1985; Beck and Housen, 2003). This change in far-field stresses resulting from plate reorganization was interpreted to mark the beginning of orogenic gold formation in the Sierra Nevada foothills (Taylor et al., 2015). This same tectonic trigger is seemingly responsible for initial orogenic gold formation in the Klamath Mountains (Schroeder deposit, 159.58 ± 0.58 Ma). Sinistral movement along the Wolf Creek fault zone, one of the major faults in the Sierra Nevada foothills, was coincident with the next major gold event in the Sierra Nevada foothills province at ~152 Ma (Snow et al., 2008; Taylor et al., 2015). This event is noted as far south as the ~151 Ma syntectonic emplacement of the Guadalupe igneous complex in the southernmost Sierra Nevada foothills (Vernon et al., 1989) and about 80 km further south in the Owens Mountain area as sinistrally-sheared syntectonic dike swarms that were emplaced between 155 and 148 Ma (Wolf and Saleeby, 1992), all of which indicates an extensive area that was affected by this strike-slip movement. This change from a more compressional regime to a more transcurrent deformation in the Late Jurassic may also reflect a broader change in far-field stresses. The second orogenic gold hydrothermal event in the Sierra Nevada foothills also coincided with the second event in the Klamath Mountains with the formation of the Yankee John deposit at 152.5 ± 1.7 Ma. Prior to this study, the only other known age constraints on orogenic gold mineralization in the Klamath Mountains were the 147.7 ± 2.8 to 145.8 ± 3.0 Ma K-Ar ages obtained on hydrothermal sericite from the Morrison-Carlock deposit in the Oro Fino district (Elder and Cashman, 1992). These ages match up within uncertainty of the Quartz Hill deposit of the Scott

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Figure 4.7: Diagram showing a timeline of events from the Middle Jurassic through Early Cretaceous of California. The apparent intrusive magma flux of the Sierra Nevada is shown as the red line (from Ducea, 2001). The light yellow boxes represent periods during which gold mineralization is permissible. The darker yellow boxes represent times when alteration products related to gold mineralization have been dated. The age range that is permissible for gold mineralization in the Klamath Mountains ranges from the beginning of orogenic gold mineralization in California at 160 Ma until the lateral offset of the Klamath Mountains from the active arc beginning around 140 Ma.

Bar district dated in this study at 144.7 ± 0.7 Ma. Mineralization at the McKeen deposit in the Callahan district closely followed these events at 141.4 ± 0.8 Ma. Sinistral movement along the terrane bounding faults in the Klamath Mountains and the Sierra Nevada foothills continued during this period. All of the dated deposits in the Klamath Mountains fall within the range of initial development of these hydrothermal systems and pre-date the Early Cretaceous separation of the Klamath Mountains from the Sierra Nevada (Fig. 4.8). This window lasted from approximately 160 Ma to 140 Ma. At the lower time limit of this range, the Klamath Mountains were laterally translated westward away from the tectonically active arc. Thus, magmatism and orogenic gold formation was terminated along with deep-seated tectonic movement along the terrane bounding faults as they were translated away from the active arc. Although the Klamath Mountains were translated westward away from the active arc and the Sierra Nevada, sinistral movement along the terrane-bounding faults of the Sierra Nevada foothills continued until the switch to dextral movement was initiated in the Early Cretaceous

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Figure 4.8: Time slices of northern California and southern Oregon. Pluton emplacement within the accreted terranes occurs mostly between ~160-140 Ma. Gold mineralization within the Klamath Mountains occurs between ~160-140 Ma and within the Sierra Nevada between ~160- 115 Ma. The lateral offset of the Klamath Mountains away from the Sierra Nevada and the active arc began around ~140-135 Ma. Age data used for the gold deposits is listed in Table 4.4. Locations for the Klamath Mountains deposits are shown in Figure 4.2. Enlarged geology maps are given in Figures 4.1 and 4.2. The outline of the current border of northern California in relation to the location of the Sierra Nevada foothills is used for reference. Maps are modified from Irwin (2003) and Ernst et al. (2008).

