Download Lower crustal deformation beneath the central Transverse Ranges

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 107, NO. B7, 10.1029/2001JB000354, 2002
Lower crustal deformation beneath the central Transverse Ranges,
southern California:
Results from the Los Angeles Region Seismic Experiment
Nicola J. Godfrey,1,2 Gary S. Fuis,3 Victoria Langenheim,3 David A. Okaya,1
and Thomas M. Brocher3
Received 22 February 2000; revised 3 October 2001; accepted 8 October 2001; published 25 July 2002.
[1] We present a P wave velocity model derived from active source seismic data
collected during the 1994 Los Angeles Region Seismic Experiment. Our model extends
previously published upper crustal velocity models to mantle depths. Our model was
developed by both ray tracing through a layered model and calculating travel times
through a gridded model. It includes an 8-km-thick crustal root centered beneath the
surface trace of the San Andreas fault, north of the highest topography in the San
Gabriel Mountains. A simple mass balance calculation suggests that 36 km of northsouth shortening across the San Andreas fault in the central Transverse Ranges could
have formed this root. If north-south compression began when the ‘‘Big Bend’’ in the
San Andreas fault formed at 5 Ma, 36 km of shortening implies a north-south
contraction rate of 7.1 mm/yr across the central Transverse Ranges. If, instead, northsouth compression began when the Transverse Ranges formed at 3.4–3.9 Ma, 36 km of
shortening implies a contraction rate of 9.2–10.6 mm/yr. North of the San Andreas
fault, the Mojave Desert crust has a low-velocity (6.3 km/s) mid and lower crust and a
28-km-deep Moho. South of the San Andreas fault, beneath the Los Angeles and San
Gabriel Valley basins, there is a fast (6.6–6.8 km/s), thick (10–12 km) lower crust with
a 27-km-deep Moho. Farther south still, the lower crust of the Continental Borderland is
INDEX TERMS:
fast (6.6–6.8 km/s) and thin (5 km) with a shallow (22 km deep) Moho.
9350 Information Related to Geographic Region: North America; 7205 Seismology: Continental crust
(1242); 8122 Tectonophysics: Dynamics, gravity and tectonics; 8150 Tectonophysics: Plate boundary—
general (3040); KEYWORDS: Crustal root, P wave velocity, density, flexure, thermal
1. Introduction
[2] Southern California is actively deforming. One region
of interest is the Big Bend region of the San Andreas fault,
where there is a left step in the right-lateral transform
system (Figure 1a). This restraining bend, in principle,
results in approximately north-south compression across
the San Andreas fault [e.g., Atwater, 1970; Bird and
Rosenstock, 1984], although Weldon and Humphreys
[1986] point out that little contraction is actually currently
occurring across the San Andreas fault itself. Nevertheless,
the Transverse Ranges, located on both sides of the San
Andreas fault in the region of the Big Bend, responded, and
are still responding, to north-south compression by thrust
faulting, folding, and block rotations [Bailey and Jahns,
1954; Luyendyk et al., 1980; Jackson and Molnar, 1990;
Luyendyk, 1991]. Deformation in the upper crust is believed
1
Earth Science Department, University of Southern California, Los
Angeles, California, USA.
2
Now at Landmark EAME, Leatherhead, UK.
3
U.S. Geological Survey, Menlo Park, California, USA.
Copyright 2002 by the American Geophysical Union.
0148-0227/02/2001JB000354$09.00
ETG
to be decoupled from that of the lower crust and mantle
[Yeats, 1981; Nicholson et al., 1994; Shen et al., 1996;
Ryberg and Fuis, 1998; Kohler, 1999; Fuis et al., 2001a].
Until recently, it was thought that the San Gabriel Mountains (central Transverse Ranges) had little or no crustal
root [Hadley and Kanamori, 1977; Hearn, 1984; Fuis and
Mooney, 1990; Shearer and Richards-Dinger, 1997]. Kohler and Davis [1997], however, modeled a significant
crustal root beneath the surface trace of the San Andreas
fault in the San Gabriel Mountains using relatively highresolution teleseismic earthquake data collected during the
passive phase of the (1994) Los Angeles Region Seismic
Experiment. The Transverse Ranges are also associated
with a vertical, tabular body of anomalously fast mantle
velocities 60– 80 km wide that extend 100– 250 km into the
mantle lithosphere [Hadley and Kanamori, 1977; Raikes,
1980; Humphreys et al., 1984; Humphreys and Clayton,
1990; Zhao et al., 1996; Kohler, 1999]. Kohler [1999]
mapped the high-velocity mantle anomaly as a near vertical
feature directly beneath the crustal root. A high-velocity,
and presumably high-density, mantle anomaly may be
gravitationally unstable. It may sink into the lithosphere,
drawing lower crustal material down with it, thereby
enhancing crustal contraction [Fleitout and Froidevaux,
8-1
ETG
8-2
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
1982; Humphreys and Hager, 1990; Jackson and Molnar,
1990; Kohler, 1999].
[3] This paper presents results for the lower crust and
mantle beneath line 1 of the 1994 active phase of the Los
Angeles Region Seismic Experiment (LARSE) (Figure 1).
In a companion paper [Fuis et al., 2001a], we discuss the
upper and middle crust. In this paper we investigate what
the geometry of the lower crust and Moho can tell us about
lower crustal deformation in this active region, as well as
how it may relate to the previously documented mantle
anomaly. We use active source data from LARSE line 1, and
gravity data from the same profile, to model velocity and
density variations in the crust along a transect that extends
from the offshore Continental Borderland, across the Los
Angeles basin and San Gabriel Mountains, into the Mojave
Desert. We use our velocity and density models to determine how the transition from the offshore Pacific plate to
the undeformed North American plate is accommodated in
this complexly deforming region. The lower crust has
deformed by thickening into an 80-km-wide root centered
beneath the surface trace of the San Andreas fault. We
believe that this thickening in the lower crust is most likely
a ductile response to north-south compression across the
Transverse Ranges, where ductile lower crustal deformation
is decoupled from brittle upper crustal deformation.
2. Tectonic History
[4] At 30 Ma the Pacific and North American plates
came into contact, when the Pacific-Farallon spreading
ridge collided with the North American plate margin
[Atwater, 1970; Atwater and Molnar, 1973; Atwater,
1989]. Two triple junctions formed as a result of the
collision [Atwater, 1970, 1989], and they migrated away
from one another due to their geometry [McKenzie and
Morgan, 1969], allowing the San Andreas transform fault
system to develop between them. Between 24 and 18 Ma, a
period of extensional tectonics began. Extension manifested
itself in mid-Miocene rifting, widespread volcanism and
high heat flow [Henyey, 1976; Dokka, 1989; Tennyson,
1989; Wright, 1991]. Crustal extension coincided with
clockwise block rotations that may have been related to
the shallowing of Farallon subduction [Luyendyk, 1991]
and/or microplate capture [Nicholson et al., 1994].
[5] At 5 Ma, the opening of the Gulf of California
caused the Pacific-North American plate boundary to
migrate from offshore southern California onshore to the
east side of the Los Angeles basin [Atwater, 1970, 1989;
ETG
8-3
Nicholson et al., 1994]. The right-lateral San Andreas
transform fault either inherited the ‘‘Big Bend’’, or developed and/or exaggerated this shape, after the migration of
the plate boundary [Cox and Engebretson, 1985; Atwater,
1989]. Uplift of the Transverse Ranges began between 3.4
and 3.9 Ma in the vicinity of the Big Bend in the San
Andreas fault [Woodford et al., 1954; Wright, 1991].
