Tectonophysics 429 (2007) 1 – 20
www.elsevier.com/locate/tecto
Continuation of the San Andreas fault system into the upper mantle:
Evidence from spinel peridotite xenoliths in the Coyote
Lake basalt, central California
Sarah J. Titus ⁎, L. Gordon Medaris Jr., Herbert F. Wang, Basil Tikoff
Department of Geology and Geophysics, University of Wisconsin-Madison, 1215 W. Dayton, Madison, WI 53706, USA
Received 28 July 2005; received in revised form 12 June 2006; accepted 12 July 2006
Available online 16 November 2006
Abstract
The Coyote Lake basalt, located near the intersection of the Hayward and Calaveras faults in central California, contains spinel
peridotite xenoliths from the mantle beneath the San Andreas fault system. Six upper mantle xenoliths were studied in detail by a
combination of petrologic techniques. Temperature estimates, obtained from three two-pyroxene geothermometers and the Al-inorthopyroxene geothermometer, indicate that the xenoliths equilibrated at 970–1100 °C. A thermal model was used to estimate the
corresponding depth of equilibration for these xenoliths, resulting in depths between 38 and 43 km. The lattice preferred orientation
of olivine measured in five of the xenolith samples show strong point distributions of olivine crystallographic axes suggesting that
fabrics formed under high-temperature conditions. Calculated seismic anisotropy values indicate an average shear wave anisotropy
of 6%, higher than the anisotropy calculated from xenoliths from other tectonic environments. Using this value, the anisotropic
layer responsible for fault-parallel shear wave splitting in central California is less than 100 km thick. The strong fabric preserved in
the xenoliths suggests that a mantle shear zone exists below the Calaveras fault to a depth of at least 40 km, and combining xenolith
petrofabrics with shear wave splitting studies helps distinguish between different models for deformation at depth beneath the San
Andrea fault system.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Peridotite xenoliths; San Andreas fault; Calaveras fault; Geothermometry; Lattice preferred orientation; Seismic anisotropy
1. Introduction
Collision of the East Pacific Rise with the western
margin of North America at ∼ 29 Ma generated the
Mendocino triple junction, a transform–transform–trench
intersection, whose subsequent northwesterly migration
along coastal California produced the San Andreas
⁎ Corresponding author. Department of Geology, Carleton College,
Northfield, MN 55057, USA.
E-mail address: [email protected] (S.J. Titus).
0040-1951/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2006.07.004
transform system (Atwater, 1970; Dickinson and Snyder,
1979). Migration of the Mendocino triple junction
generated an expanding slab window (Thorkelson and
Taylor, 1989; Severinghaus and Atwater, 1990), which
was accompanied by several pulses of volcanism, thought
to have been triggered by decompression melting as
upwelling mantle flowed into the slab window (Johnson
and O'Neil, 1984; Dickinson, 1997). These volcanic
rocks are exposed along a 280-km long belt in the eastern
Coast Ranges and are generally older in the south and
younger in the north (e.g. Fox et al., 1985; Dickinson,
2
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
1997). The migratory volcanic centers are predominantly
intermediate to silicic in composition and have relatively
high δ18O values and 87Sr/86Sr ratios, indicating a
significant degree of crustal anatexis in their petrogenesis
(Johnson and O'Neil, 1984).
Among the eastern Coast Range volcanic fields, the
Coyote Lake basalt, which is located near the intersection of the Hayward and Calaveras faults about 100 km
southeast of San Francisco (Fig. 1), is noteworthy for
several reasons. First, the Coyote Lake basalt represents
a clear exception to the northward younging trend of the
volcanic fields. In contrast to other eastern Coast Range
volcanic fields whose eruptions followed closely after
passage of the Mendocino triple junction, the Coyote
Lake basalt was erupted at 2.5–3.6 Ma (Nakata et al.,
1993), well after passage of the triple junction at this
latitude (Dickinson, 1997). Instead, eruption of the
Coyote Lake basalt volcanic rocks has been ascribed to
near-fault normal extension along the Calaveras fault in
Pliocene time (Jové and Coleman, 1998). Second, in
contrast to other siliceous volcanic centers of the eastern
Coast Ranges (Johnson and O'Neil, 1984), this small
volcanic field covering an area of ∼ 11 km2 consists of a
sequence of basaltic tuffs, breccias and flows of alkaline
affinity (Nakata, 1977). Lastly, the Coyote Lake basalt is
unique among Cenozoic volcanic rocks in the Coast
Ranges because of the inclusion of mantle xenoliths,
providing the only known samples from the mantle
beneath the San Andreas fault system.
This study examines six mantle xenoliths in detail
from the Coyote Lake basalt to document the mineralogy, petrology and microstructures preserved in these
samples. Geothermometry coupled with thermal modeling allows us to estimate the temperatures, pressures
and depths at which these xenoliths equilibrated in the
mantle. Petrofabric analysis, using the lattice preferred
orientation of olivine and the modal composition of each
xenolith, allows calculation of the seismic properties.
The xenolith petrofabrics and the associated seismic
anisotropy calculations place constraints on the thickness of the anisotropic layer observed in modern shear
wave splitting studies in central California.
2. Description of mantle xenoliths
2.1. Occurrence and lithology
Small (1–3 cm diameter), subangular to subrounded
xenoliths of lower crustal and upper mantle material are
abundant in massive basalt flows and feeders of the
Fig. 1. Map of central California showing the locations of Tertiary volcanic fields. Dark arrows indicate the position of the Mendocino triple junction
through time in Ma. Upper inset shows the location within the state of California. Lower inset shows a more detailed map of the Coyote Lake basalt
with a star denoting the xenolith sample site. Modified from McLaughlin et al. (1996) and Jové and Coleman (1998).
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
Coyote Lake basalt southwest of Coyote Lake. Lower
crustal xenoliths include metagabbronorite, which consists of plagioclase, olivine, orthopyroxene, clinopyroxene and Fe–Ti oxides, and andesine megacrysts. Upper
mantle xenoliths include Cr-diopside lherzolite, harzburgite, wehrlite and dunite, and Al-augite clinopyroxenite and clinopyroxene megacrysts (following the
classification for mantle xenoliths proposed by Wilshire
and Shervais, 1975). Spinel is the characteristic
aluminous phase in the ultramafic rocks, although
plagioclase is present in some clinopyroxenite and has
been reported in some lherzolite (Nakata, 1977). Garnet
is absent in these xenoliths.
Six spinel-bearing ultramafic xenoliths were selected
for detailed investigation, including three samples of Crdiopside lherzolite (3a, 7a, 12a), two of harzburgite (2a,
7c) and one Al-augite clinopyroxenite-wehrlite composite (6a); mineral modes for the analyzed samples are
summarized in Table 1.
Because of the uncertainty in determining the identity
of each silicate mineral grain in these ultramafic rocks
during conventional point-counting for modal analysis,
modes were obtained from false-color images prepared
from Al, Ca and Fe X-ray maps, which provide
unequivocal discrimination among olivine, orthopyroxene
and clinopyroxene, and yield a relative precision from
b 1% for major phases to 4% for minor phases. The five
Cr-diopside peridotite samples contain a four-phase
assemblage of olivine (64–83%), orthopyroxene (14–
23%), clinopyroxene (3–15%) and spinel (0.2–1.9%). The
variation in modes (and mineral compositions described
below) among the five samples most likely reflects
different degrees of partial fusion ranging from 3.0% to
8.5%, as estimated from the compositions of constituent
spinel (Table 1; Hellebrand et al., 2001). Composite
sample 6a consists of Al-augite olivine-spinel clinopyrox-
3
enite, which includes feldspar (plagioclase, 8%) and
apatite (2%), and adjoining spinel wehrlite, which also
includes feldspar (alkali feldspar, 4%) and apatite (1%).
2.2. Microstructures
The grain shapes and grain boundary textures observed
in all xenoliths are quite similar and are described using
the terminology of Passchier and Trouw (1998; Fig. 2).