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(Glazner, 1991). The majority of the gold deposits in the Mother Lode belt in the southern Sierra Nevada foothills mostly formed between ~130-125 Ma and may be associated with the transition from sinistral to dextral movement on the terrane bounding faults at ~125 Ma (Table 4.3 and Fig. 4.8; Marsh et al., 2008). This ca. 125 Ma event has been correlated with a major shift in far-field stresses due to the formation of the Ontong-Java plateau and the subsequent reorganization of plates in the Pacific basin (Goldfarb et al., 2007). The youngest (ca. 115 Ma) dated orogenic gold deposits in California occur in the Bagby and Alleghany districts of the Sierra Nevada foothills province. At this time, another change in plate motion occurred with the Pacific plate changing from northwest to west-southwest movement, the Farallon plate changing from south-southeast to north movement, and North America staying relatively constant (Engebretson et al., 1985). The age range from ~160-115 Ma for formation of orogenic gold deposits in the Sierra Nevada foothills coincides with the range in age for ductile deformation along the terrane bounding faults of the Sierra Nevada foothills. The age range from ~160-140 Ma represents the permissible time for orogenic gold formation in the Klamath Mountains during this ductile deformation along the regional scale faults prior to the westward translation of the Klamath Mountains away from the active arc (Fig. 4.8). Tobisch et al. (1989) found that ductile deformation along these faults was active from roughly 160 to at least 123 Ma, with some secondary structures possibly being younger. However, the episodic mineralization within this time period indicates that specific tectonic triggers are required for the formation of orogenic gold deposits within a favorable structural regime. Taking all of the gold deposits in the Klamath Mountains and the Sierra Nevada foothills into account, it is clear that gold formation is directly linked to changes in far-field stress that acted as triggers for fault activation and reactivation, fluid flow, and gold precipitation (Fig. 4.7). Orogenic gold deposition is related to these tectonic triggers and not genetically related to magmatism. The observed general temporal overlap between gold mineralization and magmatic activity is related to the fact that the area was undergoing subduction and was typified by a high geothermal gradient leading to regional metamorphism and partial melting of the crust. 4.6 Conclusions Hydrothermal white mica that is an alteration product associated with the fluid responsible for gold mineralization was separated from multiple lode gold deposits in the Klamath

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Mountains, California. These white mica separates are of the muscovite endmember with variable amounts of minor biotite and celadonite substitution. Of the eight deposits analyzed during this study, three had disrupted or invalid age spectra resulting from excess argon due to subsequent deformation or abundant mineral inclusions. One other sample did not provide a plateau, but provided a reasonable integrated age within the permissible age range for mineralization. The other four samples provided plateau ages between ~160-140 Ma. This new data indicates that this 20 Myr window is the range of permissible ages for gold mineralization in the Klamath Mountains. The age of 160 Ma represents the oldest orogenic gold deposits formed in California, both in the Klamath Mountains (this study) and the Sierra Nevada foothills (Grass Valley; Taylor et al., 2015). The youngest dated deposit in the Klamath Mountains, the ~141 Ma McKeen deposit, formed just before the lateral translation of the Klamath Mountains away from the active arc between ~140-136 Ma. This marks the time when magmatism and gold formation was terminated. There is an overall discrepancy in age of individual deposits and that of local intrusive phases. All but one of the deposits from this study have substantial age differences with local plutons or dikes, as can be shown with published ages and geologic relationships. The McKeen deposit and the hosting Craggy Peak pluton are indistinguishable in age, both having formed near the end of when the Klamath Mountains were a part of the active arc. However, this age overlap does not indicate a genetic relationship between the two; instead, an elevated geothermal gradient led to hydrothermal fluid circulation and magmatism in a localized environment which has been documented in other isolated environments of California and in other orogenic gold provinces. No clear correlation between magmatism and formation of these lode gold deposits exists, whether considered at the local or regional scale. Alternatively, a correlation exists with far-field stress changes, deformation and fault movement, and ages of deposits in the Sierra Nevada foothills gold province. Initial gold-forming hydrothermal activity in both the Klamath Mountains and the Sierra Nevada foothills occurred at ~160 Ma, coincident with a major change in far-field stresses that resulted in changes in motion of the Farallon and North American plates. The second gold-forming event occurred 8 Myr later at ~152 Ma in both the Klamath Mountains and the Sierra Nevada foothills. This event and the remainder of the gold deposits in the Klamath Mountains formed during a transcurrent tectonic regime and sinistral shearing of the

