SESolid EarthSESolid Earth1869-9529Copernicus PublicationsGöttingen, Germany10.5194/se-8-379-2017Structural and rheological evolution of the Laramide subduction channel in
southern CaliforniaXiaHaoranhaoranxi@usc.eduPlattJohn P.https://orcid.org/0000-0002-4459-1864Department of Earth Sciences, University of Southern California, Los Angeles, CA 90089-0740, USAHaoran Xia (haoranxi@usc.edu)30March20178237940322October201626October20162March20178March2017This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://se.copernicus.org/articles/8/379/2017/se-8-379-2017.htmlThe full text article is available as a PDF file from https://se.copernicus.org/articles/8/379/2017/se-8-379-2017.pdf
The Pelona Schist in the San Gabriel Mountains, southern California, formed
in the Laramide subduction channel, exhibits multiple phases of
deformation/metamorphism and provides valuable insights into the rheological
properties of the subduction channel. Petrological and microstructural
analysis indicates that the Pelona Schist has undergone three major
deformational/metamorphic events. Subduction of volcanic and sedimentary
protoliths during D1 was recorded by aligned mineral inclusions in albite and
epidote porphyroblasts. Metamorphic temperature and pressure at the end of
subduction yielded by Raman spectroscopy of carbonaceous material and
phengite barometry were 519 ± 20 ∘C and
10.5 ± 0.4 kbar, respectively. During D1 the dominant deformation
mechanism was quartz pressure solution, and the estimated shear stress at the
end of D1 was less than 10 MPa. D2, the first stage exhumation of the Pelona
Schist along the upper section of the subduction channel during return flow,
was recorded by retrogressive metamorphism, isoclinal folding, and a
pervasive schistosity that wraps around earlier porphyroblasts. Metagreywacke
was deformed mainly by quartz pressure solution and metachert was deformed
dominantly by dislocation creep during D2. The shear stress in metagreywacke
was less than 10 MPa and that in metachert was between
8.3 + 2.7/- 1.5 and 12.9 + .9/- 2.3 MPa, resulting in a strain
rate of 1.4 × 10-13 to 5.5 × 10-13 s-1. A
topography driven model is proposed as the main driving force of D2
exhumation. D3 records normal-sense movement on the Vincent Fault, which
separates the schist from overlying arc and continental basement. This
resulted in the second stage of exhumation, creating a major synform and
associated mylonitic fabric in the upper section of the Pelona Schist.
Conditions at the beginning of D3 were 390 ± 13 ∘C and
5.8 ± 0.8 kbar given by the TitaniQ thermometer and phengite
geobarometer. The deformation was dominated by quartz dislocation creep with
a strain rate of 4.5 ± 1.2 × 10-13 s-1 at a shear
stress of 20.1 + 7.3/- 4.0 MPa.
Introduction
Sedimentary and volcanic materials lying on subducting plates are commonly
carried down and deformed within a subduction channel between the two plates
(Cloos and Shreve, 1988; Gerya and Stockhert, 2002).
Some of this material may be exhumed after experiencing high-pressure, low-temperature metamorphism. Various models have been proposed to interpret the
burial and exhumation cycle of high-pressure subduction complexes in
accretionary settings (e.g., Brandon
et al., 1998; Brun and Faccenna, 2008; Platt, 1986; Ring et al., 2007). The
best way to test these models is to examine the deformational and
metamorphic history of rocks exhumed from the subduction channel.
The rheological properties of subduction channels may place controls on the
interaction of the converging plates, such as depth of seismicity
(e.g., Ruff and Tichelaar, 1996) and coupling of the
plates (e.g., Stöckhert, 2002). Geophysical
observations cannot easily yield the mechanical properties of the plate
interface in active subduction zones (Grigull et al., 2012;
but see Houston, 2015). The only directly available information on small-scale deformation mechanisms, stress states and material strength along the
plate interface comes from the rocks exhumed from a subduction zone
(e.g., Wassmann and Stöckhert, 2013).
The Pelona Schist in the Transverse Ranges, southern California, was formed
during the Late Cretaceous–Paleocene Laramide subduction event, and it is an
ideal case to scrutinize the subduction zone. In this paper we first
construct the deformational and metamorphic history of Pelona Schist by
employing multiple thermobarometers and zircon fission track analysis, then
decipher the exhumation mechanism of the Pelona Schist, and finally we infer
deformation mechanisms and rheological properties including stresses and
temperatures along the subduction channel during the different deformation
stages.
Geologic background
Laramide subduction caused eastward migration of magmatic activity in western
North America from Late Cretaceous to early Tertiary time; this, together
with basement-involved thrusting in the Laramide foreland, led to the idea
that this was a period of flat-slab subduction (Coney and Reynolds, 1977;
Dickinson and Snyder, 1978). Although debates exist (e.g., Barth and
Schneiderman, 1996; Ehlig, 1981), it is widely accepted that the Pelona,
Orocopia, and Rand schists and the schists of Portal Ridge and Sierra de
Salinas in southern California and adjacent areas are products of this
subduction event (e.g., Burchfiel and Davis, 1981; Crowell, 1981; Grove et
al., 2003; Hamilton, 1988; Jacobson et al., 2007), and that their protolith
ages decrease from northwest to southeast (Grove et al., 2003). These
ocean-affiliated schists were metamorphosed under moderately high-pressure
conditions from blueschist facies to amphibolite facies beneath the
Cordilleran Mesozoic magmatic arc, and generally preserve retrograde
metamorphic textures (e.g., Chapman et al., 2010; Jacobson and Dawson, 1995;
Kidder and Ducea, 2006).
Nowadays the schists are separated from the upper plate of Precambrian to
Mesozoic igneous and metamorphic rocks by normal faults. For the Rand Schist
and the schist of Sierra de Salinas, Chapman et al. (2010) suggested that the
contact between the schists and the upper plate was remobilized as a normal
fault during exhumation even though it was proximal to the original
subduction megathrust. In the Orocopia Mountains, Jacobson and Dawson (1995)
interpreted the upper boundary of the Orocopia schist as an early Miocene
low-angle detachment fault, underlain by a thin zone of mylonitized and
retrogressed schist. However, the Vincent Fault above the Pelona Schist,
which is also known as the “Vincent thrust”, was argued to be the original
subduction zone megathrust (Ehlig, 1981; Jacobson et al., 1996).
The Pelona Schist occurs in the Sierra Pelona and in the eastern San Gabriel
Mountains of the Transverse Ranges (Fig. 1); we focus on the latter in this
paper because of the outstanding exposures along the East Fork of the San
Gabriel River. The Pelona Schist in the East Fork area is separated by the
regionally gently SW-dipping Vincent Fault from a Meso-Proterozoic gneiss
complex and Mesozoic granitoid rocks in the hanging wall (Fig. 2), which
formed part of the Mesozoic arc in California. A striking feature of the
Pelona Schist is the complete lack of intrusive rocks related to the Late
Cretaceous arc that directly overlies the Vincent Fault. A mylonite zone is
developed in the lower 500 to 1000 m of the hanging wall. The Pelona Schist
in the East Fork area occurs in the core of a NW-trending post-metamorphic
antiform (Jacobson, 1997), and is truncated to the NE by the Punchbowl and
San Andreas faults (Dibblee, 1967). Through the approximately 4 km thick
transect in the East Fork, the Pelona Schist consists of ∼ 90 %
metagreywacke, ∼ 10 % mafic greenschist and a minor amount of
metachert (Jacobson, 1997), all of which were metamorphosed under
high-pressure greenschist facies conditions. The peak pressure and
temperature were constrained in the Sierra Pelona at up to 10 ± 1 kbar
and 620–650 ∘C (Graham and Powell, 1984). An inverted thermal
gradient in the Pelona Schist was reported both in the Sierra Pelona (Graham
and Powell, 1984) and in the East Fork of the San Gabriel Mountains
(Jacobson, 1983a, 1995). The protolith of the Pelona Schist has a maximum age
of 68 Ma based on detrital zircon 206Pb/238U dating (Grove et al.,
2003). Amphibole 40Ar/39Ar ages from mafic schist in the East Fork
are 60.3 ± 2.6 and 58.9 ± 2.5 Ma (Jacobson, 1990), and white
mica 40Ar/39Ar ages cluster between 55.5 and 60.8 Ma (Grove et
al., 2003; Jacobson, 1990) though one metagreywacke sample from the deepest
structural level yields a white mica 40Ar/39Ar age of
31.7 ± 0.2 Ma. To improve the precision of cooling history of the
Pelona Schist, in this study we also performed zircon fission track analyses
as presented in Sect. 5.4.
Regional tectonic map of southern California showing
outcrops of Laramide subduction related schists, Mesozoic magmatic arc, and
major faults. Modified from Jennings (1977).
(a) Geologic map of eastern San Gabriel Mountains. Modified
from Dibblee (2002a, b, c, d). The background topographic map is adapted from
the 1 : 100 000-scale Metric Topographic Map of San Bernardino,
California, compiled and published by the United States Geological Survey in 1982.
Contour interval of 200 m. (b) Cross section of A-A′ in
panel (a).
A phase of Miocene magmatic activity in the East Fork area is characterized
by intermediate dykes and sills in both the upper plate and the Pelona
Schist, as well as a dacite sill complex with a cumulative thickness of a few tens
of meters parallel to the main schistosity of the Pelona Schist near Iron
Fork, which was dated at 14–16 Ma by the K/Ar method (Miller and
Morton, 1977).
MethodsThermobarometersRaman spectroscopy on carbonaceous material (RSCM)
Organic material in rocks undergoes irreversible graphitization during
prograde metamorphism, and the degree of this transformation is a function
of temperature and is not subject to retrogression, so the temperature
measured by this approach approximates the maximum temperature
(Beyssac et al., 2002). The degree of organization of
carbonaceous material can be measured using Raman microspectroscopy by
calculating ratios of the peak areas at bands of 1350, 1580, and 1620 cm-1. The calibration used here
(Beyssac et al., 2002) comes with a maximum error of
±50 ∘C, stemming from the uncertainties in the
thermobarometric methods used to calibrate it. The precision of the method,
however, is ∼ 10 ∘C (Beyssac
et al., 2002). RSCM measurements of carbon-rich metagreywacke samples were
completed with a 514 nm argon laser in the Mineral Microspectroscopy
Facility at the California Institute of Technology, and the spectra were
processed with the software Peakfit following the procedure outlined in
Beyssac et al. (2002).
Quartz c axis fabric opening-angle thermometer
The empirical quartz c axis fabric opening-angle thermometer
(Kruhl, 1998) utilizes the linear relationship
between the opening angle of quartz c axis fabric and the temperature during
fabric formation. This deformation-related thermometer has an uncertainty of
±50 ∘C, and is applicable for rocks deformed in the
∼ 300–650 ∘C range. Kruhl (1998) noticed that at given temperatures,
the opening angles of quartz c axis fabrics from experimentally deformed
rocks at high strain rates are smaller than those from naturally deformed
rocks, but he argued that strain rate should not affect the thermometer when
applied to rocks deformed at natural conditions because the thermometer was
established on rocks deformed at natural strain rates. By promoting prism
[c] slip, high water content may result in larger opening angles of fabrics,
but the direct link between water content and opening angle of quartz fabric
has not yet been demonstrated experimentally (Law, 2014).
Quartz fabrics, or crystallographic preferred orientations (CPOs), were
measured and analyzed using a Hikari electron backscatter diffraction (EBSD)
detector mounted on a JEOL-7001F scanning electron microscope at the Center for
Electron Microscopy and Microanalysis, University of Southern California, and
associated OIM collection/analysis software developed by EDAX. The
acceleration voltage was 25 kV, the working distance 15 mm, and the step
size half of the pre-estimated average grain size.
Titanium-in-quartz thermobarometer (TitaniQ)
The substitution of Si by Ti in quartz is P and T dependent, and it enables
P–T estimation by measuring titanium content in quartz
(Thomas et al., 2010; Wark and Watson, 2006) and
combining the results with an independent thermobarometer. In this study we
combined TitaniQ with the phengite geobarometer. Calibrations were performed
at constant pressure (Wark and Watson, 2006) and varying
pressure conditions (Thomas et al., 2010) by synthesizing
quartz and rutile from silica-saturated aqueous fluids. The uncertainty in
the temperature determined using the calibration of Thomas et al. (2010) is approximately ±20 ∘C if the error of the independently constrained pressure is
±1 kbar. Huang and Audétat (2012) determined
an alternative calibration by assuming that their lowest Ti concentration
measurements were closest to equilibrium. Thomas
et al. (2015) evaluated the published calibrations by synthesizing and then
recrystallizing quartz at different P–T conditions, and confirmed the
validity of the calibration of Thomas et al. (2010),
which we used in this study. Samples were first examined by
cathodoluminescence imaging using a Tescan Vega-3 XMU scanning electron
microscope at University of California, Los Angeles, and were then measured
on a Cameca IMS 6f secondary ion mass spectrometer (SIMS) at Arizona State
University. SIMS settings and analytical procedures followed those given by
Behr et al. (2011).