3. Geology
[6] LARSE line 1 crosses several geological terranes and
tectonic features and passes close to the epicenters of three
major earthquakes. A magnitude 6.3 earthquake (Long
Beach earthquake) occurred on the Newport-Inglewood
fault in 1933 [Richter, 1958; Hauksson, 1987], a magnitude
5.9 earthquake (Whittier Narrows earthquake) occurred on a
blind thrust in 1987 [Hauksson et al., 1988], and a magnitude 5.8 earthquake (Sierra Madre earthquake) occurred in
1991 [Hauksson, 1994] (Figure 1c). Line 1 crosses the Los
Angeles region approximately perpendicular to the major
mapped faults [Jennings, 1977; Murphy et al., 1996]. Below
we briefly describe the major tectonic elements (from north
to south) relevant to LARSE line 1.
3.1. Mojave Desert
[7] The pre-Tertiary basement of the Mojave Desert
consists mainly of Mesozoic plutonic and volcanic rocks,
and Paleozoic sedimentary rocks [see Jennings, 1977]. The
plutonic rocks are the remnant of a continental Mesozoic
magmatic arc, continuous with the batholiths of the Sierra
Nevada and Peninsular Ranges. Generally, shallow basins
filled with Tertiary and Quaternary sediments overlie the
Mesozoic batholith [Dibblee, 1968; Jennings, 1977].
3.2. San Andreas Fault
[8] The modern San Andreas fault in southern California
became active at 5 Ma (see summary of Powell [1993, pp.
49– 50]. Between the Mendocino triple junction and the
Gulf of California, the San Andreas fault in most places has
a trend of 330° (Figure 1a). This is approximately parallel
to relative plate motion between the Pacific and North
American plates. In the Transverse Ranges in southern
California, however, the trend of the San Andreas fault
changes to 295°, a more east-west trend, which is oblique to
the relative plate motion. In this section of the San Andreas
fault, known as the ‘‘Big Bend’’, the San Andreas fault
system consists of a broad belt of northwest-trending strike
slip faults, of which the San Andreas fault itself is the
Figure 1. (opposite) (a) Index map of California showing the location of the LARSE experiment (black box). CA is
California, MX is Mexico, NV is Nevada, GF is the Garlock fault, MTJ is the Mendocino triple junction, and SAF is the
San Andreas fault. (b) Enlargement of black box shown in Figure 1a showing the location of the three LARSE profiles.
Shading is topography; white is sea level. (c) Detailed map of LARSE line 1 showing the sources and receivers used for the
model presented in this study. Geological features are labeled. Shading is topography; white is sea level. EF is the Elsinore
fault, GF is the Garlock fault, HF is the Hollywood fault, NIF is the Newport-Inglewood fault, PF is the Punchbowl fault,
PH are the Puente Hills, PVF is the Palos Verde fault, RF is the Raymond fault, SAF is the San Andreas fault, SaMoF is the
Santa Monica fault, SFV is the San Fernando Valley, SGF is the San Gabriel fault, SGV is the San Gabriel Valley, SJF is the
San Jacinto fault, SMF is the Sierra Madre fault, SMM are the Santa Monica Mountains, VT is the Vincent thrust fault, and
WF is the Whittier fault. Earthquakes close to line 1 are shown by black stars where 1 is the 1991 M5.8 Sierra Madre
earthquake, 2 is the 1987 M5.9 Whittier Narrows, 3 is the 1933 M6.3 Long Beach earthquake, 4 is the 1971 M6.7 San
Fernando earthquake, and 5 is the 1994 M6.7 Northridge earthquake.
ETG
8-4
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
northeasternmost [e.g., Powell, 1993] (Figure 1). Older
strands include the Punchbowl and San Gabriel faults
(Figure 1c) [Crowell, 1952; Ehlig, 1981; Powell, 1993].
3.3. San Gabriel Mountains (Central Transverse
Ranges)
[9] The Transverse Ranges are a late Cenozoic feature
resulting from a combination of compression across the
left-stepping bend in the San Andreas fault [Bailey and
Jahns, 1954], and upper crustal block rotations [Luyendyk
et al., 1980; Jackson and Molnar, 1990; Luyendyk, 1991].
The Transverse Ranges include the western Transverse
Ranges, the San Gabriel Mountains or central Transverse
Ranges, and the San Bernardino and Little San Bernardino Mountains or eastern Transverse Ranges. They are
east-west trending mountain ranges that cut across the
predominantly northwest-southeast trending structures of
the surrounding provinces [Bailey and Jahns, 1954]. The
Transverse Ranges are undergoing convergence as is
evident, (1) from faulting [e.g., Crowell, 1968; Morton
and Yerkes, 1987], (2) from the occurrence of moderate
historical reverse-fault earthquakes (the 1971 M6.7 San
Fernando, 1991 M5.8 Sierra Madre, and 1994 M6.7
Northridge earthquakes) (Figure 1c), and (3) from modern
north-south compression measured from geodetic data
[Feigl et al., 1993; Shen et al., 1996]. The country rocks
making up the San Gabriel Mountains south of the San
Andreas fault are divided into two plates by the Vincent
thrust fault (Figure 1c). The upper plate consists of
Precambrian metasedimentary rocks, and anorthosite and
Mesozoic plutonic rocks [Ehlig, 1981]. The lower plate
consists of Pelona Schist, a chiefly mica-quartz-albite
schist [Ehlig, 1968, 1981]. For a discussion of the
geometry of these two plates, the reader is referred to
Fuis et al. [2001b].
3.4. Los Angeles and San Gabriel Valley Basins
[10] The Los Angeles basin is a fault-bounded basin
whose bounding faults include the Palos Verde, Santa Monica, Hollywood, Raymond, and Whittier faults (Figure 1c).
The San Gabriel Valley basin, a subbasin of the Los Angeles
basin, is bounded on the southeast by the Puente Hills and on
the north by the Sierra Madre fault and the San Gabriel
Mountains (Figure 1c). The present-day Los Angeles basin
began evolving in the Late Miocene as a result of subsidence
between the right-oblique Whittier and Palos Verde fault
zones and the left-oblique Santa Monica fault system
[Wright, 1991]. Since the mid-Pliocene, the Los Angeles
ETG
8-5
basin has been deforming by crustal shortening accommodated by blind thrusts beneath the basin, such as the thrust
fault that moved in the 1987 Whittier narrows earthquake
[Hauksson et al., 1988]. The sediments in the Los Angeles
basin are more than 10 km thick [Yerkes et al., 1965; Fuis et
al., 2001a]. They form a northwest-southeast elongated
synclinorium, with its flanks folded and cut by Quaternary
active faults such as the Whittier-Elsinore fault system, and
the Newport-Inglewood and Palos Verde faults [Ziony and
Yerkes, 1985; Wright, 1991].
3.5. Peninsular Ranges
[11] The Peninsular Ranges batholith is exposed southeast
of the Los Angeles basin (Figure 1c), and it may extend
northwest to floor at least part of the Los Angeles basin. The
Peninsular Ranges batholith consists of Mesozoic plutons
(tonalite, granodiorite and gabbro), and Mesozoic and older
metamorphic rocks [e.g., Jennings, 1977; Silver et al., 1979].
3.6. Continental Borderland
[12] The Continental Borderland region is broken up into
northwest-trending blocks separated by strike slip faults
[e.g., Gibson et al., 1985]. The northeasternmost block,
which underlies LARSE line 1, consists of Cenozoic
volcanic and sedimentary rocks, and Catalina Schist, which
is thought to be coeval and cogenetic with the Franciscan
accretionary prism assemblage of the Coast Ranges and
Pelona Schist [Sorensen, 1984].