Grain shapes vary from polygonal to interlobate, although
polygonal shapes with relatively straight grain boundaries
meeting in 120° triple junctions are more common. All
xenoliths, except 6a, are inequigranular, where the matrix
(primarily olivine) has an average grain size of ∼ 1 mm
and larger grains (both olivine and pyroxene) are ∼ 4 mm
in diameter. Sample 6a, in contrast, has a bimodal
distribution of grain sizes (0.5 mm and 1 mm) and no
larger porphyroclasts. Sample 6a is also noteworthy
because of the abundance of fine-grained, subhedral
spinel grains.
Undulose extinction is common in many olivine
grains and, more rarely, deformation bands and
deformation lamellae are present. Pyroxenes vary in
size; generally, they are the same size as the olivine
matrix, although some larger grains (3–4 mm) are
present in sample 3a. Extremely fine-grained material,
which is localized along the xenolith margins and grain
boundaries, especially those of clinopyroxene, is the
result of the early stages of partial fusion during
entrainment and ascent of the xenoliths in basalt. No
shape preferred orientation of olivine or pyroxene grains
is observed in any of the xenoliths studied in detail.
This, however, may be related to the orientation of the
xenolith thin sections, which were created to maximize
sizes and not relative to structural fabrics within the
xenolith.
Table 1
Modal analysis of CLB mantle xenoliths
Sample
2a
3a
6a
6a⁎
7a
7c
12a
Vol.%
Harzburgite
Lherzolite
Wehrlite
Pyroxenite
Lherzolite
Harzburgite
Lherzolite
Olivine
Orthopyroxene
Clinopyroxene
Spinel
Apatite
Alkali feldspar
Plagioclase
Sum
F1
81
15
4
0.5
–
–
–
100
6.4
64
23
12
0.5
–
–
–
100
3.5
68
11
16
0.9
1
4
–
100
–
68
24
7
2
–
–
–
100
3.0
83
14
3
0.2
–
–
–
100
8.5
66
19
15
0.4
–
–
–
100
3.1
6
–
70
14
2
–
8
100
–
Modal analyses of xenoliths based on image analysis of X-ray maps, described in more detail in the text. F denotes the percent of partial melting of
each sample where F = 10 × ln(Cr#)spinel + 24 (Hellebrand et al., 2001). Sample 6a, the composite xenolith, has two entries in this table for the wehrlite
and pyroxenite portions of the xenolith.
4
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
2.3. Mineral compositions
Mineral compositions were determined by means of
a Cameca SX 50 electron microprobe, using an
accelerating voltage of 15 kV, a beam current of
20 nA, a suite of analyzed natural minerals as standards
and the ϕ(ρz) data reduction program (Armstrong,
1988). Mineral grains were analyzed in several domains
in each sample, and no significant compositional
differences were found in the peridotites on the interor intragranular scale, except locally at the rims of some
clinopyroxene and spinel grains near the xenolith
borders, reflecting changes due to xenolith excavation
and entrainment. In contrast to the peridotites, the
composite sample 6a shows a marked compositional
variation in minerals between clinopyroxenite and
adjoining wehrlite, as described below. Representative
analyses of constituent minerals are summarized in
Table 2.
Mineral compositions in the Coyote Lake basalt Crdiopside xenoliths are typical of those in Cr-diopside
xenoliths elsewhere. Olivine is magnesian, with Mg#
being restricted to values between 89.3 and 90.4, where
Mg# is defined as 100 × molecular Mg/(Mg + Fe). Spinel
is rich in MgAl2O4 component (Fig. 3a), with Mg#
ranging from 78.5 to 84.8 and Cr# from 12.4 to 21.2,
where Cr# is defined as 100 × molecular Cr/(Cr + Al).
The most Cr-rich spinel occurs in the two samples of
harzburgite, for which the highest degrees of partial
melting, 6.4% and 8.5%, are calculated (Fig. 3a).
Orthopyroxene and clinopyroxene are also magnesian
and contain relatively large R2O3 contents (0.42–
0.62 wt.% Cr2O3 and 4.23–5.05 wt.% Al2O3 in
orthopyroxene; 0.76–1.19 wt.% Cr2O3 and 5.45–
6.94 wt.% Al2O3 in clinopyroxene). The R2O3 contents
of coexisting ortho- and clinopyroxene and Cr#'s in
associated spinel are shown in Fig. 3b, demonstrating
the positive correlation between Cr# in spinel and Cr/Al
ratio in ortho- and clinopyroxene (i.e. an increase in
spinel Cr# is accompanied by a continuous shift in
position and increase in slope of the R2O3 tie lines in
pyroxene). The two harzburgite samples contain
pyroxene with the highest Cr/Al ratio.
Minerals in composite xenolith 6a are less magnesian
and chromiferous than those in the Cr-diopside
peridotites. In addition, mineral compositions differ
between the two lithologies of the composite xenolith,
with an increase in Mg/Fe and Cr/Al ratios and a
decrease in TiO2 content from clinopyroxenite to
wehrlite. For example, the Mg# in clinopyroxene
increases from 83.4 in clinopyroxenite to 88.5 in
wehrlite, TiO2 decreases from 1.68 to 0.94 wt.% and
Cr2O3 increases from 0 to 0.68 wt.%, with a corresponding decrease in Al2O3 from 8.62 to 6.58 wt.%. A
concomitant variation occurs in spinel, which is Cr-free
in clinopyroxenite and Cr-bearing (Cr# = 6.3) in wehrlite. Such variations between mineral compositions in
the clinopyroxenite and wehrlite of sample 6a are illustrated graphically in Fig. 3, where these differences may
be compared with the mineral compositions in the five
peridotite xenoliths that cluster more closely on each
graph.
3. Geothermometry
Equilibration temperatures for the Coyote Lake
mantle xenoliths were determined by application of
three different calibrations of the two-pyroxene
geothermometer (Bertrand and Mercier, 1985; Brey
and Köhler, 1990; Taylor, 1998) and by the Al-inorthopyroxene geothermometer (Witt-Eickschen and
Seck, 1991). Temperatures were calculated at an
assumed pressure of 10 kbar, because of the uncertainty
in estimating pressures for spinel peridotites (see below).
Results from the three two-pyroxene geothermometers are in excellent agreement (Table 3, Fig. 4),
although the Brey and Köhler (1990) calibration yields
temperatures that are consistently about 50 °C higher
than those from the Bertrand and Mercier (1985) and
Taylor (1998) calibrations. Least squares analysis of
temperatures from the three geothermometers, taking
the Taylor calibration as the independent variable,
yields:
TBertrandMercier ð-CÞ ¼ 17:5 þ 0:985 TTaylor ð-CÞ
R2 ¼ 0:977
TBreyKPhler ð-CÞ ¼ 151 þ 0:896 TTaylor ð-CÞ
R2 ¼ 0:969
Using the Bertrand–Mercier and Taylor calibrations,
the mean temperatures for the six analyzed samples
range from 970 to 1060 °C (Table 3).
Temperatures were also calculated for three additional peridotite xenoliths, which were analyzed in grain
mounts prepared from crushed samples (shown in
Fig. 4, but not included in Table 3). Results from two
of these xenoliths, 1030 and 1045 °CTaylor, lie within the
range of the previous six samples, and the third, 1090 °CTaylor, extends the range to a slightly higher temperature. Recalculation of temperature estimates from three
other Coyote Lake basalt peridotite xenoliths, based on
mineral compositions reported in Jové and Coleman
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
(1998), yield 925, 965 and 1085 °CTaylor. Based on all
available results, the Coyote Lake basalt mantle suite
appear to have been derived from a relatively restricted
temperature range of 175 °C.
Results from the Al-in-orthopyroxene geothermometer are comparable to those from the two-pyroxene
geothermometers (Table 3), being within 40 °C of the
temperatures obtained from the Bertrand–Mercier and
Taylor calibrations. Such correspondence between the
5
two types of geothermometers and the regular distribution of sub-parallel pyroxene tie lines in Fig. 3b indicates
the attainment and preservation of equilibrium compositions in pyroxene with respect to Ca, Mg, Fe, Al and Cr.
