Hydrothermal Circulation During Slip on the Mohave Wash Fault, Chemehuevi
Mountains, SE CA: Oxygen Isotope Constraints
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
Cody J. MacDonald
August 2014
© 2014 Cody J. MacDonald. All Rights Reserved. 2 This thesis titled
Hydrothermal Circulation During Slip on the Mohave Wash Fault, Chemehuevi
Mountains, SE CA: Oxygen Isotope Constraints
by
CODY J. MACDONALD
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
3 ABSTRACT
MACDONALD, CODY J., M.S., August 2014, Geological Sciences
Hydrothermal Circulation During Slip on the Mohave Wash Fault, Chemehuevi
Mountains, SE CA: Oxygen Isotope Constraints
Director of Thesis: Craig B. Grimes
Fluids are likely significant during the life-cycle of low-angle normal faults
(LANFs) as well as other fault systems, but the role of those fluids and their source at fault initiation are unclear. The Mohave Wash Fault (MWF), a LANF situated within the
Chemehuevi Mountains core complex (SE CA), offers a well-exposed site to evaluate this question. The MWF slipped 1-2 km during the Miocene before being denuded passively to the surface by extension localized on the higher-level Chemehuevi
Detachment Fault. To evaluate fluid-rock interactions during the early slip on this fault
18 system, δ O values of whole rocks, quartz, and epidote were measured by CO2-laser fluorination and interpreted along with field and microscopic observations of fault rocks from this area.
The MWF damage zone is variable in thickness and characterized by cracked granitic rocks hosting mineralized fractures, cohesive cataclasites, thin foliated shear zones, and rare pseudotachylite. δ18O of quartz hosted by undeformed granite ranges from 9.0-10.3‰, defining predeformation values. Foliated shear zones and quartz veins
18 18 extend to lower δ OQtz from 10.1-6.1‰, while cataclasites record the lowest δ OQtz values down to 1.1‰. The δ18O values of epidote (from all types) ranges from 5.3‰ to -
0.4‰; the lowest values are generally in cataclasites. The shifts to lower δ18O are
18 explained by interaction with heated, low δ O fluids from an external source (evolved
4 meteoric fluids or basin brines). Apparent temperatures from stable isotope thermometry on coexisting quartz and epidote (from 0.5 cc of rock) from the footwall are typically 50-
150˚C higher than ambient footwall temperatures at 23 Ma (fault initiation) determined using 40Ar/39Ar closure temperatures (John and Foster, 1993). Temperatures defined by both methods increase in the paleodip direction. The temperature difference across the
18 footwall estimated from Δ O(Qtz-Ep) versus Ar/Ar closure temperatures either indicates the mineralization occurred prior to Ar/Ar closure or reflects localized upwelling of hot,
18 deep-seated fluids during slip along the MWF. Calculated δ OH2O in equilibrium with mineral pairs decreases with lower temperature, consistent with influx of progressively lower δ18O fluids with decreasing temperature and depth.
5 DEDICATION
To my family, you are my foundation.
6 ACKNOWLEDGMENTS
This work would not have been possible without the help and support from Craig
Grimes, Barbra John, Justin LaForge, and James Brown.
7 TABLE OF CONTENTS
Page
Abstract ...... 3 Dedication ...... 5 Acknowledgments ...... 6 List of Tables ...... 9 List of Figures ...... 10 Chapter 1: Introduction ...... 11 Chapter 2: Geologic Background ...... 15 Regional Geologic Framework ...... 15 Cretaceous Granites – Chemehuevi Plutonic Suite ...... 17 Proterozoic Gneiss ...... 18 Mohave Wash Fault Footwall Geology ...... 18 Chemehuevi Dike Swarm ...... 19 Existing Thermal Structure of the Footwall ...... 20 Constraints on Hydrothermal Fluid Circulation along LANFs ...... 21 Brief Overview of δ18O Studies on Selected Core Complexes ...... 23 Chapter 3: Methods ...... 28 δ18O System ...... 28 Fractionation ...... 29 Water-Rock Interactions and Oxygen Isotope Thermometry ...... 30 Formation of Hydrothermal Epidote ...... 32 Analytical Techniques ...... 33 Chapter 4: Results ...... 36 Terminology ...... 36 Field Relations ...... 36 Petrographic Observations from Damage Zones and Hydrothermal Products ...... 38 Monomineralic Quartz Veins ...... 38 Foliated Shear Bands ...... 38 Injection Vein - Pseudotachylite ...... 39 Shear Band – Green Cataclasites ...... 39
8 Shear Band – Black Cataclasite ...... 40 Quartz + Epidote Veins ...... 40 Oxygen Isotopes ...... 41 Oxygen Isotope Signature of Cretaceous Granitoids ...... 41 Oxygen Isotope Signature of Precambrian Gneiss ...... 41 Oxygen Isotope Signature of the Chemehuevi Dike Swarm ...... 42 Oxygen Isotope Signature of Deformation Structures ...... 42 Whole Rock Data ...... 43 Chapter 5: Discussion and Conclusions ...... 44 Characteristics of the Mohave Wash Fault ...... 44 Constraints on Fluid-Rock Interactions from Oxygen Isotope Ratios ...... 44 Origin of Epidote: Magmatic Fluids or Externally-Derived Hydrothermal Fluids? .... 46 Isotopic Equilibrium and Oxygen Isotope Thermometry ...... 47 Temperature of Fluid-Rock Interactions Along the Mohave Wash Fault ...... 49 Chemehuevi Mountain Isotopic Signature – The Story of Two Fluid Regimes ...... 50 Footwall Isotope Thermometry Variation from Mineral Closure Temperatures ...... 52 Fluids Role in the Development of the Chemehuevi Metamorphic Core Complex ..... 56 Conclusions ...... 58 Tables ...... 60 Figures ...... 63 References ...... 85
9 LIST OF TABLES
Page
Table 1: Overview of fluid source and isotopic signature ...... 60
Table 2: Overview of δ18O studies on metamorphic core complexes ...... 60
Table 3: Laser fluorination δ18O data for quartz, epidote, garnet and whole rocks from the
Chemehuevi Mountains, SE CA ...... 61
10 LIST OF FIGURES
Page
Figure 1. Active low-angle normal fault and possible fluid flow pathways...... 63
Figure 2. Geologic map of the Chemehuevi Mountains...... 64
Figure 3. Cross section from A to A’ ...... 65
Figure 4: Mohave wash fault lithological profile ...... 66
Figure 5: Mohave wash fault outcrop characteristics ...... 67
Figure 6: δ18O vs. structural rock type ...... 69
Figure 7: Whole rock δ18O vs. structural rock type ...... 70
Figure 8: Deformation microstructures ...... 71
Figure 9: Whole rock δ18O vs. vertical distance from damage zone ...... 73
Figure 10: δ18O vs. distance in the slip direction ...... 76
18 18 Figure 11: δ OQtz vs. δ OEp ...... 75
Figure 12: Mineral pairs and associated calculated temperatures ...... 76
Figure 13: Temperature vs. distance in the slip direction ...... 77
18 18 Figure 14: Temperature, δ OH2O, and δ OQtz vs. vertical distance from damage zone 78
Figure 15: Geographic location of calculated temperatures ...... 79
18 18 Figure 16: Temperature vs. δ OH2O, and δ OQtz ...... 81
Figure 17: Model for fluid circulation during the development of the Chemehuevi Mountains metamorphic core complex ...... 82
11 CHAPTER 1: INTRODUCTION
Fluids are likely to be significant during the evolution of low-angle normal faults, but the role of those fluids and their source at fault initiation are unclear (Figure 1). Low
Angle Normal Faults (LANFs) initiate at <30° below horizontal and are an unusual class of geologic faults because of their orientation, yet they appear frequently in the rock record and in geophysical surveys over western North America and other extensional environments around the globe. The mechanics of formation and slip on LANFs remains an ongoing debate, similar to the debate over strength of the San Andreas fault (Faulkner et al, 2006) Classic rock mechanics theory (Anderson, 1951) dictates that frictional forces are too high to allow fault slip at such low angles, and that a fault dipping ~60° should break to accommodate strain. The standard friction coefficient values determined for faults in the laboratory, 0.6<µs<0.85 (Byerlee, 1978), predict that normal faults should cease to slip at dips of <30°-40° (Collettini et al., 2009). Despite the mechanical theory, many natural examples of LANFs are known from a variety of geologic settings, including mid-ocean ridges, rifted margins, and metamorphic core complexes. Some examples of (microseismically) active LANFs include the Dixie Valley fault in Nevada and the Alto Tiberina fault in Italy (Collettini, 2011), yet their origin remains a topic of continual debate. Also, there has been much discussion on the earthquake paradox and seismic slip along LANFs (i.e., Anderson, 1951; Sibson, 1985; Arbasz & Julander, 1986;
Jackson, 1987; Jackson and White, 1989; Collettini & Sibson, 2001; Axen 2004). The earthquake paradox is that LANFs accommodate large amounts of displacement and yet they are not associated with large magnitude earthquakes. Only a handful of examples of seismic slip on LANFs have been attributed to historic intermediate and large magnitude
12 earthquakes (Abers, 1991, 2001; Wernicke, 1995; Axen, 1999). The rarity of significant earthquakes (greater than magnitude 3) in the historic record has led to the suggestion by some that continental detachments were simply not initiated in a low angle
(<30°) orientation, and/or are anomalously weak (e.g., Collettini, 2011). Alternatively, large magnitude seismic events may occur on said faults, but slip intervals are long relative to the period of observation (see Abers, 2009).
Several hypotheses have been offered to account for slip on LANFs: 1) Passive rotation of an originally higher-angle fault surface; 2) Rotation of the principal stress from vertical such that adjacent to the fault zone the slip surface is oriented consistent with Andersonian theory (Spencer and Chase, 1989; Yin, 1989; Chery, 2002); 3) The planar region along which the fault forms is weakened by alignment of low-friction minerals formed during hydrothermal alteration; 4) Increased pore-fluid pressure may aid in continued slip after the fault has formed; over time the permeability decreases and compaction of the fault leads to overpressure and promotes slip (Byerlee, 1990; Rice,
1992; Faulkner and Rutter, 2001); 5) The fault may slip on a ‘rolling hinge’, where high- angle faults grade listrically into a deeper detachment rooted near the brittle-ductile transition zone, and then roll to shallower dips through isostatic uplift (e.g., Wernicke and
Axen, 1988); 6) Once formed, fault weakening and strain localization may be promoted by ‘refrigeration’ of the footwall by circulating hydrothermal fluids (e.g., Morrison and
Anderson, 1998). Models 3, 4, and 6 are directly related to the circulation of hydrothermal fluids to mid-crustal depths.
The Chemehuevi Mountains (Figure 2) offer an excellent location to study the relationship between hydrothermal fluids, and early deformation along LANF’s owing to
13 the exposure of two stacked, low-angle normal faults that have been previously mapped, and for which detailed thermochronologic data is available that constrain the ambient thermal structure of the footwall at initiation. In the Chemehuevi Mountains, field relations show the LANFs were the most recent faults to slip, precluding the possibility of the faults being rotated from a higher dip to their present orientation
(hypothesis 1 above). Unrotated dikes emplaced prior to/during faulting and thermochronologic constraints on the initiation dip-angle of the Chemehuevi detachment fault (CDF) and Mohave Wash fault (MWF) system support a shallow initial dip of <30°
(John and Foster, 1993). Extensive Miocene-related alignment of phyllosilicates or other low-friction minerals are not observed in the damage zone of the MWF, precluding hypothesis 3 above. Models involving crustal-scale fluid flow and fault weakening seems to be more plausible for the development of this (and likely other) detachment normal fault systems (e.g., Sibson, 2000; Collettini, 2009), but the transport mechanisms and sources of fluids at depth during the incipient stages of faulting remain unclear.
The goal of these investigations is to quantify the imprint of fluids along the
Mohave Wash Fault using stable isotope geochemistry and petrology, in an effort to constrain the timing and source of fluids available during faulting. The MWF slipped during the initial phase of deformation related to the regional-scale CDF system, but was later trapped and denuded as part of the footwall to the higher-level Chemehuevi Fault
(Figure 3). Three questions in particular will be addressed by this study: (1) what role, if any, did hydrothermal fluids play in early microstructure development? (2) Is there evidence for surface-derived fluids during deformation, or were any fluids present magmatic/metamorphic in origin? (3) What is the temperature of mineralization, fluid
14 flow, and deformation across the study area as defined by stable isotope thermometry – is there evidence for rapid cooling along the fault driven by convection of hydrothermal fluids as previously proposed by Morrison and Anderson (1998)?
15 CHAPTER 2: GEOLOGIC BACKGROUND
Regional Geologic Framework
The Chemehuevi Mountains lie within the Colorado River Extensional Corridor
(CREC) (Figure 2 and 3). The CREC underwent up to 100% crustal extension between
23 and 12 Ma (John and Foster, 1993), and is a product of crustal relaxation in conjunction with Basin and Range extension. Extension was accomplished along a series of northeast-dipping, detachment fault systems. Detachment faults related to Basin and
Range extension are also exposed around domal core complexes in the central part of the
CREC throughout the Whipple and Buckskin Mountains, Chemehuevi Mountains and the
Dead Mountains. Past field and microstructural studies indicate motion of the upper plates of each fault was to the northeast (Davis et. al., 1980, John, 1982; Howard et. al,
1982; Spencer, 1985). Northeast slip-direction indicators are present on all scales from outcrop to regional. Slickenlines, lineations, offset markers, preserved striae, drag folds, minor faults within related cataclasites, and the southwest dip of syntectonic strata above the CDF are observed (Yin and Dunn, 1992). On the regional to local scale (amplitude
100s of meters), the CDF and MWF exhibit corrugations parallel to the slip direction.