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terrane-bounding faults after the more compressional tectonic regime found during the first gold event. At ~140 Ma, orogenic hydrothermal activity in the Klamath Mountains ceased and the province was translated westward. A hiatus of gold mineralization in the Sierra Nevada foothills also commenced at this time, but only lasted ~5 Myr. Additional gold mineralization lasted from ~135-115 Ma, with formation of the Mother Lode belt coincident with the switch from sinistral to dextral motion along the terrane-bounding faults at ~125 Ma. The youngest orogenic gold deposits in California, in the Alleghany and Bagby districts, formed during a period of far-field stress changes of the motions of the Pacific and Farallon plates. 4.7 References Albers, J.P., 1965, Economic geology of the French Gulch quadrangle, Shasta, Trinity Counties, California: California Division of Mines and Geology, Special Report 85, 43 p. Albers, J.P., Kistler, R.W., and Kwak, L., 1981, The Mule Mountain Stock, an early Middle Devonian pluton in northern California: Isochron/West, v. 31, p. 17. Allen, C.M., and Barnes, C.G., 2006, Ages and some cryptic sources of Mesozoic plutonic rocks in the Klamath Mountains, California and Oregon, in Snoke, A.W., and Barnes, C.G., eds., Geological studies in the Klamath Mountains province, California and Oregon: A volume in honor of William P. Irwin: Geological Society of America Special Paper 410, p. 223– 245. Averill, C.V., 1931, Redding Field Division: California Journal of Mines and Geology, v. 27, 582 p. Averill, C.V., 1933, Gold Deposits of the Redding and Weaverville Quadrangles: California Journal of Mines and Geology, v. 29, p. 57–58. Averill, C.V., 1935, Mines and Mineral Resources of Siskiyou County: California Journal of Mines and Geology, v. 35, p. 154–159. Averill, C.V., 1939, Mineral Resources of Shasta County, Califonria: California Journal of Mines and Geology, v. 31, p. 287–288. Barnes, C.G., Petersen, S.W., Kistler, R.W., Murray, R., and Kays, M.A., 1996, Source and tectonic implications of tonalite-trondhjemite magmatism in the Klamath Mountains: Contributions to Mineralogy and Petrology, v. 123, p. 40–60. Barrow, W.M., and Metcalf, R.V., 2006, A reevaluation of the paleotectonic significance of the Paleozoic Central Metamorphic terrane, eastern Klamath Mountains, California: New constraints from trace element geochemistry and 40Ar/39Ar thermochronology, in Snoke, A.W., and Barnes, C.G., eds., Geological studies in the Klamath Mountains province, California and Oregon: A volume in honor of William P. Irwin: Geological Society of America Special Paper 410, p. 393–410. Beck, M.E., Jr., and Housen, B.A., 2003, Absolute velocity of North America during the Mesozoic from paleomagnetic data: Tectonophysics, v. 377, p. 33–54.