4. Data
[13] In 1994, the first active source phase of the Los
Angeles Region Seismic Experiment was carried out
[Brocher et al., 1995; Okaya et al., 1995; Fuis et al.,
1996; Murphy et al., 1996; ten Brink et al., 1996]. Three
multichannel seismic (MCS) reflection lines were recorded:
two (lines 1 and 2) trending approximately north to northeast and one trending approximately east-west (line 3)
(Figure 1b). Onshore receivers collinear with the MCS
profiles recorded air gun shots generated offshore by the
R/V Ewing (Figures 1 and 2) resulting in an onshoreoffshore reflection/refraction data set. Lines 1 and 2 also
included ocean bottom seismometers (OBSs) beneath the
MCS profiles. Line 1 was augmented by an onshore
reflection/refraction survey utilizing explosive sources
recorded by closely spaced receivers. The onshore explosions were not recorded by the OBSs. We present a model
along LARSE line 1 that incorporates all three data sets:
Figure 2. (opposite) (a) Example of onshore-offshore data from station 39 at model coordinate 165 km with picked
phases labeled. The data have been band-pass filtered and are shown as true amplitude plots. Example of OBS data is
presented in Figure 4 of ten Brink et al. [2000]. Example of onshore data is presented in Figures 2 and 3 of Fuis et al.
[2001b]. (b) –(k) Travel time plots for a selection of stations used in the model. Solid lines are model arrival times and
dashed lines are times picked from the data. (b) Ocean bottom seismometer C4 (model coordinate 17 km) recording Ewing
air gun shots. (c) Ocean bottom seismometer A1 (model coordinate 76 km) recording Ewing air gun shots. (d) Onshoreoffshore data recorded at station 21 (model coordinate 132 km). (e) Onshore-offshore data recorded at station 39 (model
coordinate 165 km). Shaded box shows extent of seismic data shown in Figure 2a. (f) Onshore-offshore data recorded at
station 47 (model coordinate 182 km). (g) Onshore-offshore data recorded at station 64 (model coordinate 216 km). (h)
Explosion 9450 (model coordinate 103 km) recorded on onshore stations. (i) Explosion 8050 (model coordinate 152 km)
recorded on onshore stations. (j) Explosion 8302 (model coordinate 178 km) recorded on onshore stations. (k) Explosion
8360 (model coordinate 184 km) recorded on onshore stations.
ETG
8-6
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
OBSs recording air gun shots, onshore recorders recording
air gun shots, and onshore recorders recording land shots.
[14] The offshore OBS data set recorded along line 1
consists of eight OBSs located beneath the line 1 MCS
profile [ten Brink et al., 1996] (Figure 1c). These OBSs
recorded air gun shots from the 8470 cubic inch tuned air
gun array of the R/V Ewing. These data show upper crustal
refractions (Pg) out to maximum offsets of 60 km. The
southernmost two instruments also recorded Moho reflections (PmP). Examples of the OBS data are given by ten
Brink et al. [2000, Figure 4]. These data primarily provide
information about the upper crust beneath the offshore
region.
[15] The onshore explosion data set recorded along line 1
consists of shot points every kilometer across the northern
San Gabriel Valley and San Gabriel Mountains, and shot
points spaced at 5 –50 km elsewhere [Murphy et al., 1996]
(Figure 1c). Shot sizes generally ranged from 5 to 907 kg,
although a 2722-kg shot was used at the northernmost end
of line 1. Seismic recorders were closely spaced (100 m) in
the region of dense shot locations and more sparsely spaced
(250 – 1000 m) elsewhere [Murphy et al., 1996]. A total of
649 instruments were used. Examples of the onshore
explosion data are given by Fuis et al. [2001b, Figures 2
and 3]. These data show crustal refractions out to maximum
offsets of 160 km, and crustal and Moho reflections. This
study is mainly concerned with lower crustal structure, and
we have not included the crustal reflections in the model
presented here. A detailed upper crustal model, including
the crustal reflections, is given by Fuis et al. [2001a].
[16] The onshore-offshore data set recorded along line 1
consists of the R/V Ewing air gun shots recorded by 84
onshore seismometers [Okaya et al., 1995]. The R/V
Ewing made multiple traverses along each MCS profile.
Some traverses were carried out at night to minimize
cultural noise, some had large shot intervals (60 s)
appropriate for OBS recording, while others had short
shot intervals (20 s) appropriate for higher-fold MCS
acquisition. We examined all the data and picked the best
traverse in terms of signal-to-noise ratio for each station,
trying where possible to use the 20-s shot interval data.
Signal-to-noise ratios on the line 1 onshore-offshore gathers are very low in most of the Los Angeles basin and San
Gabriel Valley, and are also low in parts of the Mojave
Desert, where offsets are large. We were only able to
correlate phases with confidence on 29 out of the 84
stations (Figures 1c and 2a). Data from these stations
and the land shots are used to create the model of the
lower crust presented in this study. The line 1 onshoreoffshore data set contains crustal refractions (Pg) out to
maximum crossover distances of 170 km. Pn is observed
out to offsets of 215 km, and strong PmP is observed on
most of the 29 gathers used.
5. Modeling
[17] We modeled line 1 as a 270-km-long, 2-D model,
extending from San Clemente Island (model coordinate
0 km) to the Mojave Desert (model coordinate 270 km).
We use a forward modeling approach that utilizes both
conventional ray tracing (using MacRay [Luetgert, 1992]),
and travel time calculations through a gridded model (based
on code developed by Hole [1992]). A single model of the
upper crust along the entire line 1 transect was obtained as
follows. For the onshore region, the model of Fuis et al.
[2001a] was used. For the offshore region, a starting model
was taken from ten Brink et al. [2000] and T. M. Brocher
(written communication). The composite model was checked
to make sure it was consistent with all three data sets
(offshore OBS data, onshore explosion data, and onshoreoffshore data) by interactive ray tracing, resulting in slight
adjustment to the offshore part of the model. The range of
uncertainties for the travel time picks are ±20 ms for Pg in
the onshore explosion data and offshore OBS data, ±30 ms
for PmP in the onshore-offshore data, ±40 ms for Pn in the
onshore-offshore data, ±40 ms for PmP in the onshore
explosion data and offshore OBS data, and ±75 ms for Pg
in the onshore-offshore data.
[18] We extended the upper crustal model to upper mantle
depths by including travel times (Pg, Pn and PmP) picked
from data recorded by the 29 best stations in the onshoreoffshore data set. Moho reflections (PmP) from all three
Figure 3. (opposite) Velocity model. Solid lines are velocity boundaries, dashed lines are velocity contours. Numbers are
velocity in km/s. Vertical exaggeration is 2, except in Figure 3b, which has a vertical exaggeration of 4. LAB is the Los
Angeles Basin, NIF is the Newport-Inglewood fault, PVF is the Palos Verde fault, SAF is the San Andreas fault, SGF is the
San Gabriel fault, SGM are the San Gabriel Mountains, SGV is the San Gabriel Valley, SMF is the Sierra Madre fault, VT is
the Vincent Thrust, and WF is the Whittier fault. (a) Upper 10 km of the final model constrained by crustal refractions
generated by offshore air gun shots recorded by ocean bottom instruments, and onshore stations recording onshore
explosions. Heavy lines are bathymetry, topography and the ocean surface. The upper crust between model coordinates 0
and 100 km is modified from ten Brink et al. [2000]. The upper crust between 100 and 270 km is slightly modified from
Fuis et al. [2001a]. Circles are ocean bottom seismometer locations, triangles are onshore-offshore receiver locations; stars
are explosion locations for gathers recording PmP arrivals. Thick shaded lines represent reflectors modeled from wideangle reflections in the explosion data set [Fuis et al., 2001a]. (b) Forward ‘‘layered’’ model. Light and medium shaded
regions in the upper crust are constrained by the Pg phase in the OBS data (I), the onshore-offshore data (II) and the
explosion data (III). The dark shaded region (IV) shows the constrained part of the mantle (onshore-offshore Pn data). The
heavy line on the Moho shows the parts of the Moho constrained by PmP bounce points from all data sets. Box with dashed
outline shows region of poorly constrained velocities. (c) Velocity model showing the lower crust (shaded) divided into five
regions (A– E), which are discussed in the text. The decoupling zone between brittle and ductile lower crustal deformation
is shown with a heavyweight dashed line. Vertical dashed line at model coordinate 50 is the boundary between the Inner
and Outer Continental Borderland. (d) Ray diagram for PmP from the two OBSs that have this phase. (e) Ray diagram for
PmP from the onshore-offshore data. (f) Ray diagram for PmP from the onshore explosion data.