4. Estimated depths
In general, spinel peridotites are assigned to an
intermediate pressure range from ∼ 8 to ∼ 18 kbar
Fig. 2. Photomicrographs of xenoliths taken under crossed-polarizers. Inset in each figure taken under plane-polarized light illustrates the texture of
spinel. Width of each photograph is 1 cm. Numbers indicate temperature (°C) and depth (km) estimates, based on the thermal model in Fig. 5.
6
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
Table 2
Representative mineral analyses for Coyote Lake basalt mantle xenoliths
Wt.%
SiO2
TiO2
Al2O3
Cr2O3
V2O3
FeO
MnO
MgO
ZnO
NiO
CaO
Na2O
Sum
Si
Ti
Al
Cr
V
Fe
Mn
Mg
Zn
Ni
Ca
Na
Sum
Mg#
Cr#
Olivine
Spinel
2a
3a
6a
7a
7c
12a
40.31
40.52
39.87
40.07
40.37
40.87
3a
0.07
54.42
12.54
0.08
11.21
0.11
20.96
0.08
0.33
0.22
60.42
5.79
0.08
12.22
0.09
21.08
0.08
0.30
99.80
100.28
10.37
0.15
48.33
9.91
0.14
48.48
12.03
0.16
47.07
9.37
0.16
49.31
9.61
0.14
48.97
9.71
0.13
48.78
0.32
0.36
0.27
0.37
0.36
0.34
0.14
50.66
15.54
0.10
12.90
0.13
20.12
0.08
0.27
99.48
99.41
99.40
99.28
99.45
99.82
99.94
Cations per 4 oxygen atoms
0.997
1.000
0.995
6a⁎
2a
6a
7a
7c
12a
0.26
65.25
0.00
0.10
14.93
0.11
20.13
0.07
0.17
0.13
56.22
11.93
0.08
9.07
0.11
21.91
0.07
0.33
0.16
46.38
18.56
0.10
14.72
0.13
19.36
0.07
0.30
0.09
54.08
11.44
0.07
12.71
0.10
20.76
0.09
0.36
101.02
99.85
99.78
99.70
Cations per 4 oxygen atoms
0.990
0.996
1.003
0.214
0.003
1.782
0.205
0.003
1.784
0.251
0.003
1.751
0.194
0.003
1.816
0.198
0.003
1.800
0.199
0.003
1.785
0.006
0.007
0.005
0.007
0.007
0.007
3.003
89.3
3.000
89.7
3.005
87.5
3.010
90.3
3.004
90.1
2.997
90
0.003
1.598
0.329
0.002
0.289
0.003
0.803
0.002
0.006
0.001
1.684
0.260
0.002
0.246
0.002
0.821
0.002
0.007
0.004
1.824
0.117
0.002
0.262
0.002
0.805
0.002
0.006
0.005
1.939
0.000
0.002
0.315
0.002
0.757
0.001
0.003
0.003
1.716
0.244
0.002
0.196
0.002
0.846
0.001
0.007
0.003
1.498
0.402
0.002
0.337
0.003
0.791
0.001
0.007
0.002
1.684
0.239
0.001
0.281
0.002
0.818
0.002
0.008
3.033
73.5
17.1
3.025
76.9
13.4
3.024
75.4
6.0
3.025
70.6
0.0
3.017
81.2
12.4
3.045
70.1
21.2
3.036
74.4
12.4
Upper
tabulation shows weight percent oxides of representative analyses
of olivine, spinel, orthopyroxene and clinopyroxene from
Orthopyroxene
Clinopyroxene
*
each
xenolith
sample.
6a
has
two
different
lithologies:
6a
denotes
the
wehrlite
portion
portion.
2a
3a
6a
7a
7c
12a
2a
3a
6a and 6a
6a⁎ denotes
7a the pyroxenite
7c
12a
Lower tabulation shows the number of cations per 4 oxygen atoms for olivine and spinel, and per 6 oxygen atoms for
orthopyroxene and clinopyroxene. The bottom two rows show the Mg# and Cr# for each mineral. The Mg# is defined for silicates
as the molecular percent Mg/(Mg + Fetotal) and for spinel as Mg/(Mg + Fe2+), where Fe2+ is calculated by stoichiometry. Cr# is
similarly defined as the molecular percent Cr/(Cr + Al).
(depending on bulk-rock Mg# and Cr#), between the
stability fields of plagioclase peridotite and garnet
peridotite, and maximum pressures are estimated for
individual samples from the Cr content of spinel
(O'Neill, 1981). However, no viable procedure exists
at present for specific pressure determinations of spinel
peridotites. Although a geobarometer based on the Ca
content of olivine was proposed by Köhler and Brey
(1990), and accurate and precise determinations of trace
amounts of Ca in olivine can be made by electron
microprobe under appropriate operating conditions,
application of the geobarometer produces spurious
results, especially for lower temperature xenoliths
(O'Reilly et al., 1997; Medaris et al., 1999).
Despite the present lack of a geobarometer suitable
for spinel peridotites, depth estimates for spinel
peridotite xenoliths can be obtained by combining
equilibration temperatures with a geotherm associated
with the eruptive site at the time of xenolith transferal to
the surface. The method and results of a numerical
model for thermal evolution of the Coyote Lake basalt
site are presented in the following section.
5. Thermal modeling and evolution
The thermal evolution of lithosphere beneath the
eastern Coast Ranges was estimated from a thermal
model, which simulates upwelling of 1300 °C asthenosphere that fills in the slabless window following the
northward movement of the Mendocino triple junction
(Lachenbruch and Sass, 1980). The one-dimensional,
finite difference model extended from the surface to a
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
Orthopyroxene
2a
3a
7
Clinopyroxene
6a
7a
12a
2a
3a
6a
6a⁎
7a
7c
12a
53.90
0.01
4.49
0.48
53.59
0.07
5.05
0.50
53.21
0.30
5.17
0.26
54.02
0.13
4.94
0.43
54.31
0.06
4.23
0.62
54.19
0.12
5.05
0.42
51.73
0.19
5.55
1.19
51.35
0.33
5.93
0.76
50.04
0.94
6.58
0.68
48.02
1.68
8.62
0.00
51.56
0.47
6.04
0.78
51.75
0.27
5.45
1.08
51.05
0.36
6.94
0.85
6.87
0.15
33.15
6.39
0.14
32.91
7.97
0.19
32.21
6.20
0.14
33.49
6.35
0.14
33.37
6.39
0.14
32.84
3.34
0.12
16.69
3.12
0.10
16.25
3.66
0.12
15.84
5.20
0.16
14.67
2.66
0.07
16.11
3.07
0.10
16.56
3.25
0.10
16.15
0.94
0.08
100.07
1.51
0.13
100.29
0.84
0.07
100.22
1.06
0.07
100.48
1.07
0.08
100.23
1.07
0.08
100.30
20.06
1.04
99.91
20.76
0.94
99.54
20.73
0.91
99.50
19.90
0.98
99.23
20.97
1.04
99.70
20.39
1.12
99.79
19.78
1.16
99.62
Cations per 6 oxygen atoms
1.873
1.858
1.855
0.000
0.002
0.008
0.184
0.206
0.212
0.013
0.014
0.007
7c
1.864
0.003
0.201
0.012
1.880
0.002
0.173
0.017
1.873
0.003
0.206
0.011
Cations per 6 oxygen atoms
1.878
1.871
1.833
0.005
0.009
0.026
0.238
0.255
0.284
0.034
0.022
0.020
1.774
0.047
0.375
0.000
1.873
0.013
0.259
0.022
1.881
0.007
0.233
0.031
1.855
0.010
0.297
0.024
0.200
0.004
1.717
0.185
0.004
1.701
0.232
0.006
1.674
0.179
0.004
1.722
0.184
0.004
1.722
0.185
0.004
1.692
0.101
0.004
0.903
0.095
0.003
0.883
0.112
0.004
0.865
0.161
0.005
0.808
0.081
0.002
0.872
0.093
0.003
0.897
0.099
0.003
0.875
0.035
0.005
4.031
89.6
6.6
0.056
0.009
4.035
90.2
6.4
0.031
0.005
4.030
87.8
3.2
0.039
0.005
4.029
90.6
5.6
0.040
0.005
4.026
90.3
8.9
0.040
0.005
4.018
90.1
5.1
0.780
0.073
4.017
89.9
12.5
0.811
0.066
4.015
90.3
7.9
0.814
0.065
4.022
88.5
6.6
0.788
0.070
4.027
83.4
0.0
0.816
0.073
4.011
91.5
7.8
0.794
0.079
4.019
90.6
11.7
0.770
0.081
4.015
89.8
7.5
depth of 81 km. The initial condition consisted of a
steady-state crustal geotherm with exponential decay of
heat production from the surface to a depth of 30 km.