The amount of slip on the fault system that accommodates extension in the CREC increases to the northeast and totals an estimated 40-75 km (Howard and John, 1987;
Davis 1988; Spencer and Reynolds, 1991; John and Foster, 1993). Initiation of crustal stretching in the CREC began ~23 Ma based on K-Ar ages of volcanic rocks within synextensional basins, the crystallization ages of the oldest syntectonic plutons within core complexes, and 40Ar/39Ar cooling ages within footwall rocks (Howard and John,
1987; Spencer and Reynolds, 1991; Anderson et al., 1988). The period of extension
16 recorded in the Chemehuevi footwall rocks is interpreted to have ended by ~13-14 Ma, based in large part on fission track (apatite) and 40Ar/39Ar (biotite, K-feldspar) ages on footwall rocks (Foster and John, 1993).
John and Foster (1993) suggest that the fault system containing the CDF and
MWF crops out in the Old Woman Mountains that are ~45 km east of the Chemehuevi
Mountains, and the Sacramento Mountains (~10 km to the northwest). The Chemehuevi
Detachment Fault is the youngest and is of greatest regional significance, and has been equated with the nearby Whipple detachment fault exposed in the Whipple Mountains 40 km to the south (John, 1987a; John and Foster, 1993).
Within the Chemehuevi Mountains, the hanging-wall rocks of the CDF are crosscut by many high-angle faults that rotated to gentler dips with progressive slip
(John, 1987a). These high-angle faults do not cut through the regional low-angle detachment fault in any known location, but the opposite relationship is reported. For example, the more steeply dipping Devil’s Elbow fault within the hanging-wall block is truncated by the younger low-angle CDF (Figure 3). These field-based observations indicate that no significant passive rotation could have occurred to accommodate the present-day low-angles measured for the CDF and MWF. Instead, all indications are that the slip occurs at this low angle.
The Mohave Wash fault is the lowest exposed fault in the Chemehuevi Mountains
(Figure 3), exposed as a sinuous trace over more than 350 km2 (John and Foster, 1993).
The MWF fault accommodated 1-2 km of slip, but then ceased and was denuded to the near surface by the overlying CDF, preserving initial faulting structures and mineralization developed at initiation depth. Both faults dip gently (10-15˚) SW in the
17 west, and (2-15˚) in the east, reflecting doming in response to isostatic uplift driven by hot, buoyant footwall rocks (John, 1988) At outcrop scale the fault is roughly planar, but at map scale it features corrugation parallel to slip (John, 1987). At structurally shallow levels (1-5 km depth at initiation), the MWF is characterized by principle slip zones < 1 m thick comprised of chloritized cataclasites, with an asymmetric damage zone 10s m thick. At the deepest exposed structural levels (12-15 km paleodepth at initiation) the
MWF features similar asymmetric distribution of damage zone thickness with the addition of Miocene syn-extensional dikes featuring mylonitic deformation.
Cretaceous Granites – Chemehuevi Plutonic Suite
The Chemehuevi Plutonic suite is composed of five distinct granitic lithologies covering the southwest to north central portion of the Chemehuevi Mountains. The suite is metaluminous to peraluminous and exhibits crude normal, vertical, and temporal zonation (John, 1988). The zonation features a decrease in age and an increase in silica content toward the center, away from the roof and walls (John, 1988). In decreasing age the older outer portion is made up of hornblende and sphene bearing granodiorite (Kga ).
Two bands of biotite-bearing granodiorite (Kg) trend southwest to northeast on the northern and southern portions of the suite. Porphyritic biotite granodiorite (Kpg) is the most extensively exposed in the range making a concentric pattern around the youngest granites. The youngest, most evolved members of this suite are (mostly) undeformed muscovite-bearing, locally garnetiferous granite and granodiorite (Kgg) (John, 1988).
18 Proterozoic Gneiss
The northern portion of the Chemehuevi Mountains is composed primarily of
Proterozoic layered gneisses and migmatites cut by a dense swarm of Cretaceous(?) dikes of variable composition (Figure 2). The gneisses form a coherent gently SW dipping
(15˚) unit in the northeastern part of the range and a steeply dipping (60-90˚) unit in the northern portion of the range. The gneissic rock assemblages are composed of strongly foliated, variably mylonitized, layered orthogneisses and paragneisses. The mineralogy of these gneisses is upper greenschist- to lower amphibolite-facies and feature elongated granitic leucosome pods. The gneisses are cut by high-angle (~60˚) shear bands containing hydrothermally-derived greenschist facies mineralization. The Mohave Wash fault cuts the Cretaceous granites (Kg) in the southwest and central portions of the range and cuts the Precambrian gneisses in the northeast portion of the Chemehuevi Mountains.
Mohave Wash Fault Footwall Geology
The footwall to the Mohave Wash and Chemehuevi Fault system is dominated by isotropic granitic rocks of Cretaceous age, originally emplaced at 12-20 km depth based on fluid inclusion studies (John, 1988), and Proterozoic layered gneisses and migmatites
(John, 1988; John and Mukasa, 1990; John and Foster, 1993). The Chemehuevi
Mountains pluton intrudes the layered gneiss and underlies most of the southern and central parts of the range. The youngest intrusions in the range make up a dense swarm of
Miocene mafic to silicic dikes.
The Chemehuevi Mountains core complex lies near the center of the CREC in the region of maximum crustal stretching. As much as 20 km of slip has been estimated
19 along the CDF-MWF system (John, 1987a). The majority of this extension was accommodated along the CDF (18 km) while, the MWF is believed to have accommodated only 1-2 km of extension during the early stage of formation of the extensional complex. This off set is based on geologic mapping of John (1987) showing a
1-2 km offset of the different lithologies in the granites between the footwall and hanging-wall of the MWF. The MWF therefore represents structures associated with the earliest stages of faulting. Extension transported Cenozoic and Proterozoic rocks to the surface from estimated paleodepths of at least 10-15 km, juxtaposing them against much
Tertiary volcanic rocks (John and Foster, 1993).
In contrast with most other detachment systems, including that of the nearby
Whipple Mountains, the footwall of the Chemehuevi Fault system does not include thick, impermeable mylonitic rocks formed during Miocene extension, although field investigations in March, 2013 revealed numerous, thin (up to 10 cm thick) mylonitic shear zones, and felsic dikes with strong mylonitic fabrics in the northeastern part of the field area, which were not known previously. Thus, although the fault system is interpreted to have initiated near the brittle-ductile transition zone (John and Cheadle,
2010), and to have slipped almost entirely in the brittle (seismogenic) regime, there is evidence for localized deformation by plastic deformation processes
Chemehuevi Dike Swarm
The Cretaceous Chemehuevi plutonic suite is intruded by Tertiary dike swarms above, below and into the Chemehuevi detachment fault in western, southern, and central parts of the Chemehuevi Mountains. The dikes range in composition from basalt and
20 microdiorite to rhyolitic and are of several generations (John and Foster, 1993). The exact timing and duration of the dike intrusion in the Chemehuevi Mountains is poorly constrained, but a K-Ar hornblende age from a NW-striking dacite dike of 20.7 + 1.3 Ma has been reported (John and Foster, 1993). This implies that at least some intrusions occurred during regional extension. Field observations during the 2013 and 2014 seasons revealed strong internal lineations oriented parallel to the established extension direction, mylonitic fabrics, and chloritized shear zones at the margin of several dikes, primarily in the northeastern part of the field area. The deformation fabrics indicate that the dikes were deformed plastically, likely before or during Miocene extension. The youngest dikes dated in the area include an undeformed basalt plug at 11.1 Ma (John, 1986, 1987a,
John and Foster 1993); therefore, the dike intrusions spanned from at least ~21 to 11 Ma.
Existing Thermal Structure of the Footwall
John and Foster (1993) constrained the initial dip of the Chemehuevi detachment fault by comparing estimated regional thermal gradients (Foster et al., 1991) with thermochronologic data collected on minerals sampled along the slip surface of the fault.
Thermochronometric dating utilizes the principle that some isotopic systems record the timing that a rock cooled their closure temperature, which corresponds to the time when radiometric daughter products stop diffusing out of the crystal lattice. John and Foster
(1993) used 40Ar/39Ar dating versus sampling distance along the fault to estimate the slip rate to be 7.7± 1.8 mm/yr between 19 and 15 Ma, with a slower rate before 20 Ma. This rate is consistent with the nearby Harcuvar and Buckskin-Rawhide Mountains (7-9 mm/yr) (Spencer and Reynolds, 1991; Foster et al., 1993) and the Whipple Mountains
21 (~8 mm/yr) (Davis, 1988). Plausible paleodepth estimates at ~23 Ma were estimated from 40Ar/39Ar and fission track cooling ages and a range of geothermal gradients from
20-50 °C/km (John and Foster, 1993). Howard and Foster (1990) used 40Ar/39Ar fission track results to estimate a paleothermal gradient of between 30 and 50°C/km for the
Mohave Mountains that lie directly east, and structurally above, the Chemehuevi detachment fault (Howard et al., 1982b; Howard and John, 1987; John, 1987a). Hurlow et al. (1991) estimated a 45 °C/km gradient for the Ruby Mountains and East Humboldt
Range in northeastern Nevada. Modern heat flow data from extensional terranes suggest geothermal gradients in excess of 25 °C/km, with higher gradients in areas of active, rapid extension (Lachenbruch and Sass, 1978). Considering these estimated thermal gradients, the initial dip on the Chemehuevi fault would have been between 12° and 32°, yielding depth estimates of 5 to 15 km for the southwest to northeastern most exposed portions of the footwall at the time of initiation. John and Foster (1993) suggest that the geothermal gradient during the early stages of extension was above 30 °C/km, constraining the initial dip to <20°.
Constraints on Hydrothermal Fluid Circulation along LANFs
Numerous previous workers have evaluated the role of fluid flow along LANFs using stable isotope geochemistry, though none have examined the well-mapped
Chemehuevi Mountains and few have been able to focus on early-slip microstructures that have not been overprinted by subsequent, down-temperature faulting and continued fluid flow. The sources of fluids that may permeate faulted crustal sections are reviewed here (Table 1), and include magmatic, metamorphic, meteoric, and evolved-meteoric
22 (brine) fluids. The constraints discussed below are based on representative studies summarized in Table 2.
The δ18O of water derived from magmatic systems is expected to be 6 to 10‰ depending on the parent material of the magma (Hoefs, 2009). The lower limit is defined by water in equilibrium with mantle-derived melts; higher values are achieved through the assimilation/contamination of supracrustal material that had previously been hydrothermally altered at low temperatures (Hoefs 2009). Water-rich fluids might also be derived from dehydration reactions of the country rocks, such as the mica-rich
Precambrian biotite gneisses found in the northeast portion of the Chemehuevi Mountains today. Kohn et al. (1997) studied high grade metamorphosed muscovite and biotite from the Bethlehem Gneiss, located in New Hampshire, USA, and found that the expected
18 δ OH2O values in equilibrium with metamorphic mineral assemblages formed at P ≥ 4 kbar at T ≤ ~650˚C would be from 5 to >7‰ (e.g., Kohn et al., 1997). The specific δ18O for water derived through dehydration will depend on the temperature at which dehydration occurs (>>650˚C; Graphchikov et al., 1999). Thus, magmatic and metamorphic fluids will have quite similar δ18O values that are difficult to distinguish from one another. .
Meteoric δ18O values vary significantly depending on latitude. For the latitude of the Chemehuevi and Whipple Mountains (34˚ 1186 – 34˚ 6467 N), the modern average regional meteoric precipitation is -5‰ (Bowen and Wilkinson, 2002). Fricke et al.,
(1992) determined the composition of pristine Tertiary meteoric waters to be -4.9‰ from a fluid inclusion in an unstrained quartz breccia in the near by Ruby Mountains (40˚
2820). The meteoric δ18O value can be considered the lower limit for fluids that could
23 interact during faulting. Evolved meteoric fluids may produce saline brines. Saline waters (brines) may form during evaporation of surface-waters without recharge or when surface waters interact with salt-rich rocks (Roddy et al., 1988). In the southwestern
United States such conditions may be promoted during the initial stages of metamorphic core complex-formation when tilting of hanging-wall blocks creates a series of parallel isolated basins that fill with water, which subsequently evaporates in this arid setting.
Brines develop higher δ18O values through evaporation (owing to preferential evaporation of 16O relative to 18O), and show a wide range in composition. Clayton
(1966) analyzed brines from the Michigan and Illinois Basins, and found δ18O ranges of -
9 to 3‰ and -8 to 5‰ respectively. Roddy et al., (1988) concluded that brines with δ18O values of -5 to +3‰ were responsible for hydrothermal alteration in the nearby Harcuvar
Mountains (AZ). Constraints from Roddy et al., (1988) are taken to be representative of brine waters that could have infiltrated the Chemehuevi system. In conclusion, fluids originating as magmatic/metamorphic, meteoric, and saline brine waters are readily distinguished on the basis of δ18O (Table 1).
