Brown, James accepted thesis 11-24-15 Fa 15-3

Ion Microprobe δ18
O-contraints on Fluid Mobility and Thermal Structure During Early
Slip on a Low-angle Normal Fault, Chemehuevi Mountains, SE California
A thesis presented to
the faculty of
the College of Arts and Sciences of Ohio University
In partial fulfillment
of the requirements for the degree
Master of Science
James E. Brown
December 2015
© 2015 James E. Brown. All Rights Reserved.

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This thesis titled
Ion Microprobe δ18
by
JAMES E. BROWN
has been approved for
the Department of Geological Sciences
and the College of Arts and Sciences by
Craig B. Grimes
Assistant Professor of Geological Sciences
Robert Frank
Dean, College of Arts and Sciences

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Abstract
BROWN, JAMES E., M.S., December 2015, Geological Sciences
Ion Microprobe δ18
Director of Thesis: Craig B. Grimes
The Mohave Wash fault (MWF), a low angle normal fault (~2 km of slip)
initiated near the brittle-ductile transition in crystalline rocks, is associated with the
regionally developed Chemehuevi detachment system. To address the role of water on
initiation and early slip, δ18
O of quartz/epidote pairs from thin shear zones and vein-fill
were analyzed in situ using a 10 μm ion microprobe spot (precision ±0.3‰, 2 SD). 480
analyses were made on 317 grains in 23 samples collected from three vertical transects
from the footwall and through the damage zone, distributed over 17 km down-dip. Quartz
from undeformed hosts defines pre-faulting δ18
O = 9.0–10.4‰ VSMOW. δ18
O values
decrease within damage zone microstructures down to -1.0‰ for quartz and -5.3‰ for
epidote. Such low-δ18
O values at the structurally deepest exposures are interpreted to
reflect influx of surface-derived fluids to depths of > 10 km.
Syn- and post-deformation mineralization in ~25% of the shear zones record
heterogeneous δ18
O(mineral) on the scale of < 100 mm2
. Inter- and intra-crystalline
variability in δ18
O is greatest in the damage zone. Host clasts are often preserved, but
textural relations also signify heterogeneity in new mineral growth within discrete shear
zones. Of 123 grains analyzed with multiple spots, 36% are zoned in δ18
O; single-grain
gradients reach 8.7‰ (over 500 μm) for quartz and 2.1‰ (over 300 μm) for epidote.

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Differences in Δ18
O(Qtz-Ep) from adjacent rims over < 100 mm2
range from 0.2–8.0‰ (in
damage zone) and 0.6–2.2‰ (below damage zone). Large variability in measured
Δ18
O(Qtz-Ep) is consistent with variable oxygen isotope exchange, and sub mm-scale
heterogeneities in permeability. Despite the intrasample-variability, overall trends in
Δ18
O(Qtz-Ep) from rims on adjacent grains (and thus temperature, assuming rims
equilibrated) vs. vertical position are resolved. Δ18
O(Qtz-Ep) generally increases (=
decreasing temperature) over ~30–100 m vertical transects from the footwall into the
damage zone at structurally deep exposures, consistent with footwall refrigeration.
Temperature defined at shallow exposures is relatively high, and implies significant heat
transfer up the fault. These results are interpreted to reflect surface-derived fluid
infiltration at the onset of slip followed by fluid recirculation likely driven by syntectonic
dike emplacement.

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Acknowledgements
I would like to thank my advisor Craig Grimes for his support and enthusiasm
since we met during my career as an undergraduate at Mississippi State University. His
excitement about my project has been extremely encouraging, especially during the
challenging times. I would like to acknowledge Dr. Barbara John and Justin LaForge for
their assistance on this project. Research and technical staff of the WiscSIMS lab at the
University of Wisconsin as well as the electron microprobe lab at the University of
Tennessee, Dr. John Valley, Dr. Kouki Kitajima, Jim Kern, and Alan Patchen assisted me
in analyses or discussions. Thanks are due to my committee at Ohio University, Drs.
Gregory Nadon and Damian Nance. I would like to thank the faculty, staff, and students
of Clippinger Laboratories for their friendship and encouragement during my time here.
Especially of note are my fellow advisees of the past two years Cody MacDonald and
Cody Strack for providing support and helping to alleviate stress. Funding came from
NSF (EAR-1145183), the Ohio University Department of Geological Sciences, and the
Geological Society of America (GSA). I would like to thank my family who has shown
me enormous support not only during my graduate work, but also throughout my life. I
want to end by thanking my partner Jen for being completely understanding and
supportive of all my ideas and eccentricities. She has given me an abundance of support
scientifically and emotionally.

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Table of Contents
Page
Abstract................................................................................................................................3
Acknowledgements..............................................................................................................5
List of Tables .......................................................................................................................8
List of Figures......................................................................................................................9
1. Introduction....................................................................................................................12
2. Background....................................................................................................................16
2.1 Low-angle detachment normal faults.......................................................................16
2.2 Detachment fault related mineralization..................................................................17
2.3 Stable Isotopes and thermal structure of detachment shear zones...........................18
2.3.2 Oxygen isotope studies on detachment faults.................................................. 20
2.4 Geologic setting .......................................................................................................23
2.4.1 Mohave Wash fault.......................................................................................... 26
2.4.2 Previous thermal structure studies ................................................................... 27
3. Methods..........................................................................................................................32
3.1 Sampling strategy.....................................................................................................32
3.2 Analytical techniques...............................................................................................33
3.2.1 Microscopy ...................................................................................................... 33
3.2.2 Electron probe microanalysis........................................................................... 34
3.2.3 Ion microprobe analysis................................................................................... 35
3.2.3.1 Sample preparation ................................................................................... 35
3.2.3.2 SIMS oxygen isotope analysis.................................................................. 36
3.2.3.3 Post-SIMS imaging................................................................................... 37
3.3 Oxygen-isotope thermometry ..................................................................................37
4. Results............................................................................................................................45
4.1. The Saddle Section: Generalized outcrop and sample description.........................45
4.1.1 The Saddle Section: Petrographic and Microstructural description ................ 46
4.2 The Bat Cave Wash Section: Generalized outcrop and sample description............48
4.2.1 The Bat Cave Wash Section: Petrographic and Microstructural description .. 49
4.3 Vertical transect summary .......................................................................................51
4.4 Additional samples...................................................................................................52
4.5 Electron probe microanalysis results .......................................................................54

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4.6 Oxygen isotope results.............................................................................................55
4.6.1 Oxygen isotope composition of The Saddle.................................................... 56
4.6.2 Oxygen isotope composition of Bat Cave Wash ............................................. 58
4.6.3 Oxygen isotope composition of additional MWF samples.............................. 62
4.6.4 Intercrystalline variability in oxygen isotope composition.............................. 64
4.6.4.1 Heterogeneity: shear zones versus veins................................................... 65
4.6.4.2 Oxygen isotope zonation within mineral grains ....................................... 67
4.3.2.3 Mineral pair variability within microstructural domains.......................... 68
5. Discussion......................................................................................................................96
5.1 Evidence for early fluid-infiltration along the Mohave Wash fault.........................96
5.2 Miocene fluid-rock interaction ................................................................................97
5.3 Grain-scale oxygen isotope variability ..................................................................100
5.3.1 Ion microprobe data verses conventional analyses of isotope composition .. 101
5.4 Calculated temperatures of in situ mineral pairs....................................................102
5.4.1 Vertical isotopic and thermal characteristics through the Mohave Wash fault
................................................................................................................................. 106
5.4.1.1 The Saddle .............................................................................................. 106
5.4.1.2 Mohave Wash ......................................................................................... 107
5.4.1.2 Bat Cave Wash........................................................................................ 107
5.4.2 Summary of vertical transect trends .............................................................. 108
5.5 Surface-derived fluids and the Mohave Wash fault...............................................110
5.6 Stable isotopic constraints on lateral variations along the Mohave Wash fault ....114
6. Conclusions..................................................................................................................131
References........................................................................................................................134
Appendix – Additional elemental and stable isotope data...............................................143

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List of Tables
Page
Table 3.1: Samples from the Mohave Wash fault analyzed for hydrothermal
minerals….….….….….….….….….….….….….….….….….….….….….….….….…..40
Table 4.1: Results of petrographic analysis of Mohave Wash fault samples associated ..69
Table 4.2: Data from electron microprobe analysis Mohave Wash fault samples……....71
Table 4.3: Summary of oxygen isotope compositions…………………………………...72
Table 4.4: Summary of intercrystalline homogeneity of analyzed minerals in δ18
O.……76
Table 4.5: Summary of intracrystalline zonation patterns in δ18
O……...…………….…79
Table 5.1: Summary of calculated temperatures from the Mohave Wash fault…..…....117
Table 5.2: Summary of calculated temperatures of the Mohave Wash fault from samples
with heterogeneous microstructural domains……..……….………………………...…120
Table A1: Weight percent oxide data for epidote from electron microprobe analysis....143
Table A2: Epidote number of ions data from electron microprobe analysis…………...152
Table A3: Feldspar weight percent oxide data from electron microprobe analysis……161
Table A4: Feldspar number of ions data from electron microprobe analysis………….164
Table A5: Ion microprobe data for analysis of quartz, epidote, and K-feldspar……….165

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List of Figures
Page
Figure 2.1: Idealized low-angle normal fault with the effects of extensional shearing and
footwall heating on geothermal gradient ……………………………………………..…29
Figure 2.2: Simplified geologic map showing sample locations and cross-section of the
Chemehuevi Mountains, California…………………………………….………….…….30
Figure 3.1: Field characteristics of the Mohave Wash fault damage zone at The Saddle
vertical transect……………………………..………………………………………...….42
Figure 3.2: Field characteristics of the Mohave Wash fault damage zone at the vertical
transect located at Bat Cave Wash………………………………..……………..………43
Figure 3.3: Example analysis pits by ion microprobe…………….……………..………44
Figure 4.1: Hand sample example of a cataclasite from the Mohave Wash fault at The
Saddle vertical transect…………………………………………………..……………....80
Figure 4.2: Annotated photographs, X-ray maps and backscattered electron images of
samples characteristic of the Mohave Wash fault at The Saddle vertical transect……....81
Figure 4.3: Annotated photograph, X-ray maps, and backscattered electron images of
sample CG-14CH-126 from the top of the Mohave Wash fault damage zone at the mouth
of Bat Cave Wash……………………...........………………………………...…………82
Figure 4.4: Annotated photograph, X-ray maps, and backscattered electron images of
sample CG-14CH-128 taken from the bottom of the Mohave Wash fault damage at the
mouth of Bat Cave Wash ……………............................………………………………..83
Figure 4.5: Annotated photograph and X-ray maps of sample CG-14CH-124 from 1 m
below the main Mohave Wash fault damage zone at the mouth of Bat Cave Wash ……84
Figure 4.6:Annotated backscattered electron images of sample CG-14CH-135 from 10 m
below the main Mohave Wash fault damage zone at the mouth of Bat Cave Wash.........85
Figure 4.7: Annotated photograph and backscattered electron images of sample CG-
13CH-4 from the Studio Spring sampling area……………………………….....……….85
Figure 4.8: Annotated photograph and backscattered electron image of sample CG-
13CH-24 from the Trampas Wash sampling area……………………………..…………86

