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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
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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 orehosting 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 orehosting 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 brittleductile 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.
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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
2.6
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
40
2.5.7
Hornblende Geobarometry and Plagioclase-Amphibole Geothermometry ......25
Ar/39Ar Geochronology ..................................................................................24
Results .........................................................................................................................25
2.6.1
Petrography and Geochemistry of the Grass Valley Granodiorite ....................25
v
2.7
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
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
3.5
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
Discussion....................................................................................................................88
3.5.1
Deformation of Vein Minerals ..........................................................................88
3.5.2
Paragenesis ........................................................................................................90
3.5.3
Timing of Gold Deposition ...............................................................................94
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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
4.1
4.2
40
AR/39AR GEOCHRONOLOGY OF HYDROTHERMAL ACTIVITY
RELATED TO GOLD MINERALIZATION IN THE KLAMATH
MOUNTAINS, CALIFORNIA ...........................................................................106
Background................................................................................................................106
4.1.1
Tectonic Setting...............................................................................................107
4.1.2
Klamath Mountains Geology ..........................................................................109
4.1.3
Gold Deposits ..................................................................................................110
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
vii
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
40
Ar/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
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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
x
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
xi
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
xii
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
xiii
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
xiv
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 NNWtrending 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: ~170140 Ma and ~120-80 Ma.
1
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).
2
3
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 JuraTriassic 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.
4
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
5
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-inhornblende 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 magmatichydrothermal 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
6
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-inhornblende 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.
7
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: CircumPacific 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.
8
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.
9
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.
10
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. Goldbearing 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 modernday 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).
11
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
12
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; SPFSpencerville 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.
13
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 JuraTriassic 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 ~170140 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
14
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
15
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.
16
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).
17
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
18
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 JuraTriassic 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 NS 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.
19
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 orehosting 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
20
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
21
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, ICPAES, and ICP-MS were monitored using the georeference material GSP-2. Repeated analysis
22
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.
23
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.
25
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 207Pbcorrected 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
26
27
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.
28
29
Figure 2.7: U-Pb isotope data for zircon from the Grass Valley granodiorite. (A) TeraWasserburg 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
207
Pb-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
206
Pb/238U weighted average ages for 204Pb-corrected, 207Pb-corrected, and 208Pb-corrected ages
are statistically indistinguishable.
30
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
(~6σ) outlier. The main group of 206Pb/238U 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
31
32
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.
33
34
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).
35
36
37
38
Figure 2.10: U-Pb isotope data for xenotime from the Empire vein in Grass Valley (A) TeraWasserburg 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
39
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 Tiin-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-
40
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
41
42
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).
Analyses of crystals with Al#>0.21 (Al#=[6]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
43
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).
44
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,
45
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
46
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 magmatichydrothermal 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
47
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 arcrelated 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
48
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 arcrelated 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
49
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
50
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 magmatichydrothermal 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 NS 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.
51
Increased compressive stress caused fracturing and vein formation. The second vein set formed
subsequently when the far-field stresses changed again, to a more transcurrent regime.
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58
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 quartzcarbonate 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
59
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) Crosspolarized photomicrograph of undulose extinction in quartz from the Schroeder deposit, YrekaFort 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 NNW60
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 DevonianPermian 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).
61
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 quartzcarbonate 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
62
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 JuraTriassic 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
63
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.
64
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 NS 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).
65
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
40
Ar/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.
66
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 ablationinductively 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,
67
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 highresolution 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.
68
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
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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).
70
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 BSEbright 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.%.
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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.
73
Figure 3.5: Continued.
74
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. <DL =
chemical values are below the detection limit.
75
Figure 3.6: Continued.
76
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.
77
78
79
80
81
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 BSEdark 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.
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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-ICPMS 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
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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,
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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 NS 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.
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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.
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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 BSEbright 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.
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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
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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
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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.
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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
91
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. Potassiumrich 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;
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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.
93
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
94
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 BSEbrighter 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 BSEbright 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
97
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
98
faults. Based upon paragenesis and textures, a sequence of events is proposed within a fixedspace 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 H2OCO2 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.
100
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.
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105
CHAPTER 4
40
AR/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 goldbearing 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
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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 170140 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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116
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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119
120
121
122
123
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
39
Ar 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,
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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
129
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.
131
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 magmatichydrothermal 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,
132
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
133
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
134
135
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
136
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 ~160115 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).
137
(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
138
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
139
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.
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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 162160 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 magmatichydrothermal 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
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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 farfield 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
Geological Society of Nevada 2010 Symposium Poster
Society of Economic Geology 2010 Keystone
Conference, Keystone, Colorado – Extended Abstract
Society of Economic Geology 2010 Keystone
Conference, Keystone, Colorado - Poster
Geological Society of America 2011, Minneapolis,
Minnesota - Abstract
Geological Society of America 2011, Minneapolis,
Minnesota - Poster
Society of Economic Geologist 2012 Conference,
Lima, Peru - Abstract
Society of Economic Geologist 2012 Conference,
Lima, Peru - Poster
Society of Economic Geology 2013 Conference,
Whistler, BC - Abstract
Society of Economic Geology 2013 Conference,
Whistler, B.C. - Poster
Geological Society of America 2014, Vancouver, B.C.
- Abstract
Mineral resources in a sustainable world, SGA 2015
Proceedings, Nancy, France – Extended Abstract
GSN 2010 poster.pdf
SEG 2010 ext abs.doc
SEG 2010 poster.pdf
GSA 2011 abstract.doc
GSA 2011 poster.pdf
SEG 2012 abstract.docx
SEG 2012 poster.pdf
SEG 2013 abstract.docx
SEG 2013 poster.pdf
GSA 2014 abstract.docx
SGA 2015 ext abs.doc
149