Brief Overview of δ18O Studies on Selected Core Complexes
Table 2 presents a brief overview of relevant fluid δ18O studies of metamorphic core complexes exposed in the Whipple Mountains, Harcuvar Mountains, Ruby-East
Humboldt Mountains, and the Valhalla Mountains, British Columbia. The Whipple
Mountain metamorphic core complex lies 30 km south of the Chemehuevi core complex and was also formed during Miocene extension and is the most proximal to the present study area. Both areas feature regionally extensive low-angle detachment faults, high-
24 angle normal faults in the hanging-wall that are truncated by, or merge downward into the detachment fault, and a domal core of uplifted crystalline rocks in the central complex
(Morrison, 1994). The major difference between the two is the presence of extensive mylonites in the Whipple Mountains that are believed (although not confirmed) to be
Miocene. Oxygen isotope studies of quartz-epidote pairs associated with detachment faulting in the Whipple Metamorphic core complex (Morrison and Anderson, 1998;
Anderson, 1994) show a low- δ18O shift along the fault zone that indicates surface- derived fluids circulated to a paleodepth of up to 8 km during extension (Morrison,
1994). Morrison and Anderson (1998) report oxygen isotope values from coexisting quartz and epidote sampled along a transect in the footwall of the fault. The δ18O of quartz ranges from 7.9‰ to 10‰ and the δ18O of epidote ranges from 2.2‰ to 5.5‰.
Calculated δ18O values for water in equilibrium with quartz-epidote pairs ranges from
2.4‰ (12 m below the fault) to 6.5‰ (50 m below the fault). Since magmatic values should have δ18O values ~6-10 per mil, the fluids involved are unlikely to be igneous in
18 origin. Instead, the calculated δ OH2O was attributed to the influx of surface-derived, low-δ18O fluids to depth soon after fault initiation (Morrison and Anderson, 1998).
Apparent oxygen isotope equilibrium temperatures were calculated to be from 320˚C to
519˚C (using the fractionation factor of Matthews (1994) for pure epidote and quartz), and to decrease closer to the fault. The extreme thermal gradient implied of 82˚C over 38 meters (2160˚C per kilometer) was taken as evidence for extreme cooling along the fault, which would facilitate brittle fracturing and strain localization.
The Bullard detachment fault is a regional low-angle normal fault located within the Harcuvar Mountains of west-central Arizona, and is located 155 km southeast of the
25 Chemehuevi Mountains. The Bullard detachment fault separates footwall mylonites and chlorite breccia from upper-plate volcanic and sedimentary rocks. Hydrothermal fluids circulating along this fault have led to aerially extensive K-metasomatism and converted upper-plate mafic flows and felsic ash-flow tuffs into rocks with simple K- feldspar-hematite-quartz mineralogy. Cu-Au-Ag mineralization is concentrated in the damage zone of the Bullard detachment fault, and serves as further evidence for migration of hydrothermal fluids. Fluid inclusion studies of the fault show that the fluids were saline, and had minimum temperatures of 290 – 330˚C along the detachment fault and 240 – 290˚C in mafic flows within the hanging-wall. The δ18O of the mineralizing fluids were +3‰ for high-temperature quartz-sulfide veins, and -5‰ for lower- temperature barite-calcite-Mn-oxide veins. The high salinity of these fluids, along with their oxygen isotope values, indicates that they were basinal brines (Roddy et al., 1988).
Roddy et al. (1988) used oxygen isotope, fluid inclusion, and geochemical data to show that multiple fluid regimes were active during the lifecycle of the detachment fault in the
Harcuvar Mountains. Basin-brine-dominated mineralization has overprinted older lower- plate mylonitic rocks and breccias that were previously thought to be equilibrated with igneous or metamorphic fluids (δ18O = 6-10‰).
The Ruby-East Humboldt metamorphic core complex lies in northeast Nevada,
650 km north-northwest of the Chemehuevi Mountains. The Ruby-East Humboldt metamorphic core complex was exhumed during the Tertiary along low-angle a normal fault system. Like the Whipple Mountains, the Ruby-East Humbolt Mountains feature extensive mylonites in the lower-plate. Low whole rock oxygen isotope values were measured on quartzites and pelites, and were interpreted as evidence that meteoric waters
26 were not restricted to brittle deformation zones in the upper plate rocks, but may have penetrated (50-100 m) into the lower-plate footwall rocks that were still undergoing plastic deformation (Wickham et. al., 1993). However, it is also possible that the low values represent a later, unrecognized, low-temperature hydrothermal overprint. Whole rock δ18O values down to +2‰ were measured in lower-plate quartzites, much lower
δ18O values than those of quartzite from deeper structural levels (2‰ at the detachment fault, 13‰ 1,300 m below the detachment fault), interpreted as a consequence of interaction with meteoric waters. The measured 18O-depletion is restricted to the upper
100m of the lower plate. Hydrogen isotopes were also measured, and provide evidence that surface-derived fluids actually penetrated several kilometers into the lower-plate but were reequilibrated with formation waters within 50 – 100m of the damage zone
(Wickham et. al., 1993).
The Valhalla metamorphic core complex lies within the Valhalla Mountains,
British Columbia, 1,700 km north of the Chemehuevi Mountains. Holk and Taylor (2007) measured whole rock, feldspar, calcite, sphene, quartz, K-feldspar, amphibole, and epidote δ18O values within the Valhalla metamorphic complex to determine the presence of contrasting hydrothermal regimes. According to their study the Valhalla metamorphic core complex preserves evidence for three distinct stages of hydrothermal activity from
δ18O data: (1) fluids expelled from the lower plate (~10‰) moving upward along the damage zone and in to the upper plate (2) deep fluids mixed with surface-derived waters as brittle deformation became the dominate tectonic process. (3) Meteoric-hydrothermal activity associated with the emplacement of the Coryell plutons where evidence of long-
27 lived hydrothermal activity in the late stages of the development of the Valhalla metamorphic core complex (Holk and Taylor, 2007).
28 CHAPTER 3: METHODS
δ18O System
The reason δ18O is used to investigate water-rock interactions is because partitioning of 18O is (1) temperature dependent and (2) the abundance of oxygen is approximately twice as high in water than it is in rocks, giving leverage to change δ18O of rock derived from hydrothermal alteration. As discussed in the previous section (3) natural waters carry a range of diagnostic δ18O values. Water-rock interactions can occur through diffusive exchange of oxygen at grain boundaries, or through reactions that result in the formation of new minerals through metamorphism or hydrothermal precipitation.
Oxygen is the most abundant element in the Earth’s crust, playing a major role in
!" !" almost all geochemical reactions. There are three stable isotopes of oxygen: !O, !O,
!" !O, with natural abundances 99.756%, .039%, and .205%, respectively (Faure and
Mensing, 2009). For natural systems involving water and water-rock interactions 16O and
18O are compared to assess the source(s) of water. The difference between the oxygen isotopes is represented in a normalized fraction called delta (δ), and is represented by:
! δ18O = !"# − 1 � 10!‰ (1) !!"#$
Where R= 18O/16O
SMOW = standard mean ocean water
A positive δ18O means that the sample has a higher 18O/16O than the standard and a negative δ18O means that sample has a lower 18O/16O than the standard.
29
Fractionation
Isotopes of light elements can be fractionated during reactions, phase changes, and biological processes. Heavy Isotopes have stronger bond energy because of the larger mass of the nucleus - this means as a reaction occurs the lighter isotope will react first, enriching the products in the lighter isotope. The product of this variation in reaction energy is an unequal distribution of isotopes. The inequality is called isotope fractionation and can be quantified by the fractionation factor (α, defined below). There are two type of isotope fractionation: equilibrium isotope partitioning and kinetic fractionation (Faure and Mensing, 2009).
Equilibrium fractionation involves the separation of multiple substances in chemical equilibrium by difference in vibrational energy of the atoms. The heavier isotope will vibrate at a lower frequency than the lighter isotope causing the lighter isotope to have a greater potential energy for a reaction to occur. In phase changes the heavy isotope will occupy the phase with the lower energy (e.g. 18O will be concentrated in a solid over a liquid given that the two phases are in equilibrium) because of the relationship.
Kinetic fractionation is a mass dependent process that does not require equilibrium. A common example of kinetic fractionation is the chemical process of photosynthesis where the reaction is sensitive to atomic mass of the reactants. In mass fractionation the isotopes are physically separated based on differential rates of bond breaking. Lighter isotopes tend to move more quickly though a medium and react at faster rate than heavier isotopes. Phase changes, diffusion and biological processes
30 involve mass fractionation. During the phase change of water to vapor the different masses of the isotopic molecules of water have different ranges of vapor pressures. This range in vapor pressure causes the lighter molecule (16O) to preferentially evaporate
(Faure and Mensing, 2009).
Water-Rock Interactions and Oxygen Isotope Thermometry
Oxygen isotope thermometry is based on the premise that the magnitude of the partitioning of light stable isotopes between two phases is temperature-dependent.
Temperature estimations can be made by measuring δ18O on two or more isotopically-
18 equilibrated phases of interest and comparing the difference (defined as Δ Omin1-min2) to experimentally-determined values. The temperature calculation is based on temperature- dependent fractionation where the difference in δ18O values between two minerals decreases with increasing temperature at a constant rate. The difference is the fractionation factor (α) and is defined by the relation:
! !! �! = (2) !!
Where (a) and (b) represent the two phases of interest. The experimentally-constrained relationship between temperature and isotope fractionation between two phases is defined as:
! 10!���! = ! ! !" + � (3) ! !!
Where T is temperature in Kelvin and A and B are experimentally-defined constants
(Faure and Mensing, 2009). Quartz and epidote were chosen for this study because of their slow diffusion rate of oxygen (Bird and Helgeson, 1980), and thus likely retention
31 of initial isotopic compositions upon formation. Therefore samples of deformation structures related to Miocene extension are interpreted to preserve Miocene isotopic signatures.
The fractionation of 18O between epidote and quartz is also dependent on the composition of the epidote. Epidote is actually a mineral group, represented by a solid solution between end-members. The pistachite (Ps) component of epidote is the Fe-rich
+3 end-member, having the composition Ca2Fe3 [Si2O7][SiO4]O(OH). The proportion of the Ps component is represented as Ps =[100 mol Fe3+/( mol Fe3+ + mol Al)] (Deer et. al.,
1997). For oxygen isotope fractionation between quartz and epidote with Ps0, a value of
A=2.00 is defined by Matthews (1994). For more intermediate composition of Ps30, a value of A=2.225 has been defined by Matthews (1994). These values were determined from high-pressure mineral-water exchange experiments and constrain the end members in Fe variation for hydrothermal epidote, therefore, constraining the error range of calculated temperatures for epidote whose composition is not otherwise known. For this study, Ps30 is considered an appropriate epidote composition, based on mineral properties such as light green color in hand sample, and optically negative character in thin section. Thus, all temperature calculations were made using the experimental constant A=2.225 and B = 0. This is also consistent with the epidote composition used by Morrison and Anderson (1998) for samples from the nearby Whipple Mountains, and used for oxygen isotope thermometry calculations.
32 Formation of Hydrothermal Epidote
In the Chemehuevi Mountains epidote is abundant in hydrothermal veins and deformation structures and is rare and dispersed as only a very minor primary mineral in the plutonic suite (John, 1988). Coupled with its retention of formation δ18O at temperatures below 500°C (Ferreira et al., 2003), epidote is a key mineral for assessing the involvement of fluids during deformation in the Chemehuevi Mountains.
The disilicate epidote group consists of four minerals: Zoisite/clinozoisite
3+ (Ca2AL2O.AlOh[Si2O7][SiO4]), Epidote (Ca2AL2O.(Al,Fe )OH[Si2O7][SiO4]),
3+ 3+ Piemontite (Ca2(Mn ,Fe ,Al)3O.OH[Si2O7][SiO4]), and Allanite
2+ 3+ 3+ (Ca,Mn,Ce,La,Y,Th)2(Fe ,Fe ,Ti)(Al,Fe )2O.OH[Si2O7][SiO4]). For this study only zoisite/clinozoisite and epidote are present based on initial SEM analysis. Epidote present in the Chemehuevi Mountains is a product of metasomatic alteration where it occurs in quartz – epidote veins and monomineralic epidote veins. The presence of hydrothermal vein occurrence of epidote (zoisite) is a result of the alteration of feldspars in granitic rocks and can form by either of the following idealized reactions:
Anorthite (component of plagioclase) + orthoclase (component of K-feldspar) (4) + H2O = muscovite + zoisite + quartz
Anorthite + Orthoclase + H2O + CO2 = muscovite + calcite + quartz (5)
Experimental work by Bird and Helgeson (1980) showed at higher temperatures (up to
650˚C) epidote forms for the following equation:
Anorthite + Garnet + Hematite + Quartz + H2O 2Epidote (6)
33 Where in this equation epidote is formed from the reaction of water and breakdown of garnet, anorthite, quartz, and hematite. but is rare in the Chemehuevi Mountains (Deer et. al., 1997)
Analytical Techniques
To assess the role of hydrothermal fluids during deformation along the Mohave
Wash system, petrographic description and oxygen isotope analysis were carried out.