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Figure 4.9: The XFe of 17 samples taken from within and outlying the main damage zone
of the Mohave Wash fault………………………………….………………………....….87
Figure 4.10: Summary showing petrographic relations of analyzed textures for all 503
measurements of δ18
O (‰, VSMOW) of quartz, epidote, and K-feldspar in 23 analyzed
samples…………………………………………………………………..………….……88
Figure 4.11: Secondary electron images of sample CG-14CH-127 from Bat Cave Wash
showing significant intracrystalline and grain-to-grain variation in δ18
O values.....…….89
Figure 4.12: All oxygen isotope analyses of quartz and epidote sampled by field site
plotted versus distance along the Mohave Wash fault……….…………………………..89
Figure 4.13: All ion microprobe measurements arbitrarily arranged in order of increasing
δ18
O (‰) of quartz and epidote in samples taken from three vertical transects of the
Mohave Wash fault damage zone……………………………………..…..……………..90
Figure 4.14: All ion microprobe measurements arbitrarily arranged in order of increasing
δ18
O (‰) of quartz and epidote in additional samples taken along the Mohave Wash fault
damage zone.…………………………………...…………………………….…………..91
Figure 4.15: Summary showing petrographic relations of all 116 measurements of δ18
O
(‰) of quartz, epidote, and K-feldspar in six analyzed samples from the Mohave Wash
fault vertical transect at The Saddle…………………………………….……………..…92
Figure 4.16: Summary showing petrographic relations of all 235 measurements of δ18
O
(‰) of quartz and epidote in nine analyzed samples from Mohave Wash fault vertical
transects at Bat Cave Wash…………………….……………………………...…………93
Figure 4.17: Annotated photograph and backscattered electron images of sample CG-
13CH-RF from the Range Front sampling area………………….……………...……….94
Figure 4.18: Backscattered electron image of sample CG-14CH-133 from Bat Cave
Wash…………………………………………………...……..………………...………..95
Figure 5.1: Comparison of stable isotope compositions of δ18
O and elemental iron
composition of epidote for a given sample ………………………………………..…...122
Figure 5.2: Comparison of stable isotope compositions of δ18
O (‰) of quartz and epidote
for a given sample ………………………………….………………………….…..…...123
Figure 5.3: Summary of measured δ18
O (‰, VSMOW) of quartz-epidote mineral pairs in
fault rocks analyzed by ion microprobe……………………………………………..….124

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Figure 5.4: (a) Quartz and epidote δ18
O values (‰, VSMOW) from a given vertical
transect plotted versus Mohave Wash fault (MWF) position. (b) Apparent temperatures
calculated using the oxygen isotope fractionation plotted versus the MWF position.…125
Figure 5.5: Measured δ18
O(Qtz) values plotted versus respective calculated δ18
O of fluids.
Measured δ18
O(Ep) values plotted versus respective calculated δ18
O of fluids……….…126
Figure 5.6: (a) Calculated apparent temperatures of quartz-epidote mineral pairs from a
given field site plotted versus distance along the Mohave Wash fault (MWF). (b) The
effect of rapid advection of heat transport along the MWF relative to the overall
geothermal gradient…………………………………………………………………….127
Figure 5.7: (a) Summary cartoon of the Mohave Wash fault with results from this study.
(b) Modeling by Gottardi et al. (2013) showing colder temperatures within a detachment
recharge zone and hotter temperatures within a detachment discharge zone…………..129

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1. Introduction
Despite recognition of regionally developed, large-slip, low-angle normal faults
(LANFs) globally and in various geotectonic settings, controversy remains regarding
their initiation and protracted slip at shallow dips through the seismogenic crust (review
by John and Cheadle, 2010; Whitney et al., 2013). Hydrothermal fluid circulation, heat
flow, and the behavior of actively slipping geologic faults are most likely intimately
linked, and fluids may contribute to early fracture development and later strain
localization of low-angle faults through weakening processes involving reaction
softening, elevated pore pressure and/or reduced frictional coefficients, which are often
invoked to explain fault movement (e.g., Lachenbruch, 1980; Famin et al., 2004;
Collettini, 2011). Low-angle normal fault systems are widely recognized as major
conduits for fluid migration (e.g., Kerrich and Rehrig, 1987; Fricke et al., 1992;
Wickham et al., 1993; Nesbitt and Muehlenbachs, 1995; Losh et al., 1997; Morrison and
Anderson, 1998; Holk and Taylor, 2007). Speculations on the source of fluids moving
through these faults vary widely and are based largely on stable isotope data. The
spectrum of inferred fluids include shallow level meteoric water (Kerrich and Hyndman,
1986; Glazner and Bartley, 1991), basinal brines (Spencer and Welty, 1986; Roddy et al.,
1988), deep magmatic or metamorphic sources (Smith et al., 1991; Axen, 1992; Smith et
al., 2008), or mixing of multiple sources (Spencer and Welty, 1986). Some authors have
suggested that surface-derived fluids penetrate to 10-15 km depths (Wickham et al.,
1993; Fricke et al., 1992; Kerrich and Rehrig, 1987). Other workers favor the concept
that downward fluid penetration is restricted to the upper, brittle sections of detachment
faults, whereas the release of metamorphic or deeply seated magmatic fluids account for

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the alteration of ductile portions of the system (e.g., Axen et al., 2001). Once initiated,
cataclasis associated with faulting increases permeability, channeling fluids into a fault
zone, allowing fluid-assisted deformation processes to enhance break-down reactions of
feldspar to form weaker phyllosilicates.
Past research of LANFs has focused in large part on fault breccias and gouges
related to late slip that occurred after fault initiation since these are often readily
preserved (reviewed by Collettini, 2011). Such studies have consistently suggested that
increased fluid pore pressure and development of aligned phyllosilicate-rich networks
contribute to fault slip based on laboratory evidence of fault zone fabrics (reviewed by
Collettini et al., 2009). However, field evidence is elusive, and it is not clear when these
weakening mechanisms develop or how the fault initially breaks.
Oxygen isotope geochemistry can be an effective monitor of fluid rock
interactions, and the fractionation of 18
O between minerals is temperature sensitive. If
equilibrated, two co-existing minerals formed from the same fluid can be used to monitor
the temperature of formation based on comparisons between their δ18
O values and
established experimental oxygen isotope fractionation factors (e.g., Valley, 2001). Most
stable isotope studies on LANFs have been conducted using bulk measurements on whole
rocks or mineral separates. Such measurements effectively constrain integrated fluid
histories, but likely obscure fluid-rock interaction associated with early slip along faults.
For example, Morrison (1994) demonstrated that mylonitic footwall rocks to the Whipple
detachment (California) had low-δ18
O caused by a secondary overprint (on feldspar)
related to late circulation of meteoric water at ~350°C, rather than infiltration of fluids
while the rocks experienced ductile deformation. Overprinting of isotopic signatures of

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microstructures by later fluid flow may be quite common. Morrison and Anderson (1998)
found spatially varying Δ18
O(Qtz-Ep) (δ18
O(Qtz) - δ18
O(Ep)) fractionations in minerals
separated from chlorite breccias in the Whipple detachment fault footwall within
gneisses. They showed Δ18
O(Qtz-Ep) increased from 4.54 ± 0.46 ‰ (yielding an oxygen
isotope temperature of 458°C) 50 m below the fault to 5.81 ± 0.52 ‰ (~350°C) 12 m
below the fault. They attributed this extreme geothermal gradient (82°C over 38 m or
2160°C/km) to convection of cool surface-derived fluid down high-angle faults in the
upper plate. More recent studies have reported similar transient vertical geotherms in
detachment faults of ~2000°C/km in mylonitic micaschists and marbles of the Tinos
detachment in the Aegean (Famin et al., 2004), and 140°C over 100 m in quartzite fault
rocks of the Raft River detachment in Utah (Gottardi et al., 2011). Similarly, McCaig and
Harris (2012) suggest upward fluid and heat migration along oceanic detachment faults
where a high temperature heat source (melt lens) occurs at depth. If common, this process
would lead to cooling and strain localization along brittle structures due to the rapid
advection of heat by infiltrating surface-derived fluids.
Past research considered, the goal of this study is to evaluate the role of footwall
refrigeration (or heating) during initiation of the extinct limited-slip, low-angle Mohave
Wash fault (MWF) seated in the footwall to the regional Chemehuevi fault system
located in SE California using oxygen isotope geochemistry. The MWF is thought to
have limited slip history (~2 km of displacement and lacking the development of a gouge
zone common in mature faults) partly in isotropic granites (no preexisting fabrics to help
localize deformation), preserving conditions shortly after fault initiation near the brittle-
ductile transition zone (John and Foster, 1993). Thus, limited fluid-rock interactions

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during the pre- and post-faulting history allow isotopic signatures reflecting fluid flow
within MWF microstructures to be constrained directly to the early slip history. Sampling
of vertical transects through the MWF damage zone, and laterally over ~17 km in the
down-dip direction allow characterization of the stable isotopic composition on various
scales. The Δ18
O(Qtz-Ep) of mineral pairs have been determined in situ by ion microprobe
using a 10 μm spot. The principal advantage of this technique is the ability to relate
specific textures or zones/domains within single grains identified by optical microscope
and Scanning Electron Microscope (SEM) to stable isotope compositions. The
assumption of stable isotope equilibrium can be evaluated more effectively when
adjacent rims on two grains are analyzed, allowing mineral zoning and mineralization
related to MWF deformation and fluid flow to be recognized and resulting in more
geologically meaningful temperature calculations. The isotopic data are used to address:
1) the extent to which heat and mass transfer along a LANF creates a locally steep
vertical gradient that may help facilitate strain localization; 2) the likely source of fluids
at fault initiation and with progressive slip; and 3) lateral variations in fluid-rock
interactions over 17 km in the down-dip direction, reflecting paleodepths ranging from
~5-11 km.

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2. Background
2.1 Low-angle detachment normal faults
Detachment faults, or low-angle normal faults (LANFs), are gently dipping (30°
or less) regional features showing domed topography with offsets of 10–50 km (Figure
2.1a; Axen, 2004). These features have been recognized in a wide variety of settings such
as the Basin and Range province of the Western US, rifted continental margins, and mid-
ocean ridge spreading centers (Figure 2.2; John and Cheadle, 2010) and are considered
important structures along which extreme lithospheric extension is accommodated.
Although appreciation of large offset detachment faults in continental and oceanic crustal
settings has expanded in recent decades, discussion with regard to their initiation and
early slip mechanisms remains controversial (John and Foster, 1993; Axen, 2004; Famin
et al., 2004; Collettini, 2011; Gottardi et al., 2015) since Andersonian fault mechanical
theory does not predict the development of normal faults at such low-angles to horizontal
(Anderson, 1951; Collettini and Sibson, 2001). Contrary to theoretical predictions, field
observations from detachment faults accompanied by thermochronometric and
paleomagnetic data indicate both initiation and kilometer-scale displacement within the
brittle crust (John and Foster, 1993; Axen, 2007). Detachment faults are suggested by
some to initiate within the brittle zone and without conventional stick-slip behavior by
providing considerable extension through aseismic creep since the accommodation of
large amounts of displacement found with detachment faults is anomalous due to the lack
of observed large magnitude earthquakes (Howard and John, 1987; Abers, 1991; Axen et
al., 1999; Collettini and Holdworth, 2004; Abers, 2009). Detachment faults are