The steady-state geotherm was based on a reduced heat
flow of 20 mW/m2 and a heat production contribution of
25 mW/m2, yielding an initial surface heat flow of
45 mW/m2. The reduced heat flow reflects the insulating
effect of the subducted Farallon plate underlying the
North American plate (Furlong, 1984) and the initial
surface heat flow is comparable to that used by Zandt
and Furlong (1982) and Furlong (1984). The asthenospheric upwelling was represented by a step temperature
discontinuity to 1300 °C followed by an adiabatic
gradient of 1 °C/km to 81-km depth. The boundary
conditions were specified as temperatures of 0 °C at the
surface and 1351 °C at 81-km depth. Other parameters
used were a thermal conductivity of 2.5 W/m K and
thermal diffusivity of 10− 6 m2/s.
The calculated geotherms at 3.0 m.y. after upwelling
(Fig. 5) match closely those shown by Furlong (1984) in
his Fig. 3. The predicted surface heat flow changes
through time, with 73 mW/m2 at a model time of 6 m.y.,
a peak value of 78 mW/m2 at 8 m.y. and 77 mW/m2 at
9 m.y. (0 Ma), compared to a present-day observed
surface heat flow of 80 mW/m2. These results for the
one-dimensional thermal model lie between the convective model I and conductive model II of Zandt and
Furlong (1982, their Fig. 3c).
Depths for the six Coyote Lake basalt xenoliths were
estimated by intersecting their temperatures (using
results from the Taylor two-pyroxene geothermometer
and taking into account its pressure dependence) with the
8
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
the Coyote Lake basalt xenoliths occurs at depths
between 37.8 and 42.5 km. Thus, the Coyote Lake
basalt spinel peridotite xenoliths were likely derived
from a relatively narrow depth interval, ∼ 5 km, in the
upper part of the lithospheric mantle. This estimate is
consistent with their phase petrology and well within the
maximum pressures allowed by the compositions of
constituent spinel (Table 2). If the higher temperatures
from the Brey and Köhler two-pyroxene geothermometer are used, the depth interval remains constant at
∼ 5 km, but the depth estimates increase by about 3 km
representing depths of approximately 40–45 km.
Thus, the spinel peridotite xenoliths from the Coyote
Lake basalt allow us to evaluate the physical properties
of the mantle associated with the San Andreas fault
system, from a depth of ∼ 40 km at 3 Ma, approximately
6 m.y. after passage of the Mendocino triple junction.
6. Olivine petrofabrics
Fig. 3. Graphical representations of geochemical data. (a) Compositions of spinel in mantle xenoliths in the Coyote Lake basalt. Such Mgrich and Cr-poor spinel compositions are typical for relatively
undepleted mantle peridotite. The numbers adjacent to the spinel
data points indicate the Mg#'s of coexisting olivine. Note the relatively
Fe-rich and Cr-poor compositions of spinel in the composite sample,
6a: pd, peridotite (wehrlite); pxite, pyroxenite. (b) R2O3 contents of
orthopyroxene (opx) and clinopyroxene (cpx) in mantle xenoliths in
the Coyote Lake basalt. The sub-parallel tie lines between opx and cpx
indicate chemical equilibrium among pairs of pyroxenes. Cr#'s of
coexisting spinels are indicated by the italicized numbers adjacent to
clinopyroxene compositional points. Higher Cr/Al ratios in pyroxenes
(and corresponding spinels) reflect higher degrees of partial melting.
Note the lower Cr/Al ratios of pyroxene in the composite sample; pd,
peridotite; pxite, pyroxenite.
geotherm calculated at a model age of 6 m.y. (Table 3,
Fig. 5). This geotherm corresponds to the eruption of the
Coyote Lake basalt at 3 Ma. The range of temperatures of
The lattice preferred orientation (LPO) of olivine was
measured in the five peridotite xenolith samples (but not
the composite pyroxenite/wehrlite sample 6a) using a
five-axis universal stage microscope following the
method of Emmons (1943). For each thin section, the
orientation of each individual olivine grains was
measured. Due to variations in the physical size of
each xenolith, between 70 and 165 grain orientations
were determined in each thin section. The LPO was not
measured for the orthopyroxene and clinopyroxene
grains in thin section since the small volume fraction
of pyroxene represented in each sample (Table 1) makes
it impossible to determine a statistically significant LPO
pattern for these minerals.
Olivine LPO data are presented in Fig. 6a. On each
lower hemisphere, equal-area projection, olivine a-axes,
[100], are oriented horizontal and east–west, while the caxes, [001], are approximately vertical. This orientation
was chosen for several reasons. First, LPO data are
typically plotted with a vertical east–west foliation and a
horizontal lineation, where the foliation and lineation are
determined independently based on the shape preferred
orientation of the constituent minerals. Because no shape
preferred orientation was apparent in these samples, the
selected orientation provides a common framework to
compare the LPO patterns of each sample in the suite,
and to compare the patterns with those from other studies
of mantle xenoliths lacking a strong shape preferred
orientation (e.g. Christensen et al., 2001; Vauchez et al.,
2005). Second, this orientation is reasonable based on
the bulk kinematics of the San Andreas fault system,
where a vertical foliation with a horizontal lineation is
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
9
Table 3
Estimated temperatures, maximum pressures, depths and calculated seismic anisotropy values of Coyote Lake basalt mantle xenoliths
Sample
6a
7a
3a
7c
2a
12a
Average
T at 10 kbar
2-Px, T
2-Px, BM
2-Px, BK
Al-in-Opx
970
975
1000
1005
1040
1060
1016
975
980
1005
1010
1040
1060
1019
1025
1020
1045
1055
1080
1100
1060
930
990
1005
1045
1010
1030
1016
Pmax
(kbar)
Temp.
(°C)
Depth
(km)
Vp
max
(km/
s)
Vp
min
(km/
s)
AVp
(%)
Vs1
max
(km/
s)
Vs1
min
(km/
s)
Vs 2
max
(km/
s)
Vs2
min
(km/
s)
AVs
(%)
δt
(s)
16.1
18.5
21.2
19
20.5
19.6
20
975
980
1015
1020
1055
1070
1028
37.8
37.9
39.6
39.8
41.7
42.5
40.3
–
8.71
8.61
8.84
8.80
8.53
8.70
–
7.96
8.03
8.01
7.97
8.17
8.03
–
9.0
7.0
9.9
9.9
4.3
8.0
–
5.01
4.95
5.01
5.02
4.94
4.99
–
4.83
4.84
4.87
4.82
4.83
4.84
–
4.87
4.88
4.92
4.90
4.87
4.89
–
4.68
4.70
4.64
4.68
4.74
4.69
–
6.8
5.2
7.6
7.0
4.1
6.1
–
1.4
1.1
1.6
1.4
0.9
1.3
The left section of the table shows estimated temperatures, maximum pressures and depths based on geothermometry and thermal modeling. Four
thermometers were used to estimate temperatures at an assumed pressure of 10 kbar: three two pyroxene thermometers: T—Taylor (1998), BM—
Bertrand and Mercier (1985), BK—Brey and Köhler (1990) and one Al-in-opx geothermometer from Witt-Eickschen and Seck (1991). Maximum
pressures are calculated from the composition of spinel (O'Neill, 1981). Temperatures and depths are based on a combination of two-pyroxene
geothermometry and the thermal model presented in Fig. 5. The right section of the table shows the calculated seismic anisotropy values, based on the
LPO fabrics presented in Fig. 6a and computed using software by Mainprice (1990). The anisotropy of P-waves, AVp, is defined as (Vp max − Vp
min) / (average Vp) from Mainprice and Silver (1993); the shear wave anisotropy, AVs, is similarly defined. Shear wave delay times (δt) are reported
for a 100-km path defined as AVs normalized by the average S-wave velocity.
anticipated due to dominantly wrench deformation at the
plate boundary. Although no shape preferred orientation
was visible in these samples, olivine a-axes in many
naturally and experimentally deformed peridotites are
sub-parallel to the lineation direction and a- and c-axes
define a plane parallel to the foliation (e.g. Zhang and
Karato, 1995; Ben-Ismail and Mainprice, 1998; Tommasi et al., 2004). Third, this orientation produces
maximum values of shear wave splitting for vertically
Fig. 4. Comparison of temperature estimates (at 10 kbar) for the Coyote
Lake basalt mantle xenoliths from the T—Taylor (1998), BM—
Bertrand and Mercier (1985) and BK—Brey and Köhler (1990) twopyroxene geothermometers.
incident shear waves, described in more detail in the
following section.