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Böhlke, J.K., Kirshbaum, C., and Irwin, J.J., 1989, Simultaneous analyses of noble gas isotopes and halogens in fluid inclusions in neutron-irradiated quartz veins using a laser microprobe noble gas mass spectrometer: U.S. Geological Survey Bulletin 1890, p. 61–80. Boucot, A.J., Dunkle, D.H., Potter, A., Savage, N.M., and Rohr, D., 1974, Middle Devonian orogeny in western North America: A fish and other fossils: Journal of Geology, v. 82, p. 691–708. Cathles, L.M., Erendi, A.H.J., and Barrie, T., 1997, How long can a hydrothermal system be sustained by a single intrusive event?: ECONOMIC GEOLOGY, v. 92, p. 766–771. Clark, W.B., 1970, Gold districts of California: California Division of Mines and Geology, Bulletin 193, 186 p. Craig, J.R., and Rimstidt, J.D., 1998, Gold production history of the United States: Ore Geology Reviews, v. 13, p. 407–464. Cumbest, R.J., Johnson, E.L., and Onstott, T.C., 1994, Argon composition of metamorphic fluids: Implications for 40Ar/39Ar geochronology: Geological Society of America Bulletin, v. 106, p. 942–951. Dickinson, W.R., 2000, Geodynamic interpretation of Paleozoic tectonic trends oriented oblique to the Mesozoic Klamath-Sierran continental margin in California, in Soreghan, M.J., and Gehrels, G.E., eds., Paleozoic and Triassic paleogeography and tectonics of western Nevada and northern California: Geological Society of America Special Papers 347, p. 209–246. ——, 2004, Evolution of the North American Cordillera: Annual Review of Earth and Planetary Sciences, v. 32, p. 13–45. ——, 2008, Accretionary Mesozoic-Cenozoic expansion of the Cordilleran continental margin in California and adjacent Oregon: Geosphere, v. 4, p. 329–353. Ducea, M.N., 2001, The California arc: Thick granitic batholiths, eclogitic residues, lithospheric- scale thrusting, and magmatic flare-ups: Geological Society of America Today, v. 11, p. 4– 10. Elder, D., and Cashman, S.M., 1992, Tectonic control and fluid evolution in the Quartz Hill, California, lode gold deposits: ECONOMIC GEOLOGY, v. 87, p. 1795–1812. Engebretson, D.C., Cox, A., and Gordon, R.G., 1985, Relative motions between oceanic and continental plates in the Pacific Basin: Geological Society of America Special Paper 206, p. 1–60. Ernst, W.G., 2013, Earliest Cretaceous Pacificward offset of the Klamath Mountains salient, NW California-SW Oregon: Lithosphere, v. 5, p. 151–159. Ernst, W.G., Snow, C.A., and Scherer, H.H., 2008, Contrasting early and late Mesozoic petrotectonic evolution of northern California: Geological Society of America Bulletin, v. 120, p. 179–194. Ferrero, T., 1990, The Liberty gold mining district: California Geology, v. 43, p. 123–133. Glazner, A.F., 1991, Plutonism, oblique subduction, and continental growth: An example from the Mesozoic of California: Geology, v. 19, p. 784–786.

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Goldfarb, R.J., Groves, D.I., and Gardoll, S., 2001, Orogenic gold and geologic time: a global synthesis: Ore Geology Reviews, v. 18, p. 1–75. Goldfarb, R.J., Baker, T., Dubé, B., Groves, D.I., Hart, C.J.R., and Gosselin, P., 2005, Distribution, character, and genesis of gold deposits in metamorphic terranes: ECONOMIC GEOLOGY 100TH ANNIVERSARY VOLUME, p. 407–450. Goldfarb, R.J., Hart, C., Davis, G., and Groves, D., 2007, East Asian gold: Deciphering the anomaly of Phanerozoic gold in Precambrian cratons: ECONOMIC GEOLOGY, v. 102. p. 341–345. Hacker, B.R., Donato, M.M., Barnes, C.G., McWilliams, M.O., and Ernst, W.G., 1995, Timescales of orogeny: Jurassic construction of the Klamath Mountains: Tectonics, v. 14, p. 677–703. Helper, M.A., Walker, N.W., and McDowell, F.W., 1989, Early Cretaceous metamorphic ages and Middle Jurassic U-Pb zircon protolith ages for the Condrey Mountain Schist, Klamath Mtns [abs.]: Geological Society of America Abstracts with Programs, Spokane, WA, v. 21, no. 5, p. 92. Hotz, P.E., 1971a, Geology of lode gold districts in the Klamath Mountains, California and Oregon: U.S. Geological Survey Bulletin 1290, 91 p. Hotz, P.E., 1971b, Plutonic rocks of the Klamath Mountains, California and Oregon: U.S. Geological Survey Professional Paper 684-B, 19 p. Irwin, W.P., 1985, Age and tectonics of plutonic belts in accreted terranes of the Klamath Mountains, California and Oregon, in Howell, D.G. (ed.), Tectonostratigraphic Terranes of the Circum-Pacific Region, Earth Science Series, Number 1: Circum-Pacific Council for Energy and Mineral Resources, p. 187–199. Irwin, W.P., 2003, Correlation of the Klamath Mountains and the Sierra Nevada: U.S. Geological Survey Open-File Report 02-490, 2 sheets. Irwin, W.P., and Wooden, J.L., 1999, Plutons and accretionary episodes of the Klamath Mountains, California and Oregon: U.S. Geological Survey, Open-File Report 99-374, 1 sheet. Irwin, W.P., and Wooden, J.L., 2001, Map showing plutons and accreted terranes of the Sierra Nevada, California, with a tabulation of U/Pb isotopic ages: U.S. Geological Survey, Open-File Report 01-229, 1 sheet. Jenks, J., and Tregaskis, S., 2007, Technical report: The Washington-Niagara mine French Gulch project: Written for Bullion River Gold Corp., 69 p. Kendrick, M.A., Burgess, R., Pattrick, R.A.D., and Turner, G., 2001a, Fluid inclusion noble gas and halogen evidence on the origin of Cu-porphyry mineralising fluids: Geochimica et Cosmochimica Acta, v. 65, p. 2651–2668. Kendrick, M.A., Burgess, R., Pattrick, R.A.D., and Turner, G., 2001b, Halogen and Ar-Ar age determinations of inclusions within quartz veins from porphyry copper deposits using complementary noble gas extraction techniques: Chemical Geology, v. 177, p. 351–370.