17
influenced by extension of uplifted core complexes that form domal geometries, or
upwarping, parallel to the extensional direction (Yin and Dunn, 1992). This domed
detachment geometry may be the product of various processes: isostatic response from
past tectonic events (Rehrig and Reynolds, 1980); reverse drag from a deeper underlying
detachment fault (Davis and Lister, 1988); formation of shear zones in the lower plate of
the detachment (Reynolds and Lister, 1990); movement initiated by a flat fault surface
(John, 1987). Analysis on the origin of domal detachment zones has focused on the link
between detachment faults and their observed lower-plate structures (John, 1987), dikes
(Spencer et al., 1986), and mylonitic-zones (Davis, 1988).
2.2 Detachment fault related mineralization
Significant evidence for fluid migration along detachment faults, which may
promote reaction-weakening processes and facilitate slip, comes from field observations
(Spencer and Welty, 1986; Roddy et al., 1988; Spencer and Reynolds, 1989). Greenschist
facies minerals including epidote, chlorite, and calcite are typically found throughout the
damage zone in early-slip portions of many detachment faults of the Colorado River
extensional corridor (CREC), including the Mohave Wash fault (MWF) (John, 1987;
Lister and Davis, 1989). Distinct features of detachment-fault-related mineralization in
general are:
1. Mineralization is controlled by structures formed during detachment faulting.
Structures include the low-angle detachment-fault system, high-angle faults in the

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lower-plate just below the detachment fault, and low- to high-angle normal faults
in the upper-plate.
2. Mineralization localized in zones that have been brecciated or deformed by
movement along or above the detachment fault.
3. Chlorite-epidote-calcite alteration along and below the detachment fault.
Late mineralization consists of iron and copper oxides, principally specular to earthy
hematite. Common gangue minerals are quartz, barite, fluorite and manganese oxides.
Lower-temperature clay gouge mineralization is also common in faults active to low
temperatures.
2.3 Stable Isotopes and thermal structure of detachment shear zones
The fractionation of 18
O between water and minerals provides a sensitive
indicator of fluid-rock interactions (O’Neil, 1986; Chacko et al., 2001). The fractionation
of 18
O between two phases is also temperature dependent. However, oxygen isotope
thermometry has proven more difficult in large part due to uncertainties about
equilibration between mineral assemblages and an altering fluid (Valley, 2001). Mineral
pairs may be equilibrated through coprecipitation from a single fluid, reequilibration
during crystal plastic deformation, or from bulk diffusive exchange between preexisting
minerals, although the latter is typically only expected at adjacent grain boundaries
(O’Neil, 1986). To determine geologically meaningful temperatures using stable isotopes,
mineral pairs must be equilibrated and must not have experienced differential exchange
or resetting during cooling or later fluid-rock interactions (O’Neil, 1986; Valley, 2001).

19
This constraint can become problematic at the grain scale where growth zoning,
recrystallization, grain boundary diffusion, and exchange with a hydrothermal fluid occur
(Valley, 2001; Valley and Kita, 2009; Ferry et al., 2014).
As a rock cools, minerals will continue to exchange oxygen isotopes with
surrounding minerals of different δ18
O values as part of a closed system exchange. High
oxygen diffusivity minerals (e.g., K-feldspar) will exchange oxygen isotopes during
cooling to low temperatures (< 300°C). Low oxygen diffusivity minerals (e.g., epidote)
will exchange oxygen isotopes only at high temperatures (> 800°C) and δ18
O values
should not be affected by cooling. Minerals with medium oxygen diffusivity (e.g., quartz)
will restrict exchanging oxygen isotopes below ~550°C (Cole and Chakraborty, 2001).
Open system exchange occurs when a fluid moves through a rock allowing
minerals to exchange oxygen isotopes with the fluid, however the roles of fault
permeability and deformation mechanisms in oxygen isotope transport and exchange
during fluid flow are poorly understood (Bowman et al., 1994; Person et al., 2007;
Gottardi et al., 2013). Inter-grain fluids may be preferentially incorporated into one
mineral relative to another. Oxygen isotope disequilibrium is frequently interpreted to be
present in shear zones, where kinetic fractionation (physical separation of isotopes)
surpasses equilibrium fractionation (thermodynamic separation). Disequilibrium
exchange is unlikely to have affected δ18
O values from the samples containing quartz and
epidote from the MWF due to relatively high oxygen diffusivity properties (e.g.,
Morrison and Anderson, 1998).

20
2.3.2 Oxygen isotope studies on detachment faults
In an attempt to constrain fluid-rock interactions during faulting of the
detachment zone, many previous stable isotope studies used laser fluorination
measurements from whole rock and mineral separates consisting of 2–3 milligrams of
material (Losh, 1989; Fricke et al., 1992; Wickham et al., 1993; Morrison, 1994;
Morrison and Anderson, 1998; Holk and Taylor, 2007; Gottardi et al., 2011; MacDonald,
2014). A common finding was lowered δ18
O along fault rocks, consistent with the influx
of low-δ18
O surface-derived fluids. Morrison (1994), found the Whipple detachment fault
had low absolute δ18
O values of quartz and K-feldspar associated with late circulating
surface-derived fluids overprinting feldspar found in the mylonitic footwall; the low
absolute δ18
O values were interpreted to post-date the mylonite-forming event. Studies
using quartz-feldspar mineral pairs often find oxygen isotope exchange trajectories
showing δ18
O(Qtz) – δ18
O(Kfs) plots with a vertical slope (e.g., Morrison, 1994; Holk and
Taylor, 2007). Feldspars are particularly sensitive to low temperature oxygen exchange
with fluids as well as hydrolysis reactions that produce secondary phyllosilicates (i.e.,
clays; Valley, 2001). Through careful sampling of adjacent quartz and epidote grains in
the footwall to the Whipple detachment, Morrison and Anderson (1998) reported
systematic variations in mean Δ18
O(Qtz-Ep) (δ18
O(Qtz) - δ18
O(Ep)) within the footwall, and
based on oxygen isotope thermometry interpreted them to reflect an extreme geothermal
gradient (82°C over 38 m) from 50–12 m below the damage zone (Figure 2.1c). Based on
a systematic change in Δ18
O (Qtz-Ms), Gottardi et al. (2011) subsequently suggested that a
thermal gradient of 140°C, found over a 100 m thick shear zone of the Raft River
detachment, forms near the brittle-ductile transition zone to account for the formation of

21
shearing and convection of fluids (i.e., Figure 2.1b). Similarly, Famin et al. (2004)
reported a thermal gradient of ~2000°C/km in the Tinos detachment using Δ18
O (Qtz-Cc)
fractionations from quartz-calcite mineral pairs. MacDonald (2014) provided evidence
for a downward shift in whole rock and mineral δ18
O values from quartz and epidote
within shear zones along the MWF relative to undeformed granitic host rocks, indicating
that infiltration of low-δ18
O fluid (surface-derived) permeated the MWF zone early in the
development of shear zones.
In an attempt assess the hydrologic and thermal controls on fluid-rock isotopic
exchange and transport along idealized detachment faults, modeling by Person et al.
(2007) and Gottardi et al. (2013) has been used to show that “domino” or “book shelf”
thinning effects of the upper brittle crust (e.g., Lister and Davis, 1989) allows infiltrating
fluid to channelize as its base and transfer heat (e.g., Lopez and Smith, 1995). Modeling
by Person et al. (2007) and Gottardi et al. (2013) show that considerable oxygen isotope
and heat distributions resulting from low-δ18
O fluid flow at mid-crustal depths is highly
dependent upon permeability (i.e., detachment fault damage zone). The transfer of heat
along permeable fault systems at depth would promote a steep geothermal gradient in the
footwall, which is supported by existing oxygen isotope thermometry constraints (e.g.,
Morrison and Anderson, 1998; Famin et al., 2004; Gottardi et al., 2011).
Field and thin section observations from studied detachment faults indicate that
individual shear zones experienced several episodes of deformation which in some cases
included early semi-brittle deformation followed by cataclasis and subsequent
hydrothermal alteration of feldspars and along fractures (Morrison, 1994). Conventional
analytical techniques (laser fluorination) using millimeter-scale sample size may

22
homogenize multiple events and obscure heterogeneities such as mineral zoning due to
extended growth events, inclusions of other minerals, or hydrothermal alteration
overprinting formation compositions (i.e., Morrison, 1994; Morrison and Anderson,
1998; Gottardi et al., 2011; MacDonald, 2014). These factors make it difficult to relate
bulk geochemical compositions to specific microstructures. In contrast to conventional
techniques, analysis by ion microprobe provides improved spatial resolution and the
ability to correlate geochemistry directly to specific microstructures. Valley and Graham
(1996) found regular variations of 3–13‰ over 200–400 m extensional shear zones,
respectively, in δ18
O using single quartz grains with ion microprobe analysis. The only
application of secondary ion mass spectrometry (SIMS) techniques to LANFs known to
the author was conducted by Famin et al. (2004) produced results providing support of
footwall refrigeration from surface-derived fluids (absolute δ18
O values < 5‰) circulating
along a LANF with a geotherm of > 100ºC/50 m using quartz-calcite oxygen isotope
fractionations. The cumulative results of these studies establish that circulating fluids
along faults is a complex system and requires in situ oxygen isotope geochemistry by
spatially resolved ion microprobe analysis to better understanding these fluid interactions.
Late low-temperature overprinting of isotopic signatures from infiltrating fluids is
common in evolving fault systems (Fricke et al., 1992; Morrison 1994; Morrison and
Anderson, 1998; Famin et al., 2004, 2005; Holk and Taylor; 2007; Gottardi et al., 2011).
However, Sharp et al. (1991) demonstrated that quartz will be significantly less altered
than whole rock or feldspars due to extremely slow bulk diffusion at temperatures <
500ºC. Coexisting quartz and epidote have been shown to be effectively closed to
subsolidus oxygen isotope diffusion at temperatures below ~550ºC (Sharp et al., 1991;

23
Ferreira et al., 2003). Thus, these minerals are expected to preserve δ18
O values inherited
during original formation or later recrystallization that are unaffected by the subsequent
uplift, extension, and hydrothermal fluid flow. Mathews (1994) found that the
equilibrium fractionation for quartz and epidote (Δ18
OQtz-Ep) to varied by 4‰ over 250–
450ºC and established that the quartz-epidote thermometer is reasonably sensitive in this
temperature range given typical analytical uncertainties (±0.1‰ for laser fluorination;
±0.3‰ for ion microprobe; Valley and Kita, 2009).
2.4 Geologic setting
The Chemehuevi Mountains and Whipple Mountains are centrally located
features of the Colorado River extensional corridor (CREC), which underwent crustal
extension from 23–12 Ma, accommodated an estimated 40–75 km of motion thought to
be caused by crustal relaxation and Basin and Range extension (Figure 2.2; Davis et al.,
1980; Howard and John, 1987). The CREC stretches from southeastern California and
western Arizona to southern Nevada and lies within the curved boundary of Cordilleran
core complexes containing the Whipple, Buckskin, Dead, and Chemehuevi Mountains
(Coney, 1980). These mountains represent metamorphic core complexes comprising
upper to mid-crustal rocks denuded by regional detachment faults. The corridor is
initiated along a rooted asymmetric zone of crustal extension. Seismic refraction and
structural data has suggested Chemehuevi detachment system is connected in the
subsurface to the Whipple Mountains detachment fault, lies < 3 km beneath the Mohave