The LPO patterns in all the Coyote Lake basalt
xenoliths display strong point distributions of olivine a-
Fig. 5. Calculated thermal evolution of the Coyote Lake area, showing
the estimated temperatures and depths for the Coyote Lake basalt
mantle xenoliths. Samples symbols as in Fig. 4.
10
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
axes except 12a, in which a-axes form a girdle. Olivine
b-axes are characterized by slightly weaker point
distributions (fewer contours) and in 7c, a weak girdle
is developed. The c-axes show the most diffuse LPO
patterns, which is typical in both naturally deformed
peridotites from ophiolite massifs (e.g. Boudier and
Coleman, 1981; Salisbury and Christensen, 1985),
xenoliths (e.g. Soedjatmiko and Crhistensen, 2000),
experimentally deformed peridotites (e.g. Carter and
Avé Lallement, 1970; Nicolas et al., 1973) and in
modeled olivine LPO patterns from many different
kinematic environments (e.g. Tommasi et al., 1999).
There is no systematic change in olivine LPO patterns as
a function of depth, such as that observed in other
xenolith suites (Christensen et al., 2001; Kennedy et al.,
2002; Kobussen, 2005).
The olivine LPO patterns can be compared to predicted
LPO patterns under a variety of imposed kinematic
constraints and based on numerical modeling. The pattern
observed in the Coyote Lake basalt xenoliths is more
consistent with deformation in simple shear or in pure
shear, both of which predict a-axis point maxima
Fig. 6. (a) Lower hemisphere, equal area projections of olivine a-, b- and c-axes measured on a universal stage microscope; contours are multiples of
uniform distribution with the lowest contour indicated by the dashed line. (b) Seismic anisotropy values calculated from the olivine LPO patterns.
Compressional wave velocity (Vp) in km/s, shear wave splitting (AVs) in % and polarization directions for shear waves with different incidence
angles. The black square indicates the maximum velocity or anisotropy and the white circle indicates the minimum; CI stands for contour interval.
The numerical valves for Vp and Vs for each xenolith sample are included in Table 3.
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
(Tommasi et al., 1999). The LPO patterns developed in
transpression, which might be expected given the oblique
convergence along the San Andreas fault system, are
characterized by a-axis girdles (Tommasi et al., 1999),
which are only observed in sample 12a.
Because an independent measure of foliation and
lineation is not available in these xenoliths, we cannot
definitively determine the predominant slip system.
However, the tendency towards orthorhombic point
distributions with strong a-axes concentrations suggests
that deformation may best be described by dislocation
creep at relatively high temperatures, N1000 °C, on the
slip system (010)[100] (e.g. Nicolas and Poirier, 1976;
Carter and Avé Lallement, 1970). This is one of the most
common slip systems for olivine from xenolith suites
(e.g. Ben-Ismail and Mainprice, 1998; Tommasi et al.,
2000) and is consistent with the temperatures determined by geothermometry.
7. Seismic anisotropy calculations
The small sample sizes of these xenoliths require
calculation, rather than direct measurement, of seismic
properties. Seismic velocities are calculated using the
LPO, density and the elastic stiffness coefficient for each
mineral in the rock (e.g. Crossen and Lin, 1971;
Mainprice and Silver, 1993). These calculations assume
samples are free from cracks and alteration, and use the
modal composition of each xenolith (Table 1) for the
volume fraction of each phases. We used published
elastic stiffness coefficients for olivine (Abramson et al.,
1997), enstatite (Duffy and Vaughan, 1998) and diopside
(Collins and Brown, 1998), and our calculations were
made using updated software by Mainprice (based on
Mainprice, 1990; Mainprice and Humbert, 1994)
utilizing the Voigt averaging technique. Choice of this
average is based on the agreement between laboratory
measurements of seismic velocities and calculated
values from LPO patterns for the Twin Sisters dunite
(Crossen and Lin, 1971), and is useful for comparison
with other xenolith petrofabric studies, the majority of
which also use the Voigt average (Table 4). The
averaging technique affects the absolute velocities but
not the anisotropy values (Mainprice and Silver, 1993).
Because pyroxene LPOs were not measured in our
samples, but their LPOs tend to decrease the overall
anisotropy (Christensen and Lundquist, 1982), we
approximated this contribution by assuming that pyroxene LPO is uniformly distributed (i.e. isotropic to seismic
waves). The small contribution from minor phases like
spinel (b 5% of the total volume) was ignored in the
calculations. Delay times were computed for a 100-km
11
thick slab with an anisotropy equal to the aggregate
anisotropy of each sample. These assumptions, which are
common in many petrofabric studies (e.g. Mainprice and
Silver, 1993), generally produce anisotropies that are
∼ 1% higher than for seismic anisotropy measured in
rocks in situ (e.g. Christensen, 2002).
The results of the seismic anisotropy calculations are
presented in Fig. 6b. The orientation of each sample is
the same as in Fig. 6a. Regardless of the LPO fabric
strength (or total number of grains), the seismic
anisotropy values are similar for all five xenoliths
(Table 3). The average values for the compressional
wave velocity (Vp) and anisotropy (AVp) are 8.4 km/s
and 8.0%; the average shear wave velocity (Vs) and
anisotropy (AVs) are 4.9 km/s and 6.1%; and the average
delay time (δt) is 1.3 s for a 100-km path. It is important
to note that the maximum shear wave splitting is near
vertical in Fig. 6b, suggesting that with this particular
orientation, delay times for vertically incident seismic
waves will be close to the maximum.
These results can be compared with data from similar
studies of mantle xenoliths from different tectonic
environments, which have been compiled in Table 4.
For this compilation, we recompute seismic anisotropy
parameters from each paper, where necessary, to create a
common framework for comparison between studies. For
example, we weight each xenolith with normal mantle
composition equally when calculating the average seismic
anisotropy parameters. We do not use reported seismic
anisotropy values calculated from aggregate LPO data
(e.g. Ji et al., 1994; Saruwatari et al., 2001); thus, the
values in this table may not agree with stated values in the
original source.
For completeness, Table 4 also summarizes the
modal variations of major minerals, as well as
temperatures and pressures provided for each xenolith
suite. The modal composition of each individual
xenolith or the regional modal average was used in
the seismic anisotropy calculations except in the
studies by Pera et al. (2003), Christensen et al.
(2001) and Kobussen (2005), where seismic anisotropy
calculations assumed xenoliths were composed of
100% olivine and therefore represent maximum values.
Additionally, the reported temperature ranges in some
instances are based on older geothermometers (e.g.
Boyd, 1973; Wells, 1977) that have been superseded
by more recent geothermometers (e.g. Bertrand and
Mercier, 1985; Brey and Köhler, 1990; Taylor, 1998).
Depth estimates are commonly based on the inferred
geothermal gradients in each tectonic setting, since
estimation of pressures from spinel peridotites is
problematic (e.g. Medaris et al., 1999).