Phengite geobarometer
The composition of the muscovite–celadonite solid solution series is a
function of pressure and temperature. Substitution of celadonite for
muscovite, resulting in higher Si content, is favored by increasing pressure
and decreasing temperature, based on which Massonne and
Schreyer (1987) established a phengite geobarometer. The calibration is
based on independent P–T estimates from coexisting mineral assemblages, and
is subject to large uncertainties. Mineral analyses were performed on a JEOL
JXA-8200 electron microprobe at University of California, Los Angeles. The
acceleration voltage was 15 kV, the beam current 10 nA, and the beam diameter
5–8 µm.
Recrystallized grain size piezometry for quartzGrain size measurement
Grain boundaries were manually traced on the optical photomicrographs and
then the grain size was measured in ImageJ. The grain size is defined as the
diameter of a circle with the same area as the grain in thin section. The
average 2-D grain size of recrystallized quartz grains in a sample is
calculated as the root mean square (RMS) diameter of all the measured grains
in this sample. No stereological correction was applied in order to keep
consistent with the piezometer calibration (Stipp and Tullis,
2003). The grain sizes are reported with 1 standard deviation. Thin sections
used for grain size measurement were cut normal to foliation and parallel to
lineation.
Stress calculation
Differential stress was calculated using quartz recrystallized grain size
piezometer D=103.56±0.27×σ-1.26±0.13, where D is the grain size of recrystallized quartz in microns and
σ is the differential stress in megapascals (Stipp and Tullis,
2003). Differential stress calculated from the previous equation was
multiplied by a factor of 0.73 to correct for possible errors in the stress
measurements on the Griggs-type apparatus used for the original calibration
(Holyoke and Kronenberg, 2010). The piezometer can be
applied to quartz recrystallized grain size up to 120 µm
(Stipp et al., 2010). Within uncertainties, the
piezometer is independent of the water content of the quartz, temperature,
and the quartz alpha–beta phase transition (Stipp et al.,
2006). Shear stress in a plane stress setting was calculated from
differential stress by dividing by 3 (Behr and
Platt, 2013).
Quartz shape preferred orientation
Digital polarized-light photomicrographs were taken and then the grain
boundaries of quartz were traced. The graphs of quartz grain boundaries were
processed by ImageJ to measure the orientation of grain long axes. The
results were presented in rose diagrams plotted by MATLAB.
Outcrops of Pelona Schist. Photo locations can be found in Fig. 2a
and Table S1. (a) D2 folds in metagreywacke. (b) Mafic
greenschist with boudinaged quartz veins. (c) D2 folds in metachert.
(d) Sheath folds in metachert and mafic greenschist.
(e) Minor D3 folds near the hinge of the Narrows synform.
(f) Vincent Fault as shown by the arrows; viewing SE.
Detrital zircon fission track analysis
Detrital zircon fission track analysis was performed on six metagreywacke
samples by Apatite to Zircon, Inc. The zircon grains were mounted, etched and
polished for optical analysis following the procedure outlined in Moore et
al. (2015). The U content was measured using Agilent 7700x LA-ICP-MS mounted
with a Resonetics RESOlution M-50 laser ablation system.
Deformation history of the Pelona Schist in the eastern San Gabriel
Mountains
The Pelona Schist in the San Gabriel Mountains consists mainly of
metagreywacke (Fig. 3a) with minor amounts of mafic greenschist (Fig. 3b)
and metachert (Fig. 3c), all of which were metamorphosed up to high-pressure
greenschist facies. The typical mineral assemblage of metagreywacke is
albite + quartz + white mica + chlorite + stilpnomelane + epidote
+ sphene with garnet grains and graphitized carbonaceous material
preserved within the albite porphyroblasts. The composition of garnet
changes from spessartine rich in the core to grossular rich in the rim
(Jacobson, 1983a). Mafic greenschist is made up of amphibole +
albite + chlorite + epidote + quartz + sphene. Amphibole here is
actinolite to actinolitic hornblende with high Na / Al ratio
(Jacobson, 1997). Metachert consists of quartz and minor amounts
of white mica, stilpnomelane, albite, garnet, sodic amphibole, and
piemontite. Calcite can be found in all three rock types. Boudinaged quartz
veins are common in greenschist.
S1 fabric in metagreywacke and greenschist. See Fig. 2a and Table S1
for sample locations. (a) The mineral inclusions in albite outline
S1. Note the differentiated cleavage. Sample PS52. (b) S1 in mafic
greenschist shown by the mineral inclusions in the epidote porphyroblast.
Sample 95. (c) CPO of quartz inclusions in one albite porphyroblast
in Sample PS52. (d) S1 in albite grains is crenulated as shown by
the trails of graphite. Sample PS109.
The earliest fabric (S1) in the Pelona Schist is preserved as straight
inclusion trails in albite porphyroblasts in metagreywacke, and in albite
and epidote porphyroblasts in greenschist (see Sect. 4.2, below). The
dominant deformational structure in the Pelona Schist is folding of the
compositional layering (Fig. 3a, c, d and e). Two types of the folds have
been recognized in the field: (1) pervasive tight isoclinal folds (Fig. 3a, c
and d) and (2) the large-scale Narrows synform and associated minor folds
(Fig. 3e). Isoclinal folds occur pervasively in the East Fork transect, and
are characterized by their tightly closed limbs with a length varying from
less than a meter to a few tens of meters. SW-dipping axial-planar
schistosity (S2) is well developed. Sheath folds (Fig. 3d) and axis-parallel
stretching lineation indicate SE–NW stretching. The Narrows synform and
associated minor folds occurred in the upper 700 m section of the Pelona Schist and refold the isoclinal folds. The Narrows synform is more open,
but its hinge is generally parallel to those of the isoclinal folds.
Moderate SW-dipping axial-planar crenulation cleavage (S3) is developed in
the Narrows synform. The overturned upper limb of the Narrows synform is cut
by the Vincent Fault. Both S2 and S3 are moderately SW dipping, and the
pervasive sub-horizontal stretching lineation trends SE.
S2 fabric in metagreywacke. Sample locations are shown in Fig. 2a
and listed in Table S1. (a) Quartz filled between two pulled-apart
ablite porphyroblasts in the center of the image. Sample PS107.
(b) Quartz precipitated in the pressure shadows between albite
porphyroblasts. Sample PS185. (c)Q domain in Sample PS190.
(d) Undulose extinction and bulging in quartz grains precipitated in
pressure shadows in Sample PS185. (e) Shear bands and asymmetric
albite porphyroblast tails show a top-to-NW sense of shear. Sample PS185.
(f) Conjugate shear bands outlined by white mica and chlorite in
Sample PS139.
MicrostructuresD1 fabrics
S1 is preserved by the trails of mineral inclusions in albite porphyroblasts
in metagreywacke (Fig. 4a), and in epidote and albite porphyroblasts in
greenschist (Fig. 4b). The mineral inclusions are white mica, epidote,
quartz, garnet, zircon and abundant carbonaceous material. Some of the
trails are rich in carbonaceous material or mica, while others are rich in
quartz, which defines a differentiated cleavage. Orientation of S1 varies
from one porphyroblast to another. Quartz in S1 has a grain size ranging from
10 to 50 µm and shows a random CPO (Fig. 4c). No undulose extinction or
recrystallization was observed in quartz inclusions.
D2 fabrics
The Pelona Schist exhibits ductile deformation during D2, and no
pseudotachylite concordant with S2 was found in the Pelona Schist.
Metagreywacke
Rims of some albite porphyroblasts are free of
inclusions, which probably resulted from post-D1 overgrowth. Albite does not
show recrystallization or undulose extinction. A few albite grains were
broken apart with quartz deposited in between (Fig. 5a). White mica and
chlorite are strongly differentiated into phyllosilicate-rich domains, or
P domains (Shelley, 1993), and define S2 (at top of Fig. 5b),
which wraps around the albite porphyroblasts. In the hinges of isoclinal
folds, the trails of minerals included in albite (S1) were crenulated (Fig. 4d), which might represent an early stage in the development of S2, prior to
the growth of albite.
Quartz crystals in quartz-rich domains, or Q domains (Shelley,
1993), and in strain shadows around albite and epidote porphyroblasts, are
equant and coarse-grained (100 to 200 µm), with trails of fluid
inclusions and relatively straight grain boundaries, though bulging occurs
in some crystals (Fig. 5b, c and d). They show undulose extinction. Quartz
crystals in P domains have a large aspect ratio (more than 2) and a
grain size of ∼ 30 to 100 µm (Fig. 5b and c).
(a) Metachert from Prairie Fork (Sample PS186) and
(c) its quartz CPO. The sense of shear cannot be inferred from
panel (a) due to the lack of indicators. The quartz c axis pole
figure in panel (d) shows a symmetrical skeleton, but there are
variations in density suggestive of top-to-SE sense of shear.
(b) Metachert from Iron Fork (Sample PS103) and (d) its
quartz CPO. The diagram in panel (b) illustrates the shape
orientation of recrystallized quartz grains; quartz grains pinned by other
minerals were not counted. Both quartz recrystallized grain fabric in
panel (b) and quartz c axis pole figure in panel (d) show
a top-to-SE sense of shear. Note the grain size difference between
panels (a) and (b). (e) Metachert with greenish
needle-shaped amphibole and brown stilpnomelane. Sample PS131 from Fish Fork.
Shear bands defined by amphiboles exhibit a top-to-NW sense of shear.
(f) Metachert sample (PS100) from Iron Fork. The diagram in
panel (f) illustrates the shape orientation of recrystallized quartz
grains; quartz grains pinned by other minerals were not counted. The quartz
recrystallized grain shape fabric and tilted prismatic subgrain walls
indicate that the sense of shear was top-to-SE. Sample locations are
available in Fig. 2a and Table S1.
Shear sense indicators are not common in the metagreywacke. Shear bands and
asymmetric albite porphyroblasts are the only two types of shear sense
indicators found in the samples from north of Iron Fork, in the structurally
deepest part of the schist. They show a top-to-NW sense of shear (Fig. 5e).
In the Iron Fork area, conjugate shear bands are observed and no consistent
shear sense can be inferred (Fig. 5f). In samples south of Iron Fork, S2 has
been overprinted by S3 and its sense of shear cannot be inferred.
Metachert
Quartz grains exhibit equant grain shape and relatively
straight grain boundaries, which may be pinned by oriented white mica. Most
quartz grains show weak undulose extinction, and deformation lamellae occur
in a few cases. In the structurally lowest level near Prairie Fork, the
grain size ranges from 100 to 300 µm and the average is 136 ± 47 µm (Fig. 6a). In samples from Fish Fork and Iron Fork that were not
strongly annealed, quartz shows evidence of recrystallization by subgrain
rotation and/or grain boundary migration (Fig. 6b). The average
recrystallized grain size varies between 47 ± 14 and 82 ± 25 µm (Table 1). Quartz c axis pole figures of samples close to Prairie
Fork exhibit maxima around the z axis and a sub-maximum near the y axis
(Fig. 6c), while samples near Iron Fork show discontinuous to continuous
crossed girdles (Fig. 6d).
Sense of shear in metachert can be inferred from quartz c axis pole figures
and shear bands. For rocks north of (structurally below) Iron Fork, the
quartz CPO of a Prairie Fork sample appears to be symmetrical (Fig. 6c) and
the shear bands of a Fish Fork sample show a top-to-NW sense of shear (Fig. 6e). For rocks south of (structurally above) Iron Fork, quartz
recrystallized grain shape fabric and CPOs show a consistent top-to-SE
sense of shear (e.g., Fig. 6b, d and f).
Mafic greenschist (sample PS193) showing S2 fabric under
(a) plane light and (b) crossed polarized light. Ab:
albite; Act: actinolite; Chl: chlorite; Ep: epidote; Qtz: quartz.