24
Mountains, and can be rooted as far east as the Hualapai Mountains, ~80 km (Howard
and John, 1987; John, 1987).
The rock types exposed in the Chemehuevi Mountains core complex include
Cretaceous granitic lithologies, Proterozoic layered gneisses, and Tertiary basaltic to
rhyolitic dike swarms (Figure 2.2). The Chemehuevi Plutonic Suite makes up the central
and southwestern portions of the area and is exposed primarily as granodiorite (Kpg)
showing a zonation of increasing silica content toward the center of the pluton.
Proterozoic gneiss makes up the northeast portion of the area composed of layered
orthogneiss and paragneiss with common leucosome pods. The gneisses contain
subvertical veins with greenschist mineralization (i.e., epidote) and typically show
alteration of biotite to chlorite. Basaltic to rhyolitic Tertiary dike swarms intrude the
Chemehuevi plutonic suite in southwest and central parts of Chemehuevi Mountains.
Dikes are of several generations, but a K-Ar age of 20.7 ± 1.3 Ma implies some of the
intrusions occurred during regional extension (John and Foster, 1993). Dikes in the
northeastern portion of the area show strong internal lineations oriented parallel to the
established extension direction. Mineralized shear zones are observed at the margin of
several dikes.
Field studies of the Chemehuevi Mountains reveal that the Cretaceous granitoids
and Proterozoic gneiss country rocks were exposed by a series of at least two stacked
faults that formed at the time of detachment with > 23 km of displacement in the original
dip direction (Howard and John, 1987; John and Foster, 1993). Of the two major low-
angle normal faults recognized previously (John, 1987), the Chemehuevi detachment
fault (CDF) is the shallowest structurally and accommodated the majority of the

25
extension at the Chemehuevi Mountains. The CDF is associated with the neighboring
Whipple detachment fault ~30 km SE and the Sacramento detachment fault ~20 km NW
(Figure 2.2; John, 1987).
Field observations from the Chemehuevi and Whipple detachment faults have
shown displacements of more than ~8 km along the Chemehuevi fault (Miller and John,
1988) and ~40 km along the Whipple fault system (Davis and Lister, 1988). Slip-
direction indicators such as slickenlines, lineations, offset markers, preserved striae, drag
folds, minor faults within related cataclasites, and the southwest dip of syntectonic strata
above each detachment fault show motion of the upper plates was to the northeast at 050
(John, 1987; Yin and Dunn, 1992). Age of the CREC initiation has been determined by
crystallization ages of syntectonic plutons, 40
Ar/39
Ar footwall cooling ages, and K-Ar
ages from synextensional volcanic rocks to be ~23 Ma (Spencer and Reynolds, 1991;
Anderson et al., 1988; Howard and John, 1987).
The hanging-wall of the CDF contains many high-angle normal faults that have
rotated over time to shallower dips but which do not cut the detachment providing
evidence for detachment fault emplacement without passive rotation (Howard and John,
1987). The faults cut across large portions of isotropic plutonic rocks in the southwestern
portion and gneisses similar to those found in the Whipple Mountains in the northeastern
portion of the mountains (Howard and John, 1987). The gneisses are the structurally
deepest fault rocks and contain thin (1–10 cm) shear zones. The faults are thought to have
served as fluid pathways based on previous oxygen isotope studies. Even though fluid
source and infiltration mechanisms for low permeability crystalline rock within
continental crust to depths of the brittle-ductile transition remain problematic (Fricke et

26
al., 1992; Morrison, 1994; Morrison and Anderson, 1998; Famin et al., 2004; Holk and
Taylor, 2007), numerical modeling of oxygen isotope transport and exchange has proven
useful in constraining parameters allowing meteoric fluid to circulate to these depths
(Bowman et al., 1994; Person et al., 2007; Gottardi et al., 2013).
2.4.1 Mohave Wash fault
The Mohave Wash fault (MWF) is a relatively small-displacement (1–2 km) low-
angle fault outcropping as a sinuous trace over 350 km2
that was denuded to near the
surface within the CDF footwall and exposed through erosion (John and Foster, 1993).
The lack of fault gouge on the MWF indicates that the fault did not reactivate at
shallower depths and it is considered to preserve the initial faulting structures and
mineralization associated with detachment fault initiation at depth (John and Foster,
1993). Previous studies of the MWF describe a damage zone varying from 10 to ~200 m
thick, represented by cracked granite/gneiss, chlorite-rich breccia/cataclasite, and
cohesive cataclasite with indication of sequential fracturing and fluid flow (John, 1987;
LaForge et al., 2014). The metamorphic minerals epidote, chlorite, and calcite are found
hosted throughout the damage zone of the MWF, but are scarce away from the fault. John
(1987) determined that cataclasis was the primary deformation microstructure during
early slip history producing the thick chlorite-rich cataclasite/breccia zones with little
evidence of mylonitization. In the southwest region the MWF cuts isotropic granodiorite
and is dominated by brittle deformation. The MWF in the structurally deepest northeast
region cuts gneissic fabric with plastically deformed mafic and felsic dikes intruding the

27
damage zone with foliations parallel to slip direction (John and Foster, 1993; LaForge et
al., 2014).
2.4.2 Previous thermal structure studies
Both structural and thermochronologic data from the Chemehuevi Mountains
show that low-angle normal faulting began 22–24 Ma (John and Foster, 1993). Using
multiple thermochronometric systems (40
Ar/39
Ar on hornblende [closure temperature of
490ºC (Harrison, (1982)] and biotite [closure temperature of 373ºC (Berger and York,
(1981)], and fission-track on apatite), John and Foster (1993) defined a southwest to the
northeast trend of decreasing cooling ages in samples from the lower plate footwall rocks.
The trend of younger biotite 40
Ar/39
Ar ages toward the northeast is consistent with deeper
structural levels at the time of detachment-fault activity and is interpreted as
demonstrating rapid cooling associated with detachment initiation (John and Foster,
1993). Based on the closure temperatures and ages of minerals from samples collected
over 16 km in the spreading direction, they determined a continuous increase in
temperature of < 200ºC in the southwest to > 450ºC in the northeast at the time of fault
initiation (~23 Ma). MacDonald et al. (2014) found apparent temperatures using oxygen
isotope thermometry on coexisting quartz and epidote from the MWF footwall to be
typically 50–150ºC higher than ambient footwall temperatures found by John and Foster
(1993) at fault initiation. John and Foster (1993) and MacDonald et al. (2014) both found
that temperatures increased along fault with paleodip. Using these data along with
estimated thermal gradients of 30–50ºC/km, the fault system was modeled to root at a
minimum depth of ~10–12 km with a paleodip ≤30º, and an estimated slip rate of ~8

28
mm/yr (John and Foster, 1993). However, circulation of surface-derived fluids along the
fault (i.e., footwall refrigeration) could locally perturb geothermal gradients by creating
lower temperatures deeper than expected that resulted in closure of thermochronometers
prior to substantial uplift (Figure 2.1b,c). Carter et al. (2004) found anomalous young
ages from the Chemehuevi Mountains among the consistent age decrease along the slip
direction using the apatite (U-Th)/He (closure temperature of ~40-80ºC)
thermochronometer indicative of localized heat flow, possibly due to syntectonic dike
emplacement.

29
Figure 2.1: (a) Schematic cross section of an idealized low-angle normal fault shortly
after initiation, with possible fluid flow paths and channelized fluid flow (blue arrows)
along high-angle faults in the upper plate and along the main detachment. (b) Two
models for the thermal structure along a fault showing effect extensional shearing (left)
resulting in localized footwall heating of a given detachment fault (red dashed lines) on
geothermal gradient, and (right) the effects of fluid flow penetrating a given detachment
fault (blue dashed lines) on geothermal gradient with grey dashed box highlights the
region most affected by an extreme thermal gradient (Gottardi et al., 2011; 2013). (c)
Measured difference in δ18
O of quartz and epidote (Δ18
OQtz-Ep) in the footwall to the
Whipple detachment fault, and corresponding oxygen isotope temperatures showing a
geothermal gradient of 82ºC over 30 m within the uppermost 50 m of footwall of the
nearby Whipple detachment fault (Morrison and Anderson, 1998).

31
Figure 2.2: Simplified geologic map and cross section of the Chemehuevi Mountains,
California (after John and Foster, 1993) showing sample locations for this study. Yellow
stars identify locations where vertical transects were made. Notched lines show major
faults. Bold lines represent the thermal structure of the footwall at 23 Ma, the inferred
timing of initiation (John and Foster, 1993).

32
3. Methods
3.1 Sampling strategy
Fieldwork and sampling for this investigation took place during December 2013
and March 2014. To characterize vertical gradients across the fault using stable isotope
geochemistry, two locations were targeted for sampling along the Mohave Wash fault
(MWF) separated by 17 km in the slip direction. The two sites selected include The
Saddle section, located near the W-SW margin of the exposed footwall (shallower at
initiation), and the Bat Cave Wash located at the far NE (deeper at initiation) portion of
the footwall (Figures 2.2, 3.1, 3.2).
A transect of 10 samples at The Saddle covered ~120 m of continuous vertical
section (Figure 3.1). Two transects of approximately 30 m and which were perpendicular
to the fault were made at the Bat Cave Wash site due to poor MWF footwall exposure.
One site was near the mouth of the wash and another 1.75 km to the SW with 10 total
samples collected from both locations (Figure 3.2). The two sites were surveyed to
increase the vertical distance relative to the fault that was accessible for sampling.
Combined, both Bat Cave Wash transects cover ~61 meters extending from 40 meters
below the main damage zone through the intensely-fractured interval (~10 m thick) of the
MWF.
During Spring of 2013, 113 samples were collected along the MWF at four sites
known as Range Front, Studio Springs, Trampas Wash, and Mohave Wash, spanning ~15
km in the slip direction (Figure 2.2). Many of these samples were originally analyzed by
laser fluorination and reported by MacDonald (2014). Eight of 113 samples were selected

33
for additional secondary ion mass spectrometry (SIMS) analysis and are incorporated in
this study for comparison with vertical transects and to constrain lateral variations in the
fault-slip direction. The rock types sampled include granitoids of the Chemehuevi
Plutonic Suite (Cretaceous), Precambrian gneiss, breccias featuring greenschist facies
mineralization, quartz-epidote cataclasite shear zones containing brittle deformation, and
veins mineralized with epidote and quartz. Table 3.1 summaries the location, sample
type, and structural orientations of all samples incorporated during this study.
3.2 Analytical techniques
Oxygen isotope values were determined in situ in thin section or rock chips by ion
microprobe to characterize fluid rock interactions, heat, and mass transfer during early
slip on the Mohave Wash fault system. Examination of samples using optical
petrography, scanning electron microscopy (SEM), energy dispersive X-ray spectroscopy
(EDS), and electron probe microanalysis (EPMA) was carried out to characterize
microstructures, mineralogy, and geochemistry along the fault zone and to guide ion
microprobe analysis.
3.2.1 Microscopy
In order to characterize microstructures and map minerals for oxygen isotope
measurements, 19 standard thin sections were prepared. Additionally, five rock chips
from representative microstructures were cut, mounted in epoxy, and polished.
Mineralogy was determined, and generations of deformation zone/vein formation were

34
established by cross-cutting relationships. Deformed rock type name was determined
using the following classification modified from Davis et al. (1996):
Breccia: 200–500 μm angular clasts in a finer grained matrix.
Cataclasite: < 200 μm angular clasts in a finer grained matrix.
Ultracataclasite: < 200 μm angular clasts in a glassy matrix.
Mylonite: ductile deformation of feldspar clasts in quartz ribbons.
Thin sections were examined by secondary electron and backscattered electron
(BSE) imaging with the Hitachi S-2460N natural scanning electron microscope in the
Department of Physics and Astronomy at Ohio University to assess texturally complex
areas and characterize mineralogy. EDS confirmed identified minerals. BSE techniques
were used to prepare maps of thin sections and polished rock chips and to identify
adjacent rims on appropriate minerals interpreted to be in textural equilibrium. Sharp
grain boundaries between minerals in distinct textural domains were used as evidence for
textural equilibrium.
3.2.2 Electron probe microanalysis
Once sample mineralogy and deformation textures were documented, samples
were mounted and prepared for in situ geochemical analysis by electron probe
microanalysis (EPMA). Four samples were cut into ~1 cm3
rock chips and mounted in
~2.5 cm diameter epoxy rounds along with an oxygen-isotope standard UWQ-1 (Kelly et
al., 2007) at the center of each round. Thin sections of 17 samples were cut down to ~2.5
cm diameter circular thin sections with UWQ-1 quartz standard mounted in epoxy at the
center of each thin section.