Petrographic observation is essential to determine the presence of equilibrium textures and compositions of minerals. Imaging allows for the determination of mineralogy and the degree of deformation within samples that cannot be seen on the outcrop scale.
Mineral pairs in isotopic equilibrium can be used to determine temperature of formation and provide information on the δ18O of fluids they were in equilibrium with. In an effort to sample equilibrated mineral pairs, minerals were handpicked from small volumes
(<0.5cm3) of crushed rock or thin (~2mm) slices of rock for application of isotope- exchange thermometry along with duplicate analysis for each pair (following methods of
Morrison & Anderson 1998). The quartz separates were then purified using 49% hydrofluoric acid (HF) rinse treatment. The HF reacts with any plagioclase in the separate and creates a powdery film over the grain making it readily distinguishable from quartz.
A mass of 2-3 milligrams of material was then picked for analysis, weighed and placed in a sample holder along with the standard, UWG-2. Separates were then pretreated overnight in BrF5 in preparation for analysis the following day. For each replicate analyses of quartz and epidote pairs were analyzed to assess the homogeneity within the crushed volume.
34 Laser fluorination is a technique that liberates oxygen from mineral separates by heating a sample with a laser, reacting the released gas with carbon to form CO2 and measuring the different isotopes released with a mass-spectrometer. Quantitative extraction of oxygen from silicates and oxides is accomplished best by reaction with strongly oxidizing reagents such as fluorides (BrF5, ClF3). The laser fluorination lab at the University of Wisconsin use BrF5 because of it is an excellent reagent and physical properties that are easy to handle (e.g. 438 hPa (mbar) at 20˚ C and freezing point at -
61˚C) (Kusakabe and Nakamura, 2004). Once the sample has been heated and reacted with the oxidizing agent it is then reacted with a carbon rod, converted to CO2, and ran though a mass-spectrometer to yield a ratio of the two isotopes.
Oxygen isotope ratios of minerals and whole rocks used in this study were analyzed at the University of Wisconsin by laser fluorination during three consecutive days. During analysis conducted in fall of 2013, oxygen was liberated by heating from a
10 µm CO2 laser reacted in the BrF5, then reacted with heated graphite, and converted in
18 to CO2 gas (Valley, 1995). Values of δ O were determined in a Finnigan MAT 251 mass-spectrometer. On each day UWG-2 garnet standards were analyzed at the beginning, middle, and end of each day. UWG-2 or UW-Gore Mountain Garnet #2 was prepared in 1994 from a single 2 kg block of garnet from the center of a single crystal.
Valley (1995) reported 1081 analyses of UWG-2 from laser fluorination with an average value of 5.74‰ ±0.005‰. Small amount of systematic drift was reported from the long- term average with precision on most days better than ±0.10‰. This drift is accounted for by adding or subtracting the difference of the measured-accepted value of 5.80‰ SMOW based on Valley et al. (1995) yielding corrected δ18O results. Daily precision ranged
35 from5.86 ± 0.07‰ (1ST. DEV) to 5.93 ± 0.12‰ (1ST. DEV). Whole rock samples were analyzed separately from quartz and epidote mineral separates. Whole rock samples were powdered, and then placed in an air-lock container, treated with BrF5, and analyzed immediately to avoid reaction and loss of highly-reactive phases present.
36 CHAPTER 4: RESULTS
Terminology
In this paper the term cataclasite will be used to describe rocks as a “..cohesive fault rock composed of broken, crushed, or rolled grains. Unlike breccia, it is a solid rock that does not disintegrate when struck with a hammer.” (Van der Pluijm and Marshak,
2004). The term ultracataclasite follows the same definition but is defined as 90-100% matrix. A mylonite will be used to describe rocks in which “…the matrix minerals undergo extensive dynamic recrystallization and strong patterns of preferred crystallographic orientation typically develop.” (Lister and Snoke, 1984). In this definition, mylonites are not considered as cataclastic rocks. While cataclasite processes may be involved during mylonitization, crystal-plastic deformation is the dominant mechanism by which strain is accommodated (Fricke et al., 1992) (Figure 6, Figure 7).
Field Relations
Field investigations reveal that the MWF is defined by a variable-thickness damage zone (~ 60 meter section shown in Figure 4) that ranges from 1 – 250 m across in the field area, and is defined by cohesive cataclasites, semi-brittle fault rocks, and thin semi-brittle to plastic shear bands (Figure 5). In one outcrop at the Pool Flats site, a pseudotachylite (Figure 5.2F) up to a few mm thick was identified, implying that the fault was seismically active during at least part of its slip history. From the field observations, ten lithologies were identified for the purpose of stable isotope characterization: Host granitoids (Kg), Precambrian gneiss, monomineralic quartz veins (possibly all plastically deformed), pseudotachylite, shear bands (foliated and unfoliated), quartz-epidote veins,
37 and monomineralic epidote veins. The exposed fault rocks with the shallowest paleodepths (Range Front, Figure 2), based on John and Foster’s (1993) thermochronologic constraints, are in a 220 m zone of coherent and incoherent cataclasites in Cretaceous granitoid. These cataclasites feature grain-size reduction and often greenschist facies mineralization (e.g. chlorite and epidote) (Figure 5).
Structural logs across the damage zone were documented for the Studio Springs section (Figure 4), a 220m section in which fracture density, orientation, mineralization present, and dike abundance were recorded. Fracture density ranged from 11 fractures per meter (section 0-5m, below damage zone) to 280 (section 185-195, upper end of damage zone) with lineation trends ranging from 20° NE to 85°E. Mineralization was sparse, consisting of mm-cm thick veins of epidote, calcite, and hematite. The dikes present in the damage zone were intensely fractured with areas of offset up to 2m along mm-thick shear bands. Moving down paleodepth, coherent cataclasites and (rare) ultracataclasites become more abundant as localized planar bands.
In the deepest regions of the fault zone, Miocene deformation was recorded in mylonites cutting the fabric of Precambrian gneisses. Siliceous dikes intruding the fault zone show a strong foliation and lineation orientation of approximately 050˚, consistent with the documented Miocene slip direction. Many dike margins featured reworked gneissic rocks, along with chlorite schist in the more mafic dike compositions.
Monomineralic quartz veins are present in the siliceous dikes, and often cut the foliation
(Figure 5.4B). High-angle (~60˚) shear bands that cut the Precambrian gneiss foliation often include quartz and epidote mineralization. In several of the high-angle shear bands the Precambrian fabric was deflected parallel to the shear band.
38
Petrographic Observations From Damage Zones and Hydrothermal Products
Monomineralic Quartz Veins
Monomineralic quartz veins within red cataclasites, felsic dikes, Cretaceous granite, and leucosomes associated with the Precambrian gneiss were collected throughout the
Chemehuevi Mountains, and include samples CH-35, CH-36, CH-37, CH-38, CH-41,
CH-44, CH-45, CH-46, CH-47, CH-51, CH-79. Quartz veins collected within mylonitic
Miocene dikes (CH-38, CH-41, CH-42, CH-46) appeared to cut the foliation of the dikes in some cases (CH-41, CH-42), but internally exhibit foliation parallel to the dikes
(Figure 8.5). The foliation within the quartz veins is defined by elongate, flattened grains with undulose extinction and sutured grain boundaries (Figure 8.4). In summary, all of the veins for which thin sections were examined show textural evidence of plastic deformation (Figure 8.5). No thin sections were made for samples CH-44 and CH-45.
Foliated Shear Bands
Foliated Shear bands, defined on the basis of macroscopic fabric, are 1-10 cm thick, and often featured outcrop structural indicators of offset. The foliated shear bands were collected across the Chemehuevi Mountains and were sampled within both granitoids and Precambrian gneiss. Within the granitic host rocks, the shear band mineralogy is primarily quartz + epidote; plagioclase in CH-42 features undulose extinction and elongated grains in areas. Within the gneiss, in which primary mineralogy includes biotite, hornblende, plagioclase, and quartz, the shear band mineralogy was
39 more diverse. Epidote and quartz mineralization occurs in bands cutting the fabric of the gneiss and feature sharp mineralogical and grain (e.g. shape, size and deformation intensity) boundaries between the shear bands and the host rock (Figure 8.4). All foliated shear bands feature evidence for plastic deformation of quartz, and in some cases plagioclase (CH-11; Pool Flats field site).
Injection Vein - Pseudotachylite
Sample CH-19 collected from the Pool Flats field site, featured what is interpreted as pseudotachylite. Macroscopically, the sample features a rounded injection plume into a plastically deformed quartz lens (Figure 8.2). The injection plume features quartz clasts
(from the host granite) and layering with the finest material parallel to the edges of the plume and repeated dark-light layering toward the center. The grains of the granite country rock are >1 cm long while the majority of the grains in the pseudotachylite are
<1 mm. The matrix appears glassy in thin section, supporting the interpretation that this feature is a solidified frictional melt.
Shear Band – Green Cataclasites
The most frequently observed deformation structure associated with the MWF were green cataclasites, found throughout the granitic and gneissic footwall rocks. These shear bands range from 1mm up to 10 cm thick, and were observed to accommodate up to 10 cm of offset. They commonly feature epidote+quartz mineralization, and occasional crack-seal veins of calcite (e.g., CH-4, CH-RF) and chlorite (CH-69). Epidote occurs as 1 mm fractured euhedral crystals (e.g., CH-24, Figure 8.4) and 15 µm grains in a
40 microcrystalline matrix (e.g., CH-RF; Figure 8.2). The largest grains of quartz feature dense cracks. The matrix epidote grains in the green cataclasites were ~100 times smaller than the grains in the wall rocks (Kg and Xgn). Figure 8.1.B shows a typical green cataclasite zone with an epidote-dominated matrix featuring large sub-rounded clasts of quartz and rock fragments.
Shear Band – Black Cataclasite
Black cataclasites were significantly less common. These were defined on the basis of having a black Fe-rich matrix between sub-angular clasts of host granitoids, with the clasts ranging from 0.5 to 3 mm. CH-30 featured small (~1mm wide) veins of epidote cutting the cataclasite and evidence of grain size reduction, with grains up to 100x smaller than the wall rock near the edges of the cataclasite.
Quartz + Epidote Veins
Hydrothermal veins containing primarily quartz and epidote were defined on the basis of being undeformed and having sharp boundaries cutting the host rock fabric. They characteristically exhibit blocky euhedral crystals. Examples include samples CH-22 and
CH-23, which cut granitoids and feature large (~3 mm) euhedral epidote grains in a 2 cm thick vein. The vein also includes quartz grains ranging from 0.25 mm to 0.5 mm; grain boundaries between quartz and epidote are sharp. Some of the quartz grains are elongated near the edge of the vein. Samples CH-28, CH-78, and CH-79 cut gneiss; no thin sections were made for these samples. On outcrop scale, these veins cut across the Precambrian gneiss banding at angles up to 70˚ in the Trampas Wash and Mohave Wash field sites.
41 The veins are offset by hematite- and calcite-filled fractures and cataclasites throughout the Chemehuevi Mountains.
Oxygen Isotopes
Results from laser fluorination analysis of oxygen isotope ratios in minerals
(quartz, epidote) and whole rocks are presented in Table 3, and summarized in the following section.
Oxygen Isotope Signature of Cretaceous Granitoids
To permit meaningful comparison and interpretation of deformed, hydrothermally-altered products, primary magmatic δ18O values of the footwall rocks were defined for samples that appeared free of deformation microstructures related to
Miocene extension. The δ18O of quartz from undeformed granite sampled at Range Front,
Studio Springs, and Mohave Wash sites define a narrow range of 9 to 10.3‰ (Table 3,
Figure 6). These constraints indicate generally uniform oxygen isotopic composition for undeformed rocks throughout the Chemehuevi Metamorphic core complex.
Oxygen Isotope Signature of Precambrian Gneiss
Quartz collected from a leucosome within the Precambrian gneiss yielded a value of 9.1‰, representing equilibrium values of quartz within the gneiss. One whole rock sample was analyzed from the Mohave Wash field site with a value of 5‰.
42 Oxygen Isotope Signature of the Chemehuevi Dike Swarm
Three samples from the Chemehuevi dike swarm were analyzed for whole rock
δ18O values. CH-5 a mafic dike in the Studio Springs area yielded the lowest value of
δ18O = 1.0‰. CH-37 and CH-47 were both collected in the Mohave Wash area and have values of δ18O = 7.5‰ and δ18O = 6.9‰ respectively. Quartz collected within the same dike as CH-37 yielded a value of δ18O = 10.4‰ .
Oxygen Isotope Signature of Deformation Structures
Quartz from monomineralic quartz veins showing microstructures indicative of plastic deformation give δ18O values from 5.5 to10.3‰; 4 of the 7 samples in this group fall below the δ18O range defined by quartz from of host granites (Figure 6). Within the foliated shear bands, the δ18O of the quartz ranged from 6.1 to 10.1‰, while the epidote ranged from 3.9 to 5.0‰ (Figure 6). Two analyses run of the pseudotachylite yielded a
18 δ OQtz replicate average of 9.8‰, i.e, no apparent shift away from the range of the host granites. The quartz analyzed within the green cataclasites ranged from 0.1 to 10.3‰.
The majority of the quartz values are below the host granite range with the exception of a
18 sample from the foliated shear band in CH-RF (Figure 8.1C) at δ OQtz = 10.3‰ and CH-
18 15 at δ OQtz = 9.3‰. Epidote values from all sample types ranged from -1.5 to 5.3‰.