(c) Crystallographically continuous overgrowths on an albite
porphyroblast in mafic greenschist. Sample PS97. (d) Fibrous quartz
strain shadow around a magnetite crystal in mafic greenschist. Sample PS101.
Locations of samples can be found in Fig. 2a and Table S1.
(a) D3 crenulations in Sample PS32 from the hinge of the
Narrows synform. (b) Recrystallized quartz close to the hinge of the
Narrows synform. Sample PS52. (c) Quartz in a mylonitized
metagreywacke sample (PS36). (d) Lineation-perpendicular view of the
mylonitized metagreywacke. S3 defined by the crenulated white mica is
sub-vertical, and so is the quartz shape preferred orientation. Sample PS35.
See Fig. 2a and Table S1 for sample locations.
Greenschist
Chlorite, actinolite, tabular albite and epidote
define S2 (Fig. 7a and b). As in the metagreywacke, albite in greenschist
also exhibits evidence of overgrowth (Fig. 7c). Quartz is concentrated in
strain shadows around epidote porphyroblasts or magnetite (Fig. 7d). Most of
the quartz is recrystallized by grain boundary migration or subgrain
rotation.
D3 fabric
S3 is best preserved in the metagreywacke in the hinge and upper limb of the
Narrows synform; the S2 foliation was strongly crenulated, and the axial
planes of the crenulation define S3 (Fig. 8a).
In the upper limb of the Narrows synform, quartz is completely
recrystallized by subgrain rotation (SGR), and the recrystallized grain size
clusters around 30 µm. Noticeably, some of the grain boundaries of the
recrystallized quartz grains show an irregular shape due to bulging, forming
fine-grained quartz crystals (Fig. 8b). Albite remains intact and does not
show any evidence of plastic deformation.
The top ∼ 100 m of the Pelona Schist immediately beneath the
Vincent Fault exhibits strong mylonitic microstructure. Quartz grains have
been completely recrystallized by SGR, showing a recrystallized grain shape
fabric and indicating a top-to-SE sense of shear. The grain size of
recrystallized quartz is around 18 µm (Fig. 8c). In thin sections cut
perpendicular to the lineation, axial planes of white mica crenulations are
parallel to the shape preferred orientation of recrystallized quartz (Fig. 8d). This suggests that crenulation of white mica due to the Narrows synform
was coeval with mylonitization along the Vincent Fault.
Sense of shear during the deformation history
Here we summarize information on the sense of shear in the Pelona Schist.
It is not possible to determine the sense of shear during D1, as its
original orientation was strongly modified during later deformation.
The sense of shear during D2 varies depending on the structural level. For
rocks structurally deeper than Iron Fork, though exceptions exist, the shear
bands in metagreywacke and metachert show a top-to-NW sense of shear (Figs. 5e and 6e). The metachert sample PS186 (Fig. 6a) does not show an
asymmetric geometry in its quartz c axis pole figure (Fig. 6c) and therefore
its sense of shear cannot be determined. For rocks structurally shallower
than Iron Fork (south of Iron Fork), quartz pole figures of metachert
samples show a top-to-SE sense of shear. There is evidence of both senses of
shear in the Iron Fork area. Metachert samples from the Iron Fork and
structurally above all exhibit a top-to-SE sense of shear, whereas
boudinaged quartz veins in greenschist that is a few tens of meters
structurally higher than Iron Fork exhibit an apparent top-to-NW sense of
shear (Fig. 3b). There are indications that structures showing top-to-NW
sense of shear were formed prior to the top-to-SE ones.
The quartz recrystallized grain shape fabric developed during D3 implies a
top-to-SE sense of shear.
P–T–t path of Pelona SchistPeak metamorphic temperature
Raman spectroscopy on carbonaceous material (RSCM) and the opening angle of
the quartz c axis fabric thermometer were used to quantify the peak
metamorphic temperature and/or to detect any possible inverted thermal
gradient. A series of carbon-rich metagreywacke samples from a 4 km thick
transect of Pelona Schist yield an average peak temperature of
519 ± 20 ∘C (see Table S1 in the Supplement for sample
locations and Table S2 for detailed analysis results). All but a couple of
the measurements lie within uncertainty of ∼ 512 ∘C, and there
is no obvious systematic trend of peak temperature across the East Fork cross
section (Fig. 9). The peak temperatures of one metagreywacke and two
metachert samples given by the quartz c axis fabric opening-angle
thermometer are 554 (PS185), 562 (PS186), and 527 ∘C (PS187).
Considering that this thermometer has a calibration uncertainty of
±50 ∘C, the peak temperatures given by this thermometer are
consistent with those from RSCM.
Peak temperature of Pelona Schist constrained by laser Raman
spectroscopy. See Tables S1 and S2 for sample locations and detailed
results.
Peak metamorphic pressure
Within the isoclinal folds, white mica wrapping albite in metagreywacke
exhibits straight grain shapes parallel to the axial plane (S2) and is
optically strain-free. The lack of large white mica grains in albite
porphyroblasts indicates that the white mica outside albite presumably
crystallized during S2 at or after peak conditions. The phengite barometer
requires a mineral assemblage of phengite, K-feldspar, phlogopite and quartz,
whereas it will give a minimum pressure if phengite only coexists with a (Mg
Fe) silicate (Massonne and Schreyer, 1987). Samples analyzed were all
metagreywacke with a mineral assemblage of phengite + quartz +
stilpnomelane/biotite and do not contain K-feldspar, so the results given by
this barometer are minimum pressures. The average Si content is
3.39 ± 0.03 atoms per formula unit (a.p.f.u.) of white mica, i.e.,
K(Mg, Fe2+)xAl2-x[Si3+xAl1-xO10](OH)2, where
x is the number of Mg and Fe2+ atoms in the octahedral site (see
Tables S1 and S3 for sample locations and analysis results). These results
are consistent with the compositions of old white mica grains reported by
Jacobson (1983a, 1984). Combined with the peak temperature, the peak pressure
is 10.5 ± 0.4 kbar. The result is similar to 10 ± 1 kbar, the
peak pressure of the Pelona Schist in the Sierra Pelona (Graham and Powell,
1984). The peak pressure corresponds to a depth of approximately 39 km with
an assumed upper plate density of 2750 ± 50 kg m-3.
P–T conditions of mylonitization of Pelona Schist
In mylonitized metagreywacke, coarse-grained white mica grains defining S3
are curved and show undulose extinction, whilst fine-grained white micas are
usually free of lattice distortion. Some grains show compositional zonation
in backscattered electron images. Electron probe results indicate that the
cores of the zoned white mica have Si content (∼ 3.35 a.p.f.u.) similar to that of S2 white mica, and the rims of the zoned white
mica have Si content similar to that of fine-grained white mica
(∼ 3.25 a.p.f.u.; see Table S1 for sample locations and Table S3 for analysis results). Quartz in one metagreywacke sample collected close
to the Vincent Fault, which has a mylonitic microstructure as described
previously, reveals a titanium concentration of 1.55 ± 0.17 ppm (see
Table S1 for sample location and Table S4 for analysis results). Combination
of phengite barometry and TitaniQ thermobarometry yields an average P of 5.8 ± 0.8 kbar and T of 390 ± 13 ∘C, which can be
attributed to this stage in the structural development of the schist.
Timing of exhumation of Pelona Schist
Several geochronological studies of Pelona Schist have been reported. The
youngest 206Pb/238U age of the detrital zircon from the Pelona Schist is 68 Ma (Grove et al., 2003), marking the earliest
possible time for the subduction of the protolith of the Pelona Schist.
Three amphibole 40Ar/39Ar isochron ages from the greenschist in
the East Fork area reported are 58.9 ± 2.5, 60.3 ± 2.6 and
73.4 ± 3.0 Ma (Jacobson, 1990). Jacobson (1990) noted that
the sample giving 73.4 ± 3.0 Ma age had an older isochron age than the
plateau age, indicating an unusual case that the initial 40Ar/39Ar
is less radiogenic than atmospheric argon, and this age should be treated
with caution. Given that the youngest detrital zircon 206Pb/238U
age suggests that underthrusting of the Pelona Schist had not started until
68 Ma, the 73.4 Ma amphibole 40Ar/39Ar age is unlikely to be
reliable. The other two samples giving ages of 58.9 ± 2.5 and 60.3 ± 2.6 Ma do not have well-developed plateaus and Jacobson (1990)
concluded that this could be due to excess radiogenic argon. However, the
consistency of the amphibole and white mica 40Ar/39Ar ages
suggests that they are reliable. The amphibole grains in the mafic schist
grew at the beginning of D2 at ∼ 500 ∘C, which is at
or below the closure temperature for the Ar system in Ca amphibole, and they
were not affected by D3, so we interpret the amphibole 40Ar/39Ar
ages as crystallization ages.
Jacobson (1990) reported white mica isochron ages of three metagreywacke
samples of 60.8 ± 0.6, 58.1 ± 0.8 and 55.9 ± 0.2 Ma,
which decrease with structural depth. White micas of two Pelona Schist
samples (SG69 from right beneath the Vincent Fault and SG532A from Prairie
Fork, the structurally deepest level) from the East Fork area were analyzed
by Grove et al. (2003). No excess 40Ar was noticed
except for the initial step of Sample SG532A. The 40Ar/39Ar total
gas ages for SG69 and SG532A are 57.8 ± 0.1 and 34.2 ± 0.2 Ma, respectively. These results fit a trend that the ages decreases with
increasing structural depth. Grove et al. (2003) suggest that the closure
temperature for white mica in the Pelona Schist is ∼ 400 ∘C, which is the same as the temperature during D3 within
uncertainty. We therefore interpret the age of SG69, which was affected by
recrystallization during D3 mylonitization, as a crystallization age, and
the age of SG532A, which contains D2 micas unaffected by D3, as a cooling
age.
To better constrain the time when the Pelona Schist entered the brittle
regime, we performed zircon fission track analysis on six metagreywacke
samples from East Fork in this study. The zircon fission track age is as old
as ∼ 46.9 Ma immediately beneath the Vincent Fault, and
becomes progressively younger toward the north except one outlier
(∼ 13.5 Ma), which may have been affected by nearly Miocene
dikes (see Fig. 2b for sample locations in the cross section, Table S1 for
detailed sample locations and Table S5 for analysis results). The sample
from Prairie Fork has the youngest age of ∼ 23.5 Ma, which is
consistent with the white mica 40Ar/39Ar age of 34.2 ± 0.2 Ma (Grove et al., 2003) from the same area. The results above
imply that the top section of the Pelona Schist cooled through the zircon
fission track closure temperature of ∼ 237 ∘C
around ∼ 46.9 Ma (see Appendix A for the calculation of the
closure temperature), and that the structurally deeper section of the schist
experienced slow cooling over the next 20 million years.
Implications of the pressure, temperature and time estimatesDeformation history
We make the case here that the main foliation (S2) of the Pelona Schist in
the San Gabriel Mountains preserves mineral assemblages and microstructures
of retrograde metamorphism based on the following observations: (1) biotite
rarely appears in the Pelona Schist in the San Gabriel Mountains, whereas
biotite is more common in the Pelona Schist in Sierra Pelona
(Graham and England, 1976; Graham and Powell, 1984);
(2) garnet occurs exclusively as inclusions in albite porphyroblasts, where
it was protected from retrogression; and (3) the compositional zonation of white
mica indicates decompression and hence exhumation during its formation. A
plausible explanation for these observations is that the Pelona Schist in
the San Gabriel Mountains experienced retrograde metamorphism during which
biotite and garnet were broken down to chlorite and white mica.
If S2 was formed during exhumation, it is likely that S1 formed during
subduction, while S3 is related to later reactivation and mylonitization. As
discussed in Sect. 4.2.3, S3 is characterized by crenulation of S2 and
mylonitization, and therefore is inferred to postdate S2. Mylonitization in
the Pelona Schist shows similar stress, pressure and temperature conditions
as the mylonites immediately above the Vincent Fault in the upper plate
(Xia and Platt, 2017), and therefore may have
occurred at the same time.
It should be noted that the subduction zone and the Vincent Fault originally
dipped E and have been rotated and tilted afterwards into their present
orientations. Paleomagnetic study of Neogene volcanic rocks in the San
Gabriel block bounded by the San Gabriel Fault and the San Andreas Fault
yielded a net clockwise rotation of 37.1∘± 12.2∘ since the early Miocene (Terres and Luyendyk,
1985). As a result, the subduction channel was rotated to a SE dip, and the
top level of the subduction channel shows top-to-SE sense of shear. The
Vincent Fault and the Pelona Schist now dip SW in the limb of a faulted
NW-trending antiform that was produced during Neogene motion on the San
Andreas and Punchbowl faults.