35
Analyses were made using a Cameca SX-100 electron microprobe equipped with
four wavelength-dispersive spectrometers, natural and synthetic silicate standards, and
ZAF (ZAF refers to atomic number, absorption, and fluorescence) correction in the
Department of Earth & Planetary Sciences, University of Tennessee at Knoxville. All
epidote analyses were conducted with a spot size of ~1 μm, 20 kV accelerating voltage,
and 20 nA probe current over a two day analytical session. All K-feldspar analyses were
conducted with a spot size of 5 μm, a 10 kV accelerating voltage, and a 20 nA probe
current over the same two day analytical session. Backscattered electron imaging and
quantitative wavelength dispersive spectrometry (WDS) were conducted for in situ
chemical analysis of epidote (17 samples; 259 points), K-feldspar (2 samples; 4 points),
and plagioclase (6 samples; 74 points). Cation-oxide proportions in epidote were
calculated on the basis of 25 oxygens, and the pistacite composition of epidote, defined
by XFe (molecular iron / iron + aluminum), was found on the basis of 25 oxygen atoms.
Cation-oxide proportions in feldspars were found on the basis of 8 oxygen atoms.
Orthoclase (Or) composition of K-feldspar was determined as K / (K + Na + Ca), and the
anorthite content (An#) of plagioclase was calculated as Ca / Ca + Na.
3.2.3 Ion microprobe analysis
3.2.3.1 Sample preparation
Thin sections of four samples were cut into ~1 cm3
chips and mounted in one ~2.5
cm diameter epoxy round along with an oxygen-isotope standard UWQ-1 (Kelly et al.,
2007) mounted at the center of the round. All epoxy rounds and thin sections were

36
polished using 3 μm diamond suspension to minimize topographic effects that can cause
isotopic fractionation across the analysis area (i.e., Kita et al., 2009). All samples were
cleaned by sonicating in deionized water and ethanol multiple times, and then dried in a
vacuum oven. After drying, a ~30 nm Au coat was applied to each sample mount.
Detailed mineral maps were made prior to analyses based on SEM, EPMA, and optical
microscope imaging to guide spot positioning.
3.2.3.2 SIMS oxygen isotope analysis
Oxygen isotope analyses were made on selected minerals using a CAMECA ims-
1280 ion microprobe at the University of Wisconsin-Madison WiscSIMS Laboratory.
Oxygen isotope analyses of quartz, epidote, and k-feldspar in all prepared samples were
acquired in six consecutive 12-hour analytical sessions with a primary ion beam diameter
of 10–12 μm and depth of ~1 μm. Operating and analytical conditions followed those
described by Kita et al. (2009). Primary ion beam current ranged from 1.7–2.6 nA. The
working standard for all samples was UWQ-1 (12.33‰, Vienna Standard Mean Ocean
Water; VSMOW). The working standard in each sample was measured four times before
and after each 10–12 unknown analyses. The difference between the measured δ18
O
values (δ18
Oraw) of the quartz standard and the true δ18
O defines the instrumental mass
fractionation (IMF) for each bracket, which was then used to correct δ18
Oraw of the
unknowns to their true δ18
O (VSMOW). The IMF is known to vary systematically with
composition in minerals that show solid solution, known as the ‘matrix effect’, and so an
additional correction was applied to δ18
Oraw of epidote and K-feldspar using an in-house

37
calibration curve defined by analyzing compositionally-variable standards. Standards
used for matrix effects corrections were Tz-1 (epidote), Tz-3 (epidote), Corse1 (epidote),
CD23 (epidote), B28 (epidote), C30 (epidote), MES-4 (K-feldspar; Pollington, 2013),
FCS (K-feldspar; Pollington, 2013), and Gem28 (K-feldspar; Pollington, 2013).
Coexisting quartz, epidote, and K-feldspar were targeted for oxygen isotope thermometry
at spots within 6 mm of the center of the sample to avoid spatial variations in the IMF
(i.e., Kita et al., 2009).
3.2.3.3 Post-SIMS imaging
Following ion microprobe analysis, an additional thin coating of Au was applied
to each sample to minimize electron-charging effects in analysis pits. Every analysis pit
was then examined by secondary electron and BSE imaging with the Hitachi S3400N
SEM in the Department of Geoscience at the University of Wisconsin to assess pit
location, verify the mineral analyzed, and inspect for any cracks or other irregular pit
features (Figure 3.3). Energy-dispersive spectrometry was used to analyze pits placed in
extremely fine-grained and texturally-complex areas to verify mineral composition. A
total of 85 pits were observed to overlap grain boundaries or cracks, in which case they
were excluded from the following results and discussion.
3.3 Oxygen-isotope thermometry
Temperatures can be calculated provided that (1) experimentally-determined
values for the temperature-dependent oxygen isotope fractionation have been measured

38
and, (2) minerals have reached oxygen isotope equilibrium without retrograde change to
that composition. The latter constraint is often difficult to exclude in natural systems, and
variations in mineral δ18
O have been shown to be common over small areas (~1 cm2
) in
thin sections from low-grade metamorphic rocks (up to 7.4‰ in silicates; Ferry et al.,
2014). Various stable isotope studies of detachment zones have concentrated on the
oxygen isotope values from quartz-feldspar (e.g., Fricke et al., 1992; Morrison, 1994;
Holk and Taylor, 2000), quartz-muscovite (e.g., Mulch et al., 2007; Gottardi et al., 2011),
and quartz-epidote (e.g., Morrison and Anderson, 1998; MacDonald, 2014). These
studies find systematic variations in Δ18
O(m-n) = δ18
O(m) - δ18
O(n) (i.e., the thermal
gradient) between two minerals phases m and n across the damage zone of a given
detachment fault using laser fluorination measurements from whole rock and mineral
separates. Determining that two mineral phases are in isotopic equilibrium at the scale
required from laser fluorination is improbable and thus analysis by SIMS is appropriate
for useful oxygen-isotope thermometry.
Temperature calculations come from measuring the difference in δ18
O of two or
more isotopically equilibrated mineral phases (Δ18
O(m-n)) and a temperature-dependent
experimentally determined fractionation factor (α) between two isotopically equilibrated
minerals. The temperature-dependent oxygen isotope fractionation between two minerals
phases m and n is expressed as:
1000𝑙𝑛α 𝑛
𝑚
=
𝐴×106
𝑇2 + 𝐵 + C
where A, B, and C are experimental determined constants and T is temperature in Kelvin
(O’Neil et al., 1969). Temperature calculations between quartz and epidote in the study

39
used constants calibrated experimentally by Matthews (1994) of A = 2.180, B = 0.000,
and C = 0.000 for epidote having a composition of XFe = 24 for the temperature range of
0 – 1200°C. Temperature calculations between quartz and K-feldspar in the study used
constants calibrated experimentally by Zheng (1993) of A = 0.16 and B = 1.50, and C = -
0.62 for the temperature range of 0 – 1200°C.

40
Table 3.1. Samples from the Chemehuevi Mountains (SE CA) associated with the Mohave Wash fault used in this study.
Sample Site Sample type Host rock
Structural
position
relative to
MWF (m)a
Latitude Longitude Strike Dip
CG-13CH-24 Trampas Vein Gneiss 34° 35.996 114° 30.101
CG-13CH-30 Mohave Wash Chlorite breccia Granodiorite 0 34° 39.939 114° 30.327
CG-13CH-4 Studio Springs Vein Granodiorite 34° 34.850 114° 32.767
CG-13CH-60 Range Front Vein Granodiorite 34° 34.020 114° 35.183
CG-13CH-78 Mohave Wash Vein Gneiss -20 34° 40.039 114° 30.174 011°
CG-13CH-RF Range Front Undeformed Granodiorite
CG-13CH-RF Range Front Shear zone Granodiorite
CG-13CH-RF Range Front Ductile shear zone Granodiorite
13JL-7 Range Front Shear zone Granodiorite -5 34° 41.218 114° 35.230
13JL-8 Range Front Shear zone Granodiorite -5 34° 41.218 114° 35.230
CG-14CH-104 Saddle Chlorite breccia Granodiorite 1 34° 34.137 114° 34.524
CG-14CH-105 Saddle Undeformed Granodiorite 0 34° 34.137 114° 34.524
CG-14CH-105 Saddle Shear zone Granodiorite 0 34° 34.137 114° 34.524 135° 52°NE
CG-14CH-106 Saddle Undeformed Granodiorite -1 34° 34.137 114° 34.524
CG-14CH-106 Saddle Shear zone Granodiorite -1 34° 34.137 114° 34.524 132° 31°NE
CG-14CH-109 Saddle Shear zone Granodiorite -27 34° 34.13 114° 34.527 355° 22°E
CG-14CH-110 Saddle Chlorite breccia Granodiorite -28 34° 34.129 114° 34.539 355° 22°E
CG-14CH-111 Saddle Vein Granodiorite -88 34° 34.129 114° 34.539 025° 54°W
CG-14CH-112 Saddle Vein Granodiorite -97 34° 34.119 114° 34.649 060° 25°S
CG-14CH-113 Saddle Chlorite breccia Granodiorite -98 34° 34.119 114° 34.649 110° 54°NE
CG-14CH-124 Mouth Bat Cave Wash Vein Gneiss -4 34° 42.358 114° 29.669 165° 30°E
CG-14CH-125 Mouth Bat Cave Wash Vein Gneiss 0 34° 42.351 114° 29.648 Subvertical

41
Table 3.1. (continued)
Sample Site Sample type Host rock
Structural
position
relative to
MWF (m)a
Latitude Longitude Strike Dip
CG-14CH-126 Mouth Bat Cave Wash Vein Gneiss 4 34° 42.35 114° 29.641 165°
CG-14CH-127 Mouth Bat Cave Wash Undeformed Gneiss 24 34° 42.336 114° 29.612
CG-14CH-127 Mouth Bat Cave Wash Syntaxial vein Gneiss 24 34° 42.336 114° 29.612 110° Subvertical
CG-14CH-128 Mouth Bat Cave Wash Crack-seal vein Gneiss -3 34° 42.358 114° 29.669
CG-14CH-133 Bat Cave Wash Vein Gneiss -32 34° 41.59 114° 30.345 042° 24°SE
CG-14CH-134 Bat Cave Wash Vein Gneiss -31 34° 41.59 114° 30.345
CG-14CH-135 Bat Cave Wash Vein Hornblende-diorite -10 34° 41.577 114° 30.331 128° Subvertical
CG-14CH-136 Bat Cave Wash Chlorite breccia Gneiss -33 34° 41.59 114° 30.345
CG-14CH-137 Bat Cave Wash Vein Gneiss -37 34° 41.59 114° 30.345 135° Subvertical
a
Vertical structural position of samples taken from the Mohave Wash fault where “0 m” marks the base of the heavily fractured damage zone.

42
Figure 3.1: (a) Google Earth view to the east of the ~40 m thick Mohave Wash fault
(MWF) damage zone at The Saddle with sample locations shown. (b) View at the base of
the MWF main damage zone at The Saddle containing highly fractured granodiorite –
samples CG-14CH-104, CG-14CH-105, and CG-14CH-106 originate from this outcrop.
(c) Implosion breccia formed between crack-seal veins mineralized with epidote and
quartz within the MWF damage zone at The Saddle (sample CG-14CH-106). The CSZ
contains both angular fragments of host rock and epidote-rich veins. (d) Low-angle
fracture (< 5 mm thick) offsetting pegmatite by ~30 cm towards North, located ~25 m
below the MWF damage zone of The Saddle – samples CG-14CH-109 and CG-14CH-
110 originate from this outcrop. (Right and below) Structural profile schematically
depicting variations in fracture density from the footwall through the main damage zone
of the MWF. Fracture orientation data of a given sample represented in a pole stereonet
with Kamb density contouring (Kamb, 1959).