12
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
Table 4
Comparison of mantle xenolith compositions and calculated seismic anisotropy values from different tectonic settings
Location
Lithology
Cratonic environments
Kaapvaal craton Peridotite ±
kimberlites,
garnet
South Africa
Kaapvaal craton Peridotite ±
kimberlites,
garnet
South Africa
Kaapvaal craton Peridotite ±
kimberlites,
spinel/
South Africa
garnet
Labait volcano,
Spinel/
Tanzania
garnet
peridotite
Convergent setting
N. British
Spinel
Columbia,
peridotite
Canada and
Alaska, USA
S. British
Spinel
Columbia,
peridotite
Canada
Peridotite
Torre Alfina,
Northern
Apennines,
Italy
Method ol
opx cpx sp gt Ave. Vp
Vs
AVp AVs
(%) (%) (%) (%) (%)
(km/s) (km/s) (%) (%)
5
U
56– 22– –
80 44
–
3– V
8
8.4
NA
5.4
3.7
∼ 1 900–
1050
120–
170
6
U
33– 10– b4
60 32
–
7– H
12
8.3
4.8
5.4
4.4
0.9
980
150
48 U/
EBSD
NA NA NA NA NA V
8.2
4.7
3.0
2.6
0.6
1200– 170–
1400 200
Mainprice
and Silver
(1993)
Long and
Christensen
(2000)
Ben-Ismail
et al. (2001)
13 EBSD
60– 4–
94 28
1 – b15 VRH 8.3
3
NA
6.5
5.1
N1
1000– 70–
1400 140
Vauchez et
al. (2005)
13 U
64– 21– 6–
68 26 11
1– –
2
12 U
54– 25– 12– b3
61 34 21
15 EBSD
Extension setting (rifts)
Vitim, Baikal
Peridotite + 4 U
region, Russia spinel/garnet
Cima volcanic
Spinel
15 U
field, California, peridotite
USA
Kozákov volcano,
Czech Republic
As Shamah
volcanic field,
Syria
Hotspot track
Polynesian
hotspots,
South Pacific
1–
7
NA
6.3
4.5
1.0
900–
1100
45–
59
Ji et al.
(1994)
–
V
8.0
4.6
7.9
6.6
1.5
900–
1100
45–
67
Saruwatari et
al. (2001)
N90 NA NA NA –
V
8.2
4.6
13.5 11.5
1.8
1000– 50–
1080 60⁎
8.2
4.4⁎
5.8
4.4⁎
0.4
8.3
NA
7.4
5.2
40–
90
30–
45
680–
1065
900–
1100
57–
86
54–
93
10– FIX b3 b6 NA
26
6– 1– ≤1 – H
38 7
49–
82
65–
90
0–
34
0–
25
Kern et al.
(1996)
Soedjatmiko
and
Crhistensen,
2000
Christensen
et al. (2001)
Kobussen
(2005)
1– –
6
2– –
4
NA
8.6
4.9
NA
8.5
4.9
4–8 2 . 4 – 1.0
5.0
8.5 6.3
1.3
4–
37
NA –
VRH 8.3
4.8
6.0
4.2
0.9
b1100 NA
Tommasi et
al. (2004)
b2
V
4.9
8.0
6.1
1.3
970–
1100
This study
20 EBSD
45– 0–
100 30
5
64– 11– 4–
83 23 16
U
925–
1200
NA 955–
1060
Pera et al.
(2003)
5 –
13
5–
8
Spinel
peridotite
U
Depth Citation
(km)
8.0
15 U
9
T
(°C)
V
Spinel
lherzolite
Spinel
peridotite
Transform setting
Calaveras fault,
Spinel
California, USA peridotite
δt
(s)
N
–
8.4
32–
70
NA
38–
43
Compilation of xenolith composition data and calculated seismic anisotropy values based on LPO measurements from different tectonic
environments. N represents the total number of xenoliths studied. The method of LPO measurement is either by using a U-stage (U) or electron
backscatter diffraction patterns (EBSD). Range of modal percentages is reported for olivine (ol), orthopyroxene (opx), clinopyroxene (cpx), spinel
(sp) and garnet (gt). The averaging technique for seismic anisotropy calculations (Ave.) is abbreviated as Voigt (V), Hill (H) or Voigt–Reuss–Hill
(VRH). Average seismic anisotropy values are reported for P-wave velocity (Vp), S-wave velocity (Vs), P-wave anisotropy (AVp), S-wave anisotropy
(AVs) and shear wave delay time (δt) for a 100-km ray path. The range of temperatures (T) and estimated depths are reported for each xenolith suite.
The S-wave velocity and anisotropy that are starred (Kern et al., 1996) are from experimentally determined values (not calculated from the xenoliths'
LPOs). NA indicates that specific data were not available.
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
The seismic velocities of compressional (Vp) and shear
waves (Vs) in the Coyote Lake basalt xenoliths are similar
to those observed from other tectonic environments, while
the anisotropy of compressional (AVp) and shear waves
(AVs) is generally higher. In particular, the 6% shear wave
anisotropy from the Coyote Lake basalt xenoliths is
higher than the 4–5% anisotropies calculated for
xenoliths from cratons and rift environments but similar
to values from orogenic belts that range from 4.5% to
11%. This higher anisotropy significantly affects the
interpretations of modern shear wave splitting studies,
which generally assume 4% shear wave anisotropy for all
tectonic environments, based primarily on the LPOs of
kimberlite xenoliths (e.g. Silver and Chan, 1991; Silver
and Savage, 1994). The implications of the strong LPO
patterns and associated seismic anisotropy values of the
Coyote Lake basalt xenoliths on the tectonics of the San
Andreas fault system are discussed below.
8. Discussion
The xenoliths from the Coyote Lake basalt represent
the only known samples of the mantle beneath the San
Andreas fault system. The variation in mineral modes
and compositions of the xenoliths (Tables 1 and 2)
reflects compositional variability within a vertically
restricted (∼ 5 km) region in the mantle. Note, however,
that the xenoliths investigated here do not represent the
full suite of mantle xenoliths from the Coyote Lake
basalt; they were chosen because geothermometry was
possible based on their coexisting mineral phases.
The seismic anisotropy values calculated from petrofabrics in the Coyote Lake basalt peridotite xenoliths are
significantly higher than those calculated from mantle
xenoliths in most other tectonic settings, especially from
xenoliths that sample the mantle beneath cratonic
environments (Table 4). This suggests that the LPO
strength is related to deformation associated with the San
Andreas fault system; thus the Coyote Lake basalt
xenoliths represent the first non-geophysical documentation of a shear zone in the mantle beneath the San Andreas
fault. Because geothermometry and thermal modeling of
the mantle xenoliths entrained in the Coyote Lake basalt
indicate that xenoliths originated from depths of approximately 40 km, the San Andreas fault system presumably
continues into the mantle as a shear zone to at least 40-km
depth. This depth is approximately 10–15 km below the
crust–mantle boundary in west–central California, based
on seismic refraction profiles across the San Andreas fault
system (Fuis and Mooney, 1990 and references therein).
Because of their location and history within the
deforming plate boundary system, the Coyote Lake
13
basalt xenoliths provide important information about the
lithospheric structure along the San Andreas fault
system. The LPO patterns aid the interpretation of
shear wave splitting values from central California, and
together these data sets shed light on the patterns of
deformation in the mantle beneath the San Andreas fault
system, as discussed in more detail below.
8.1. Comparison with shear wave splitting values
While analysis of petrofabrics from xenoliths represents one of the few ways to directly study continental
mantle LPO, albeit with small sample sizes, shear wave
splitting measurements provide a large-scale picture that
is interpreted to reflect LPO patterns in the mantle (e.g.
Vinnik et al., 1989; Silver and Chan, 1991). If the mantle
has a consistent anisotropy over a wide region due to the
alignment of anisotropic crystals (primarily olivine),
shear waves traveling through the medium will be split
into two orthogonal polarized waves traveling at
different speeds. Shear wave splitting studies generally
use this birefringence in SKS and SKKS waveforms,
which have near-vertical incidence, and measure the
difference between arrival times of the fast and slow
directions of these waves (e.g. Silver, 1996). The delay
time between S-wave arrivals is indicative of the degree
of anisotropy sampled by the waveform as well as the
thickness of the anisotropic layer.