18 These values are all below the magmatic δ Oep range derived from Matthews (1994).
18 One sample (CH-69) is within the magmatic δ Oep range based on the fractionation calculated by Ferreira et al., (2003) but below the magmatic range based on Matthews
(1994). Quartz within the black cataclasites ranged from 7.3 to 9.8‰. Quartz in quartz/epidote veins ranged from 6.5 (replicate average) to 9.2 (replicate average) ‰.
43 The epidote analyzed from green cataclasites yielded values from -0.1 to 5.6‰. A single sample (CH-24) of undeformed epidote was collected in the northern portion of the range cutting a leucosome within the gneiss. The epidote vein had an average value of
5.5‰.
Whole Rock Data
Separate airlock whole rock laser fluorination analysis was run on 23 samples.
These samples define a large range of values from -2.6 to 8.6‰ (Figure 7). Whole rock host granite range was constrained by two samples (CH-18 and CH-80) to be between 7.6 and 8.6‰, while deformation structures ranged from -2.5 to 8.1‰. One sample (JL-15) of the Precambrian gneiss was analyzed and had a value of 4.96 ‰. This value is significantly lower than whole rock values from granite, and can largely be attributed to the gneiss containing more low- δ18O minerals like biotite and amphiboles.
44 CHAPTER 5: DISCUSSION AND CONCLUSIONS
Characteristics of the Mohave Wash Fault
Field investigations reveal that the MWF is composed of cohesive cataclasites, semi-brittle fault rocks, and thin semi-brittle to plastic shear bands within a variable- thickness damage zone (Figure 4) that ranges from 1 – 250 m across in the field area.
Locally, rare veins of pseudotachylite (Figure 5.2e, Figure 8.2) up to a few mm thick were identified, implying that the fault was seismically active during part of its slip history. Mineralogical changes along the Mohave Wash fault zone are defined primarily by the formation of quartz, epidote and chlorite, accompanied by grain-size reduction of primary country rock mineralogy. The formation of hydrous mineral assemblages indicate that fluid-rock interactions were an integral, and perhaps significant influence during the earliest stages of deformation recorded during detachment faulting in the
Chemehuevi Mountains. Chlorite is a common greenschist facies mineral that forms from
<200˚C to 450˚C, at pressures up to 1.5 GPa, while epidote forms from 250˚C up to magmatic temperatures and pressures (Winter, 2010). The wide range in temperatures and pressures over which these mineral form means that their presence alone is not enough to constrain metamorphic conditions during faulting. Therefore, oxygen isotopes are critical in evaluating fluid and temperature constrains.
Constraints on Fluid-Rock Interactions from Oxygen Isotope Ratios
Laser fluorination oxygen isotope values for 76 samples of quartz, epidote, garnet, and whole rock collected in the in the Chemehuevi Mountains are listed in Table 3.
Measured δ18O for quartz within the Chemehuevi Mountains ranged from 1.1‰ to
45 10.3‰. Undeformed host granite quartz ranges from 8.9‰ to 10.3‰, while quartz samples from deformed structures show variably lower δ18O from 1.1‰ to 10.3‰.
Epidote varies from -1.5‰ to 5.6‰.
From the samples collected along the Mohave Wash fault, ten structural groupings were designated for the purpose of stable isotope characterization:
Undeformed host granite, leucosomes in Precambrian gneiss, quartz veins, pseudotachylites, hydrothermally-altered granite, shear bands (foliated, green cataclasites), quartz and epidote veins, and epidote veins. These sample types are plotted vs. their δ18O value in figure 6. The quartz samples collected within the undeformed
Chemehuevi plutonic suite have a narrow range of values (8.9-10.3‰) (Figure 6; 10).
Any quartz values below the grey bar are interpreted to represent values not in equilibrium with the Chemehuevi plutonic suite, and thus require interactions with an
18 externally-derived fluid. This interpretation is consistent with the δ Oquartz = 9.2‰ from undeformed granite (sample number: BJ81-CH5), which preserves near-magmatic
18 quartz-garnet fractionation of Δ O(qtz-gt) = 3.8‰ (both minerals exhibit extremely slow volume diffusion of oxygen at submagmatic temperatures). This would be equivalent to an equilibrium temperature of 600°C (using calibration of Bottinga and Javoy, 1975), which has been interpreted as the closure temperature for oxygen isotope diffusion in igneous quartz (Ferreira et al., 2003).
Whole rock (WR) analysis of samples from a structural transect was made at the
Pool Flats field area (Figure 9, Figure 2 marked measured section) to document fluid- rock interactions along a structurally-well constrained vertical transect across the Mohave
Wash Fault from the lower to upper plate. WR δ18O values of undeformed granite ranged
46 from 7.7 to 8.5‰, whereas deformation structures ranged from 4.5 to 8‰. No apparent trend in δ18O values across the damage zone is present, although there is a marked downward shift to lower δ18O along thin (1-50 cm thick) green cataclasite bands (e.g.,
Figure 2C-D). The downward shift relative to undeformed country rock is interpreted to
18 reflect oxygen isotope exchange with a low-δ O, externally-derived fluid. These observations are consistent with open exchange of fluids at, or immediately after, initiation of the Mohave Wash Fault.
Origin of Epidote: Magmatic Fluids or Externally-Derived Hydrothermal Fluids?
18 To provide a framework for interpreting the δ O values for epidote sampled from deformation microstructures, the isotopic composition of epidote in magmatic equilibrium with the Chemehuevi plutonic suite was determined. Using the calibration from Matthews (1994), fractionation of oxygen isotopes between quartz and epidote gives an expected δ18O range for epidote (at 600˚C) between 6.4 and 7.6‰. This range was calculated using A=2.225 for Ps30 in equilibrium with quartz from undeformed samples of the Chemehuevi plutonic suite determined to range from δ18O values of 9.0 and 10.3‰. Ferreira et al. (2003) established a different calibration for oxygen isotope fractionation between quartz-epidote (A = 3.1), based on empirical calibrations determined from natural isotopic variations between quartz and epidote in peraluminous granites from the Emas and Sao Rafael plutons in NE Brazil. The range from their calibration is 4.9 to 6.2‰ for magmatic epidote (Figure 6). However, this calibration is only applicable at temperatures ≥600°C (range of igneous crystallization for the granites sampled by Ferreira et al., 2003), and thus the calibration of Matthews (1994) is
47 preferred. Only epidote from samples CH-3, CH-69, CH-78, and CH-24 fall within the magmatic epidote range of Ferreira et al. (2003) and none fall within the range defined by the preferred calibration of Matthews (1994) (Figure 6). Therefore, the majority of epidote from this study is resolvably shifted to lower δ18O than would be expected for fluids in magmatic equilibrium with the undeformed country rocks. This observation is interpreted to indicate epidote formed during interaction with hydrothermal fluids derived from an external source, associated with early high-temperature slip along the Mohave
Wash Fault zone.
Isotopic Equilibrium and Oxygen Isotope Thermometry
Oxygen isotope thermometry can offer a powerful constraint on the conditions of fluid-rock interactions. However, before accurate temperature calculations can be made from measured mineral-mineral fractionations, several assumptions must be evaluated.
Three of the most basic assumptions that need to be met for accurate temperature calculations are: (1) the sample was isotopically equilibrated during a specific event and those compositions are preserved during later, down-temperature processes. Minerals are in textural equilibrium (sharp grain boundaries), and do not contain growth zonation or retrograde alteration. (2) Analysis is accurate at the appropriate scale. (3) Isotope fractionation is sufficiently sensitive to temperature and accurately calibrated, including the effects of solid solution, mixed fluids, and pressure (Valley, 2001). For assumption
(2), the laser fluorination techniques outlined above meet the criteria for accurate analysis on the single-grain scale. Assumption 3 is met by using the internally-consistent calibration for quartz-epidote oxygen isotope fractionation of Matthews (1994) calibrated
48 for temperatures between 0-1200°C, which accounts for solid solution in epidote and has been used previously in the literature for similar geologic settings (e.g., Morrison and
Anderson, 1998). The first assumption listed above is more difficult to constrain. To evaluate whether the first assumption listed can be reasonably made, petrographic observations, SEM observations, and analytical approaches were taken. These are discussed below.
On the thin section scale, textural equilibrium was assessed by looking at sharpness of grain boundaries, optical evidence of growth zoning, retrograde reactions, and partial recrystallization (e.g., Valley and Graham, 1996; Valley, 2001: Famin et al.,
2004; Bindeman, 2008). In samples from which apparent oxygen isotope temperatures are estimated, coexisting pairs of quartz and epidote were selected from discrete microstructural domains (Figure 8). Grain boundaries between epidote and quartz are characteristically sharp, indicating textural equilibrium. In samples CH-RF, CH-22, CH-
23, Ch-24 where euhedral epidote is noted (Figure 8.3), no optical evidence for zoning was observed that would indicate variable composition. Lack of zoning was also present in SEM backscatter images obtained from sample CH-RF (Figure 8.6). Analytically, replicate analysis of minerals of interest can be made to evaluate whether they are homogeneous. A necessary, but not sufficient criteria for proving isotopic equilibrium between coexisting minerals, is that replicate analyses of the same mineral from a small volume of rock must be homogeneous in δ18O – this is known as the modified outcrop test described by Valley (2001). Replicate analysis varied up to 0.7‰ (1 STD) for the quartz and up to 1.3‰ (1 STD) for the epidote, any value greater than the standard
(UWG-2) variation of up to 0.1‰ (1 STD) is considered heterogeneous. Epidote from
49 CH-28, CH-77, and JL-7 are heterogeneous, indicating that temperature estimates based on these samples are less accurate (see Table 3).
For chemical equilibrium to be achieved in natural samples, at least one of the following must have occurred: (1) both quartz and epidote were co-precipitated from the same fluid, and no later diffusion / alteration of δ18O occurred; (2) one of the minerals precipitated after the other, and the original mineral was reequilibrated during formation of the second either by deformation enhanced diffusion of thermal diffusion (the latter is considered unlikely at the relevant temperature range for these minerals); (3) a rim of reequilibration developed on two preexisting minerals by diffusion exchange with a new fluid (Famin et al., 2004). In cases (2) and (3) zoning may be apparent in CL and possibly in petrographic imaging, however it is difficult to control texture of quartz by CL, therefore these types of isotopic equilibrium were not used as equilibrium pairs in this study.
Temperature of Fluid-Rock Interactions Along the Mohave Wash Fault
18 Thirteen apparent temperatures were calculated from Δ Oquartz-epidote measured on coexisting mineral pairs (Table 3, Figure 12) from deformation microstructures interpreted to be associated with Miocene deformation. Calculated apparent temperatures ranged from 248˚C to 542˚C showing a general increase in the down-dip direction within the footwall of the MWF (Figure 13). The highest and lowest apparent temperatures observed came from the Mohave Wash field site, (Figure 2, Figure 9) and show a significant shift in temperature and δ18O of quartz and epidote from samples collected at and below the main damage zone relative to those collected above the main damage zone.
50 This shift implies a change in fluid source accompanying the decrease in temperature
(Figure 16). Calculated apparent footwall temperatures average 452˚C while the average temperature of the hanging-wall is 295˚C. This approximately 150˚C difference suggests infiltration of cold surface-derived fluids at shallower depths and within the hanging-wall of the MWF. The large temperature difference also indicates that there was not an exchange of fluids between the hanging-wall and footwall at the Mohave Wash site, which is inconsistent with the findings of Morrison and Anderson (1998) from the nearby
Whipple Mountains. Based on a pattern of decreasing oxygen isotope temperatures from
0 to 60 meters below the fault, they inferred significant footwall refrigeration resulting from the infiltration of surface-derived fluids. The temperatures reported in the Whipple
Mountains ranged form 321˚C (15 m below the fault) to 519˚C (40 m below the fault). In the footwall of the Mohave Wash fault the temperatures show the opposite relationship, decreasing from 511˚C (20 m below the main damage zone) to 464˚C (45 m below the main damage zone) (Figure 14)
Chemehuevi Mountain Isotopic Signature – The Story of Two Fluid Regimes
18 Once the temperature has been constrained, the δ OH2O in equilibrium with
18 mineral pairs can also be calculated using δ Omineral-H2O fractionations. The values are
18 then applicable to evaluating possible fluid sources. Calculated δ OH2O values from the
MWF range from -1.4‰ to 6.7‰, and are presented in Table 3 and Figure 16.