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
ETG
8-7
ETG
8-8
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
data sets (southernmost two OBSs, 13 of the onshoreoffshore stations and data generated by 13 of the onshore
explosions) were used to constrain the depth to the Moho
and, to some extent, the velocity of the lower crust.
[19] The upper crust is quite complex, with large lateral
velocity variations beneath the entire length of the model
(Figure 3a) resulting from the fact that major faults juxtapose rocks of significantly different velocities. In the offshore region, substantial bathymetry results in large velocity
variations in the uppermost 2 –3 km of the model. Using
MacRay [Luetgert, 1992], it was not always possible to
generate rays through the complex layered model that
reached the surface at all the offsets represented in the data.
By rasterizing the layered model into a grid with 100 m by
100 m cells, however, we were able to generate travel times
through the gridded model using code modified from Vidale
[1988] and Hole [1992] to compare with the observed travel
times from the three data sets.
[20] Our final model for the whole crust is shown in
Figure 3b. The onshore explosion data constrain the uppermost 10 km between model coordinates 100 and 270 km.
This upper crustal part of the model is changed little from
that of Fuis et al. [2001a]. At the base of this constrained
region velocities reach 6.2 km/s in the Mojave Desert and
about 6.0 km/s in the San Gabriel Mountains. In the
offshore region (model coordinates 0 – 100 km), the eight
OBSs constrain the uppermost 8 km of the offshore region,
and velocities of 6.0 km/s are reached at 8 km depth. The
onshore-offshore data provide midcrustal constraints from
crustal refractions (Pg), and require velocities of 6.6 km/s at
15 km depth beneath the center of the model (model
coordinates 90 –130 km). Examples of model fit to the data
are shown in Figure 2.
[21] Lower crustal velocity information is obtained indirectly from mantle refractions (Pn) in the onshore-offshore
data set, and Moho reflections (PmP) in all three data sets.
Arrivals through the mantle and to the Moho have to travel
through the lower crust on their downward and upward
paths. This does not constrain velocity in the way that rays
turning in the lower crust would, but it does provide some
measure of average lower crustal velocity. For simplicity,
we assume that there are no velocity inversions with depth
below the regions constrained by the upper crustal refractions. We estimate the error in lower crustal velocity by
perturbating the lower crustal velocities until the misfit
between the modeled travel times and the picked travel
times exceeds the picking error.
[22] Beneath the northern part of the model (model
coordinates 180 to 270 km), velocities are relatively low.
We assume velocities in the mid and lower crust in this
region to be greater than or equal to 6.2 km/s, the deepestconstrained upper crustal velocity. Velocities producing the
best fit to Pn are 6.3±0.1 km/s for the lower crust, and
7.8±0.1 km/s for the uppermost mantle (Figure 3b). In the
center of the model (model coordinates 80 to 160 km), we
assume lower crustal velocities are at least 6.6 km/s, the
deepest-constrained midcrustal velocities. The best fitting
velocity at the base of the lower crust in this region is
6.8±0.05 km/s. The southern part of our model shows low
to intermediate (6.0 –6.6 km/s) velocities for the upper and
midcrust, similar to that of ten Brink et al. [2000]. We
model a thin (5 km) layer of intermediate-to high-velocity
(6.6±0.05 –6.8±0.05 km/s) material beneath the entire Continental Borderland. In contrast, ten Brink et al. [2000]
model high velocities (6.75 km/s) at the base of the crust
only beneath Santa Catalina Island, with lower (6.5 km/s)
velocities at the base of the crust elsewhere. Our model is
constrained by more data, including the onshore-offshore
data, than that of ten Brink et al. [2000]. The lower crustal
region between model coordinates 160 and 180 km (dashed
box, Figure 3b) is poorly constrained, as there are no PmP
arrivals from the Moho in this region (Figures 3e and 3f ).
Extending the uppermost-constrained velocity of 6.0 km/s
down to the Moho in this region results in the same model
travel times (within picking errors) as those generated using
a velocity of 6.3 km/s extrapolated from the lower crust to
the north, or even 6.5 km/s. We assign this region a velocity
of 6.25±0.25 km/s (dashed box, Figure 3b) This part of the
model is discussed in more detail in section 6.
[23] Using the best fitting lower crustal velocities of
6.3 km/s in the northern part of the model, and 6.6 –
6.8 km/s in the central and southern part of the model, we
modeled the depth to the Moho using the travel times of
Moho reflections (PmP). We modeled the PmP arrivals
from all three data sets. The data require the lower crust to
have a ‘‘root’’ between model coordinates 160 and 240 km,
which reaches a maximum depth of 35 km in a region
centered beneath the surface trace of the San Andreas fault
(Figure 3). This root represents a deepening of 6– 8 km
compared with crustal thickness to the north and south. In
addition, the Moho becomes shallower still south of model
coordinate 100 km, reaching a depth of 22 km beneath the
offshore region southwest of Santa Catalina Island (model
coordinate 50 km) (Figure 3). The exact depth to the Moho
is, or course, dependent on the lower crustal velocities used,
which, as discussed above, are not well resolved. Lower
crustal velocities as much as 0.2 km/s greater than modeled,
the maximum velocities allowed by the data, would result in
a 1-km-increase in the depth to the Moho. The general
shape of the Moho, however, is fairly robust.
[24] Using the Pn arrivals from the onshore-offshore
data set, and the PmP arrivals from all three data sets,
we can obtain a reasonably well constrained model with
regards to the lower crust and Moho. Despite the uncertainties in our modeling, there are two persistent features of
the lower crust and Moho. First, there is the crustal root,
which is located beneath the surface trace of the San
Andreas fault. Second, there is a difference in lower crustal
velocities and thickness between the southern (fast, thin),
central (fast, thick) and northern (slow, thick) parts of the
model (Figure 3b).
[25] In the onshore region (model coordinates 100 –
270 km), our velocity model is based on the upper and
midcrustal parts of the model of Fuis et al. [2001a]. It does
not, therefore, differ significantly from this model. Like the
model of Fuis et al. [2001a], the model of Lutter et al.
[1999] was derived from the onshore LARSE explosion
data, and is also quite similar to our model (Figures 4k to
4o). In the offshore region (model coordinates 0 – 100 km),
the mid and upper crust of our model is generally similar to
that of ten Brink et al. [2000], except in the Outer Continental Borderland, where it is somewhat faster (Figure 4b),
and beneath Santa Catalina Island, where it is somewhat
slower. Velocity models used by Ryberg and Fuis [1998],
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
ETG
8-9
Figure 4. Velocity model showing the locations (shaded bars and arrows) of velocity-depth profiles
shown in Figures 4b– 4p. ICB is the Inner Continental Borderland, LAB is the Los Angeles Basin, NIF is
the Newport-Inglewood fault, OCB is the Outer Continental Borderland, PVF is the Palos Verdes fault,
SAF is the San Andreas fault, SGF is the San Gabriel fault, SGM are the San Gabriel Mountains, SGV is
the San Gabriel Valley, SMF is the Sierra Madre fault, and WF is the Whittier fault. Dashed lines are
velocity contours (see Figure 3 for velocity values). (b) – (i) Velocity-depth curves for our model
compared to previously published models that show a poor match (mostly models earlier than 1977).
( j) –( p) Velocity-depth curves for our model compared to previously published models that show a good
match. Numbers on velocity depth curves are additional identifiers for model authors (see legend).