43
Figure 3.2: (a) Google Earth view to the NE of the ~10 m thick Mohave Wash fault
(MWF) damage zone at the mouth of Bat Cave Wash with sample locations shown. (b) A
representative ~1 cm thick vein mineralized with epidote and quartz within the MWF
damage zone (breccia in gneiss) at Bat Cave Wash. (c) Google Earth view to the NE of
the MWF damage zone and splay up Bat Cave Wash with sample locations shown. (d)
View of the 1 m thick MWF splay located up Bat Cave Wash containing altered gneiss
with cross-cutting epidote veins and the location of sample CG-14CH-133 shown. (Right)
Structural profile with estimated fracture density of the MWF measured at the two sites
along Bat Cave Wash with sample locations shown. Fracture orientation data of a given
sample represented in a pole stereonet with Kamb density contouring (Kamb, 1959).

44
Figure 3.3: Secondary electron image showing quartz analysis pits by ion microprobe
measuring 10 μm in diameter and ~1 μm in depth. The lower spot shows debris from ion
sputtering process ablating quartz of the allowing analysis (above).

45
4. Results
4.1. The Saddle Section: Generalized outcrop and sample description
The country rock in this section was primarily granodiorite intruded by minor
mafic dikes. The Mohave Wash fault (MWF) was recognized by a variable damage zone
up to 40 m in thickness consisting of cracked granodiorite, chlorite breccia, and cohesive
cataclasite. Figure 3.1 shows the 138 m vertical transect sampled from the base of the
most intensely fractured zone.
Chlorite breccia (CG-14CH-108, CG-14CH-107) makes up the top 10 meters of
The Saddle vertical transect with only very minor cataclasis found in localized 0.5–2 mm
thick shear zones (CG-14CH-109, CG-14CH-110, CG-14CH-113) in the lower 70
meters. Samples CG-14CH-104, CG-14CH-105, and CG-14CH-106 were taken from the
main deformation zone (0 m on fault column). Sample CG-14CH-104 was taken from a
representative chlorite breccia of the MWF damage zone containing thin (< 1 mm)
cataclasite shear zones. Sample CG-14CH-105 features a 1 cm thick shear zone
containing crack-seal veins striking 135° and dipping 52°NE. Sample CG-14CH-106
features a 2.5 cm thick cataclasite shear zone containing crack-seal veins striking 132°
and dipping 31°NE. Samples CG-14CH-109 and CG-14CH-110 were taken from slightly
altered-deformed host granodiorite at -27 m on the fault column featuring 1 cm thick
cataclasite shear zones striking 355° and dipping 22°E. Sample CG-14CH-111 was taken
at -88 m on the fault column and features a ~2 cm thick quartz + epidote vein in
undeformed granodiorite striking 025° and dipping 54°W. Sample CG-14CH-112 was
taken at -97 m on the fault column and features a ~1 cm thick cracked zone in

46
undeformed granodiorite containing quartz + epidote veins striking 060° and dipping
25°S. Sample CG-14CH-113 was taken at -98 m on the fault column from a 1–2 meter
thick isolated chlorite breccia containing thin (~1 mm) cataclasite zones striking 110° and
dipping 54°NE.
4.1.1 The Saddle Section: Petrographic and Microstructural description
The mineralogy and microstructural character of the footwall and MWF damage
zone at The Saddle section is based on 10 thin sections sampled within and below the 40
m thick MWF damage zone. Throughout the interval, primary igneous minerals in the
host granodiorite include quartz, k-feldspar, plagioclase, muscovite, and biotite. The
grains are typically 0.25–2 mm in diameter. The quartz shows weak undulatory
extinction, whereas the feldspars are undeformed. Plagioclase shows signs of incipient
alteration to fine-grained phyllosilicate. Biotite is altered to chlorite at grain boundaries.
In addition to clasts of the primary minerals, shear zones contain greenschist facies
mineral assemblages of chlorite + minor epidote +/- calcite. Quartz microstructures
within shear zones include microfractures and undulose extinction, but no evidence for
subgrain formation. Average grain size within shear zones is visually estimated to be <
0.5 mm, and bands up to 1 cm thick of fine-grained (< 10 μm) K-feldspar + quartz +
epidote are observed in some samples (CG-14CH-105, CG-14CH-106, CG-14CH-109;
Figure 4.2).
Samples CG-14CH-108, CG-14CH-107, and CG-14CH-104 feature thin
cataclasite zones within chlorite breccia ranging 100–500 μm in thickness of fine-grained
(< 10 μm) quartz, K-feldspar, albite, and epidote. Samples CG-14CH-107 and CG-14CH-

47
104 also contain thin calcite-rich veins (0.25–0.5 mm thick) cross-cutting the brecciated
granodiorite. Samples CG-14CH-108, CG-14CH-107, and CG-14CH-104 were not
targeted for δ18
O analysis due to the fine grain size (< 10 μm). The crack-seal veins found
at the margins of samples CG-14CH-105, CG-14CH-106, and CG-14CH-109 are also
found within the feldspar-rich shear zone as broken angular fragments and are evidence
of formation during slip along the MWF (Figures 4.1, 4.2). Sample CG-14CH-105 shows
a 1 cm thick crack-seal vein containing broken fragments (0.5–1 cm long segments) of
fine-grained (< 10 μm) quartz, K-feldspar, and albite cemented within a matrix of the
same composition having an average grain size of ~100 μm (Figure 4.2). Epidote found
within the shear zone sample CG-14CH-105 has an average grain size of ~5 μm. Sample
CG-14CH-106 features a 2.5 cm thick cataclasite shear zone of angular host granodiorite
fragments 1 mm to 1 cm in size cemented within a matrix of albite, K-feldspar, and
quartz grains 100–500 μm in size surrounded by 2 mm thick localized crack-seal veins of
fine-grained (10 μm) quartz and epidote (Figures 4.1, 4.2). SamplesCG-14CH-109 and
CG-14CH-110 feature a similar mineralogy between the host rock and shear zone of each
respective sample, consisting of fine-grained (< 10 μm) K-feldspar-albite and quartz
matrix with minor calcite and epidote (Figure 4.2). Samples CG-14CH-111 and CG-
14CH-112 contain a 0.1–2 mm thick vein infill by undeformed quartz and epidote.
Sample CG-14CH-113 contains multiple cataclasite zones within chlorite breccia ranging
0.25–1 mm in thickness of fine-grained (< 10 μm) quartz, K-feldspar, and epidote.

48
4.2 The Bat Cave Wash Section: Generalized outcrop and sample description
The country rock in this section was quartz-biotite gneiss with a prominent
foliation containing upper greenschist- to lower amphibolite-facies mineralogy and quartz
leucosomes. Greenschist facies shear zones related to Miocene deformation typically cut
the gneissic fabric, although slip along folia is also likely based on field observations.
Syntectonic dikes of felsic to mafic composition are commonly found within gneiss, and
mineral lineations at 050° were measured on several examples. The MWF was
recognized by a damage zone ~10 m in thickness and is represented by cracked gneiss,
chlorite breccia, and cohesive cataclasite. Mineralized fractures cutting gneiss below and
within the zone mapped as the MWF (John and Foster, 1993) were sampled. Due to poor
MWF footwall exposure, the section for Bat Cave Wash is a composite of two localities
separated by 1.75 km. The second transect was sampled deeper in the footwall and
crossed a sharp fault with evidence for substantial greenschist facies mineralization and
fluid flow (Figure 3.3). The sharp fault is interpreted as a deeper splay off the main
MWF, although no crosscutting relations were observed and the exact structural relations
are not clear. Figures 3.2 and 3.3 each show ~30 m vertical transects from the base of the
most intensely fractured zone.
Sample CG-14CH-127 is a 1 cm epidote-rich vein taken 24 m above the base of
the MWF. Sample CG-14CH-126 taken from the MWF damage zone (4 m on fault
column) contains cross-cutting veins striking 110° and with a subvertical dip. Sample
CG-14CH-125 was taken from a chlorite breccia zone containing cross-cutting veins at
the bottom of the MWF damage zone (0 m on fault column) with a subvertical dip.
Sample CG-14CH-128 features a ~1 cm thick zone of quartz + epidote veins taken 3 m

49
beneath the base of the MWF damage zone. Sample CG-14CH-124 features interspersed
quartz + epidote veins taken 1 m below CG-14CH-128 striking 165° and dipping 30°E (-
4 on fault column). Sample CG-14CH-135 was taken 22 m above the splay (~10 m below
the MWF) and features 5 mm thick quartz + epidote vein striking 128° and with a
subvertical dip. The host rock for sample CG-14CH-135 is a hornblende-biotite diorite
dike and is the only undeformed host rock found in the Bat Cave Wash section. Sample
CG-14CH-134 was taken < 0.5 m above the splay and features a quartz + epidote shear
zone. Sample CG-14CH-133 is from the middle of the 1 m thick MWF splay (~32 m
below the MWF) containing a quartz + epidote shear zone striking 042° and dipping
24°SE and many subvertical cross-cutting veins. Sample CG-14CH-136 was taken 1 m
below the splay. Sample CG-14CH-137, featuring epidote-rich veins striking 135° and
with a subvertical dip, was taken 5 m below the splay (~37 m below the MWF).
4.2.1 The Bat Cave Wash Section: Petrographic and Microstructural description
The mineralogy and microstructural character of the footwall and MWF damage
zone at Bat Cave Wash section is based on nine thin sections. Throughout the interval,
primary igneous minerals in the host quartz-biotite gneiss were dominated by quartz,
plagioclase, and biotite. Primary igneous minerals in the undeformed hornblende-biotite
diorite dike include plagioclase, hornblende, and biotite. The grains are typically 0.1–1
mm in diameter. The quartz grains have undulatory extinction. Biotite has been altered to
chlorite within and surrounding the MWF damage zone. Cross-cutting veins sampled
throughout the transect exhibit greenschist facies mineral assemblages of chlorite +
epidote + calcite + titanite. Surrounding cross-veins, gneissic foliation containing primary

50
biotite and plagioclase is typically heavily altered to fine-grained (< 10–100 μm) epidote
(Ep) and chlorite intergrowth among quartz (Qtz) ribbons (CG-14CH-126, CG-14CH-
134, CG-14CH-133). Thin calcite veins occasionally are observed cutting veins (CG-
14CH-135; Figure 4.6).
Sample CG-14CH-127 contains a syntaxial vein of coarse undeformed
(subhedral) quartz and epidote grains (~300 μm) cemented by calcite and bounded by 0.5
mm thick zones of fine-grained quartz and epidote (< 50 μm) cutting gneissic fabric.
Sample CG-14CH-126 features fine-grained (< 10–100 μm) epidote and chlorite
intergrowth among quartz ribbons containing subgrains parallel to the gneissic fabric as
well as a cross-cutting epidote vein containing quartz showing undulatory extinction
(Figure 4.3). Sample CG-14CH-125 features fine-grained (50 μm) quartz + epidote veins
cross-cutting gneissic fabric. Sample CG-14CH-128 features a 1 cm thick zone of quartz
+ epidote crack-seal veins showing sharp contacts (up to 1 mm in thickness) of reduction
in grain size with quartz-epidote grain sizes decreasing from 100 μm to 10 μm (Figure
4.4). Sample CG-14CH-124 features interspersed veins of epidote + quartz 0.5–3 mm
thick cutting gneissic fabric (Figure 4.5). Sample CG-14CH-135 features a 5 mm thick
zone of quartz, epidote, and calcite veins cutting undeformed host (Figure 4.6). Samples
CG-14CH-134 and CG-14CH-133 show abundant epidote intergrowth (10–200 μm in
size) among quartz ribbons (50–1000 μm in size) containing subgrains parallel to the
gneissic fabric with minor K-feldspar, Ca-plagioclase, and titanite. Sample CG-14CH-
133 contains undeformed cross-cutting veins of quartz and epidote with a grain size of
10–100 μm (Figure 3.1). Sample CG-14CH-136 features zones of chlorite and epidote
marked by a brecciated contact without identifiable gneissic fabric. Sample CG-14CH-