Using the quartz-H2O oxygen isotope fractionation of Clayton and O’Neil (1972), (see above), water in magmatic equilibrium with undeformed granite sample BJ81-CH5 would have a δ18O of ~7.7‰. This δ18O value serves as an upper bound for fluids
51 expected in this system, and is consistent with typical magmatic/metamorphic fluid compositions from other settings (e.g., Hoefs, 2009). Undeformed quartz and epidote within a late, crack-seal vein sampled from the Sacramento Detachment fault (equivalent to the structurally higher level Chemehuevi Detachment fault) from the nearby
Sacramento Mountains (Sac09SPK-15) was also analyzed. The quartz-epidote pair gave a
18 Δ O(qtz-ep) = 8‰, yielding a calculated equilibrium temperature of 254°C. At this temperature, these minerals would be in equilibrium with water having a much lower
18 18 δ O = -3.3. This value is fairly close to the expected δ OH2O of meteoric precipitation at
18 18 his latitude (δ OH2O = -5), and defines the low-δ O end-member for fluids that could
18 theoretically infiltrate the Mohave Wash fault zone. The lowest δ OH2O in the
Chemehuevi Mountains is calculated from a quartz-epidote vein (CH-28) collected in the
18 Mohave Wash fault hanging-wall, which yielded a calculated δ OH2O = -1.4‰ (Figure
15). This value is shifted to higher values relative to meteoric values of -5‰ (Bowen and
Wilkinson, 2002; Fricke et al., 1992). The intermediate δ18O of most fluids thought to be in equilibrium with the samples from this study requires either a shift upwards relative to meteoric water, or a shift downwards from magmatic waters prior to water-rock interactions that produces quartz-epidote mineralization. Upward shifts in δ18O can be achieved by brine formation, where 16O is preferentially evaporated, enriching the remaining waters in 18O and salts leading to higher δ18O values. Roddy et al. (1988) working in the Harcuvar Mountains concluded that saline brines with δ18O values ranging from -5 to +3‰ were the likely fluid source involved with detachment-related mineralization. It is most likely that the source of low-δ18O fluids is saline brines that were coeval with the development of the Chemehuevi Mountains metamorphic core
52 complex. Magmatic fluids, or fluids equilibrated with the country rocks, range from
>7‰ to 8.9‰ at temperatures of >600C, based on host rock quartz δ18O values. These values are determined using A=2.51, B=0.00 and C=-1.46 (equation 3), based on experimental values determined for the temperature interval of 500 - 750˚C by Clayton et al., (1972). All of the calculated values for water in equilibrium with fault rocks from this
18 study fall below magmatic values and 8 of the 13 samples have δ OH2O values above the maximum δ18O expected for brines. This observation is consistent with the mixing of fluids from at least two sources; those sources are likely to be magmatic and brines.
A third possible fluid source to consider involves the dehydration of hydrous minerals (i.e., biotite) in the Precambrian gneiss that forms the country rock in the northeast portion of the field area (Figure 2). Biotite reacts with quartz and breaks down to form orthoclase, enstatite, and water at temperatures of >650˚C (Graphchikov, 1999).
18 Considering quartz from Precambrian gneiss sample JL-15 (δ Oqtz = 9‰), water formed
18 by this reaction at 650˚C would have a δ OH2O of ~7.5‰, which is effectively equivalent
18 to δ OH2O values for magmatic fluids in equilibrium with the Chemehuevi plutonic suite country rocks. Therefore, dehydration of the Precambrian gneiss cannot explain the
18 downward shifts in δ OH20 calculated for the sample in this study, even in the unlikely event that small amounts of water derived through dehydration could be expelled, and travel up the fault.
Footwall Isotope Thermometry Variation from Mineral Closure Temperatures
The ambient thermal structure of the Chemehuevi Mountains at the time of initiation and uplift (23 Ma) was determined previously using thermochronometry on the
53 footwall rocks by John and Foster (1993). These data serve as a baseline for interpreting the direction of fluid (and heat) transport along the fault as recorded in deformed, hydrothermally-mineralized deformation microstructures.
As expected, calculated oxygen isotope temperatures increase with increasing paleodepth (i.e., towards the northeast) along the MWF, similar to the temperatures determined by thermochonometric methods. Perhaps unexpectedly, oxygen isotope temperatures from samples collected in the Mohave Wash damage zone are consistently elevated 50-200˚C relative to ambient footwall temperatures expected from mineral closure temperatures at 23 Ma (Figure 13). A paleoisotherm map based on John and
Foster (1993) results from 17 samples is shown along with temperatures from this study in Figure 14. Errors associated with the position of each paleoisotherm placement vary from 1-4 km, based on constraints from the sample distribution. Even considering these uncertainties, temperatures from oxygen isotope thermometry remain elevated relative to the ambient footwall temperature defined by Ar/Ar thermochronometry (Figure 13).
The interpretation of fluids with elevated temperatures relative to the ambient footwall moving along the fault zone is seemingly at odds with the model of footwall refrigeration reported based on oxygen isotope observations from the nearby Whipple
Mountains (Morrison and Anderson, 1998). There, rapid cooling recorded in the upper 50 meters of the lower plate rocks was interpreted to result from the migration of cold, surface-derived (low δ18O) fluids down through the upper plate along high-angle normal faults. In this study, the elevated temperatures observed along the fault zone indicate instead that relatively hot fluids migrated along the fault zone, and thus presumably were derived from deeper portions of the crust.. There are several possible mechanisms for
54 generating these higher-than-predicted temperatures, including: (1) heating of fluids by dikes intruding at the same structural level as the initiating Mohave Wash Fault (2) Flash heating along the damage zone from sliding resistance (Goldsby and Tullis. 2011; Toro et. al, 2004; Lachenbruch, 1980). (3) Seismic pumping of fluids from deeper, hotter crustal regions upward along the fault zone during rupture (Sibson, 1975; Sibson, 1990;
Sibson, 2000).
McCaig et al. (2007) documented the upward migration of high-temperature fluids (>400°C) along detachment faults at mid-ocean ridges, driven by thermal buoyancy. At the TransAtlantic Geotraverse (TAG) hydrothermal field along the slow- spreading Mid-Atlantic Ridge, hot fluids are envisioned to travel up to 6 km from melt- rich zones to the sea floor, where they erupt from vents. The fluids originate as seawater, which infiltrates normal faults within the hanging-wall along the spreading ridge, becomes heated, and then travels upward along the main detachment faults to exit as black smokers on the sea floor. For this fluid circulation to take place, fault-zone permeability is at least as great, and probably greater, than that of the hanging wall. In the Chemehuevi Mountains, two possible heat sources are envisioned: 1) advection of heat by water-rich fluids traveling up the fault from greater depths, and 2) heating of water-rich fluids by synextensional, Miocene dikes. Several examples of mylonitic dikes of Miocene age intruding the Mohave Wash Fault zone were observed in the northeast portion of the field area (Mohave Wash site), and would represent a substantial heat source promoting thermally-driven upward migration of fluids along the Mohave Wash fault once it had initiated. The Miocene dike swarm and Whale Mountain thermal anomaly both occurred around 23 Ma (John and Foster, 1993).
55 Hence, one reason for the different observation in this study, versus those from the Whipple Mountains, might be the presence of the Miocene dike swarm. The dike swarm is dense and spans the entire exposed extent of the Chemehuevi Mountains, providing additional heat during extension (Figure 2). A second reason for the difference between the Whipple and Chemehuevi Mountains could be that the Mohave Wash fault ceased faulting soon after initiation and was not well connected to the surface. The
Chemehuevi detachment fault and the Whipple Mountain detachment fault are thought to be structurally related (John and Foster, 1993), therefore, the Mohave Wash fault may not represent the same feature in a metamorphic core complex as the Whipple Mountain detachment fault.
Sibson (2000) described the stress-controlled structures affecting rock permeability within normal faults to be brittle faults, mircocracks, extensional and extensional- shear fractures, and styolitic solution seams. Different combinations of these structures may contribute to the bulk permeability of the damage zone. During seismic activity dilatancy/fluid-diffusion occurs in these structural features, providing a mechanism for upward fluid transport. At the onset of dilatance, the drop in fluid pressure causes a rise in frictional resistance to shear along the fault. As fluid fills the cracks that form, fluid pressure rises again and frictional resistance decreases (Sibson, 1975). This mechanism requires multiple hydrothermal injections and episodes of extensional opening, and evidence of widespread seismic activity. Some samples from the Mohave
Wash Fault show evidence of multiple distinct mineralization events (CH-25, CH-26, Ch-
60, CH-61, CH-64, CH-RF) indicating seimic pumping of fluids from depth may be a mechanism of hot fluid migration up the MWF damage zone. The inferred aseismic
56 nature of the fault based on the lack of widespread pseudotachylites indicates that flash heating was not likely the sole mechanism of the shifted temperatures. Additionally, this heating is likely to be extremely localized, and unlikely to promote regional scale, thermally-driven fluid flow. An alternative, and perhaps the most likely mechanism is that mildly low δ18O fluids, originating from near the surface as brines? infiltrated downwards into the upper plate, and were subsequently heated by Miocene dikes intruding the permeable fault zone. Fluids from greater depths (rock-equilibrated fluids, or magmatic fluids possibly even derived from intruding magmas as they solidify) could possibly mix with surface derived fluids at the depth where the detachment initiates. This would explain the high, but still below-magmatic δ18O calculated from water in equilibrium with the highest-temperature quartz-epidote mineral pairs.
Considering the constraints from structural, mineralogical, and stable isotope observations, a model for the inferred fluid pathways and evolution during the development of the Chemehuevi metamorphic core complex is shown in Figure 17, and discussed in depth in the following section.
Fluids Role in the Development of the Chemehuevi Metamorphic Core Complex
The role of fluids in the evolution of the Chemehuevi metamorphic core complex can be highlighted in six steps (Figure 17), and incorporated into existing models for core complex formation (e.g., Reynolds and Rehrig, 1980). (1) During the Cretaceous the
Chemehuevi plutonic suite intruded into Precambrian gneissic rocks in multiple pulses.
(2) The Miocene marked Basin and Range crustal relaxation and related extension of the
CREC (initiation ~23 Ma). Early deformation products, including mylonitized Miocene
57 dikes and crosscutting quartz veins, preserve δ18O values in equilibrium with magmatic fluids (7 to 9‰). Extension was initially accommodated by brittle deformation of the upper crust along the Mohave Wash fault and possibly along the CDF. The Chemehuevi dike swarm intruded into the plutonic suite during extension and provided heat to the region. (3) Differences in response to crustal extension between different rheologies created differential strain and fragmented the rigid top portion of the crust. This created dislocation surfaces and the crust split into discrete blocks. (4) As extension continued the blocks rotated, creating sediment-filled basins. Evaporation processes led to the formation of lakes and basinal brines (δ18O = -5 to 3‰) that migrated downward. At the same time, diking provided heat to drive fluids from depth (δ18O = 3 to 7‰) up the
Mohave Wash damage zone, where permeability is highest. Brittle deformation is recorded along the shallow and intermediate depth of the MWF and plastic deformation is present along the deeper portions of the fault. These fluids may have originated as brines shifted to higher δ18O during subsequent high temperature water-rock interaction, or mixing between magmatic fluids expelled from below and surface-derived fluids. (5)
Faulting along the MWF ceased. Upper-crustal blocks continued tilting while basinal brines continued to migrate downward, moving down high-angle normal faults to become heated and driven up the MWF damage zone by thermal buoyancy. Hot, high δ18O rock- dominated fluids may have continued to move up the damage zone in the deepest portions of the fault zone, but deformation microstructures record infiltration of progressively lower δ18O fluids (δ18O = -3 to 3‰) with decreasing temperature (Figure
15). Fluid circulation at this point was dominantly localized in the hanging-wall to the
MWF, and probably along the CDF. As tectonic denudation proceeded, the area of
58 maximum thinning responded to removal of zone 1 by isostatic uplift and deformational thinning in zone 2 (Rehrig and Reynolds, 1980). Isostatic uplift and zone 2 thinning continues until ~12 Ma when extension ceased.
Conclusions
Returning to the original stated questions: (1) what role, if any, did hydrothermal fluids play in early microstructure development? The presence of greenschist facies hydrothermal minerals in deformation features indicates that fluids were present during deformation. Of the 15 samples showing plastic deformation in the quartz, only 6 show a shift below the δ18O value of the host granitoids. This observation is not consistent with a large influence of a low- δ18O fluid at depth during initiation. Quartz within deformation features showing brittle deformation and interpreted to be associated with shallow and late structures show a stronger shift away from host granite quartz 9 of the 14 samples having δ18O values below the range of the Chemehuevi Plutonic suite quartz values. This observation is consistent with low- δ18O fluids interacting with the shallow portions of the fault zone and also later in the faulting history. (2) Is there evidence for surface- derived fluids during deformation, or were any fluids present magmatic/metamorphic in origin? As stated in question 1, there is evidence for surface-derived fluids in the damage zone of the Mohave Wash fault in all areas except the deepest exposed portions. The highest temperature fluids are not in equilibrium with the country rock, yet still have high
δ18O values. With decreasing δ18O values, temperature also decreases consistent with surface-derived fluids. (3) What is the temperature of mineralization, fluid flow, and deformation across the study area defined by stable isotope thermometry – is there
59 evidence for rapid cooling along the fault driven by convection of hydrothermal fluids as previously proposed by Morrison and Anderson (1998)? Calculated temperatures ranges from 273˚C to 542˚C in the footwall and are ~50-200˚C higher than the ambient temperatures calculated by John and Foster (1993). There is no evidence of cooling along the Mohave Wash fault and the temperatures may indicate the opposite relationship is true. In summary:
• δ18O quartz values from the Mohave Wash fault zone range from 1.1‰ to 10.3 ‰
• The shift to lower δ18O in the fault zone is consistent with infiltration of low- δ18O
fluid and associated mineralization
• δ18O shift from undeformed values in the host granites provides evidence of two
fluid regimes, a surface-derived component (basinal brines) and formation water.