Kohler and Davis [1997], and Kohler [1999], which are
based on preliminary analysis of LARSE refraction data, are
superceded by our model.
[26] Velocity models developed from other data sets
include Shor and Raitt [1958] for the offshore region just
west of Santa Catalina Island (Figure 1), Hadley and
Kanamori [1977] for the San Gabriel Mountains, and
Hauksson and Haase [1997] and Hauksson [2000] for the
Los Angeles basin, San Gabriel Mountains and Mojave
Desert. The models of Hauksson and Haase [1997] and
Hauksson [2000] generally agree with our model (Figure 4).
The biggest discrepancies between our model and previously published models are seen in the models of Hadley
and Kanamori [1977] (Figures 4d to 4i) and Shor and Raitt
[1958] (Figures 4b and 4c), both of which are based on data
sets that are much sparser than the LARSE data set. Both
have significantly shallower Moho’s in the San Gabriel
Mountains and southern Mojave Desert (model coordi-
nate170 –220 km), and both are faster than our model in
the San Gabriel Mountains. The model of Hadley and
Kanamori [1977] is used for routine earthquake locations
in southern California. Their model is based on data from an
array of 120 stations within the Transverse Ranges and
Mojave Desert. They used Nevada Test Site shots and
quarry blasts to record data on three profiles, and they were
able to correlate Pg and Pn phases (no PmP), which were
used to derive a crustal and uppermost mantle velocity
model [Hadley and Kanamori, 1977]. They also created a
map of Moho depth based on P wave delay times, which
revealed a high-velocity ‘‘ridge’’ in the mantle beneath the
Transverse Ranges. Our model is derived from a much
denser 2-D array of instruments and many more sources,
enabling us to correlate Pg, Pn and PmP, and create a more
detailed velocity model in the same region.
[27] Zhu and Kanamori [2000] also derived a map of
Moho depth using teleseismic receiver functions recorded at
ETG
8 - 10
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
stations in southern California. Their method is independent
of P wave velocity, but depends on Vp/Vs ratios in the crust.
Their model shows a 37-km-deep Moho beneath the eastern
Transverse Ranges, a 31-km-deep Moho beneath the western Transverse Ranges, and a 29– 30-km-deep Moho, and
no crustal root, beneath the central Transverse Ranges,
which is the region of our model. Vp/Vs values are not well
known across the whole of southern California [Zhu and
Kanamori, 2000], and if P wave velocities vary widely in
the upper and lower crust, as demonstrated by our model, S
wave velocities may vary considerably too, creating the
potential for a larger range of Vp/Vs than that used by Zhu
and Kanamori [2000]. We believe that our data set, which is
of higher resolution than that used by Zhu and Kanamori
[2000], can be used to derive a well-constrained P wave
model for the central Transverse Ranges.
6. Density Models Along Line 1
[28] Bouguer gravity data along line 1 [Oliver et al.,
1980; Langenheim and Jachens, 1996] were modeled to
provide additional information about the nature of the lower
crust beneath line 1, in particular, beneath the San Gabriel
Mountains (model coordinates 160 to 180 km), which is the
most poorly constrained region of the velocity model.
Modeling of potential field data is particularly nonunique,
and we therefore use the velocity model as a starting point
to constrain the density models. We are chiefly concerned
with the long-wavelength features of the gravity data that
are primarily controlled by deeper density variations in the
crust. Short-wavelength misfits should be the result of upper
crustal density variations and thus be accounted for by the
velocity model. The first density model (Figure 5b) was
created by converting the velocity model to a density model
using the velocity-density functions of Nafe and Drake
[1957] and Gardner et al. [1974] for velocities <5.5 km/s.
For velocities >5.5 km/s, the velocity-density relationship of
Christensen and Mooney [1995] for ‘‘all rocks except
monomineralic rocks’’ was used. The resulting calculated
gravity does not match the observed gravity data in the
ETG
8 - 11
region between the Los Angeles Basin and the Mojave
Desert (Figure 5a). The mismatch between calculated and
observed gravity curves is as high as 50 mGal using the
Nafe-Drake function [Nafe and Drake, 1957], is even larger
using the Gardner function [Gardner et al., 1974], and
occurs above the portion of the lower crust that we wish
to constrain. Furthermore, the calculated gravity curve for
the Mojave Desert region is consistently 20 – 30 mGal
higher than the observed gravity data. The mismatch, based
on its 40-km wavelength, is primarily caused by to midcrustal rocks.
[29] We next evaluated whether the density-velocity
function used is appropriate for lithologies exposed in the
San Gabriel Mountains and Mojave Desert. McCaffreePellerin and Christensen [1998] found that the density of
the Pelona schist (2.72 g/cm3) is higher than would be
predicted from its seismic velocity (5.9 km/s; 2.68 g/cm3)
using the function of Christensen and Mooney [1995]. We
performed an additional 50 density measurements of Pelona
Schist samples, and found an average density of 2.75 g/cm3.
The crystalline rocks in the San Gabriel Mountains have an
average density of 2.73 g/cm3 [Langenheim, 1999]. For the
Mojave Desert, the average density of pre-Cenozoic rocks is
2.67 g/cm3. On the northern side of the San Andreas fault,
the Christensen and Mooney [1995] function predicts higher
densities than measured on the basis of velocity. Part of the
discrepancy in the calculated and observed values near the
San Andreas fault is resolved by recognizing that extensive
fracturing most likely causes the low velocities in the
velocity model. Fracturing does not affect density as
severely as it does velocity. A study of granitic rocks in a
well about 1 km from the San Andreas fault in the Gabilan
Range (450 km to the northwest of line 1) found that
saturated macrocracks produced 3– 7% reduction in rock
density, and a 30– 50% reduction in seismic velocity [Stierman and Kovach, 1979].
[30] We used the above information to adjust the densities
of the upper crust in the Mojave Desert and San Gabriel
Mountains. We also adjusted (1) the densities in the offshore
region to honor bathymetry, (2) the southwest boundary of
Figure 5. (opposite) (a) Gravity data plotted against gravity calculated through two density models along line 1. Solid line
is observed Bouguer gravity. Short-dashed line is the gravity calculated through a density model derived from the velocity
model using the Nafe-Drake curve [Nafe and Drake, 1957] for velocities <5.5 km/s and the Christensen and Mooney
[1995] velocity-density curve for velocities >5.5 km/s, and a density of 2.7 g/cm3 in the southern part of the crustal root,
equivalent to a velocity of 6 km/s. Long-dashed line is the gravity calculated through a density model derived from the
velocity model using the Gardner curve [Gardner et al., 1974] for velocities <5.5 km/s and the Christensen and Mooney
[1995] velocity-density curve for velocities >5.5 km/s, and a density of 2.7 g/cm3 in the southern part of the crustal root,
equivalent to a velocity of 6 km/s. (b) Density model (derived from the velocity model) for the curves shown in Figure 5a.
Dashed lines show density contours and solid lines show density boundaries. Numbers are densities in g/cm3. (c) Gravity
data plotted against gravity calculated through two density models along line 1. Solid line is observed Bouguer gravity.
Short-dashed line is the gravity calculated through a density model derived from the velocity model using the conversions
discussed in the text, and a density of 2.7 g/cm3 in the southern part of the crustal root (shaded region in Figure 5d),
equivalent to a velocity of 6 km/s. Medium-dashed line is the gravity calculated through a density model derived from the
velocity model using the conversions discussed in the text, and a density of 2.8 g/cm3 (equivalent to a velocity of 6.3 km/s)
in the uppermost part of the crustal root and 2.9 g/cm3 (equivalent to a velocity of 6.5 km/s) in the lowermost part of the
crustal root (shaded region in Figure 5d). Long-dashed line is the gravity calculated through a density model derived from
the velocity model using the conversions discussed in the text, and a density of 2.9 g/cm3 (equivalent to a velocity of 6.5
km/s) in the crustal root (shaded region in Figure 5d). (d) Density model for the curves shown in Figure 5c. Dashed lines
show density contours and solid lines show density boundaries. Numbers are densities in g/cm3. Shaded region is the region
being tested in this model.