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Kinkel., A.R.J., Hall, W.E., and Albers, J.P., 1956, Geology and base metal deposits of west Shasta copper-zinc district, Shasta County, California: U.S. Geological Survey Professional Paper 285, 156 p. Lanphere, M.A., and Jones, D.L., 1978, Cretaceous time scale from North America: The geologic time scale: AAPG Studies Geology 6, p. 259–268. Lanphere, M.A., Irwin, W.P., and Hotz, P.E., 1968, Isotopic age of the Nevadan Orogeny and older plutonic and metamorphic events in the Klamath Mountains, California: Geological Society of America Bulletin, v. 79, p. 1027–1052. Lapierre, H., Albarede, F., Albers, J., Cabanis, B., and Coulon, C., 1984, Early Devonian volcanism in the eastern Klamath Mountains, California: evidence for an immature island arc: Canadian Journal of Earth Science, v. 22, p. 214–227. Logan, C.A., 1925, Sacramento Field Division: California Journal of Mines and Geology, v. 21, no. 4, 624 p. Marsh, E.E., Goldfarb, R.J., Kunk, M.J., Groves, D.I., Bierlein, F.P., and Creaser, R.A., 2008, New constraints on the timing of gold formation in the Sierra Foothills province, central California, in Spencer, J.E., and Titley, S.R., eds., Ores and orogenesis: Circum-Pacific tectonics, geologic evolution, and ore deposits: Arizona Geological Society Digest 22, p. 369–388. Miller, M.M., 1989, Intra-arc sedimentation and tectonism: Late Paleozoic evolution of the eastern Klamath terrane, California: Geological Society of America Bulletin, v. 101, p. 170–187. Obrien, J.C., 1947, Mines and Mineral Resources of Siskiyou County, California: California Journal of Mines and Geology, v. 43, no. 4, p. 447–448. Renne, P.R., Balco, G., Ludwig, K.R., Mundil, R., and Min, K., 2011, Response to the comment by W.H. Schwarz et al. on “Joint determination of 40K decay constants and 40Ar*40K for the Fish Canyon sanidine standard, and improved accuracy for 40Ar/39Ar geochronology” by P.R. Renne et al. (2010): Geochimica et Cosmochimica Acta, v. 75, p. 5097–5100. Saleeby, J.B., 1983, Accretionary tectonics of the North American Cordillera: Annual Review of Earth and Planetary Sciences, v. 15, p. 45–73. Saleeby, J.B, and Harper, G.D., 1993, Tectonic relations between the Galice Formation and the Condrey Mountain Schist, in Dunne, G.C., and McDougall, K.A., eds., Mesozoic paleogeography of the western United States—II: Pacific Section, Society for Sedimentary Geology, Book 71, p. 197–225. Scherer, H.H., Snow, C.A., and Ernst, W.G., 2006, Geologic-petrochemical comparison of early Mesozoic mafic arc terranes—Western Paleozoic and Triassic belt, Klamath Mountains, and Jura-Triassic arc belt, Sierran Foothills, in Snoke, A.W., and Barnes, C.G., eds., Geological studies in the Klamath Mountains province, California and Oregon—A volume in honor of William P. Irwin: Geological Society of America Special Paper 410, p. 377– 392.