(a)P–T–t path and (b) thermal history of
Pelona Schist. RSCM: Raman spectroscopy on carbonaceous material; TitaniQ:
titanium-in-quartz thermobarometer.
The D3 mylonitic fabric only occurs in the top 100 m beneath the Vincent
Fault and shows a top-to-SE sense of shear, so D3 is likely to be related to
the Vincent Fault, and the motion direction of the Vincent Fault was top to
SE. After correcting for the early Miocene vertical-axis rotation
(Terres and Luyendyk, 1985), the sense of shear was initially
top-to-E. We conclude that the Vincent Fault dipped E initially and its
sense of shear was top to E. Therefore, the Vincent Fault had a normal sense
of motion.
P–T–t path of Pelona Schist
With the above data, we are able to establish the metamorphic and
deformational history of the Pelona Schist. The youngest detrital zircon
206Pb/238U age of ∼ 68 Ma (Grove et
al., 2003) places constraints on the youngest age of the source rock of the
Pelona Schist and on the earliest time of the subduction initiation. The
peak metamorphic conditions are 10.5 ± 0.4 kbar and 519 ± 20 ∘C as indicated by the phengite barometer and RSCM. The timing of
peak metamorphism is not known precisely, but it should be no later than
∼ 60.3 Ma as constrained by the amphibole 40Ar/39Ar
age of the mafic greenschist in the Narrows area (Jacobson,
1990). After reaching these conditions, part of the Pelona Schist was
underplated and exhumed, firstly by ductile flow during D2, and then by
normal sense motion on the Vincent Fault, accompanied by formation of the
Narrows synform and associated mylonite (D3). The P–T conditions during
mylonitization were 5.8 ± 0.8 kbar and 390 ± 13 ∘C
given by the phengite barometer and the TitaniQ thermometer. Mylonitization
of the top section of the schist started at ∼ 55 Ma implied by
the white mica 40Ar/39Ar ages (Grove et al.,
2003; Jacobson, 1990) and ceased by ∼ 43 Ma as constrained by
the zircon fission track ages of the schist immediately beneath the Vincent
Fault (Fig. 10).
Exhumation mechanisms of the Pelona Schist
The tectonic setting of the Pelona Schist suggests that it was formed in a
subduction channel rather than an accretionary wedge; therefore, exhumation
models based on the accretionary wedge setting (Brandon et al., 1998; Platt, 1986; Ring et
al., 2007) cannot be applied to the Pelona Schist. The flat-slab Laramide
subduction rules out the slab rollback as a possible exhumation mechanism
(Brun and Faccenna, 2008).
Various models such as upper plate normal faulting (e.g.,
Jacobson et al., 2007), passive-roof thrusting and erosion
(Yin, 2002), return flow (Oyarzabal et al.,
1997), and channelized extrusion (Chapman et al., 2010)
have been proposed to explain the exhumation of the Rand Schist, the schists
of Portal Ridge and Sierra de Salinas, and the Orocopia schist. As for the
Pelona Schist, it was still thought to preserve the original structural and
metamorphic features related to subduction (Ehlig,
1981; Jacobson, 1983a, b, 1997; Jacobson et al., 1996) and no model has
yet been proposed to account for its exhumation.
Normal faulting alone is unlikely to have brought the Pelona Schist from
∼ 39 km depth to the surface, as it cannot produce two
opposing senses of shear in the lower plate. Erosion is not plausible for
exhuming the Pelona Schist, either. The deep-water San Francisquito
Formation in southern California was deposited from latest Cretaceous
through middle Paleocene time (Kooser, 1982), indicating that the
Sierra Pelona area was in a marine environment during the subduction and
first stage exhumation of the Pelona Schist, and erosion was likely to be
minimal. The channelized extrusion model (Chapman et al.,
2010) can bring up the entire subduction assemblage as a whole. We think
this is unlikely to be the case for the Pelona Schist, for two reasons.
First, when S2 of the Pelona Schist, which is defined by the retrograde
metamorphic mineral assemblages, was formed between 60 and 58 Ma, Pacific
Ocean lithosphere was still subducting at a rate of 115 mm yr-1
(Doubrovine and Tarduno, 2008) using a recent plate
reconstruction (Müller et al., 2008), so that
schist exhumation was coeval with ongoing subduction. Return flow allows
the subducted material to be exhumed along roughly the same route as it
descended, and makes the subduction channel a “two-way street”
(Ernst, 1984). The second observation favoring
return flow is that the whole of the exposed Pelona Schist in the East Fork
shows evidence for large non-coaxial strains, which is not consistent with
the channelized extrusion model.
In the classic return flow model, the velocity of rocks in a subduction
channel results from a combination of Couette flow (laminar flow of a
viscous fluid due to viscous drag) driven by the drag of the subducting
plate and Poiseuille flow (laminar flow of a viscous fluid resulted from a
pressure gradient) caused by the buoyancy of the low-density rocks in the
subduction channel (Beaumont
et al., 2009). In this model, the maximum subducting velocity occurs at the
base of the channel, the maximum exhumation velocity occurs in the upper
part of the channel, and the velocity at the roof of the channel is zero.
When crossing the locus of the maximum exhumation velocity plane, the sense
of shear changes because the velocity gradient changes its sign.
The change in sense of shear has been observed in D2. S2 shows a top-to-NW
sense of shear at structurally deeper levels and a top-to-SE sense of shear
at structurally shallower levels. In the return flow model the sense of
shear changes from the bottom of the channel to the top, with the rocks in
the return flow moving relatively northwestward. The locus of maximum
exhumation velocity probably occurs around Iron Fork, where we observe
conflicting senses of shear. In addition, the locus of the maximum
exhumation velocity may have shifted through time, which would explain the
conflicting shear senses in the quartz veins, metacherts, and greenschist in
the Iron Fork area.
Return flow brought the Pelona Schist to a depth of ∼ 22 km,
marking the first stage of exhumation. The overprinting relationship between
S2 and S3 implies that return flow ceased to act after D2. Return flow was
immediately followed by the second-stage exhumation (Fig. 11). The
mylonitized Pelona Schist right beneath the Vincent Fault indicates that the
Vincent Fault exhumed the Pelona Schist to 8 ± 4 km depth as
constrained by the thermobarometric analysis of the upper plate
(Xia and Platt, 2017). Similar extensional
faulting of the upper plate has played an important role during the
exhumation of other Laramide subduction-related schists from middle and
lower crust to the upper crust (e.g., Chapman et al.,
2010).
Tectonic model of the Pelona Schist. (a) The Pelona Schist
was subducted no earlier than 68 Ma and reached the peak condition by
60 Ma. (b) The Pelona Schist exhumed by return flow. Close-up view
shows the velocity profile of the material in the subduction channel.
(c) The Vincent Fault exhumed the Pelona Schist to a shallower
structural level. (d) Present-day setting of the Vincent Fault and
the Pelona Schist after rotation and tilting. The scale in panel (d)
is larger than those in panels (a–c).
Rheological interpretation
Metagreywacke is the dominant rock type in the Pelona Schist, which makes up
around 90 % of the East Fork transect (Jacobson, 1983b). Albite
does not develop any crystal-plastic deformation microstructures, while
sheet silicates were mainly crenulated/kinked during S3, so quartz is likely
to account for most of the strain in metagreywacke. Metachert interlayers in
metagreywacke do not show boudinage or buckle folds without the surrounding
metagreywacke or greenschist being involved, implying that metachert and
metagreywacke were coupled during deformation. This suggests that the values
of strain rate estimated from one type of rock can be applied to another.
Deformation mechanisms evolve through time and vary among the three types of
rocks in the Pelona Schist.
D1
The absence of CPOs (Fig. 4c), presence of differentiated cleavage (Fig. 4a),
and lack of evidence of crystalline plastic deformation indicate that
pressure solution was the dominant deformation mechanism during S1.
Paleopiezometry was not applicable due to the lack of dynamic
recrystallization. All we can say about the stress conditions is that, by the
end of the subduction (transition from D1 to D2), the shear stress was less
than 10 MPa (see Sect. 7.2.1). The peak temperature at this time was 519 ± 20 ∘C given by RSCM.
D2
The broad distributive ductile deformation during D2 indicates that the
Pelona Schist was below the down-dip limit of seismicity during D2, which is
consistent with the absence of pseudotachylite.
Quartz in metagreywacke
In the metagreywacke from the lower limb of the Narrows synform, the strong
differentiation of mica and other sheet silicate minerals into P domains,
concentration of quartz in Q domains and pressure shadows, the precipitation
of quartz in the pull-aparts of albite grains, and the existence of fluid
inclusions in quartz grains indicate that pressure solution was the dominant
deformation mechanism (Fig. 5a–e). The source of the quartz is likely to
have been what are now the mica-rich P domains, and the sinks include the
pressure shadows around the albite porphyroblasts. Dynamic recrystallization
and undulose extinction in quartz imply that dislocation creep probably
occurred immediately after if not contemporaneously with pressure solution,
and also contributed to the D2 deformation (Fig. 5d and f). The average
recrystallized quartz grain size is ∼ 75 µm, indicating
the shear stress during recrystallization was less than 10 MPa (Table 1). It
should be noted that the stress estimates here may only represent the
maximum stress during pressure solution if quartz recrystallized after
pressure solution during exhumation. The peak temperature is constrained to
519 ± 20 ∘C, but retrogression during D2 suggests that
temperature decreased during deformation.
Quartz in metachert
In the lower limb of the Narrows synform, strong crystallographic preferred
orientation of quartz in metachert indicates that dislocation creep was the
dominant deformation mechanism in this rock type, though in some samples the
related microstructures were strongly modified by subsequent annealing, as
quartz shows equant grain shapes, straight grain boundaries and
120∘ triple points (Fig. 6a). From the grain size of recrystallized
quartz or quartz grains preserved in albite porphyroblasts varies between 47
and 82 µm, and the inferred shear stresses range between 8.3 + 2.7/- 1.5 and 12.9 + 3.9/- 2.3 MPa (Table 1).
D3
Crenulation cleavage in the hinge of the Narrows synform indicates that
microfolding of the S2 foliation accompanied by pressure solution was the
dominant deformation mechanism there (Fig. 8a).
In the upper limb of the Narrows synform, quartz in the metagreywacke
underwent dynamic recrystallization. Subgrain rotation was the dominant
recrystallization mechanism but bulging occurred subsequently. This suggests
that quartz was mainly deformed by climb-accommodated dislocation creep. We
estimate the temperature of recrystallization at 390 ± 13 ∘C. The average grain size of
recrystallized quartz grains is 28 ± 9 µm, indicating a shear stress of ∼ 20 MPa (Table 1).
Estimates of shear zone widths and strain rates
The width of the subduction channel is estimated as 10 ± 5 km based on
the thickness of the observed low-velocity anisotropic layer in the middle
to lower crust of the southern California (Lee et al., 2014; Li et al., 1992; Porter
et al., 2011). We assume that the 4 km thick Pelona Schist in the eastern
San Gabriel Mountains formed part of the zone of return flow during D2, and
the remaining 6 ± 5 km thickness of Pelona Schist formed the zone of
subduction flow in the channel.
The subduction of the protolith of the Pelona Schist started no earlier than
68 Ma (Grove et al., 2003) and it reached the maximum depth
by 60 Ma as discussed in Sect. 5, during which the estimated rate of
convergence between Pacific Ocean lithosphere and North America was over 110 mm yr-1 (Doubrovine and Tarduno, 2008). The
source of non-volcanic tremor in subduction zones is near the down-dip limit
of megathrust earthquakes (Ide et al., 2007), and a down-dip
limit of 25 km has been estimated for the present-day Cascadia margin
(Chapman and Melbourne, 2009). The Pelona Schist was at depths of 25–39 km and at temperatures of ∼ 500 ∘C, during at least the earlier stages of D2, so there may have
been very limited discrete slip along the megathrust at this time. Assuming
that subduction was accommodated within the 6 ± 5 km thick down-going
section of the subduction channel, the plate motion rate above corresponds
to a strain rate between 3.2 × 10-13 and 3.5 × 10-12 s-1 in this section of the subduction channel between 25 and
39 km depth. The range of this estimated strain rate was caused by the
uncertainty of the subduction channel width.