ETG
8 - 12
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
the Los Angeles basin to honor borehole data, and (3)
velocities in the mantle to reflect the high-velocity mantle
root beneath the Transverse Ranges [Hadley and Kanamori,
1977; Kohler, 1999]. Modeling the seismic data only
allowed us to determine that the velocity within the southern
part of the crustal root (model coordinates 160-180 km) is
between 6.0 and 6.5 km/s. Using the Christensen and
Mooney [1995] velocity-density conversion, we tested densities of 2.7 g/cm3 (6.0 km/s) and 2.9 g/cm3 (6.5 km/s). The
best fitting model has a density of 2.9 g/cm3 in the southern
part of the crustal root, which corresponds to a velocity of
6.5 km/s.
7. Geological Interpretation of the Velocity Model
[31] For details in interpretation of the complex upper
crust, the reader is referred to ten Brink et al. [2000] and
Fuis et al. [2001a, 2001b].
[32] In our model, the middle and lower crust have a
simpler velocity structure than the upper crust, although this
may also result from the decrease in resolving power of
seismic data with increasing depth. We suggest that beneath
the Los Angeles basin, the mid and lower crust are most
likely decoupled from the upper crust at the base of the
seismogenic zone (about 15 km depth) [Hauksson, 2000].
Beneath the San Gabriel Valley and San Gabriel Mountains
we follow the interpretation of Ryberg and Fuis [1998], that
the lower crust is decoupled from the upper crust along a
décollement, which dips northward from a depth of 15 km
beneath the southern San Gabriel Valley to 20 km depth at
the San Andreas fault, where it terminates. It is also possible
that the lower crust is decoupled from the upper crust in the
Mojave Desert at a zone of midcrustal reflectors at 20 km
depth [Fuis et al., 2001a]. We include in ‘‘lower crust’’ all
material beneath the interpreted decoupling zone (bold
dashed line, Figure 3c), and we divide the lower crust into
five regions (A – E) based on our velocity model (shaded
regions, Figure 3c).
7.1. Region A, Continental Borderland
[33] The Continental Borderland can be divided into the
Outer and Inner Continental Borderland (vertical dashed
line at model coordinate 50, Figure 3c). The Outer Continental Borderland, south of model coordinate 50, has a
low to intermediate velocity (6.0 – 6.6 km/s) midcrust (down
to 18 km depth), and a relatively thin (5 km) intermediate-to
high-velocity (6.6 – 6.8 km/s) lower crust. The Moho is at
about 22 km depth (region A, Figure 3c). The Inner
ETG
8 - 13
Continental Borderland, between model coordinates 50
and 100, includes a step in the thickness of both the middle
and lower crust at about model coordinate 60 km, where the
lower crust becomes thicker and the mid-crust becomes
thinner. The Moho starts to deepen toward the north at
model coordinate 50 km, from 22 to 27 km depth (region A,
Figure 3c). In the interpretation of ten Brink et al. [2000],
the entire Continental Borderland crust is composed of
Catalina Schist. Our lower crustal velocities are too fast,
however, to reasonably interpret as Catalina Schist (an
equivalent to Pelona Schist) (Figure 6a). Our velocity
structure for the Continental Borderland is consistent with
Catalina Schist underlain by greenschist facies basaltic
rocks at the base of the crust (Figure 6a). Such a structure
is interpretable as a fossil subduction zone, where the
basaltic rocks are tectonically underplated, similar to other
such zones imaged along the coast of California (see Fuis
and Mooney [1990]; Miller et al. [1992]; Howie et al.
[1993]; Page and Brocher [1993]; Beaudoin et al. [1996];
Fuis et al. [1996]; Henstock et al. [1996]; Holbrook et al.
[1996]; Godfrey et al. [1997], summarized by Fuis [1998];
and Ryberg and Fuis [1998]). Such structure is also interpretable as a core complex, similar to the hypothesis of ten
Brink et al. [2000], but where the basaltic rocks are
magmatically underplated.
7.2. Region B, Los Angeles and San Gabriel
Valley Basins and Their Basement
[34] The lower crust (below 15 km depth) beneath the
Los Angeles and San Gabriel Valley basins has intermediate
to high velocities (6.6 – 6.8 km/s) similar to the offshore
region, but the region of these intermediate to high velocities is greater in thickness (10 – 12 km versus 5 km)
(compare regions A and B, Figure 3c). The crust is about
27 km thick in region B. The midcrust (above 15 km depth)
beneath the Los Angeles Basin is interpretable as either
Peninsular Ranges batholithic rocks (e.g., granites, diorites
and diabase), or Catalina Schist (an equivalent to Pelona
Schist) (Figure 6b). The lower crust may be either the mafic
root of the Peninsular Ranges batholith or underplated
basaltic rocks (Figure 6b). Aeromagnetic data are consistent
with Peninsular Ranges rocks in the subsurface, which are
magnetic, rather than Catalina Schist, which is not magnetic
[Langenheim, 1999].
7.3. Regions C and D, Crustal Root
[35] The midcrust of the San Gabriel Mountains, above
the reflective zone imaged by Ryberg and Fuis [1998] at
Figure 6. (opposite) Velocity-depth profiles for our velocity model (shaded regions are envelopes enclosing these
profiles) plotted with velocity-depth profiles for various lithologies [Christensen and Mooney, 1995] (solid black lines). The
data from Christensen and Mooney [1995] are corrected for temperature with depth in an average heat flow regime. The
range of values measured for the anisotropic Pelona Schist is shown by the patterned region. Pelona Schist velocity data is
taken from Godfrey et al. [2000]. The long-dashed line shows average Pelona Schist values (two fast velocities and one
slow velocity) [Godfrey et al., 2000]. The short-dashed line shows average Pelona Schist values taken from McCaffreePellerin and Christensen [1998]. (a) Continental Borderland. (b) Los Angeles and San Gabriel Valley basins and their
basement. The light shaded region shows basin velocities. The darker shaded region shows basement velocities. (c) Mid
and lower crust of the San Gabriel Mountains and the crustal root. Velocities taken from region C using the preferred
velocity of 6.5 km/s in the crustal root (Figure 3c) are shown with medium shading. Velocities taken from region C using
the other possible velocities of 6.0 –6.5 km/s in the crustal root (Figure 3c) are shown with light shading. Velocities taken
from region D (Figure 3c) are shown with dark shading (narrow block). (d) The Mojave Block.
ETG
8 - 14
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
15– 20 km depth, south of the San Andreas fault has a
relatively low velocity (5.8 – 6.0 km/s). In contrast, the
midcrust beneath the Mojave Desert, north of the San
Andreas fault, has a low to intermediate velocity (6.2 –
6.3 km/s). Velocities within the lower crust (beneath the
reflective zone (décollement) imaged by Ryberg and Fuis
[1998]) may also change across the deep projection of the
San Andreas fault, which extends into the center of the 35km-deep crustal root. Velocities in the southern part of the
crustal root, south of the projection of the San Andreas fault,
(region C, Figure 3c) are not well constrained (between 6.0
and 6.5 km/s), although gravity data suggest densities
consistent with a faster velocity, close to 6.5 km/s. Such a
velocity is slower than the lower crust of region B. The
lower crust in the northern part of the root, north of the
projection of the San Andreas fault, (region D, Figure 3c),
has a velocity of 6.3 km/s, possibly lower than region C,
and similar to region E. The rock in region D is consistent
with intermediate to granitic rocks, or possibly Pelona
Schist (Figure 6c).