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Sharp, W.E., 1988, Pre-Cretaceous crustal evolution of the Sierra Nevada region, in Ernst, W.G., ed., Metamorphism and crustal evolution of the western United States: Englewood Cliffs, New Jersey, Prentice-Hall, p. 824–864. Shasta Gold Corp, 2014, http://www.shastagoldcorp.com/?page_id=16, accessed September, 2014. Silberman, M.L., and Danielson, J., 1991, Geologic setting, characteristics, and geochemistry of gold-bearing quartz veins in the Klamath Mountains in the Redding 1 degree x 2 degree quadrangle, northern California: U.S. Geological Survey Open-File Report 91-595, 24 p. Silberman, M.L., and Danielson, J., 1993, Gold-bearing quartz veins in the Klamath Mountains in the Redding 1 x 2 degree quadrangle northern California: California Geology, v. 46, p. 35–44 Snow, C.A., and Scherer, H., 2006, Terranes of the Western Sierra Nevada Foothills metamorphic belt, California: A critical review: International Geology Review, v. 48, p. 46–62. Snow, C.A., Bird, D.K., Metcalf, J., and McWilliams, M., 2008, Chronology of gold mineralization in the Sierra Nevada Foothills from 40Ar/39Ar dating of mariposite: International Geology Review, v. 50, p. 503–518. Stevens, C.H., Yancey, T.E., and Hanger, R.A., 1990, Significance of the provincial signature of early Permian faunas of the eastern Klamath terrane: Geological Society of America Special Papers 255, p. 210–218. Taylor, R.D., Goldfarb, R.J., Monecke, T., Fletcher, I.R., Cosca, M.A., and Kelly, N.M., 2015, Application of U-Th-Pb phosphate geochronology to young orogenic gold deposits: New age constraints on the formation of the Grass Valley gold district, Sierra Nevada Foothills province, California: ECONOMIC GEOLOGY, v. 110, p. 1313–1337. Tobisch, O.T., Paterson, S.R., Saleeby, J.B, and Geary, E.E., 1989, Nature and timing of deformation in the Foothills terrane, central Sierra Nevada, California: Its bearing on orogenesis: Geological Society of America Bulletin, v. 101, p. 401–413. Vernon, R.H., Paterson, S.R., and Geary, E.E., 1989, Evidence for syntectonic intrusion of plutons in the Bear Mountains fault zone, California: Geology, v. 17, p. 723–726. Wolf, M.B., and Saleeby, J.B., 1992, Jurassic Cordilleran dike swarms-shear zones: Implications for the Nevadan orogeny and North American plate motion: Geology, v. 20, p. 745–748. Wyld, S.J., 1991, Permo-Triassic tectonism in volcanic arc sequences of the western U.S. Cordillera and implications for the Sonoma Orogeny: Tectonics, v. 10, p. 1007–1017.