For a given thickness of anisotropic mantle composed primarily of olivine, and assuming that olivine
crystallographic axes track finite strain, the largest delay
times are observed when wave propagation is parallel to
the intermediate finite strain axis, with the fast wave
traveling parallel to the maximum finite strain axis and
the slow wave parallel to the minimum finite strain axis
(e.g. McKenzie, 1979; Ribe, 1992). For typical olivine
LPO patterns, this translates into propagation parallel to
the c-direction with the fast wave parallel to the a-axis
and the slow wave parallel to the b-axis. Anisotropy in
the crust can also contribute to the observed shear wave
splitting at the surface, but is often neglected since its
maximum contribution is estimated to be 0.1–0.2 s per
10 km depending on the composition and orientation of
bulk crustal fabrics (e.g. Barroul and Mainprice, 1993).
Shear wave splitting delay times from central California are generally interpreted to result from two different
anisotropic layers in the mantle: a deep asthenospheric
layer responsible for EW oriented splitting and a lithospheric mantle layer with splitting directions parallel to
the San Andreas fault (Silver and Savage, 1994; Ozalaybey and Savage, 1994, 1995; Hartog and Schwartz,
2001; Polet and Kanamori, 2002). The average delay
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S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
times for the lower asthenospheric layer are generally
0.85–1.7 s and 0.5–1.25 s for the upper lithospheric layer
(Hartog and Schwartz, 2001). The observed delay times
using a two-layer model are illustrated in Fig. 7a.
The deeper asthenospheric splitting is attributed to
asthenospheric flow within the slab window (Ozalaybey
and Savage, 1995; Hartog and Schwartz, 2000) or to
absolute motion of the Sierra Nevada–Great Valley
block (Hartog and Schwartz, 2001). The shallow
lithospheric fabric is generally attributed to strain in
the upper mantle along the plate boundary, the lateral
extent of which seems to extend approximately 55 km to
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
the west of the San Andreas fault and up to 80 km to the
east (Ozalaybey and Savage, 1995). The magnitude of
delay times from the upper anisotropic layer tends to
decrease towards the east (Fig. 7a; Hartog and Schwartz,
2001), although the orientation remains parallel to the
San Andreas fault. This is in marked contrast to the
rotation of fast shear wave splitting directions observed
across the Alpine fault in New Zealand (e.g. Molnar et
al., 1999). Fault-parallel fabrics are also observed in
central California, determined by Pn tomography, which
sample only the upper mantle (Hearn, 1996). The twolayer anisotropy models based on shear wave splitting
values (Ozalaybey and Savage, 1995; Polet and
Kanamori, 2002) generally assume a 4% shear wave
anisotropy in both layers of the mantle (e.g. Silver and
Chan, 1991; Silver and Savage, 1994). These models
therefore require 100–150 km of lithospheric mantle
with a fault-parallel fast polarization direction in central
California to account for the observed shear wave
splitting of the upper layer (Ozalaybey and Savage,
1995; Hartog and Schwartz, 2001).
The calculated depths of the Coyote Lake basalt
xenoliths place the xenoliths within the upper lithospheric mantle layer in the two-layer model. The
average 6% shear wave anisotropy calculated from the
xenolith LPO patterns can be used to estimate the
thickness of the upper anisotropic layer depending on
the orientation of LPO in the mantle. The fabric strength
can also be compared to the fault-parallel fast direction
and the kinematics of the fault system in order to
understand fabric development in these xenoliths.
8.2. Lithospheric thickness
If foliation (defined here by olivine a- and c-axes) is
vertical and lineation (a-axes) is horizontal (Fig. 6a), the
maximum shear wave splitting is nearly vertical and
maximum delay times will be observed at the surface for
vertically propagating waves with the fast wave parallel to
a-axes (Fig. 6b). The observed shear wave splitting delay
times require 40–95 km of anisotropic mantle with a
15
consistent orientation. If instead, foliation and lineation
are both horizontal (horizontal a- and c-axes), the
calculated shear wave anisotropies for vertical incidence
are slower due to the orthorhombic LPO symmetry in four
out of the five xenoliths. Only sample 12a predicts a
maximum vertical shear wave anisotropy for a horizontal
foliation—the four other xenoliths would have shear
wave anisotropies approximately half to two-thirds as
great, suggesting a required lithospheric mantle thickness
of 60–150 km.
Seismological data (Zandt, 1981; Hill, 1989) and
thermal modeling (Zandt and Furlong, 1982) indicate
that, although there are variations in depth to the
asthenosphere in central California, the total lithospheric
thickness is probably less than 80 km and may be as
little as 40–60 km. For such a thin lithosphere, the delay
times from shear wave splitting measurements at the
surface are more consistent with an anisotropic mantle
in which foliation is vertical and lineation is horizontal.
8.3. LPO development
Olivine crystallographic axes are interpreted to track
finite strain orientations (e.g. McKenzie, 1979; Ribe,
1992) but at high strains may rotate towards the shear
direction faster than the finite strain axis (e.g. Nicolas and
Christensen, 1987; Zhang et al., 2000). In most situations,
the olivine a-axes are aligned with the lineation direction
(see however Mizukami and Wallis, 2005; Katayama et
al., 2005); thus, the observed fast shear wave polarization
direction should parallel the maximum finite strain axis
and reflect the orientation of olivine a-axes in the
lithospheric mantle. Presently, the orientation of fast
shear waves in the lithospheric mantle layer is sub-parallel
(b 10°) to the San Andreas fault (Fig. 7a), and we can
calculate the shear strain required to produce this subparallel alignment of olivine a-axes based on the
kinematics of the San Andreas fault system.
For a simplified calculation, we assume that the plate
boundary is deforming in simple shear (instead of
intranspression) and that the width of the deforming
Fig. 7. (a) Map of central California with shear wave splitting measurements. Arrow color reflects the two-layer anisotropy model discussed in the
text, where white arrows show modeled shear wave splitting for the deeper asthenospheric mantle layer and dark arrow reflect splitting in the upper
lithospheric mantle layer. Each arrow is scaled by delay time length and oriented parallel to the modeled fast direction. Modified from Hartog and
Schwartz (2001). (b) Two cross-sections through San Francisco Bay showing possible orientations of major fault strands within the crust based on
seismic reflection studies. The first model illustrates a horizontal decollément linking the different faults of the San Andreas fault system; the second
model shows vertical continuations of all major faults. The width of the shear zones in the mantle may either be discrete or reflect distributed shearing
at depth. See text for details. Adapted from Brocher et al. (1994) and Bürgmann (1997). (c) Three-dimensional model for the lithosphere in central
California, including the preferred orientation for olivine crystallographic axes to produce fault-parallel fast shear wave directions shown in (a). A
vertical foliation with horizontal lineation produces the greatest shear wave splitting values for vertically propagating waves based on the LPO
patterns from Coyote Lake basalt mantle xenoliths, thus requiring less than 100 km of anisotropic mantle to produce the observed delay times at the
surface. Modified from Teyssier and Tikoff (1998).
16
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
zone in the mantle is approximately 130 km wide, based
on the extent of fault-parallel polarization directions
observed at the surface (Fig. 1). To align the long-axis of
finite strain (and therefore olivine a-axes) within 10° of
the San Andreas fault requires a simple shear component
of γ ≈ 5 averaged over the whole zone (e.g. Tikoff and
Fossen, 1993). This translates into 650 km of faultparallel displacement that must be accommodated after
passage of the Mendocino triple junction at ∼ 10 Ma in
central California (Fig. 1). Plate reconstructions suggest
605 ± 67 km of net transform displacement on the San
Andreas fault system since 10.9 Ma (Dickinson, 1996) in
good agreement with this simple forward calculation.
This calculation suggests that there has been
sufficient fault-parallel displacement since 10 Ma to
align or reorient olivine a-axes in the mantle parallel to
the strike of the San Andreas fault simply through the
rotation of olivine crystals. Thus, the mantle beneath the
San Andreas fault may have had a pre-existing LPO that
was reoriented after strike-slip deformation began, or the
fabric may have been erased by the high temperatures
induced in the slab window trailing the triple junction. In
either case, LPO development related to San Andreas
fault kinematics may have been aided by dynamic
recrystallization, which can speed up the process of LPO
development (e.g. Tommasi et al., 1999; Zhang et al.,
2000) so that olivine a-axes may align parallel to the San
Andreas fault at lower finite strains.