51
137 features a gneissic fabric containing chlorite as well as undeformed cross-cutting
quartz + epidote veins with a grain size of 20–100 μm cutting gneissic fabric (Figure 3.1).
All samples from the MWF damage zone located in Bat Cave Wash show well-defined
cross-cutting veins.
4.3 Vertical transect summary
Throughout both vertical transects sampled, the MWF was observed to consist of
zones < 1 m thick in which principle slip surfaces are concentrated as friable chlorite
breccia and cataclasite. The MWF damage zone was observed to be ~40 m thick at the
structurally-shallow levels of The Saddle cutting granodiorite (Figure 3.1). The fault
damage zone becomes thinner (~10 m) at Bat Cave Wash to the northeast, where it cuts
across gneissic banding within Proterozoic gneiss, and is defined again by fractured rock
with vein-fill and a zone of chlorite breccia and thin cataclasite shear bands (Figures 3.2,
3.3). Hydrothermal mineralization (especially epidote, chlorite) is observed within the
MWF damage zone at both The Saddle and Bat Cave Wash. Table 4.1 summarizes the
mineralogy and deformation of all analyzed samples.
Quartz in each of The Saddle samples commonly contains micro-fractures and
exhibits undulose extinction. Epidote was observed in seven of the ten Saddle samples
with a grain size of ~5–25 μm. Quartz, K-feldspar and albite, rather than quartz and
epidote, were the principle minerals precipitated within the most intensely fractured
section of the damage zone (0 m). Quartz-epidote mineralization was limited to veins and
shear zone margins at The Saddle. Quartz in Bat Cave Wash samples commonly feature

52
micro-fractures and subgrains in addition to exhibiting undulose extinction. Epidote was
observed in all of the Bat Cave Wash samples with a grain size of ~5–300 μm. Quartz
and epidote were found to be the principle secondary minerals precipitated, commonly
present in veins and along grain boundaries especially within the most intensely fractured
section of the MWF damage zone and splay in Bat Cave Wash.
Samples analyzed from The Saddle show structural evidence of multiple
generations of brittle deformation through cm-thick cataclasite shear zones and more
discrete, single events in the case of crack-seal veins. Samples analyzed from Bat Cave
Wash are more structurally complex relative to samples from The Saddle due to
interspersed veins and a strong preexisting fabric. Another major difference at Bat Cave
Wash is the presence of numerous plastically-deformed, 2–3 meter-wide dikes intruding
at the level of the mapped fault, and with lineations aligned with the documented regional
slip direction along the MWF of 050°, intruding in the footwall as well as into the fault
zone (Figure 3.2).
4.4 Additional samples
To further characterize the shallow portion of the WMF, samples CG-13CH-60,
CG-13CH-RF, 13JL-7, and 13JL-8 were incorporated in this study. These samples
originate from exposures of the MWF 1.25 km west of The Saddle (Figure 2.2). Sample
CG-13CH-60 is granodiorite containing a 5 mm vein of undeformed quartz and epidote
with a grain size of 50–100 μm as well as a second cross-cutting epidote vein ~1 mm
thick with undeformed quartz and epidote grains < 50 μm in size. Sample CG-13CH-RF

53
contains good evidence for strain localization and three discrete structural zones: a shear
zone containing fine-grained (< 50 μm) quartz, K-feldspar, and calcite with epidote
grains 1 mm is size; a foliated shear zone with deformed quartz, epidote, and K-feldspar;
and a cataclasite zone with large clasts (> 1 mm) of quartz with undulatory extinction and
K-feldspar set in a matrix of fine-grained (< 40 μm) epidote (Figure 4.17). Although CG-
13CH-RF is considered a well-preserved example of a shear zone related to Miocene
extension, it was found out of place as float near the MWF damage zone and its exact
location is not known. Based on the topography, it must have originated from west of
Chemehuevi Peak and thus from Range Front Wash. Sample 13JL-7 was taken at the
base of the MWF damage zone and features crack-seal fragments containing epidote
grains 100–1000 μm in size; secondary veins consist of epidote grains < 20 μm in size.
Sample 13JL-8 was taken alongside 13JL-7 and features epidote-quartz breccia with a
grain size of > 100 μm within an epidote matrix of < 10 μm cutting an andesitic dike at
the same location of 13JL-7.
Sample CG-13CH-4 was taken from an area 3 km NE of The Saddle (Figure 2.2).
This sample was found in contact with a 1 m thick, undeformed lamprophyre dike and
features a shear zone mineralized with epidote in granodiorite (Figure 4.7). The shear
zone contains angular quartz and epidote grains from 10–500 μm in size with the largest
grains located along the center and finest grains along the margins of the zone. The age of
the dike is unknown, and it is thus possible that this sample predates Miocene
deformation.
Sample CG-13CH-24 was taken from an area 7.5 km NE of The Saddle (Figure
2.2) and features an undeformed 0.5 cm thick quartz + epidote vein cutting a leucosome

54
within granodiorite of the MWF footwall. The vein that was sampled also cuts a 5 mm
thick quartz + epidote cataclasite interpreted as related to slip on the MWF. The
subhedral epidote grains of CG-13CH-24 are found up to 1 mm in length (Figure 4.8).
Samples CG-13CH-30 and CG-13CH-78 were taken from an area 13 km NE of
The Saddle (4 km south of the mouth of Bat Cave Wash; Figure 2.2). Sample CG-13CH-
30 was taken from the MWF damage zone and features a 1–2 cm thick breccia zone
consisting of quartz grains 1–2 mm in size within a matrix of principally quartz + epidote
grains 10–50 μm in size. Sample CG-13CH-78 was taken 20 m below the MWF damage
zone and features a 2 mm thick foliated quartz vein cutting a 1 mm thick epidote vein
within gneiss banding. These veins feature slicken surfaces with a lineation direction of
050°.
4.5 Electron probe microanalysis results
The (XFe) of epidote, orthoclase content (Or #) of K-feldspar, and anorthite
content (An #) of plagioclase from EPMA are presented in Table 4.2. 4.9 summarizes the
epidote composition of 17 samples taken from within and outlying the main damage zone
of the MWF. All epidote (n = 259) was found to be of intermediate XFe composition with
an average of 0.24 ± 0.03 SD (standard deviation) and ranging 0.17–0.34. This epidote
composition is consistent with compositions similar studies used for oxygen isotope
thermometry calculation (e.g., Morrison and Anderson, 1998; MacDonald, 2014).
Orthoclase content (Or#) of K-feldspar (n = 4) from two samples was found to average of
84.0 ± 14.9 SD and ranging from 62.6–95.1. Anorthite content (An#) of plagioclase (n =

55
74) from six samples was found to be relatively sodic with an average of 9.8 ± 8.8 SD
and ranging from 0–34.8. Results of a given weight percent oxide for epidote, K-feldspar,
plagioclase from EPMA are assembled in Appendix Tables A.1 and A.3 respectively.
Cation proportions for epidote and K-feldspar from EPMA are assembled in Appendix
Tables A.2 and A.4 respectively.
4.6 Oxygen isotope results
A total of 480 analyses (excluding standards and defective analyses) made of 317
mineral grains (quartz, epidote, or K-feldspar) in 23 samples were analyzed for δ18
O
(Figure 4.10). A total of 85 analyses were excluded after post-SIMS imaging revealed the
analytical spots had overlapped grain boundaries, cracks, or mineral inclusions. The
number of excluded spots are a result of the fine-grain size (< 10 μm) of many of the
targeted areas. Multiple grains were often analyzed in areas ~5 mm2
or smaller, and
always within the inner 1 cm diameter of each sample to avoid X-Y instrumental
fractionation affects (Kita et al., 2009). Analyses of multiple grains of the same mineral
were made within each microstructural domain, typically separated by 1–2 mm, in order
to evaluate intercrystalline variability of δ18
O. Analysis of cores and rims on the same
grains allows for the evaluation of homogeneity and core-rim zonation patterns of δ18
O.
Several grains within a given microstructure received one analysis at the center and two
rim analyses at opposite sides. However, grains < 40 μm in diameter commonly received
only one analysis. Examples of the SIMS spot locations are shown in Figure 4.11 (sample
CG-14CH-127).

56
Undeformed quartz from host rocks sampled away from the MWF yield δ18
O
values ranging 9.0–10.4‰ (MacDonald, 2014). Analyses from shear zones of the
associated MWF damage zone yield δ18
O values consistently lower (Figure 4.12). All
measurements by ion microprobe of δ18
O of the standard and unknowns, instrument
settings and analysis readings, corrections for instrumental mass fractionation (IMF), and
measured compositions are assembled in Appendix A; Table A.5. Table 4.3 summarizes
all 480 accepted measurements of δ18
O of quartz, epidote, and K-feldspar in 23 analyzed
samples. The average and range for ±2SD (2 standard deviations) for the working
standard (UWQ-1) over all analytical sessions were 0.30‰ and 0.13–0.46‰,
respectively. The 2SD of the bracketing standards (external error) is assigned as the
uncertainty on each unknown analysis within the respective bracket.
4.6.1 Oxygen isotope composition of The Saddle
The petrographic relation between δ18
O of quartz, epidote, K-feldspar, and the
working standard analyses locations over different microstructural domains from samples
CG-14CH-105, CG-14CH-106, CG-14CH-109, CG-14CH-111, CG-14CH-112, and CG-
14CH-113 from The Saddle is summarized in Figure 4.13a. Analyzed textures from The
Saddle were described as either crack-seal veins (contain undeformed minerals),
cataclasite shear zones, undeformed quartz + epidote veins, chlorite breccia, and
undeformed host.
Samples CG-14CH-105, CG-14CH-106, and CG-14CH-109 feature cataclasite
shear zones bounded by crack-seal veins containing fine-grained (< 10–100 μm) quartz,
epidote, and K-feldspar. Quartz grains from the host rock and 0.1 mm outside of the shear