During the first stage of exhumation, the Pelona Schist was decompressed from
10.5 ± 0.4 to 5.8 ± 0.8 kbar, which is approximately equivalent
to a vertical exhumation of 17.4 km from 60 to 58 Ma. With a channel
geometry as discussed in Appendix B, the displacement rate in the zone of
return flow was 17 to 70 mm yr-1 and hence the maximum strain rate was
1.4 × 10-13 to 5.5 × 10-13 s-1. The wide
ranges of the displacement rate and the strain rate are primarily due to the
overlapping amphibole and white mica 40Ar/39Ar ages, which
constrain the time period for the first stage of exhumation to between 1 and
4 Myr.
The width of the mylonitized Pelona Schist is ∼ 100 m, and the
lower 200 m of the mylonite zone in the hanging wall of the Vincent Fault
shows similar quartz microstructures and metamorphic conditions as those in
the Pelona Schist. Thermobarometric data (Xia and Platt,
2017) from the hanging-wall mylonites show that the mylonites were
decompressed to 2.0 ± 1.0 kbar by the end of D3. As shown in Appendix A, the mylonites crossed the 300 ∘C isotherm at around 51.1 Ma,
when ductile deformation ceased. Assuming the shear zone originally dipped
∼ 30∘, the estimated average strain rate during
mylonitization was 4.5 ± 1.2 × 10-13 s-1. The
uncertainty of the strain rate results from the uncertainty of the pressure
at the end of D3. This is slightly higher than the range of 1.3 × 10-14 to 1.1 × 10-13 s-1 given by the quartz
dislocation creep flow law (Hirth et al., 2001) for the
estimated shear stress of 20.1 + 7.3/- 4.0 MPa (differential stress of 34.8 + 12.7/- 7.0 MPa), temperature of 390 ∘C, and the maximum water
fugacity at 390 ∘C and 5.8 kbar.
Discussion
There are a number of issues arising from the data of this study, such as
the inverted thermal gradient in the Pelona Schist, the driving forces of
exhuming the Pelona Schist in return flow, the quartz pressure-solution flow
law, and the change in quartz deformation mechanisms.
Inverted thermal gradient of Pelona Schist
The Pelona Schist in the Sierra Pelona is one of the classic examples of an
inverted metamorphic gradient (e.g., Graham and England, 1976). In the Sierra
Pelona the Pelona schist was metamorphosed up to amphibolite facies and shows
an upward transition from high-pressure greenschist facies to amphibolite
facies. The temperature ranges from ∼ 480 to 620–650 ∘C
upward in a ∼ 700 m thick transect, corresponding to an inverted
thermal gradient of 170 to 250 ∘C km-1 (Graham and Powell,
1984). In contrast, the Pelona Schist in the East Fork was metamorphosed to
high-pressure greenschist facies and its peak metamorphic temperature is
lower than that in the Sierra Pelona by about 100 ∘C.
Jacobson (1995) noted the systematic composition variations in the amphibole
of the greenschist and argued that there was inverted metamorphic zonation.
However, no inverted thermal gradient can be observed from our peak
temperature estimates in the 4 km thick transect in the East Fork. It is
possible that no inverted thermal gradient was developed in East Fork area,
though it cannot be excluded that the high-temperature section of the Pelona
Schist in the San Gabriel Mountains was cut off by the Vincent Fault.
Topography- and density-driven return flow model of subduction
channels
To compare the stress and strain rate estimates of the return flow with the
geodynamic model, we used the analytical formulation given by Beaumont et al. (2009).
This one-dimensional formulation assumes that the rocks in the channel have
a linear viscous rheology (appropriate for pressure-solution creep). The
stress and strain rate profiles across the subduction channel can be
calculated for a given channel geometry, viscosity of material in the
channel, and the pressure gradient in the subduction channel.
Behr and Platt (2013) modified this formulation, so as to
relate the viscosity to the maximum exhumation rate.
The buoyancy contrast between the subducting material and the overlying
upper plate has been proposed as the driving force for return flow (e.g., England and Holland,
1979; Ernst et al., 1997). The density of the Pelona Schist, which consists
of 90 % metagreywacke, is ∼ 2700 kg m-3. The overriding
plate, after the removal of its high-density batholithic root during the
Laramide subduction, is mainly made up of felsic batholiths and felsic
plutons with gabbronorite to gabbrodiorite intrusives at the bottom of the
crustal column (Saleeby et al., 2003). The average density of
the overriding plate is assumed as 2750 ± 50 kg m-3. Thus, the
density contrast between the schist and the upper plate could be 50 ± 50 kg m-3 and might not be sufficient to drive return flow.
The topographic gradient between the trench and the arc could act as the
major driving force for return flow, as described in Appendix C. The
pressure gradient due to the topographic difference can be up to 969 Pa m-1, whereas that caused by the density contrast is
only 105 Pa m-1 with a density of 2750 kg m-3 for the upper plate. If D2 lasted
2 Myr, the Poiseuille flow driven by the topographic gradient can be up to
74.5 mm yr-1, whereas that driven by the density contrast is only 8 mm yr-1 (Fig. 12).
We have modified the pressure gradient by adding the effect of topography to
that generated by the density contrast, as shown in Appendices B and C. With
a pressure gradient caused by both topography and density contrast, the
one-dimensional model given by Beaumont et al. (2009)
yields a shear stress along the upper surface of the channel of
∼ 10 MPa, and a strain rate of 4.2 × 10-13 to 1.2 × 10-12 s-1 with the duration of D2
varying between 1 and 4 Myr. These results generally agree with our
estimates for the first stage exhumation of the Pelona Schist.
Velocity profile across the subduction channel filled by linear
viscous material. Subduction velocity is negative and exhumation velocity is
positive. Blue dash-point line: Couette flow caused by the drag of the
subducting plate; red dotted curve: Poiseuille flow driven by the buoyancy of
the material in the subduction channel; green dashed curve: topography
induced Poiseuille flow; thick black solid curve: the bulk flow of rocks in
the subduction channel. Note the sense of shear changes when crossing the
line of maximum velocity. The duration of D2 was 2 Myr for calculation. See
Appendix D for details.
(a) Stress–strain-rate curve and
(b) grain size–strain-rate curve of three pressure-solution flow
laws. The orange zones are the estimated strain rate of the Pelona Schist.
Implications for the pressure-solution flow law of quartz
Pressure solution is a common deformation mechanism in low-grade metamorphic
conditions, but it remains poorly understood and its flow law is loosely
constrained. The three primary models for pressure solution are the
thin-film model (Rutter and Elliott, 1976; Weyl, 1959),
the island channel model (Cox and Paterson,
1991; Raj, 1982), and the stress-corrosion micro-cracking model
(den Brok, 1998; Gratz,
1991). The thin-film model assumes that the transport of dissolved material
occurs by diffusion in a fluid film a few nanometers thick along
grain boundaries, and that the rate of pressure solution is governed by the
grain boundary diffusivity (Rutter and Elliott, 1976;
Weyl, 1959). In the island channel model the fluid fills channels that
surround islands where the grain boundaries are in contact. The rate of
deformation in this model is controlled by the diffusivity of solute in the
fluid-filled channel (Raj, 1982). A more sophisticated model, the
stress-corrosion micro-cracking model, assumes that continuously produced
micro-cracks along grain boundaries due to stress corrosion can cause a
rough grain-surface topography, and material is transported through the
plumbing network consisting of grain boundary film, capillary network of the
micro-cracks along grain boundaries and wide pores between grains
(Gratz, 1991). In this model the rate of deformation is
inversely proportional to the square of the island diameter instead of the
square of the average grain diameter (also see Appendix E).
Deformation mechanism maps for (a) D2 and (b) D3.
In both maps, the red line marks the boundary between the regime of quartz
pressure solution and that of quartz dislocation creep, the black dashed line
is calculated using the recrystallized quartz grain size piezometer (Stipp and
Tullis, 2003), and the orange shaded area is the Pelona Schist in the
grain size–stress space. See text for discussion.
The above three models can be tested by our estimated values of the
rheological parameters from the Pelona Schist. Deformation of metagreywacke
during D2 was dominated by pressure solution. Our strain-rate estimate for
this stage is 1.4 × 10-13 to 5.5 × 10-13 s-1
and the shear stress was less than ∼ 10 MPa. The dynamically
recrystallized grain size of quartz in metagreywacke varies between 30 and 100 µm. Comparison between the above estimates and the
stress-strain rate curves predicted by those three models for pressure
solution shows that strain-rate estimates from the Pelona Schist lie between
the predictions of the stress-corrosion micro-cracking model and the
thin-film model, whereas that forecast by the island channel model is many
orders of magnitude faster (Fig. 13).
Our strain-rate estimates are likely to be somewhat higher than those
predicted by theoretical models for pressure solution in pure quartz, as
shown in Fig. 14a. This may be because of the abundance of white mica in
Pelona Schist. It has long been observed that phyllosilicates can increase
rates of dissolution (Bukovská
et al., 2015; Heald, 1956; Rutter and Elliott, 1976; Wassmann and
Stöckhert, 2013; Weyl, 1959), and differences in electrochemical
potential between unlike minerals, e.g., quartz and mica, have been proposed
as the driving force of dissolution (Greene
et al., 2009; Kristiansen et al., 2011; Meyer et al., 2006). Probably due to
the relatively high content of white mica, metagreywacke in Pelona Schist
yielded a higher strain rate than the prediction of the thin-film model.
This also explains the distinct deformation behavior of metachert from
metagreywacke in Pelona Schist. The lack of mica in metachert resulted in
too low a strain-rate contribution from pressure solution, and therefore
dislocation creep became the dominant deformation mechanism.
An additional factor is that dislocation creep is likely to have contributed
to the bulk strain-rate estimates. Quartz shows undulose extinction and
dynamic recrystallization in metagreywacke samples that also reveal clear
evidence of pressure solution such as differentiated cleavage and
precipitation of quartz in pressure shadows. Coexistence of those two types
microstructures may indicate that the total deformation resulted from both
pressure-solution and dislocation creep.
Transition from pressure-solution to dislocation creep
When the Pelona Schist entered the D3 regime, the dominant deformation
mechanism of metagreywacke changed from pressure solution of quartz to
dislocation creep of quartz (Fig. 14b). Two differences between D2 and D3
are shear stress and temperature. Shear stresses recorded by the Pelona Schist during D2 are less than those in D3, while the temperature was
higher. When dynamic recrystallization is not dominant, so that the average
grain size is not controlled by stress, the strain rate caused by quartz
pressure solution is proportional to the first power of differential stress
(den Brok, 1998), whereas the strain rate
produced by quartz dislocation creep is proportional to the fourth power of
the differential stress (Hirth et al., 2001). That is, quartz
pressure solution is much less sensitive to stress compared to quartz
dislocation creep. During S3, shear stresses increased to a few tens of
megapascals and dramatically expedited dislocation creep. Thus, crystal
plastic deformation became the leading deformation mechanism during S3.
Increased stress also changed the scale of the shear zone. Grain size
reduction due to dynamic recrystallization caused strain localization during
S3. One clue is that the exposed width of the zone with S3 is
∼ 100 m while S2 penetrated the entire 4 km transect along the
East Fork in the San Gabriel Mountains.
Conclusions
New microstructural observations, thermobarometric analysis and
geochronologic data were used to constrain the deformation history and
rheological properties of the Pelona Schist during Late Cretaceous–Paleocene
Laramide subduction. The primary conclusions are as follows:
The Pelona Schist preserves a record of deformational and metamorphic
processes from subduction to exhumation. The Pelona Schist was subducted no
earlier than 68 Ma, and reached the peak P–T condition of 10.5 ± 0.4 kbar and 519 ± 20 ∘C by ∼ 60 Ma, after which
it underwent two stages of exhumation: first by return flow within the
channel and then along the Vincent Fault. By ∼ 43 Ma the
Pelona Schist entered the brittle regime.
The dominant deformation mechanism of the Pelona Schist during subduction
(D1) and the first stage of exhumation (D2) was pressure solution. The
presence of mica expedited quartz dissolution and may have increased the
strain rate relative to the predictions of the currently available quartz
pressure-solution flow laws based on monomineralic samples.
During the second stage of exhumation (D3), the Pelona Schist was deformed
dominantly by dislocation creep. Strain localization occurred due to
grain size reduction caused by dynamic recrystallization.