8.4. Strike-slip shear zones in the mantle
In order to understand the lithospheric structure
beneath central California, it is necessary to combine all
available geophysical and geological information.
Based on the geophysical data sets, primarily seismic
reflection/refraction surveys and gravity and magnetic
measurements, there are several models for the manner
in which plate boundary deformation of the San Andreas
fault may be accommodated across major strike-slip
faults, as summarized below. Shear wave splitting
studies (e.g. Ozalaybey and Savage, 1995; Hartog and
Schwartz, 2001; Polet and Kanamori, 2002) and
xenolith LPO patterns provide additional information
that may help distinguish between the different models.
In the first type of model (Fig. 7b), plate boundary
deformation in the crust is decoupled from the mantle
lithosphere by horizontal mid-crustal detachments (e.g.
Namson and Davis, 1988). Applying this model to
central California suggests that there is a decollément
horizon in the mid-crust linking a shallowly dipping San
Andreas fault to vertical Hayward and Calaveras faults
(e.g. Furlong et al., 1989; Brocher et al., 1994; Holbrook
et al., 1996; Bürgmann, 1997). In this model, the
Hayward and Calaveras faults represent the major plate
boundary suture, and we would expect either a vertical
or horizontal foliation in the upper mantle beneath these
faults depending on which sections of the mantle were
sampled by xenoliths. An example of this type of model
is illustrated for San Francisco Bay (Fig. 7b), slightly to
the north of the Coyote Lake basalt locale.
In the second type of model (Fig. 7b), mantle
deformation may be accommodated by discrete narrow
shear zones beneath the major crustal faults (e.g. Savage
and Burford, 1970; Thatcher, 1989), in effect blurring
plate tectonic movements in the mantle (e.g. Tapponnier
et al., 1982). Variations in crustal thickness on either
side of the fault have been used as evidence for offset of
the Moho in both northern (Henstock et al., 1997; Hole
et al., 2000) and southern California (Zhu, 2000),
implying that the plate boundary structure continues as a
discrete feature in the upper mantle. Narrow zones,
perhaps less than 5 km wide (e.g. Wilson et al., 2004),
with vertical foliation and horizontal lineation would be
predicted in this style of model beneath the major faults
in the San Andreas fault system.
A third possibility is that deformation is distributed
across a broad region in the lower crust and upper mantle
(Fig. 7b; e.g. Prescott and Nur, 1981; McKenzie and
Jackson, 1983; Lamb, 1994; Bourne et al., 1998; Molnar
et al., 1999; Wilson et al., 2004). A distributed
deformation model predicts subhorizontal anisotropy in
the lower crust (e.g. Wilson et al., 2004) and vertical
foliations with horizontal lineations in the upper lithospheric mantle reflecting the predominance of strike-slip
plate motion in the region (e.g. Molnar et al., 1999).
Applying this model to central California suggests that the
major sub-parallel faults of the San Andreas fault system
merge at depth into a wider mantle shear zone (e.g.
Parsons and Hart, 1999) and is somewhat similar to the
previous model, except that much wider interconnected
shear zones are anticipated in the mantle.
It is likely that the upper mantle beneath the San
Andreas fault has a vertical foliation and horizontal
lineation based on the strike-slip kinematics and the
necessity to create sufficiently large shear wave splitting
delay times with a relatively thin lithosphere. Vertical
foliations are predicted in all of the above models,
although the width of strongly foliated areas varies in
each model. Because consistent fault-parallel shear
wave splitting is observed across a 130-km wide zone
in central California (Fig. 1), the petrofabric and shear
wave splitting data are most consistent with the third
model of broad distributed deformation, as shown by the
schematic three-dimensional cartoon in Fig. 7c.
S.J. Titus et al. / Tectonophysics 429 (2007) 1–20
Fig. 7c illustrates the merging of the Calaveras and San
Andreas faults at depth into a wider shear zone in the
mantle. In detail, the upper crust south of San Francisco
shows the major vertical strike-slip faults flanked by foldand-thrust belts. The mid-crustal foliations are highly
interpretive (e.g. Jones et al., 1994), but are consistent
with the seismic reflection data of Parsons and Hart
(1999) and are based on an interpretation of coupling
between the mid- and upper-crust predicted by numerical
models of distributed deformation at plate boundaries
(Wilson et al., 2004).
Mantle fabrics in this schematic cartoon are based on
shear wave splitting data (Fig. 7a) and illustrate the two
layers of anisotropic mantle (Ozalaybey and Savage,
1995; Hartog and Schwartz, 2001). EW striking
lineations with horizontal foliations are inferred for the
asthenospheric mantle, similar to the model proposed by
Silver and Holt (2002). These fabrics are not only
observed along the San Andreas fault system but also
east of the Great Valley (Ozalaybey and Savage, 1995;
Polet and Kanamori, 2002). Fault-parallel vertical
foliations with horizontal lineations are shown for the
upper lithospheric mantle. The spacing of foliations in
the model indicates the fabric strength, where more
closely spaced lines reflect stronger alignment of olivine
LPO, consistent with the longer delay times observed
closer to the San Andreas fault (Fig. 7a). This model is
specifically for central California, and may not hold for
other areas of California where fault-parallel shear wave
splitting fabrics have not been observed (Ozalaybey and
Savage, 1995; Hartog and Schwartz, 2001; Polet and
Kanamori, 2002).
9. Conclusions
The Coyote Lake basalt, which is located at the
intersection of the Hayward and Calaveras faults in
central California, erupted more than 6 m.y. after the
passage of the Mendocino triple junction. This delay in
time suggests that volcanism at Coyote Lake was not
related to the slab window trailing the migrating triple
junction. Entrained in this alkali-basalt are the only
known peridotite xenoliths from the mantle beneath the
San Andreas fault system. Six upper mantle xenoliths
were studied in detail by a combination of petrologic
techniques. The compositions of the xenoliths vary and
include Cr-diopside lherzolite, harzburgite, and one
composite sample of wehrlite and clinopyroxenite.
Mineral compositions within the xenoliths are generally
similar, except for the composite sample.
Xenolith equilibration temperatures are estimated to
have been 970–1100 °C, based on two-pyroxene and Al-
17
in-orthopyroxene geothermometry. A thermal model was
used in conjunction with the results from geothermometry
to estimate depths of 38 to 43 km for the xenoliths. Olivine
LPO patterns measured in five xenolith samples show
orthorhombic symmetry with strong point distributions,
which is consistent with deformation at relatively high
temperatures (N1000 °C). Calculated seismic anisotropy
values, based on the LPO patterns and modal compositions, indicate shear wave anisotropies of approximately
6%. This value is higher than those commonly calculated
for mantle xenoliths from a variety of tectonic settings and
higher than the 4% anisotropy used to calibrate many
shear wave splitting studies.
In order to impart the strong LPO patterns onto the
mantle, the Calaveras fault probably extends to at least
40 km into the lithospheric mantle as a shear zone. To
explain the observed shear wave splitting delay times in
a two-layer model requires 40–95 km of anisotropic
upper mantle, assuming 6% shear wave anisotropy and a
vertical foliation (c-axes vertical) with a horizontal
lineation (a-axes parallel to the fault). This orientation is
consistent with the kinematics in a primarily strike-slip
plate boundary and agrees with lithospheric thickness
estimates based on seismology and thermal modeling.
The 130-km wide area in central California that displays
fault-parallel shear wave splitting values is most
consistent with distributed deformation in the mantle,
in contrast to deformation restricted to narrow horizontal
or vertical mantle shear zones.
Acknowledgments
We thank Clark Johnson for samples from the Coyote
Lake basalt and Scott Giorgis for comments on an
earlier version of this manuscript. The manuscript was
improved by the comments of two anonymous
reviewers. This material is based upon work supported
under a National Science Foundation Graduate Research
Fellowship (Titus), a Packard Fellowship (Tikoff) and
NSF-EAR0337498 (Tikoff).
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