• δ18O values and field observations indicate that fluids were present throughout the
development of the Chemehuevi metamorphic core complex
• Apparent stable isotope temperatures in the footwall of the Mohave Wash fault
increase with plaeodepth as expected, but are 50-200˚C higher than calculated
mineral closure temperatures at 23 Ma
• Low δ18O fluids from near the surface infiltrated the upper plate, were heated and
migrated up the damage zone, mixed with fluids from depth that were driven
upward by thermal buoyancy, consistent with the elevated calculated temperatures
60 TABLES
Table 1
Overview of fluid source and isotopic signature
18 H2O Source Expected δ O Range References Magmatic 6 to 10‰ Hoefs (2009) Gneiss dehydration 5 to 7‰ Kohn et. al., (1997) Meteoric -5‰ Bowen and Wilkinson (2002) Brines -5 to +3‰ Roddy et. al., (1988)
Table 2
Overview of δ18O studies on metamorphic core complexes
Location Isotopic Constraints Conclusion about fluids involvement Reference
18 18 Whipple δ OH2O = 2.4‰ just below Change in δ Oqtz-ep with distance from Morrison and Mountains the fault to 6.5‰ 50 m fault implied geothermal gradient of Anderson (1998) below the fault based on 2160˚C per kilometer. Cold, surface- measured δ18O between derived fluids interacted with the quartz and epidote inferred footwall at depths up to 8 km to be in equilibrium 18 Harcuvar δ OH2O = 3 to -5‰ along Upper-plate and damage zone Roddy et. al., Mountains the fault zone based on Overprinting of basinal brines δ18O (1988) 18 δ OWR values. signature during faulting. 18 Ruby-East δ OWR = quartzite values Infiltration of fluids from both the Wickham et. al., Humbolt near +2 ‰ within 100 m of surface of the earth (down to 15 km (1993) Mountains fault and up to +13 ‰ 1,000 paleodepth) and also deep crustal or m below the fault. The mantle levels during faulting. deepest portion of the fault exposed was in equilibrium with fluids having δ18O values of +6 per mil. 18 Valhalla δ Ofluid = 11.4 to 18.4 ‰ Three distinct stages of hydrothermal Holk and Taylor, Mountains measured in a (fluid activity: (1) fluids expelled from the (2007) inclusion) in vein siderite lower plate (~10‰) (2) deep fluids from the lower plate. mixed with surface-derived fluids as 18 δ OH2O = values ~-15 per brittle deformation became the mil were concentrated along dominant tectonic process. (3) Slocan Lake detachment meteoric-hydrothermal activity fault. associated with the emplacement of the Coryell plutons.
61
Table 3
Laser fluorination δ18O data for quartz, epidote, garnet and whole rocks from the Chemehuevi Mountains, SE CA
18 δ18O ‰, VSMOW Δ O qtz- T (˚C) Est. T error (˚C)b δ18O H O Distance (dip Sample # Structural Domaina Site 2 Quartz 1 s.d. Epidote 1 s.d. Garnet WR min calc - + calc direction) (km) Latitude Longitude BJ81-CH5 Undeformed Kg (Chemehuevi Plutonic Suite) 9.2 5.4 3.8 600 † - - 7.7 Sac09SPK-15 Shear Band - Green Cataclasite 5.9 -2.1 8.0 254 * - - -3.3 CG13- CH-1 Host Kg - undeformed (Chemehuevi Plutonic Suite) SS 9.0 34° 34.993 114° 32.746 CH-1 Shear band in Kg - foliated SS 6.1 34° 34.993 114° 32.746 CH-11 Shear band in Kg - foliated with plastically-deformed quartz clasts PF 9.7 5.0 34° 35.645 114° 34.003 CH-12 Undeformed aplite dike in Kg - 1 m below MWF PSZ PF 9.8 6.2 34° 35.645 114° 34.003 CH-13 Shear band - cohesive ultracat. 0-5 cm above lowest MWF PSZ PF 6.2 CH-14 Deformed Kg (pegmatite); 10-15 cm above CH-13 PF 9.5 8.1 34° 35.645 114° 34.003 CH-15 Host Kg - Kg quartz clasts in cohesive cataclasite PF 9.4 4.8 34° 35.645 114° 34.003 CH-15 Shear band in Kg - green cataclasite PF 9.3 34° 35.645 114° 34.003 CH-16 Shear band in Kg - green cataclasite PF 6.0 CH-16 Shear band - foliated 1 cm-thick layer at base of cataclasite PF 7.3 CH-17 Shear band in Kg - Black cataclasite with Kg quartz clasts PF 9.8 7.0 34° 35.645 114° 34.003 CH-18 Host Kg - undeformed (Chemehuevi Plutonic Suite) PF 10.1 0.1 7.6 34° 35.645 114° 34.003 CH-19 Pseudotachylite PF 9.8 0.3 7.1 34° 35.645 114° 34.003 CH-19 Deformed (cataclasite bands) Kg adjacent pseudotachylite PF 7.7 34° 35.645 114° 34.003 CH-20 Deformed Kg PF 7.0 CH-21 Shear band in Kg from MWF hanging wall - green cataclasite SS/TW 1.1 0.2 34° 35.801 114° 30.740 CH-21 Shear band in Kg from MWF hanging wall - matrix material -1.5 34° 35.801 114° 30.740 CH-22 Qtz + Ep vein - coarse-grained mineralization on fracture surface TW 6.5 0.4 3.1 0.0 3.3 542 * 51 63 4.3 10.89 34° 35.918 114° 29.794 CH-23 Qtz + Ep vein - coarse-grained mineralization on fracture surface TW 7.0 0.2 3.4 0.1 3.6 511 * 30 34 4.4 10.89 34° 35.918 114° 29.794 CH-24 Leucosome in Xmgn TW 9.1 34° 35.996 114° 30.101 CH-24 Epidote veins cutting leucosome in Xgn TW 5.5 0.1 34° 35.996 114° 30.101 CH-25 Quartz-rich band in Xmgn, cut by veins (green & black) TW 6.8 0.2 34° 35.936 114° 30.152 CH-27 Shear Zone - quartz clasts in cataclasite TW 8.8 34° 36.019 114° 30.007 CH-28 Qtz + Ep vein - cutting cataclasite zone in Kg MW 8.1 0.7 -0.1 1.2 8.2 248 * 51 72 -1.4 12.96 34° 39.832 114° 30.363 CH-3 Shear Band - qtz+epidote mylonite band 5-cm thick SS 10.1 0.0 5.0 0.3 5.1 388 * 21 24 5.3 5.95 34° 34.895 114° 32.666 CH-30 Shear Band - 1 cm black cataclasite with epidote-filled fractures MW 7.3 -0.1 1.3 7.5 273 * 41 54 -1.0 13.10 34° 39.939 114° 30.327 CH-31 Deformed Kg adacent mafic dike (with schistose margins) MW 9.5 34° 39.823 114° 30.351 CH-32 Host Kg - undeformed (Chemehuevi Plutonic Suite) MW 9.7 34° 39.823 114° 30.351 CH-35 Qtz clasts in red cataclasite above MWF MW 10.2 34° 40.700 114° 29.442 CH-36 Qtz vein (mylonitized): coarse grain size, in felsic dike MW 5.5 34° 40.409 114 29.789 CH-37 Qtz vein (mylonitized) in mylonitic felsic dike MW 10.0 34° 40.382 114° 29.849 CH-37 Felsic dike cutting Xgn - mylonitized MW 7.5 34° 40.382 114° 29.849 CH-38 Qtz vein (mylonitized) in dacitic dike MW 10.4 34° 40.369 114° 29.883 CH-4 Shear band - quartz clasts in foliated band at margin of mafic dike SS 7.4 -0.1 0.4 7.6 269 * 14 15 -1.2 5.76 34° 34.850 114 32.767 CH-4 Undeformed Kg 10 cm from mafic dike margin SS 10.3 34° 34.850 114 32.767 CH-41 Qtz Vein (mylonitized) in leucosome cutting Xmgn MW 8.4 34° 40.359 114° 30.029 CH-42 Shear Band - foliated band cutting Xmgn layering MW 7.4 0.5 3.9 0.3 3.5 525 * 72 99 5.0 13.77 34° 40.266 114°30.154 CH-44 Qtz vein in Kg (undeformed) MW 9.1 34° 40.153 114°30.014 CH-45 Qtz vein in Kg (undeformed?) MW 8.9 34° 40.153 114°30.014 CH-46 Qtz vein (undeformed?) in mylonitized silicic dike MW 9.9 0.2 34° 40.203 114°30.059 CH-47 Qtz vein in mylonitized dike MW 9.0 34° 40.203 114°30.059 CH-47 Felsic dike - mylonitic MW 6.9 34° 40.203 114°30.059 CH-5 Mafic dike (dike margin adjacent CH-4) SS 1.0 CH-51 Qtz+chl+ep pod immediately at dike / Xgn contact MW 7.9 34° 40.193 114°30.062 CH-60 Undeformed Kg (Chemehuevi Plutonic Suite) RF 10.1 34° 34.020 114° 35.183 CH-60 Altered Kg - quartz 2 cm from green cataclasite RF 4.3 0.2 34° 34.020 114° 35.183 CH-60 Shear band in Kg - green cataclasite RF 3.9 0.3 34° 34.020 114° 35.183 CH-62 Shear band in Kg - green cataclasite RF 8.7 0.1 34° 34.100 114° 34.661 CH-62 Altered Kg outside shear band RF 10.0 34° 34.100 114° 34.661 CH-62 Pegmatite dike (1 cm) cutting Kg & truncated by green cataclasite RF 9.9 34° 34.100 114° 34.661 CH-65 Undeformed Kg (Chemehuevi Plutonic Suite) RF 10.2 34° 34.100 114° 34.661 CH-66 Altered Kg (Cu-mineralized) at contact with f.g. mafic dike RF 10.3 34° 34.349 114° 35.314 CH-68 Kg (cracked) RF 10.1 34° 34.349 114° 35.314 CH-69 Shear band in Kg - green cataclasite truncated by mafic dike RF 5.3 0.0 34° 34.027 114° 35.289 CH-69 Undeformed Kg adjacent to green cataclasite RF 10.2 34° 34.027 114° 35.289 CH-71 Undeformed Kg (Chemehuevi Plutonic Suite) RF 10.1 34° 34.027 114° 35.289 CH-72 RF 10.2 34° 34.027 114° 35.289 CH-76 Undeformed Kg (Chemehuevi Plutonic Suite) MW 9.0 34° 39.983 114° 29.759 CH-76 Shear band at Kg-Xgn contact - epidote-fill, offsets 1 cm qtz veins MW 8.9 0.2 3.3 0.2 5.7 353 * 22 24 3.2 14.09 34° 39.983 114° 29.759 CH-77 Shear band in Kg - epidote vein in zone of cataclasis MW 7.1 0.2 0.9 1.0 6.2 327 * 52 70 0.1 13.93 34° 39.944 114° 29.845 62
CH-4 Shear band - quartz clasts in foliated band at margin of mafic dike SS 7.4 -0.1 0.4 7.6 269 * 14 15 -1.2 5.76 34° 34.850 114 32.767 CH-4 Undeformed Kg 10 cm from mafic dike margin SS 10.3 34° 34.850 114 32.767 CH-41 Qtz Vein (mylonitized) in leucosome cutting Xmgn MW 8.4 34° 40.359 114° 30.029 CH-42 Shear Band - foliated band cutting Xmgn layering MW 7.4 0.5 3.9 0.3 3.5 525 * 72 99 5.0 13.77 34° 40.266 114°30.154 CH-44 Qtz vein in Kg (undeformed) MW 9.1 34° 40.153 114°30.014 CH-45 Qtz vein in Kg (undeformed?) MW 8.9 34° 40.153 114°30.014 CH-46 Qtz vein (undeformed?) in mylonitized silicic dike MW 9.9 0.2 34° 40.203 114°30.059 CH-47 Qtz vein in mylonitized dike MW 9.0 34° 40.203 114°30.059 CH-47 Felsic dike - mylonitic MW 6.9 34° 40.203 114°30.059 CH-5 Mafic dike (dike margin adjacent CH-4) SS 1.0 CH-51 Qtz+chl+ep pod immediately at dike / Xgn contact MW 7.9 34° 40.193 114°30.062 CH-60 Undeformed Kg (Chemehuevi Plutonic Suite) RF 10.1 34° 34.020 114° 35.183 CH-60 Altered Kg - quartz 2 cm from green cataclasite RF 4.3 0.2 34° 34.020 114° 35.183 CH-60 Shear band in Kg - green cataclasite RF 3.9 0.3 34° 34.020 114° 35.183 CH-62 Shear band in Kg - green cataclasite RF 8.7 0.1 34° 34.100 114° 34.661 CH-62 Altered Kg outside shear band RF 10.0 34° 34.100 114° 34.661 CH-62 Pegmatite dike (1 cm) cutting Kg & truncated by green cataclasite RF 9.9 34° 34.100 114° 34.661 CH-65 Undeformed Kg (Chemehuevi Plutonic Suite) RF 10.2 34° 34.100 114° 34.661 CH-66 Altered Kg (Cu-mineralized) at contact with f.g. mafic dike RF 10.3 34° 34.349 114° 35.314 CH-68 Kg (cracked) RF 10.1 34° 34.349 114° 35.314 CH-69 Shear band in Kg - green cataclasite truncated by mafic dike RF 5.3 0.0 34° 34.027 114° 35.289 CH-69 Undeformed Kg adjacent to green cataclasite RF 10.2 34° 34.027 114° 35.289 CH-71 Undeformed Kg (Chemehuevi Plutonic Suite) RF 10.1 34° 34.027 114° 35.289 CH-72 RF 10.2 34° 34.027 114° 35.289 CH-76 Undeformed Kg (Chemehuevi Plutonic Suite) MW 9.0 34° 39.983 114° 29.759 CH-76 Shear band at Kg-Xgn contact - epidote-fill, offsets 1 cm qtz veins MW 8.9 0.2 3.3 0.2 5.7 353 * 22 24 3.2 14.09 34° 39.983 114° 29.759 CH-77 Shear band in Kg - epidote vein in zone of cataclasis MW 7.1 0.2 0.9 1.0 6.2 327 * 52 70 0.1 13.93 34° 39.944 114° 29.845 CH-78 Qtz + Ep vein (mylonitized) defining slickensurface (050 lineations) MW 9.2 0.0 5.6 3.6 519 * 5 5 6.7 13.45 34° 40.039 114° 30.174 CH-79 Qtz + Ep zone in Xgn. Euhedral ep surrounded by qtz MW 7.1 0.1 3.0 0.3 4.1 464 * 27 30 3.8 13.93 34° 40.335 114° 29.996 CH-79 Qtz vein 1 cm from qtz-ep vein MW 7.6 34° 40.335 114° 29.996 CH-8 Kg adjacent to mafic dike SS 9.8 0.5 34° 34.850 114° 32.767 CH-80 Weathered Kg SS 8.5 CH-81 Undeformed Kg (Chemehuevi Plutonic Suite) SS 9.9 34° 35.608 114° 34.165 CH-81 Deformed Kg adjacent green cataclasite SS 4.9 34° 35.608 114° 34.165 CH-81 Shear band in Kg - green cataclasite SS 3.4 0.0 4.6 34° 35.608 114° 34.165 CH-9 Shear band in Kg - green cataclasite PF 6.0 34° 35.636 114° 32.606 CH-RF Altered (cracked) Kg RF 6.3 CH-RF Shear band in Kg - quartz clasts in green cataclasite RF 10.3 4.4 0.0 5.9 341 * 1 1 4.3 1.63 JL-1 Shear band - qtz + ep mylonite; same location as CH-3 SS 9.0 0.1 34° 34.895 114° 32.666 JL-15 Shear band cutting Xgn MW 5.0 JL-19 Shear band Kg - green shear band SS 1.7 JL-19 Kg (cracked); excludes shear band SS 2.3 JL-2 Shear band in Kg 2 cm from mafic dike SS 1.4 JL-2 Shear band at dike margin - green cataclasite SS -2.6 JL-7 Shear band - green cataclasite w/ multiple generations of mineraliz. RF 8.2 3.2 1.0 5.0 396 * 57 76 3.5 1.82 34°34.218 114°35.230 aKg refers to Cretaceous granitoids of the Chemehuevi Plutonic Suite, and does not denote a particular intrusive phase; Xgn is used to refer to Precambrian gneiss; MWF = Mohave Wash Fault; PSZ = principle slip zone. *Temperature calculated using the calibration of Matthews (1994) for Ps30 (A=2.225). † Temperature calculated using the quartz-garnet calibration of Bottinga and Javoy (1975) (A=2.88). b 18 Error estimates for temperatures were calculated by incorporating 1 s.d. on replicate analyses to caluclate maximum and minimum Δ Oquartz-epidote.