7.4. Region E, Mojave Block
[36] North of the crustal root, mid-and lower crustal
velocities remain fairly low (6.3 km/s) down to the Moho,
which rises to 28 km depth. The relatively low velocity
of region E, like that of region D, is consistent with
intermediate to granitic composition, and/or Pelona Schist
(Figure 6d). It should be noted, however, that the mid-and
upper crustal velocities in the Mojave Desert are not
compatible with Pelona Schist (Figure 6d) [Lutter et al.,
1999; Fuis et al., 2001a], and models of aeromagnetic data
suggest that nonmagnetic Pelona Schist can only exist at
depths below 8 km [Langenheim, 1999].
8. Discussion
[37] The crustal root in our model is centered beneath the
surface trace of the San Andreas fault, and is therefore north
of the topographic high along line 1, the San Gabriel
Mountains. A similar crustal root is modeled by Kohler
and Davis [1997] and Kohler [1999], although the roots in
those models are deeper because those authors assumed
lower crustal velocities that were significantly different
from those presented here. Airy compensation is apparently
not occurring beneath the San Gabriel Mountains, but the
crustal velocity model is more complex than is assumed in
Airy compensation. Sheffels and McNutt [1986] modeled
southern California as consisting of two plates that partially
support the crustal load. Their model invokes a rigid northern plate (Mojave block), and a weaker southern plate
(Continental Borderland and Los Angeles Basin region).
The boundary between these plates is along the southern
side of the San Gabriel Mountains, south of the San Andreas
fault. Sheffels and McNutt [1986] propose that the Transverse Ranges are regionally compensated by the stiff northern plate, which is underthrust by the weaker southern plate.
If the southern plate is actually subducting, or sinking
vertically into the asthenosphere, it may be related to the
mantle anomalies modeled by various authors [Hadley and
Kanamori, 1977; Humphreys et al., 1984; Humphreys and
Clayton, 1990; Kohler, 1999]. In particular, Kohler [1999]
modeled a narrow, vertical anomaly that extends 200 km
into the lithosphere, centered directly beneath the surface
trace of the San Andreas fault and crustal root beneath
LARSE line 1. A gravitationally unstable sinking plate may
enhance crustal contraction at the plate boundary [Fleitout
and Froidevaux, 1982] and help form a crustal root by
drawing lower crustal material into the mantle above the
sinking Pacific plate (Figure 7). In section 9 we calculate
how much crustal contraction is implied by the formation of
the crustal root in our model.
9. Implications of the Crustal Root
[38] It has been suggested that the upper crust is deforming by rigid block rotations that are decoupled from lower
crustal and mantle deformation [Nicholson et al., 1994;
Shen et al., 1996; Ryberg and Fuis, 1998; Kohler, 1999].
We hypothesize that lower crustal deformation has been
occurring in a ductile manner based on the thickening of
the lower crust in the region of the crustal root. We
perform a simple mass balance to determine the amount
of crustal shortening required to form this root. We assume
the depth of decoupling is at the base of the seismogenic
layer, at about 15 km depth, beneath the Los Angeles
basin. We then assume it follows the décollement interpreted by Ryberg and Fuis [1998] and Fuis et al. [2001a]
beneath the southern and central San Gabriel Mountains
(Figure 3c). This décollement dips north from 15 km
depth beneath the southern San Gabriel Valley to 20 km
depth beneath the central San Gabriel Mountains. We also
assume the midcrustal reflections at 20 km depth beneath
the northern San Gabriel Mountains and Mojave Desert
form a décollement in which the relative sense of motion is
opposite, and symmetrical, to motion on the décollement
beneath the southern and central San Gabriel Mountains
[Fuis et al., 2001a]. We assume the depth of decoupling
did not alter over the time of convergence. We assume for
simplicity that all the thickening of the lower crust in
regions C and D (Figure 3c) is due to north-south compression between the Pacific and North American plates,
and that none of the thickening has resulted from out-ofplane motion, which makes our convergence estimate a
maximum. Our calculation assumes that the lower crust of
the Mojave block (region E, Figure 3c) and the mid and
lower crust beneath the Los Angeles and San Gabriel
Valley basins (region B, Figure 3c) are undeformed. We
further assume that the San Andreas fault is a vertical plate
boundary, although Fuis et al. [2001a] have modeled it
with a steep (83°) northward dip. Our simple calculation
estimates how much horizontal motion (contraction) is
required to thicken the bounding blocks (Figure 8a) into
the root seen in our model (Figure 8b).
[39] The additional rock added to the base of region C
(III, Figure 8b) amounts to an excess of 62 km2, which we
assume was derived from region B (Peninsular Ranges and
other lower crust) (I, Figure 8b), which has an average
thickness of 12 km (Figure 8b). An excess of 62 km2
derived from a starting thickness of 12 km implies 5.2 km
of shortening from the south (Figure 8). The material added
to the base of region D (IV, Figure 8b) amounts to an excess
of 238 km2, which we assume was derived from region E
(Mojave block) (V, Figure 8b). We assume the Mojave
block (region E, Figure 3c) is decoupled from the mid and
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
ETG
8 - 15
Figure 7. Schematic diagram showing the weak Pacific lithosphere sinking vertically into the
asthenosphere at the Pacific – North America plate boundary [Fleitout and Froidevaux, 1982]. The lower
crust has thickened into a crustal root (black) as a result of ductile deformation in response to horizontal
compression across the plate boundary, enhanced by the Pacific plate sinking into the lithosphere. SAF is
the San Andreas fault.
upper crust at 20 km depth, and has a 28-km-deep Moho. It
therefore has a thickness of 8 km. An excess of 238 km2
derived from starting thickness of 8 km implies 29.8 km of
shortening from the north (Figure 8). There is also an
additional amount of shortening due to the kinking of the
lower crust in region C (II, Figure 8b). The kink produces
an additional 0.6 km of shortening (Figure 8). Our simple
calculation suggests that a total of about 35.6 km (0.6 km
plus 5.2 km from the south and 29.8 km from the north) of
shortening has taken place across the San Gabriel Mountains and San Andreas fault.
[40] If the Big Bend region of the San Andreas fault
formed at 5 Ma, 35.6 km of shortening implies a northsouth convergence rate of 7.1 mm/yr in this region. If this
convergence instead began with the uplift of the Transverse
Ranges at 3.4 – 3.9 Ma, then the rate is 9.2 to 10.6 mm/yr.
The maximum estimate for the component of motion
perpendicular to the San Andreas fault based on geometrical
considerations for the Big Bend region is 15 mm/yr [Bird
and Rosenstock, 1984]. Geodetically measured rates perpendicular to the San Andreas fault in the Transverse
Ranges vary from ‘‘very little’’ to 10 mm/yr [Weldon
and Humphreys, 1986; Cheng et al., 1987; Lisowski et al.,
1991; Shen, 1991; Donnellan et al., 1993; Feigl et al., 1993;
Shen et al., 1996]. Geodetic rates, of course, measure the
amount of upper crustal shortening, and this amount might
be quite different from that in the lower crust and mantle.
Weldon and Humphreys [1986] have shown that the upper
crust accommodates compression largely by rotations and
‘‘escape’’ tectonics. There is little actual convergence thickening and uplift between individual blocks. Whether the
calculated convergence rate is 7.1 or 10.6 mm/yr, the rates
we estimate are reasonable compared to modeled and/or
measured rates. The crustal root could, therefore, result from
convergence. In our calculation, most of the crustal root was
formed from Mojave block material. This suggests that the
lower crust of the Mojave block is weaker, and more easily
able to flow into the root, than lower crustal material from
the south [cf. Sheffels and McNutt, 1986]. The inference of a
weaker lower crust beneath the Mojave block is consistent
with its relatively low velocity. Quartz-rich (low velocity)
rocks should be weaker than feldspar-and/or olivine-rich
(higher velocity) rocks.