Our estimate of shear stress at the end of during subduction and first-stage
exhumation is less than 10 MPa, while that during the second-stage exhumation
increased to 20.1 + 7.3/- 4.0 MPa.
The estimated magnitude of shear stress and strain rate in the subduction
channel agrees well with the flow-channel model for linear viscous creep.
Our estimated stresses and strain rates during the second stage of
exhumation are close to those estimated from the quartz dislocation creep
flow law of Hirth et al. (2001).
Locations of samples and field images are listed in Table
S1, RSCM results in Table S2, white mica compositions in Table S3, TitaniQ
analysis results in Table S4, and detrital zircon fission track results in
Table S5.
Closure temperature of zircon fission track analysis and cooling
rate
The relation between the closure temperature of zircon fission track and the
cooling rate can be written as
T˙=-RTc2E50%BeE50%RTc,
where T˙ is the cooling rate, Tc the closure temperature,
R the gas constant, E50% the activation energy for 50 %
annealing, and B a constant (Dodson, 1979). For α-damaged zircon,
Brandon et al. (1998) estimated 49.77 kcal mol-1 for E50%
and 3.160 × 10-22 Myr for B based on natural α-damaged zircon from Zaun and Wagner (1985) and Tagami et al. (1990).
The Pelona Schist immediately underneath the Vincent Fault cooled from the
closure temperature of white mica (400 ∘C) to the closure
temperature of zircon fission track Tc from 57.8 Ma
(Grove et al., 2003) to 46.9 Ma (this study), and the average
cooling rate can be written as
T˙=400-Tc57.8-46.9
Equations (A1) and (A2) yields that the average cooling rate was
15.0 ∘C Myr-1 between 57.8 and 46.9 Ma, and the closure
temperature for zircon fission track is ∼ 237 ∘C. By assuming
a linear cooling history between 57.8 and 46.9 Ma, the Pelona Schist was
cooled to 300 ∘C at 51.1 Ma, i.e., the upper section of the Pelona
Schist crossed the brittle-ductile transition zone at 51.1 Ma.
Dip of the Laramide subduction channel
We made the following assumptions about the dip of the subduction channel
when the Pelona Schist was subducted:
The migration of the trench was negligible from Late Cretaceous, when the
magmatic arc was active to the time when the Pelona Schist was subducted.
This allows us to estimate the trench–arc distance.
Following the previous assumption, we assume that the Mesozoic trench–arc
distance was approximately 250 km by comparison with the modern trench–arc
gap of the northwestern United States and that the magmatic arc was about
100 km wide based on the present-day outcrop of the Sierra Nevada.
The depth of the trench was ∼ 4.5 km when the Pelona Schist
was subducted. The Laramide shallow subduction was inferred to be result of
the subduction of the Hess and Shatsky conjugate oceanic plateaux (Liu et
al., 2010), while the former, which originated from the mid-ocean ridge 110
to 100 Myr before present (Vallier et al., 1983) and intersected the North
American plate about 70 Myr before present (Liu et al., 2010), is spatially
and temporally related to the formation of the Pelona Schist. Therefore, the
Pacific Ocean lithosphere adjacent to the western edge of North America was
∼ 40 Myr old when the Hess Plateau was subducted. The bottom of a
typical seafloor aged 40 Myr is ∼ 4.5 km beneath the sea level as
constrained by present-day observations and predicted by theoretical models
(Johnson and Carlson, 1992; Turcotte and Schubert, 2002). The above depth is
likely to increase by another ∼ 3 km along the trench due to buckling
and flexure, resulting in a seafloor depth of ∼ 7.5 km.
The Hess Rise in the central North Pacific Ocean is currently about 3 km
above the surrounding seafloor (Vallier et al., 1983), and if this holds true
for the Hess conjugate, the overall trench depth would be ∼ 4.5 km.
The depth of the subduction channel was 30 km underneath the western
margin of the Mesozoic arc. This allows the Pelona Schist to remain beneath
the Mesozoic magmatic arc after the first-stage exhumation, which is
constrained by the fact that the Pelona Schist and other Laramide
subduction-related schists crop out within the Mesozoic arc at present day.
This assumption gives a subduction channel dip of ∼ 3.3∘ between the trench and the
arc.
The depth of the subduction channel was more than 45 km at the east
margin of the Mesozoic arc. This fits the hypothesized ∼ 45–60 km thick crust of the Nevadaplano, which lasted well into Paleogene time
(Ernst, 2009; Wernicke et al., 1996) and requires
a minimum dip of 12.4∘ for the subduction channel underneath the
arc.
Topography-induced pressure gradient
Assume a monotonically increasing topography at constant rate over a
transect width of l (Fig. C1). The amount of increased elevation over the
interested width l is h. The slope of this transect is ω. Therefore,
one has
h=l×tanω.
A subduction zone dipping against the topographic slope lies underneath the
slope. The dip of the subduction zone is γ. The distance between a
point at the bottom of the overriding plate and the left end of the
subduction zone is x. The relationship among l, x and γ is
l=x×cosγ.
Combine Eqs. (C1) and (C2) and one has
h=x×cosγ×tanω.
The pressure P caused by the topography of the upper plate is ρgh, where
ρ is the density of the upper plate and g is the gravitational
acceleration. P can be written as
P=ρgh=ρgx×cosγ×tanω.P increases with x and works as the driving force of the exhumation of the
rocks within the channel. The gradient of P along the subduction channel is
∂P∂x=ρg×cosγ×tanω=ρghl×cosγ.
The paleotopography of southern California in Late Cretaceous–Paleocene time
could not be readily reconstructed because of the displacements associated
with the San Andreas Fault and other Cenozoic faults. However, the
paleo-geomorphology of the southern end of the Great Valley and Sierra
Nevada can shed light on the possible contemporaneous topography of southern
California.
Geometry of topography and subduction channel.
The Late Cretaceous depositional environment in the western foothills of the
southern Sierra Nevada was fluvial-deltaic (Cherven, 1983). The
elevation of the western foothills of the arc is unlikely to have been
higher than 0.5 km. To the east of the Mesozoic magmatic arc was the
“Nevadaplano”, a high and broad plain lasting from Late Cretaceous to
Eocene (DeCelles, 2004). Its
estimated average elevation was more than 3 km (DeCelles, 2004). The Sierra
Nevada magmatic arc itself, which formed the west flank of the Nevadaplano
(Henry, 2009), could have had an elevation up to 4 km at
approximately 70 Ma (House et al., 1998,
2001). Along the west slope of Sierra Nevada at that time, the elevation may
have changed from 1 km in the west to more than 4 km in the east within a
horizontal distance of ∼ 80 km (House et
al., 2001).
The above geometric setting and an assumed rock density of 2700 kg m-3
yield a pressure gradient of 969 Pa m-1. In contrast, the density
difference between the upper plate and the schist, which is assumed as
50 kg m-3, would generate a pressure gradient of 105 Pa m-1
with the same geometric setting.
Calculations of channel flow stresses and strain rates
In a parallel-sided subduction channel with a width of L, the velocity
profile in a linear viscous fluid is
v(z)=-∇PzL-z22η+V1-zL,
where z is the distance to the base of the channel, ∇P the
pressure gradient, η the viscosity of material in the channel, and V
the subducting velocity (Beaumont et al.,
2009; Behr and Platt, 2013).
Parameters used for subduction channel calculations.
ParameterDescriptionValueReference and notesLSubduction channel width (km)10 ± 5Lee et al. (2014), Porter et al. (2011)VSubduction rate (mm yr-1)115Doubrovine and Tarduno (2008)d2Displacement during D2 (km)69.8Exhumed from 39 km deep to 22 km deep along a subduction channel dipping 12.4∘ as estimated in Appendix BtDuration of D2 (Myr)1–4Sect. 5.4ρup-ρscDensity contrast between the overriding plate and the subduction channel (kg m-3)50For the calculations, 2750 kg m-3 is used for the density of the upper plate and 2700 kg m-3 for the density of the Pelona Schist.
The pressure gradient is caused by (1) the density contrast between the
rocks in the subduction channel and those in the overriding plate and (2) the
topography gradient as illustrated in Appendix B. The pressure gradient
caused by density contrast is
∇Pdensity=ρup-ρscgsinγ,
where ρup is the density of upper
plate, ρsc the density of rocks in
the subduction channel, and g the gravity acceleration (Beaumont et al.,
2009; Behr and Platt, 2013). The sum of Eqs. (D2) and (C5) is the total
pressure gradient:
∇P=ρup-ρscgsinγ+ρghl×cosγ.
The viscosity η is the lesser of the roots of Eq. (D4):
4V2L2∇Pη2-4V-2veη+L2∇P=0,
where ve is the maximum exhumation rate (Behr and Platt,
2013). ve can be approximated as the ratio of d2, displacement
during D2 to t, the duration of D2.
The strain rate profile is
v′(z)=-∇PL-2z2η-VL.
The stress profile (Behr and Platt, 2013) is
σ(z)=v′(z)×η=-∇P(L-2z)2-VηL.
Parameters used for the stress and strain rate calculations in the
subduction channel are listed in Table D1.
Parameters used for flow law calculations
Parameters used for pressure-solution calculations.
ParameterDescriptionValueReference and notesAGrain shape constant44den Brok (1998); 44 for spheric grainsVmSolid molar volume (m3 mol-1)2.269 × 10-5Berman (1988)cSolubility of solid in fluid phaseP and T dependentFournier and Potter (1982)DgbGrain boundary diffusivityT dependentFarver and Yund (2000)DfluidDiffusivity in grain boundary fluidT dependentWatson and Wark (1997)dchanWidth of island channel (µm)0.1den Brok (1998)dislIsland diameter (µm)0.5Paterson (1995)wEffective width of grain boundary (µm)0.1Joesten (1983)ρfDensity of fluid (kg m-3)923Behr and Platt (2013)ρsDensity of solid (kg m-3)2650Behr and Platt (2013)RGas constant (J K-1 mol-1))8.314dGrain sizeVaries by sample
As summarized by den Brok (1998) and Behr and Platt (2013), quartz
pressure-solution flow laws can be written as
ε˙=AVmcDgbwσρfRTd3ρs,ε˙=AVmcDfluiddchanσρfRTd3ρs, andε˙=AVmcDgbwσρfRTdisl2dρs
for the thin-film model, the island channel model, and the micro-cracking
model, respectively, where σ is differential stress in Pa and T the
absolute temperature. The descriptions and values used for the parameters in
Eqs. (E1) to (E3) are listed in Table E1.
Quartz dislocation creep flow law (Hirth et al., 2001) is
ε˙=AfH2Oσ4exp-QRT,
where A is a material parameter, fH2O water
fugacity, σ differential stress in megapascals, Q the activation energy, R the ideal gas constant, and T
the absolute temperature. As constrained by Hirth et al. (2001), A equals
to 10-11.2±0.6 MPa-4 s-1 and Q has a value of 135±15 kJ mol-1. Maximum water fugacity was assumed at given pressure and
temperature and was computed from the fugacity calculator
(https://www.esci.umn.edu/people/researchers/withe012/fugacity.htm;
Pitzer and Sterner, 1994).
The Supplement related to this article is available online at doi:10.5194/se-8-379-2017-supplement.
The authors declare that they have no conflict of
interest.
Acknowledgements
This research was funded in part by NSF grant EAR-1250128 to John P. Platt.
George Rossman, Richard Hervig and Lynda Williams, Raymond Donelick and
Margaret Donelick, Rosario Esposito, and Elizabeth Bell are thanked for
their help with Raman spectroscopy, ion probe, fission track, electron
probe, and cathodoluminescence analysis, respectively. We thank Rüdiger Kilian for handling the manuscript and his valuable comments, and appreciate
the constructive reviews from Ruth Keppler and Uwe Ring. Haoran Xia is grateful to
Jason Williams, Wenrong Cao and Haiming Tang for their help in the field.