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FIGURES
Figure 1. Schematic cross-section at of an active low-angle normal fault and possible fluid flow pathways highlighted by blue arrows.
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Figure 2. Geologic map of the Chemehuevi Mountains area, southeastern California, from John (1987) showing the location of the Chemehuevi and Mohave Wash Faults, and the distribution and orientation of the late Cretaceous (?) and middle Tertiary
Chemehuevi dike swarm. Cross section profile along line A-A’ is shown in Figure 3.
Yellow stars mark primary field description and sampling locations.
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Figure 3. Cross section of line A-A’ shown on Figure 2. Note position, and structural relations between the Mohave Wash Fault and later Chemehuevi Fault. Modified from John and Foster (1993).
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Figure 4. Structural profile of the Mohave Wash fault damage zone at the Studio Springs field site. Fractures were counted every meter; fracture density is plotted with the lithological profile. Fracture orientation data is represented in a pole stereonet, rose diagram, and histogram.
67
68
Figure 5. Outcrop photos of the Mohave Wash fault showing the variation of deformation structures moving from shallow to deep paleodepths during initiation. (1) Representative outcrops at Range Front site: (1A) mafic dike intruding granitoid, both showing intense cracking and brittle deformation; (1B) basaltic dike grading into an epidote–mineralized .
(2) Studio Springs field site: (2A) panoramic photo of the main MWF damage zone near the Pool Flats area (2B) pegmatite dikes offset by green cataclasite bands (2C) enlarged photo of green cataclasite; (2D) 50 cm-thick band of cataclasite cutting granitoid; (2E) pseudotachylite; (2F) photomicrograph from the area outlined by a black rectangle in 2E.
(3) Trampas Wash site represents mid-range paleodepths (estimated 5-8 km; Foster and
John, 1993): (3A) Fault gouge capping a cataclasite cutting Cretaceous granitoid; (3B)
Quartz-epidote vein exposed as a surface coating cutting Precambrian gneiss; the exposed quartz-epidote vein is offset by hematite-filled fractures, resulting from later-stage mineralization involving highly oxidizing fluids. (4) Mohave Wash field site represents the deepest exposed rocks in the Chemehuevi Mountains: (4A) Mylonitic Miocene dacite dike cutting Precambrian gneiss; (4B) close-up photo and photomicrograph (4C) of the dike from 4A highlighting the strong foliation present; (4d) Chlorite breccia within the
Mohave Wash fault along the MWF fault contact between Cretaceous granitoids and
Precambrian gneiss.
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Figure 6. δ18O of quartz and epidote vs. structural rock type. The grey bar represents magmatic (unaltered) values for quartz, defined by samples of undeformed rocks in the
Chemehuevi Plutonic Suite. The two green bars represent magmatic δ18O values for epidote at 600˚C calculated using two different fractionation calibrations: the dark green bar is the range calculated using A=2.225 from Matthews (1994). The light green bar is the range calculated using A= 3.1 from Ferreira et al. (2003). Both calculations assume a epidote composition of XPs = 30. Deformed rock types contain quartz and epidote that extend to lower δ18O values then undeformed host granites.
70
Figure 7. δ18O whole rock values versus structural rock type for different field sites. The grey bar represents the whole rock δ18O range of the host granites, defined by undeformed samples Ch-80, Ch-18
71
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Figure 8. Photomicrographs of deformation microstructures from the Mohave Wash fault. (1) Thin section scan (XPL) of sample CH-RF featuring multiple zones of deformation: (1A) Cracked, hydrothermally-altered Kpg granite; (1B) cataclasite zone with epidote dominated matrix and large clasts of sub-rounded quartz; (1C) Foliated shear zone with quartz, epidote, and chlorite. Sample shows an increase in intensity of deformation over a few centimeters from CàBàA. (2) SEM image of CH-RF area shown is outlined by a white box in 1. Ap = apatite, Ep = Epidote, Qtz = quartz. (3)
Photomicrograph of pseudotachylite (CH-19) plume injecting into deformed quartz pod.
(4) Photomicrograph (PPL) of quartz-epidote vein featuring large euhedral (undeformed) epidote crystals (CH-24). (5) Representative photomicrograph of a mylonitized quartz vein (CH-41). (6) Photomicrograph of shear band consisting primarily of microcrystalline epidote and quartz, with a lenticular clast of recrystallized quartz (Sample 42).
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Figure 9. δ18O whole rock values versus relative vertical height above base of the
Mohave Wash fault damage zone in the Pool Flats area. The schematic lithological profile (right) depicts fracture intensity, and was determined by counting fractures per meter and recording deformation features in the field.
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Figure 10. Distribution of quartz, epidote, and whole rock δ18O values by field site plotted versus distance along the Mohave Wash Fault in the down-dip direction. The top grey bar represents the range of primary magmatic quartz δ18O values for the
Chemehuevi plutonic suite; the darker grey bar represents the whole rock δ18O range.
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Figure 11. δ18O of quartz vs. epidote of the Mohave Wash fault zone (this study) and the nearby Whipple Mountains (Morrison & Anderson, 1998). The solid lines (preferred)
18 represent temperatures calculated based on measured Δ Oquartz-epidote and the fractionation calibration of Matthews (1994; A = 2.225) assuming an epidote composition of Ps30. The dashed lines are temperatures assuming an epidote composition of Ps0 (A= 2.0) from
Matthews (1994).
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Figure 12. δ18O of mineral pairs and calculated temperatures assuming the mineral pairs equilibrated during deformation and/or precipitation from hydrothermal fluids. Replicate analyses of minerals from the same structure are shown, where analyzed; the black circles represent the average of replicate analyses, and were used to calculate temperatures.
Uncertainties in the calculated temperatures were determined by assuming the maximum
18 and minimum possible Δ Oquartz-epidote values based on the measured data. Temperatures shown are calculated assuming an epidote composition of Ps30 using A= 2.225 from
Matthews (1994).
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Figure 13. Apparent temperature calculations from oxygen isotope thermometry vs. sampling distance in the slip direction. Temperatures calculated using equation 3 with an assumed composition for epidote of Ps30, which are considered most likely to be representative of epidote in these samples. Possible range of temperatures (including error estimates) are shown as vertical error bars. The grey bar represents mineral closure temperatures at the time of initiation (23 Ma), determined by John and Foster (1993). The width of the bar represents uncertainties in geographic locations based on their sampling.
Zero is an arbitrary starting location.
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Figure 14. Estimated vertical distance relative to the Mohave Wash Fault versus A) oxygen isotope temperature and B) measured δ18O. Zero is an arbitrary point defined within the damage zone.
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Figure 15. Map of the Chemehuevi Mountains showing simplified paleoisotherm (dashed lines) calculated by John and Foster (1993) based on mineral closure temperatures and
Ar/Ar ages. Yellow circles represent apparent oxygen isotope temperatures calculated assuming an epidote composition of Ps30 using A= 2.225 from Matthews (1994).
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Figure 16. Apparent oxygen isotope temperature vs. δ18O of quartz, epidote, and water
(calculated). The grey bars represent the δ18O range of magmatic and meteoric fluids associated with the Chemehuevi range. The black bar at the top of the figure represents the δ18O range inferred for brines responsible for alteration in the Harcuvar Mountains,
AZ (Roddy et al., 1988). Dashed arrows represent possibly rend in change of δ18O with change in temperature.
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Figure 17. Model for the evolution of Chemehuevi Mountains metamorphic core complex and fluid-rock interaction history. Adapted from Rehrig and Reynolds (1980).
(A) Concept of three crustal layers, each responding distinctly to heating and extension.
(B) Intrusion of the Chemehuevi plutonic suite into Precambrian gneiss country rocks during the Cretaceous. (C) Onset of Miocene extension and dike emplacement during the main stage of CREC formation (initiating ~23 Ma). Blue arrows define inferred movement of hydrothermal fluids. Early deformation products include mylonitized
Miocene dikes and crosscutting quartz veins from the Mohave Wash site, preserve δ18O values in equilibrium with magmatic fluids (D) Formation of the Mohave Wash fault damage zone. This is one possibility, another would be that the CDF also formed at this
83 time. Tilting of upper plate blocks create basins, allowing formation of lakes and basinal brines that migrate downward. Brittle deformation is recorded along the shallow and intermediate depth of the MWF fault and localized plastic deformation is present along the deeper portions of the fault. Diking continues throughout this time and drives rock- dominated fluids (δ18O = 3 to 7‰) up the damage zone. These fluids may originate as brines shifted to higher δ18O during subsequent high temperature water-rock interaction, or mixing between magmatic fluids expelled from below and surface-derived fluids. (E)
Initiation of the Chemehuevi detachment fault and subsequent continued tilting of upper- plate blocks, at this point slip along the MWF has ceased. Basinal brines continue to migrate downward, moving down high-angle normal faults, and become heated and driven up the MWF damage zone by thermal buoyancy. Hot, high δ18O rock-dominated fluids may continue to move up the damage zone in the deepest portions of the fault zone, but deformation microstructures record infiltration of progressively lower δ18O fluids (-3 to 3‰) with decreasing temperature. Fluid circulation is dominantly localized more in the hanging wall to the MWF, and probably along the CDF. (F) As tectonic denudation proceeds, area of maximum thinning responds to removal of zone 1 by isostatic uplift and deformational thinning in zone 2. (G) Schematic representation of the present-day configuration of Chemehuevi Mountains metamorphic core complex. portions of the fault zone, but deformation microstructures record infiltration of progressively lower d18O fluids (give d18O range) with decreasing temperature. Fluid circulation is dominantly localized more in the hanging wall to the MWF, and along the
CDF. (F) As tectonic denudation proceeds, area of maximum thinning responds to
84 removal of zone 1 by isostatic uplift and deformational thinning in zone 2. (G) Schematic representation of the present-day configuration of Chemehuevi Mountains metamorphic core complex.
85
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