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CHAPTER 5 CONCLUSIONS

This thesis represents a comprehensive examination of the Grass Valley orogenic gold district in the Sierra Nevada foothills, California. The new data constrain the timing of mineralization, the type of mineralization, the tectonic environment and evolution of mineralization, the paragenetic sequence of gold mineralization, and the relationship of gold mineralization throughout California. The Grass Valley gold district is the historically most productive gold district of the western cordillera of North America, having yielded 13 Moz of lode gold over its near 100 years of mining activity. The timing of major magmatic and hydrothermal events in Grass Valley has been constrained by various geochronometers in this study. 1) Zircon U-Pb ages of the ore-hosting Grass Valley granodiorite indicate intrusion at 162- 160 Ma. Multiple geothermobarometers indicate emplacement at temperatures of nearly 800 ° C and crystallization at paleodepths of approximately 3 km. The greenschist-facies country rocks that the granodiorite intruded were already on a retrograde cooling path by this time. 2) Overlapping U-Pb zircon and 40Ar/39Ar cooling ages of hornblende and biotite indicate a rapid cooling of the pluton to temperatures below 300 ° C. Samples for argon geochronology were taken from both the northern and the southern portion of the granodiorite, with cooling ages ranging from approximately 162-160 Ma. 3) The age of gold-bearing quartz-carbonate veins is indistinguishable from that of the intrusive age. Hydrothermal xenotime and monazite formed as a result of interaction between the hydrothermal fluid and the wall rock. Xenotime crystals provide a U-Pb formation age of 162 ± 5 Ma, indicating that the N-S veins formed at a maximum of 5 Myr after pluton solidification; they may have formed even closer to the time of granodiorite intrusion. A second vein set, the E-W veins, formed 5-10 Myr after formation of the N-S veins. 4) The chemistry of the vein-hosted phosphate crystals was also utilized to ensure that the dated crystals are hydrothermal precipitates and not xenocrysts derived from the igneous host rocks. The vein-hosted phosphates have REE and trace geochemical characteristics

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that distinguish them from magmatic phosphate phases. The U and Th contents of magmatic monazite are markedly higher than what characterizes the hydrothermal phases. Additionally, the lack of a negative Eu anomaly for both the hydrothermal monazite and xenotime contrasts with a strong negative anomaly for the magmatic monazite. The trace element geochemistry of the hydrothermal phosphates confirms that the gold-bearing quartz-carbonate veins are typical orogenic veins, and not genetically related to magmatic activity. If the vein-hosted phosphates were a product of an evolving magmatic- hydrothermal system, then they would have displayed prominent negative Eu anomalies and distinctly different REE trends. Orogenic gold veins in deposits worldwide are characteristically deformed and recrystallized because of their environment of formation. Continued fault movement typically causes both brittle and ductile deformation of the veins after their formation and extensive recrystallization of the vein minerals, especially the hydrothermal quartz. However, primary relationships in the veins at Grass Valley are notably well preserved. This atypical feature makes them ideal for microanalytical studies identifying vein paragenesis. 1) Orogenic veins typically display brittle deformation features such as cataclastic brecciation. Common ductile textures in vein quartz include evidence for grain boundary migration, undulose extinction, subgrain formation, and deformation lamellae. Softer minerals such as galena and carbonate minerals may commonly display bent and distorted cleavage. Recrystallization may cause metal redistribution and destroys chemical zoning evident in minerals such as pyrite. However, the veins of Grass Valley do not display these features. Their preservation may be a product of their formation at relatively shallow crustal depths. 2) The lack of deformation has preserved textures that allow an interpretation of the paragenetic sequence. The most important observation is that gold does not coprecipitate with quartz, and only precipitates at a specific period during pyrite crystallization. Quartz is the first mineral to form in the veins and forms during multiple generations of growth independent of any other mineral. Gold mineralization occurs broadly coevally with galena and sphalerite, all of which are related to a late growth banding in pyrite that is rich in As and elemental enrichments related to the interaction of the orogenic gold fluids with the local host rocks.