Edited by: R. Kilian
Reviewed by: R. Keppler and U. Ring
ReferencesBarth, A. P. and Schneiderman, J. S.: A comparison of structures in the
Andean orogen of northern Chile and exhumed midcrustal structures in southern
California, USA: An analogy in tectonic style?, Int. Geol. Rev., 38,
1075–1085, 10.1080/00206819709465383, 1996.Beaumont, C., Jamieson, R. A., Butler, J. P., and Warren, C. J.: Crustal
structure: A key constraint on the mechanism of ultra-high-pressure rock
exhumation, Earth Planet. Sc. Lett., 287, 116–129,
10.1016/j.epsl.2009.08.001, 2009.Behr, W. M. and Platt, J. P.: Rheological evolution of a Mediterranean
subduction complex, J. Struct. Geol., 54, 136–155,
10.1016/j.jsg.2013.07.012, 2013.Behr, W. M., Thomas, J. B., and Hervig, R. L.: Calibrating Ti concentrations
in quartz for SIMS determinations using NIST silicate glasses and application
to the TitaniQ geothermobarometer, Am. Mineral., 96, 1100–1106,
10.2138/am.2011.3702, 2011.Berman, R. G.: Internally-consistent thermodynamic data for minerals in the
system
Na2O-K2O-CaO-MgO-FeO-Fe2O3-Al2O3-SiO2-TiO2-H2O-CO2,
J. Petrol., 29, 445–522, 10.1093/petrology/29.2.445, 1988.Beyssac, O., Goffé, B., Chopin, C., and Rouzaud, J. N.: Raman spectra of
carbonaceous material in metasediments: a new geothermometer, J. Metamorph.
Geol., 20, 859–871, 10.1046/j.1525-1314.2002.00408.x, 2002.Brandon, M. T., Roden-Tice, M. K., and Garver, J. I.: Late Cenozoic
exhumation of the Cascadia accretionary wedge in the Olympic Mountains,
northwest Washington State, Geol. Soc. Am. Bull., 110, 985–1009,
10.1130/0016-7606(1998)110<0985:LCEOTC>2.3.CO;2, 1998.Brun, J.-P. and Faccenna, C.: Exhumation of high-pressure rocks driven by
slab rollback, Earth Planet. Sc. Lett., 272, 1–7,
10.1016/j.epsl.2008.02.038, 2008.Bukovská, Z., Wirth, R., and Morales, L. F. G.: Pressure solution in
rocks: focused ion beam/transmission electron microscopy study on orthogneiss
from South Armorican Shear Zone, France, Contrib. Mineral. Petr., 170, 31,
10.1007/s00410-015-1186-8, 2015.
Burchfiel, B. C. and Davis, G. A.: Mojave Desert and environs, in: The
geotectonic development of California, edited by: Ernst, W. G., Prentice
Hall, Englewood Cliffs, New Jersey, USA, 217–252, 1981.Chapman, A. D., Kidder, S., Saleeby, J. B., and Ducea, M. N.: Role of
extrusion of the Rand and Sierra de Salinas schists in Late Cretaceous
extension and rotation of the southern Sierra Nevada and vicinity, Tectonics,
29, 1–21, 10.1029/2009TC002597, 2010.Chapman, J. S. and Melbourne, T. I.: Future Cascadia megathrust rupture
delineated by episodic tremor and slip, Geophys. Res. Lett., 36, L22301,
10.1029/2009GL040465, 2009.Cherven, V. B.: A delta-slope-submarine fan model for Maestrichtian part of
Great Valley Sequence, Sacramento and San Joaquin basins, California, AAPG
Bull., 67, 772–816, 10.1306/03B5B6AA-16D1-11D7-8645000102C1865D, 1983.Cloos, M. and Shreve, R. L.: Subduction-channel model of prism accretion,
melange formation, sediment subduction, and subduction erosion at convergent
plate margins: 1. Background and description, Pure Appl. Geophys., 128,
455–500, 10.1007/BF00874548, 1988.Coney, P. J. and Reynolds, S. J.: Cordilleran Benioff Zones, Nature, 270,
403–406, 10.1038/275464a0, 1977.Cox, S. F. and Paterson, M. S.: Experimental dissolution-precipitation creep
in quartz aggregates at high temperatures, Geophys. Res. Lett., 18,
1401–1404, 10.1029/91GL01802, 1991.
Crowell, J. C.: An outline of the tectonic history of southeastern
California, in: The geotectonic development of California, edited by: Ernst,
W. G., Prentice Hall, Englewood Cliffs, New Jersey, USA, 583–600, 1981.DeCelles, P. G.: Late Jurassic to Eocene evolution of the Cordilleran thrust
belt and foreland basin system, western USA, Am. J. Sci., 304, 105–168,
10.2475/ajs.304.2.105, 2004.den Brok, S. W. J. (Bas): Effect of microcracking on pressure-solution strain
rate: The Gratz grain-boundary model, Geology, 26, 915–918,
10.1130/0091-7613(1998)026<0915:EOMOPS>2.3.CO;2, 1998.
Dibblee, T. W. J.: Areal geology of the western Mojave Desert California,
Geol. Surv. Prof. Paper, 522, 1–153, 1967.
Dibblee, T. W. J.: Geologic map of the Crystal Lake quadrangle, Dibblee
Geological Foundation Map, Santa Barbara, California, DF-87, 2002a.
Dibblee, T. W. J.: Geologic map of the Glendora quadrangle, Dibblee
Geological Foundation Map, Santa Barbara, California, DF-89, 2002b.
Dibblee, T. W. J.: Geologic map of the Mount Baldy quadrangle, Dibblee
Geological Foundation Map, Santa Barbara, California, DF-90, 2002c.
Dibblee, T. W. J.: Geologic Map of the Mount San Antonio quadrangle, Dibblee
Geological Foundation Map, Santa Barbara, California, DF-88, 2002d.
Dickinson, W. R. and Snyder, W. S.: Plate tectonics of the Laramide orogeny,
Geol. Soc. Am. Mem., 151, 355–366, 1978.
Dodson, M. H.: Theory of Cooling Ages, in: Lectures in Isotope Geology,
edited by: Jäger, E. and Hunziker, J. C., Springer-Verlag Berlin
Heidelberg, Berlin, Germany, 194–202, 1979.Doubrovine, P. V. and Tarduno, J. A.: A revised kinematic model for the
relative motion between Pacific oceanic plates and North America since the
Late Cretaceous, J. Geophys. Res., 113, B12101, 10.1029/2008JB005585,
2008.
Ehlig, P. L.: Origin and tectonic history of the basement terrane of the San
Gabriel Mountains, Central Transverse Ranges, in: The geotectonic development
of California, edited by: Ernst, W. G., Prentice Hall, Englewood Cliffs, New
Jersey, USA, 253–283, 1981.England, P. C. and Holland, T. J. B.: Archimedes and the Tauern eclogites:
the role of buoyancy in the preservation of exotic eclogite blocks, Earth
Planet. Sc. Lett., 44, 287–294, 10.1016/0012-821X(79)90177-8, 1979.Ernst, W. G.: Californian blueschists, subduction, and the significance of
tectonostratigraphic terranes, Geology, 12, 436–440,
10.1130/0091-7613(1984)12<436:CBSATS>2.0.CO;2, 1984.Ernst, W. G.: Rise and fall of the Nevadaplano, Int. Geol. Rev., 51,
583–588, 10.1080/00206810903063315, 2009.Ernst, W. G., Maruyama, S., and Wallis, S.: Buoyancy-driven, rapid exhumation
of ultrahigh-pressure metamorphosed continental crust, P. Natl. Acad. Sci.
USA, 94, 9532–9537, 10.1073/pnas.94.18.9532, 1997.Farver, J. and Yund, R.: Silicon diffusion in a natural quartz aggregate:
Constraints on solution-transfer diffusion creep, Tectonophysics, 325,
193–205, 10.1016/S0040-1951(00)00121-9, 2000.Fournier, R. O. and Potter, R. W.: An equation correlating the solubility of
quartz in water from 25∘ to 900 ∘C at pressures up to
10 000 bars, Geochim. Cosmochim. Ac., 46, 1969–1973,
10.1016/0016-7037(82)90135-1, 1982.Gerya, T. V. and Stockhert, B.: Exhumation rates of high pressure metamorphic
rocks in subduction channels: The effect of Rheology, Geophys. Res. Lett.,
29, 1261, 10.1029/2001GL014307, 2002.Graham, C. M. and England, P. C.: Thermal regimes and regional metamorphism
in the vicinity of overthrust faults: an example of shear heating and
inverted metamorphic zonation from southern California, Earth Planet. Sc.
Lett., 31, 142–152, 10.1016/0012-821X(76)90105-9, 1976.Graham, C. M. and Powell, R.: A garnet-hornblende geothermometer:
calibration, testing, and application to the Pelona Schist, Southern
California, J. Metamorph. Geol., 2, 13–31,
10.1111/j.1525-1314.1984.tb00282.x, 1984.Gratz, A. J.: Solution-transfer compaction of quartzites: progress toward a
rate law, Geology, 19, 901–904,
10.1130/0091-7613(1991)019<0901:STCOQP>2.3.CO;2,
1991.Greene, G. W., Kristiansen, K., Meyer, E. E., Boles, J. R., and
Israelachvili, J. N.: Role of electrochemical reactions in pressure solution,
Geochim. Cosmochim. Ac., 73, 2862–2874, 10.1016/j.gca.2009.02.012, 2009.Grigull, S., Krohe, A., Moos, C., Wassmann, S., and Stöckhert, B.:
“Order from chaos”: A field-based estimate on bulk rheology of tectonic
mélanges formed in subduction zones, Tectonophysics, 568–569, 86–101,
10.1016/j.tecto.2011.11.004, 2012.Grove, M., Jacobson, C. E., Barth, A. P., and Vucic, A.: Temporal and spatial
trends of Late Cretaceous – early Tertiary underplating of Pelona and
related schist beneath southern California and southwestern Arizona, Geol. S.
Am. S., 374, 381–406, 10.1130/0-8137-2374-4.381, 2003.
Hamilton, W.: Tectonic setting and variations with depth of some Cretaceous
and Cenozoic structural and magmatic systems of the western United States,
in: Metamorphism and crustal evolution of the western United States, edited
by: Ernst, W. G., Prentice Hall, Englewood Cliffs, New Jersey, USA, 1–40,
1988.Heald, M. T.: Cementation of Simpson and St. Peter Sandstones in Parts of
Oklahoma, Arkansas, and Missouri, J. Geol., 64, 16–30, 10.1086/626314,
1956.Henry, C. D.: Uplift of the Sierra Nevada, California, Geology, 37, 575–576,
10.1130/focus062009.1, 2009.Hirth, G., Teyssier, C., and Dunlap, J.: An evaluation of quartzite flow laws
based on comparisons between experimentally and naturally deformed rocks,
Int. J. Earth Sci., 90, 77–87, 10.1007/s005310000152, 2001.Holyoke, C. W. and Kronenberg, A. K.: Accurate differential stress
measurement using the molten salt cell and solid salt assemblies in the
Griggs apparatus with applications to strength, piezometers and rheology,
Tectonophysics, 494, 17–31, 10.1016/j.tecto.2010.08.001, 2010.House, M. A., Wernicke, B. P., and Farley, K. A.: Dating topography of the
Sierra Nevada, California, using apatite (U–Th)/He ages, Nature, 396,
66–69, 10.1038/23926, 1998.House, M. A., Wernicke, B. P., and Farley, K. A.: Paleo-geomorphology of the
Sierra Nevada, California, from (U-Th)/He ages in apatite, Am. J. Sci., 301,
77–102, 10.2475/ajs.301.2.77, 2001.Houston, H.: Low friction and fault weakening revealed by rising sensitivity
of tremor to tidal stress, Nat. Geosci., 8, 409–415, 10.1038/ngeo2419,
2015.Huang, R. and Audétat, A.: The titanium-in-quartz (TitaniQ)
thermobarometer: A critical examination and re-calibration, Geochim.
Cosmochim. Ac., 84, 75–89, 10.1016/j.gca.2012.01.009, 2012.Ide, S., Shelly, D. R., and Beroza, G. C.: Mechanism of deep low frequency
earthquakes: Further evidence that deep non-volcanic tremor is generated by
shear slip on the plate interface, Geophys. Res. Lett., 34, L03308,
10.1029/2006GL028890, 2007.
Jacobson, C.: Relationship of deformation and metamorphism of the Pelona
Schist to movement on the Vincent thrust, San Gabriel Mountains, southern
California, Am. J. Sci., 283, 587–604, 1983a.
Jacobson, C.: Structural geology of the Pelona Schist and Vincent thrust, San
Gabriel Mountains, California, Geol. Soc. Am. Bull., 94, 753–767, 1983b.
Jacobson, C.: Qualitative thermobarometry of inverted metamorphism in the
Pelona and Rand Schists, southern California, using calciferous amphibole in
mafic schist, J. Metamorph. Geol., 13, 79–92, 1995.