57
zone of sample CG-14CH-105 give δ18
O values of 9.0‰ at the rim and 10.0 to 10.2‰ at
the core. Quartz measured within crack-seal veins of sample CG-14CH-105 give
considerably lower δ18
O values between -1.0 to 0.7‰ (Figure 4.15). K-feldspar measured
inside the crack-seal vein of sample CG-14CH-105 give δ18
O values from -2.1 to 1.8‰.
No K-feldspar from the host rock were analyzed, but values of 8–9‰ are expected for K-
feldspar in magmatic equilibrium with δ18
OQtz = 10‰ (fractionation factor of Blattner et
al., 1974).
Quartz grains from the host rock and up to 4 mm outside of the shear zone of
sample CG-14CH-106 give δ18
O values of 1.3 and 1.5‰ at the rims of grains and 9.3 to
10.8 10.8‰ at the cores (Figure 4.15). Quartz measured within the crack-seal veins give
δ18
O values of 5.5 ‰ at the rims of clasts. Epidote measured 4 mm outside of the shear
zone give δ18
O values from -5.1 to -4.4‰ compared to δ18
O values from -4.6 to -3.5‰
measured 50 μm inside of the crack-seal veins. Epidote measured 500 μm inside the
crack-seal veins give δ18
O values of -4.1 to -3.5‰. K-feldspar measured 0.1 mm outside
of the shear zone give a δ18
O value of 0.0‰ and δ18
O values of -2.0 to -1.1‰ inside the
crack-seal vein. Only two quartz-epidote mineral pairs were measured from these
samples, yielding Δ18
O(Qtz-Ep) values of 6.3 and 6.1‰ from 4 mm outside of the shear
zone of sample CG-14CH-106.
Quartz clasts measured within the crack-seal veins of sample CG-14CH-109 give
δ18
O values of 10.1 to 10.6‰. K-feldspar measured within the crack-seal veins give δ18
O
values of -2.6 to -0.9‰ at the rims of grains and -2.7 to -0.3‰ at the cores.
Samples CG-14CH-111 and CG-14CH-112 feature cracked zones in undeformed
granodiorite containing a 0.1–2 mm thick quartz + epidote veins of possible multiple

58
generations. Quartz measured within the vein of sample CG-14CH-111 give δ18
O values
from 10.2 to 10.9‰. Epidote measured within the vein of sample CG-14CH-111 give
δ18
O values of 2.6 to 5.3‰ (Figure 4.15). Epidote measured within the vein of sample
CG-14CH-112 give δ18
O values from 1.8 to 6.6‰ with the larger of the two veins
analyzed having values from 1.8 to 5.7‰ and the smaller of the two veins analyzed
having values from 4.6 to 6.6‰ (Figure 4.15). Only two quartz-epidote mineral pairs
were measured from these samples, yielding Δ18
O(Qtz-Ep) values of 5.9 and 5.3‰ from a 1
mm thick quartz + epidote vein of sample CG-14CH-111.
Sample CG-14CH-113 was taken from a 1–2 meter thick isolated chlorite breccia
containing cataclasite zones. Quartz measured from the brittle shear zone give δ18
O
values of 5.6 to 8.5‰ at the rims of grains and 8.4 to 10.0‰ at the cores. K-feldspar
measured from the same shear zone give δ18
O values from -1.1 to 1.4‰.
4.6.2 Oxygen isotope composition of Bat Cave Wash
O of quartz, epidote, and the working
standard analyses locations over different microstructural domains from the Bat Cave
Wash is summarized in Figure 4.13b-c. Analyzed textures from Bat Cave Wash include
crack-seal veins, cataclasite shear zones, undeformed quartz + epidote veins, and host
gneiss.
Sample CG-14CH-127 contains a syntaxial vein of large undeformed, euhedral
quartz and epidote grains (~300 μm) bounded by fine-grained 0.5 mm thick undeformed
zones epidote-rich veins with minor quartz (< 50 μm) cutting a gneissic fabric (Figure
4.11). Quartz measured within the central coarse-grained zone give δ18
O values from 2.2

59
to 3.7‰ with the exception of two core δ18
O values of 8.6 and 9.1‰. Quartz measured
within the vein wall gives δ18
O values from 4.2 to 6.2‰. Epidote measured within the
central coarse-grained zone give δ18
O values of -3.1 to -2.0‰ at the rims of grains and -
3.1 to -1.6‰ at the cores. Epidote measured within the vein wall give δ18
O values from -
1.9 to 1.0‰ (Figures 4.4.1, 4.4.2, 4.4.7). Seven quartz-epidote mineral pairs were
measured from two structurally distinct zones separated by ~5 mm. Quartz-epidote
mineral pairs from the central coarse-grained zone yield Δ18
O(Qtz-Ep) values of 5.8 and
5.1‰. Quartz-epidote mineral pairs from the vein wall yield comparable Δ18
O(Qtz-Ep)
values from 5.4 to 4.9‰ (n = 3).
Sample CG-14CH-126 features fine-grained (< 10–100 μm) epidote and chlorite
intergrowth among quartz ribbons containing subgrains parallel to a gneissic fabric as
well as a cross-cutting epidote vein. Quartz measured within gneissic intergrowth give
δ18
O values of 4.4 to 6.5‰ and δ18
O values from 5.4 to 6.4‰ within the crosscutting
epidote-rich vein. Epidote measured within gneissic intergrowth yielded δ18
O values from
-2.7 to 3.3‰ and δ18
O values from -2.9 to 3.4‰ from within the crosscutting vein (Figure
4.16). Eight quartz-epidote mineral pairs were measured from two structurally distinct
zones separated by ~2–3 mm. Quartz-epidote mineral pairs from gneissic intergrowth
yielded large Δ18
O(Qtz-Ep) values from 8.1 to 6.2‰ (n = 6). A quartz-epidote mineral pair
from the epidote vein yields a Δ18
O(Qtz-Ep) value of 9.3‰.
Sample CG-14CH-125 features well-defined fine-grained (50 μm) epidote-rich
veins. Quartz measured within the host gneiss give δ18
O values from 5.8 to 7.1‰. Quartz
measured within the epidote-rich veins give similar δ18
O values from 6.6 to 7.6‰.

60
Epidote measured within the veins give δ18
O values from -0.6 to 3.4‰. Quartz-epidote
mineral pairs from the epidote vein yield Δ18
O(Qtz-Ep) values from 7.2 to 4.2‰ (n = 5).
Sample CG-14CH-128 features crack-seal veins with grain sizes decreasing from
100 μm to < 10 μm. A single coarse quartz grain analysis gives a δ18
O value of 6.0‰.
Quartz measured within the finest-grained (< 50 μm) zone gives similar δ18
O values from
4.3 to 6.3‰. Epidote measured within the coarse-grained zone gives δ18
O values of 0.5
and 1.5‰. Epidote measured within the finest-grained zone give δ18
O values of -3.8 to
1.7‰. A single quartz-epidote mineral pair of the coarse-grained epidote yields a
Δ18
O(Qtz-Ep) value of 4.5‰. Quartz-epidote mineral pairs from the finest-grained textures
yield Δ18
O(Qtz-Ep) values from 8.0 to 3.9‰ (n = 8).
Sample CG-14CH-124 contains interspersed quartz + epidote veins cutting a
gneissic fabric. Quartz measured within veins give δ18
O values from 4.3 to 4.8‰. Epidote
measured within veins give δ18
O values from -2.2 to -0.4‰. Two quartz-epidote mineral
pairs measured from veins yield Δ18
O(Qtz-Ep) values of 6.4 and 5.2‰.
Sample CG-14CH-135 features veins of fine-grained (10–50 μm) quartz and
epidote; the veins contain undeformed 100–500 μm thick calcite veins containing
undeformed quartz and epidote grains 10–100 μm in size. δ18
O values of quartz measured
within the epidote-rich veins ranged from 7.9 to 9.1‰ and from 8.1 to 9.1‰ within the
calcite veins. Epidote measured within the epidote-rich veins give δ18
O values from 0.6 to
2.0‰. Epidote measured within the calcite veins give δ18
Quartz-epidote mineral pairs from the quartz + epidote vein yield Δ18
O(Qtz-Ep) values of
8.2 to 6.5‰ (n = 4). Quartz-epidote mineral pairs from the calcite vein yield Δ18
O(Qtz-Ep)
values from 7.3 to 6.5‰ (n = 3).

61
Samples CG-14CH-133 and CG-14CH-134 feature fine-grained (< 10–100 μm)
epidote, quartz, and chlorite intergrowth among quartz ribbons surrounding cross-cutting
fine-grained (< 20 μm) epidote veins. These samples are the most structurally complex
and therefore different deformation events are difficult to distinguish. Quartz measured
within gneissic intergrowth from CG-14CH-133 gives δ18
O values from 1.1 to 7.6‰
(Figure 4.16). Epidote measured within gneissic intergrowth from CG-14CH-133 gives
δ18
O values from -5.3 to -1.7‰. Epidote measured within a single distinguishable vein
from CG-14CH-133 gives δ18
O values from -3.9 to -3.4‰ (Figure 4.18). Quartz-epidote
mineral pairs from the zone within gneissic intergrowth from CG-14CH-133 yield
Δ18
O(Qtz-Ep) values of 12.9 to 4.9‰ (n = 4). Quartz-epidote mineral pairs from the quartz
+ epidote vein within CG-14CH-133 yield Δ18
O(Qtz-Ep) values of 8.6 and 6.8‰. Quartz
measured within gneissic intergrowth from CG-14CH-134 give δ18
6.0‰. Epidote measured within gneissic intergrowth from CG-14CH-134 give δ18
O
values from -2.9 to -0.5‰. Quartz-epidote mineral pairs measured from sample CG-
14CH-134 yield Δ18
O(Qtz-Ep) values of 6.9 to 5.7‰ (n = 4).
Sample CG-14CH-137 features zones of plastic deformation fabric parallel to the
gneissic fabric containing quartz, epidote, and chlorite as well as undeformed cross-
cutting epidote-rich veins with a grain size of 20–100 μm. Primary quartz within gneissic
fabric cut by veins give δ18
O values from 4.0 to 5.9‰. Quartz measured within the
largest epidote-rich vein (1.5 mm wide) gives δ18
O values from 3.2 to 3.8‰. Quartz
measured within a thin epidote-rich vein (0.1 mm wide) gives δ18
3.3‰. Epidote measured within the largest epidote-rich vein of CG-14CH-137 gives δ18
O
values from -3.4 to -2.4‰. Epidote measured within the thin epidote-rich vein gives δ18
O

62
values of -2.3 and -2.0‰ (Figure 4.16). Quartz-epidote mineral pairs from the largest
epidote-rich vein yield Δ18
O(Qtz-Ep) values of 7.0 and 5.7‰. Quartz-epidote mineral pairs
from the thin epidote-rich vein yield Δ18
O(Qtz-Ep) values of 4.9 and 4.8‰.
4.6.3 Oxygen isotope composition of additional MWF samples
O of quartz, epidote, K-feldspar, and the
working standard analyses locations over different microstructural domains from
additional samples are summarized in Figure 4.14. Analyzed textures from additional
MWF samples include crack-seal veins, brittle deformed cataclasite shear zones,
undeformed quartz + epidote veins, foliated quartz + epidote veins, chlorite breccia, and
undeformed host.
Sample CG-13CH-60 is granodiorite containing a 0.5 cm vein of quartz and
epidote with a grain size of 50–100 μm as well as an inner epidote vein 1 mm thick with
quartz and epidote grains < 50 μm in size. Epidote measured within the inner epidote vein
gives δ18
Sample CG-13CH-RF contains three structural zones described in section 4.4
(Figure 4.17). Quartz measured within all three zones is quite similar, giving δ18
O values
of 7.9 to 9.0‰. Epidote measured within all three zones is also similar, giving values
from 4.2 to 6.1‰. Quartz-epidote mineral pairs of the epidote cataclasite zone yield
Δ18
O(Qtz-Ep) values of 3.9 and 3.7‰. A single quartz-epidote mineral pair from the ductile
zone yields a Δ18
O(Qtz-Ep) value of 3.1‰. Quartz-epidote mineral pairs from a coarse zone
containing subhedral epidote yield Δ18
O(Qtz-Ep) values of 2.6 and 2.3‰. K-feldspar
measured from the cataclasite zone give δ18

Brown, James accepted thesis 11-24-15 Fa 15-3

Brown, James accepted thesis 11-24-15 Fa 15-3

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