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3) The most important mechanisms for vein formation are pressure fluctuations and the interaction of the orogenic fluids with the wall rocks. Major fault movement opens space within the veins. This rapid drop in pressure leads to infilling and crystallization of quartz due to adiabatic decompression. Periods of increased equilibration between the fluid and the country rock are associated with gold precipitation, as this event is marked by enrichments in elements from the fluid (As, Ag, and Au) and from the local host rock (Co, Ni, and Pb). With the knowledge of how and when the veins of Grass Valley formed, it is possible to fit their formation into a temporal context over the major gold provinces of California. Gold mineralization in the Sierra Nevada foothills is now known to have formed during 160-115 Ma. The deposits in Grass Valley are the oldest deposits to have formed in this gold province. But this entire age span is not realistic for the deposits in the Klamath Mountains. New argon isotopic analyses of multiple deposits in the Klamath Mountains constrain the timing of mineralization in this important gold province and how it relates to Grass Valley and other deposits of the Sierra Nevada foothills. 1) Eight deposits from throughout the Klamath Mountains were selected that had suitably sized white mica alteration products. Of these, only four produced age plateaus with no evidence of excess argon. The other four samples contained disrupted age spectra most likely resulting from intense deformation after formation. One of these samples also contained abundant mineral inclusions that may have elevated the radiogenic argon component and resulted in an excessively old age. 2) The plateau ages ranged from ~160-140 Ma, and represent the entire span of ages that are permissible for gold mineralization in the Klamath Mountains. The oldest deposit from this study corresponds to the oldest dated deposit in the Sierra Nevada and when orogenic gold deposit formation in California was initiated. The youngest deposit from this study is just older than the interpreted age of separation of the Klamath Mountains from the Sierra Nevada. This age marks the time when magmatic and hydrothermal activity within the Klamath Mountains ceased. Although this research provides a detailed examination of the Grass Valley gold deposits and how they relate to other orogenic gold deposits in California, there is still additional work

147 that could further our understanding of the ore forming processes. Some recommendations for future work include: 1) This study has shown that quartz does not coprecipitate with gold and fluid inclusions found within quartz do not represent the fluid responsible for gold precipitation. However, sphalerite forms coevally with gold. If primary fluid inclusions can be found within sphalerite, then these would represent the fluids within the system closer to the time of gold mineralization. Examination of these primary fluid inclusions would be the first study to document the fluids that are paragenetically related to gold mineralization. 2) The deposits of Grass Valley represent minimally deformed and recrystallized orogenic gold veins with a preserved paragenetic sequence visible in the veins. A comparison with some examples of typical deposits that are visibly deformed and recrystallized will show how continued deformation may remobilize certain metals and alter the perceived paragenetic sequence. Deposits in California display a wide range of deformation and recrystallization that can be used to document these effects. 3) A portion of this study has been able to reconstruct the evolution of California during the formation of Jurassic-Cretaceous orogenic gold deposits. However, a complete tectonic reconstruction pertaining to the formation of all of the Mesozoic orogenic gold deposits around the Pacific Rim has not been undertaken. Coupling the findings from this study with the known ages of deposits in British Columbia, Alaska, Russia, China, and New Zealand would provide insight into the mechanisms of gold formation in relation to far- field stresses such as plate reorganization.

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APPENDIX SUPPLEMENTAL ELECTRONIC FILES

A list of supplemental material that supports the thesis work at Grass Valley is summarized in the table below. The files include abstracts, extended abstracts, and posters that have been presented at various conferences. They are listed in chronological order.

File Name Conference and File Type

GSN 2010 abs.doc Geological Society of Nevada 2010 Symposium - Abstract GSN 2010 poster.pdf Geological Society of Nevada 2010 Symposium - Poster SEG 2010 ext abs.doc Society of Economic Geology 2010 Keystone Conference, Keystone, Colorado – Extended Abstract SEG 2010 poster.pdf Society of Economic Geology 2010 Keystone Conference, Keystone, Colorado - Poster GSA 2011 abstract.doc Geological Society of America 2011, Minneapolis, Minnesota - Abstract GSA 2011 poster.pdf Geological Society of America 2011, Minneapolis, Minnesota - Poster SEG 2012 abstract.docx Society of Economic Geologist 2012 Conference, Lima, Peru - Abstract SEG 2012 poster.pdf Society of Economic Geologist 2012 Conference, Lima, Peru - Poster SEG 2013 abstract.docx Society of Economic Geology 2013 Conference, Whistler, BC - Abstract SEG 2013 poster.pdf Society of Economic Geology 2013 Conference, Whistler, B.C. - Poster GSA 2014 abstract.docx Geological Society of America 2014, Vancouver, B.C. - Abstract SGA 2015 ext abs.doc Mineral resources in a sustainable world, SGA 2015 Proceedings, Nancy, France – Extended Abstract

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