Jacobson, C.: Metamorphic convergence of the upper and lower plates of the
Vincent thrust, San Gabriel Mountains, southern California, USA, J.
Metamorph. Geol., 15, 155–165, 1997.
Jacobson, C., Oyarzabal, F., and Haxel, G.: Subduction and exhumation of the
Pelona-Orocopia-Rand schists, southern California, Geology, 24, 547–550,
1996.Jacobson, C. E.: Petrological evidence for the development of refolded folds
during a single deformational event, J. Struct. Geol., 6, 563–570,
10.1016/0191-8141(84)90065-8, 1984.Jacobson, C. E.: The 40Ar/39Ar geochronology of the Pelona schist
and related rocks, southern California, J. Geophys. Res., 95, 509–528,
10.1029/JB095iB01p00509, 1990.Jacobson, C. E. and Dawson, M. R.: Structural and metamorphic evolution of
the Orocopia Schist and related rocks, southern California: Evidence for late
movement on the Orocopia fault, Tectonics, 14, 933–944,
10.1029/95TC01446, 1995.Jacobson, C. E., Grove, M., Vućić, A., Pedrick, J. N., and Ebert, K.
A.: Exhumation of the Orocopia Schist and associated rocks of southeastern
California: Relative roles of erosion, synsubduction tectonic denudation, and
middle Cenozoic extension, Geol. S. Am. S., 419, 1–37,
10.1130/978-0-8137-2419-5, 2007.
Jennings, C. W.: Geologic map of California, California Division of Mines and
Geology, Sacramento, California, 1977.
Joesten, R.: Grain growth and grain-boundary diffusion in quartz from the
Christmas Mountains (Texas) contact aureole, Am. J. Sci., 283, 233–254,
1983.Johnson, H. P. and Carlson, R. L.: Variation of sea floor depth with age: A
test of models based on drilling results, Geophys. Res. Lett., 19,
1971–1974, 10.1029/92GL01946, 1992.Kidder, S. and Ducea, M. N.: High temperatures and inverted metamorphism in
the schist of Sierra de Salinas, California, Earth Planet. Sc. Lett., 241,
422–437, 10.1016/j.epsl.2005.11.037, 2006.
Kooser, M.: Stratigraphy and Sedimentology of the type San Francisquito
Formation, southern California, Pacific Section SEPM, 22, 53–61, 1982.Kristiansen, K., Valtiner, M., Greene, G. W., Boles, J. R., and
Israelachvili, J. N.: Pressure solution – The importance of the
electrochemical surface potentials, Geochim. Cosmochim. Ac., 75, 6882–6892,
10.1016/j.gca.2011.09.019, 2011.Kruhl, J. H.: Reply: Prism- and basal-plane parallel subgrain boundaries in
quartz: a microstructural geothermobaromether, J. Metamorph. Geol., 16,
142–146, 10.1046/j.1525-1314.1996.00413.x, 1998.Law, R. D.: Deformation thermometry based on quartz c axis fabrics and
recrystallization microstructures: A review, J. Struct. Geol., 66, 129–161,
10.1016/j.jsg.2014.05.023, 2014.Lee, E., Chen, P., Jordan, T. H., Maechling, P. B., Denolle, M. A. M., and
Beroza, G. C.: Full-3-D tomography for crustal structure in Southern
California based on the scattering-integral and the adjoint-wavefield
methods, J. Geophys. Res.-Sol. Ea., 119, 6421–6451,
10.1002/2014JB011346, 2014.Li, Y.-G., Henyey, T. L., and Silver, L. T.: Aspects of the crustal structure
of the western Mojave Desert, California, from seismic reflection and gravity
data, J. Geophys. Res., 97, 8805–8816, 10.1029/91JB02119, 1992.Liu, L., Gurnis, M., Seton, M., Saleeby, J., Müller, R. D., and Jackson,
J. M.: The role of oceanic plateau subduction in the Laramide orogeny, Nat.
Geosci., 3, 353–357, 10.1038/ngeo829, 2010.Massonne, H.-J. and Schreyer, W.: Phengite geobarometry based on the limiting
assemblage with K-feldspar, phlogopite, and quartz, Contrib. Mineral. Petr.,
96, 212–224, 10.1007/BF00375235, 1987.Meyer, E. E., Greene, G. W., Alcantar, N. A., Israelachvili, J. N., and
Boles, J. R.: Experimental investigation of the dissolution of quartz by a
muscovite mica surface: Implications for pressure solution, J. Geophys.
Res.-Sol. Ea., 111, 2–5, 10.1029/2005JB004010, 2006.
Miller, F. and Morton, D.: Comparison of granitic intrusions in the Pelona
and Orocopia Schists, southern California, J. Res. US Geol. Surv., 5,
643–649, 1977.Moore, T. E., O'Sullivan, P. B., Potter, C. J., and Donelick, R. A.:
Provenance and detrital zircon geochronologic evolution of lower Brookian
foreland basin deposits of the western Brooks Range, Alaska, and implications
for early Brookian tectonism, Geosphere, 11, 93–122, 10.1130/GES01043.1,
2015.Müller, R. D., Sdrolias, M., Gaina, C., and Roest, W. R.: Age, spreading
rates, and spreading asymmetry of the world's ocean crust, Geochem. Geophy.
Geosy., 9, Q04006, 10.1029/2007GC001743, 2008.Oyarzabal, F. R., Jacobson, C. E., and Haxel, G. B.: Extensional reactivation
of the Chocolate Mountains subduction thrust in the Gavilan Hills of
southeastern California, Tectonics, 16, 650–661, 10.1029/97TC01415,
1997.Paterson, M. S.: A theory for granular flow accommodated by material transfer
via an intergranular fluid, Tectonophysics, 245, 135–151,
10.1016/0040-1951(94)00231-W, 1995.Pitzer, K. S. and Sterner, S. M.: Equations of state valid continuously from
zero to extreme pressures for H2O and CO2, J. Chem. Phys., 101,
3111–3116, 10.1063/1.467624, 1994.Platt, J. P.: Dynamics of orogenic wedges and the uplift of high-pressure
metamorphic rocks, Geol. Soc. Am. Bull., 97, 1037–1053,
10.1130/0016-7606(1986)97<1037:DOOWAT>2.0.CO;2, 1986.Porter, R., Zandt, G., and McQuarrie, N.: Pervasive lower-crustal seismic
anisotropy in Southern California: Evidence for underplated schists and
active tectonics, Lithosphere, 3, 201–220, 10.1130/L126.1, 2011.Raj, R.: Creep in polycrystalline aggregates by matter transport through a
liquid phase, J. Geophys. Res., 87, 4731–4739, 10.1029/JB087iB06p04731,
1982.Ring, U., Will, T., Glodny, J., Kumerics, C., Gessner, K., Thomson, S.,
Güngör, T., Monié, P., Okrusch, M., and Drüppel, K.: Early
exhumation of high-pressure rocks in extrusion wedges: Cycladic blueschist
unit in the eastern Aegean, Greece, and Turkey, Tectonics, 26, TC2001,
10.1029/2005TC001872, 2007.
Ruff, L. J. and Tichelaar, B. W.: What Controls the Seismogenic Plate
Interface in Subduction Zones?, in: Geophysical Monograph: American
Geophysical Union, edited by: Bebout, G. E., Scholl, D. W., Kirby, S. H., and
Platt, J. P., vol. 96, 105–111, American Geophysical Union, Washington,
D.C., USA, 1996.Rutter, E. H. and Elliott, D.: The kinetics of rock deformation by pressure
solution [and Discussion], Philos. T. Roy. Soc. A, 283, 203–219,
10.1098/rsta.1976.0079, 1976.Saleeby, J., Ducea, M., and Clemens-Knott, D.: Production and loss of
high-density batholithic root, southern Sierra Nevada, California, Tectonics,
22, 1064, 10.1029/2002TC001374, 2003.
Shelley, D.: Igneous and metamorphic rocks under the microscope, Chapman and
Hall, London, UK, 1993.Stipp, M. and Tullis, J.: The recrystallized grain size piezometer for
quartz, Geophys. Res. Lett., 30, 2088, 10.1029/2003GL018444, 2003.Stipp, M., Tullis, J., and Behrens, H.: Effect of water on the dislocation
creep microstructure and flow stress of quartz and implications for the
recrystallized grain size piezometer, J. Geophys. Res., 111, B04201,
10.1029/2005JB003852, 2006.Stipp, M., Tullis, J., Scherwath, M., and Behrmann, J. H.: A new perspective
on paleopiezometry: Dynamically recrystallized grain size distributions
indicate mechanism changes, Geology, 38, 759–762, 10.1130/G31162.1,
2010.Stöckhert, B.: Stress and deformation in subduction zones: insight from
the record of exhumed metamorphic rocks, Geol. Soc. Lond. S. P., 200,
255–274, 10.1144/GSL.SP.2001.200.01.15, 2002.Tagami, T., Ito, H., and Nishimura, S.: Thermal annealing characteristics of
spontaneous fission tracks in zircon, Chem. Geol., 80, 159–169,
10.1016/0168-9622(90)90024-7, 1990.Terres, R. R. and Luyendyk, B. P.: Neogene tectonic rotation of the San
Gabriel Region, California, suggested by paleomagnetic vectors, J. Geophys.
Res., 90, 12467–12484, 10.1029/JB090iB14p12467, 1985.Thomas, J. B., Watson, E. B., Spear, F. S., Shemella, P. T., Nayak, S. K.,
and Lanzirotti, A.: TitaniQ under pressure: the effect of pressure and
temperature on the solubility of Ti in quartz, Contrib. Mineral. Petr., 160,
743–759, 10.1007/s00410-010-0505-3, 2010.Thomas, J. B., Watson, E. B., Spear, F. S., and Wark, D. A.: TitaniQ
recrystallized: experimental confirmation of the original Ti-in-quartz
calibrations, Contrib. Mineral. Petr., 169, 27,
10.1007/s00410-015-1120-0, 2015.
Turcotte, D. L. and Schubert, G.: Geodynamics, 2nd ed., Cambridge University
Press, Cambridge, UK, 2002.Vallier, T. L., Dean, W. E., Rea, D. K., and Thiede, J.: Geologic evolution
of Hess Rise, central North Pacific Ocean, Geol. Soc. Am. Bull., 94,
1289–1307,
10.1130/0016-7606(1983)94<1289:GEOHRC>2.0.CO;2, 1983.
Wark, D. A. and Watson, E. B.: TitaniQ: a titanium-in-quartz geothermometer,
Contrib. Mineral. Petr., 152, 743–754, 10.1007/s00410-006-0132-1, 2006.Wassmann, S. and Stöckhert, B.: Rheology of the plate interface –
Dissolution precipitation creep in high pressure metamorphic rocks,
Tectonophysics, 608, 1–29, 10.1016/j.tecto.2013.09.030, 2013.Watson, E. B. and Wark, D. A.: Diffusion of dissolved SiO2 in H2O
at 1 GPa, with implications for mass transport in the crust and upper
mantle, Contrib. Mineral. Petr., 130, 66–80, 10.1007/s004100050350,
1997.Wernicke, B., Clayton, R., Ducea, M., Jones, C. H., Park, S., Ruppert, S.,
Saleeby, J., Snow, J. K., Squires, L., Fliedner, M., Jiracek, G., Keller, R.,
Klemperer, S., Luetgert, J., Malin, P., Miller, K., Mooney, W., Oliver, H.,
and Phinney, R.: Origin of High Mountains in the Continents: The Southern
Sierra Nevada, Science, 271, 190–193, 10.1126/science.271.5246.190,
1996.Weyl, P. K.: Pressure solution and the force of crystallization: a
phenomenological theory, J. Geophys. Res., 64, 2001–2025,
10.1029/JZ064i011p02001, 1959.
Xia, H. and Platt, J. P.: Is the Vincent fault in southern California the
Laramide subduction zone megathrust?, in preparation, 2017.Yin, A.: Passive-roof thrust model for the emplacement of the Pelona-Orocopia
Schist in southern California, United States, Geology, 30, 183–186,
10.1130/0091-7613(2002)030<0183:PRTMFT>2.0.CO;2, 2002.Zaun, P. E. and Wagner, G. A.: Fission-track stability in zircons under
geological conditions, Nucl. Tracks Rad. Meas., 10, 303–307,
10.1016/0735-245X(85)90119-X, 1985.