10. Conclusions
[41] Data from the 1994 active phase of the Los Angeles
Region Seismic Experiment have been modeled to show an
ETG
8 - 16
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
Figure 8. Schematic diagram of regions B, C, D, and E from Figure 3c showing how the mass balance
calculation was performed. Not to scale. LAB is the Los Angeles basin, SAF is the San Andreas fault,
SGM are the San Gabriel Mountains, and SGVB is the San Gabriel Valley basin. (a) Prior to
compression. Mid and lower crust of the blocks before compression (shaded) showing the geometry used
in the calculation. The upper boundary of these blocks is the decoupling zone that separates deformation
by rigid block rotation in the upper crust (white) from ductile deformation in the lower crust (shaded). (b)
Present-day. Mid and lower crust of undeformed blocks (I and V) with additional material gained by the
lower crustal root as a result of kinking (II) and compression (III and IV). The upper boundary of these
blocks is the decoupling zone that separates deformation by rigid block rotation in the upper crust (white)
from ductile deformation (shaded) in the lower crust.
approximately 8-km-thick lower crustal root centered
beneath the surface trace of the San Andreas fault. This
root is about 80 km wide, and consists of material with
different velocities and densities on either side of the deep
extrapolation of the San Andreas fault. This may be additional evidence that the San Andreas fault is vertical or
steeply dipping, and penetrates the entire crust. We assume
that and lower crustal deformation associated with north-
south compression across the San Andreas fault in the Big
Bend region is decoupled in the midcrust. Beneath the
offshore Continental Borderland and Los Angeles basin
regions we assume decoupling occurs at the base of the
seismogenic crust (15-km depth). Beneath the San Gabriel
Valley and San Gabriel Mountains, we assume decoupling
occurs along a décollement interpreted from reflection data
[Ryberg and Fuis, 1998; Fuis et al., 2001a] at 15 –20 km
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
depth. Beneath the Mojave Desert we assume the décollement coincides with reflections at 20 km depth [Fuis et
al., 2001a]. In our model, the lower crust deforms in a
ductile manner by thickening at the plate boundary. A
simple mass balance calculation suggests about 36 km of
north-south shortening has occurred across the San Gabriel
Mountains and San Andreas fault. If shortening began with
either the uplift of the Transverse Ranges, or the formation
of the Big Bend in the San Andreas fault at 3.4– 5 Ma,
shortening of 36 km implies a convergence rate of 7.1– 10.6
mm/yr.
[42] Our model also shows three regions of dramatically
different lower crustal velocities and thickness. North of
the San Andreas fault, beneath the Mojave Desert, lower
crustal velocities are relatively low (6.3 km/s), and the
lower crust is about 8 km thick, although the lower crust is
apparently indistinguishable from the midcrust in this
region. South of the San Andreas fault, beneath the San
Gabriel Mountains, San Gabriel Valley and Los Angeles
basin, lower crustal velocities are much higher (6.6 –
6.8 km/s), and the lower crust is thicker (10 – 12 km).
Further south, beneath the Continental Borderland, lower
crustal velocities remain high (6.6 – 6.8 km/s), but the
lower crust is thin (5 km). This has implications for other
models and earthquake locations that currently assume a
constant lower crustal velocity across the entire Los
Angeles region.
[43] Acknowledgments. We would like to thank all those involved
in collecting the 1994 LARSE data sets including the U.S. Geological
Survey, the Southern California Earthquake Center (SCEC), the Geological Survey of Canada, IRIS/PASSCAL, and all the volunteers who
helped in the field. These people are listed by Murphy et al. [1996]. We
also thank R. Hawman, G. Spence and an Associate Editor of JGR for
their thoughtful and helpful reviews. The data are archived at the IRIS
Data Management Center and are available at http://www.iris.edu. Generic
Mapping Tools (GMT) software [Wessel and Smith, 1991] was used to
make some of the figures presented in this paper. This research was
supported by the Southern California Earthquake Center (SCEC). SCEC
is funded by NSF Cooperative Agreement EAR-8920136 and USGS
Cooperative Agreements 14-08-0001-A0899 and 1434-HQ-97AG01718.
SCEC contribution 536.
References
Atwater, T., Implications of plate tectonics for the Cenozoic tectonic history of western North America, Geol. Soc. Am. Bull., 81, 3513 – 3536,
1970.
Atwater, T., Plate tectonic history of the northeast Pacific and western North
America, in The Geology of North America, vol. N, The Eastern Pacific
Ocean and Hawaii, edited by E. L. Winterer, D. M. Hussong, and R. W.
Decker, pp. 21 – 72, Geol. Soc. of Am., Boulder, Colo., 1989.
Atwater, T., P. Molnar, Relative motion of the Pacific and North American
plates deduced from sea-floor spreading in the Atlantic, Indian and South
Pacific oceans, in Tectonic Problems of the San Andreas Fault System,
edited by R. L. Kovach and A. Nur, pp. 136 – 148, Stanford Univ. Press,
Palo Alto, Calif., 1973.
Bailey, T. L., R. H. Jahns, Geology of the Transverse Range province,
southern California, in Geology of Southern California, edited by R. H.
Jahns, Bull. Calif. Div. Mines, 170, 83 – 106, 1954.
Beaudoin, B. C., et al., The transition from slab to slabless: Results from the
1993 Mendocino Triple Junction Seismic Experiment, Geology, 24, 195 –
199, 1996.
Bird, P., and R. W. Rosenstock, Kinematics of present crust and mantle
flow in southern California, Geol. Soc. Am. Bull., 95, 946 – 957, 1984.
Brocher, T. M., R. W. Clayton, K. D. Klitgord, R. G. Bohannon, R. Sliter,
J. K. McRaney, J. V. Gardner, J. B. Keene, Multichannel seismicreflection profiling on the R/V Maurice Ewing during the Los Angeles
Regional Seismic Experiment (LARSE), California, U.S. Geol. Surv.
Open File Rep., 95-228, 70 pp., 1995.
Cheng, A., D. D. Jackson, and M. Matsu’ura, Aseismic crustal deformation
ETG
8 - 17
in the Transverse Ranges of southern California, Tectonophysics, 144,
159 – 180, 1987.
Christensen, N. I., and W. D. Mooney, Seismic velocity structure and
composition of the continental crust: A global review, J. Geophys.
Res., 100, 9761 – 9788, 1995.
Cox, A., and D. C. Engebretson, Change in motion of Pacific Plate at 5 myr
BP, Nature, 313, 472 – 474, 1985.
Crowell, J. C., Probable large lateral displacement on the San Gabriel fault,
southern California, AAPG Bull., 36, 2026 – 2035, 1952.
Crowell, J. C., Movement histories of faults in the Transverse Ranges and
speculations on the tectonic history of California, in Proceedings of the
Conference on Geologic Problems of the San Andreas Fault, edited by W.
R. Dickinson and A. Grantz, Stanford Univ. Publ. Geol. Sci., 11, 323 –
341, 1968.
Dibblee, T. W. J., Displacements of the San Andreas fault system in the San
Gabriel, San Bernadino and San Jacinto Mountains, southern California,
in Proceedings of the Conference on Geologic Problems of the San
Andreas Fault, edited by W. R. Dickinson and A. Grantz, Stanford Univ.
Publ. Geol. Sci., 11, 260 – 278, 1968.
Dokka, R. K., The Mojave extensional belt of southern California, Tectonics, 8, 363 – 390, 1989.
Donnellan, A., B. H. Hager, R. W. King, and T. A. Herring, Geodetic
measurement of deformation in the Ventura basin region, southern California, J. Geophys. Res., 98, 21,727 – 21,739, 1993.
Ehlig, P. L., Causes of distribution of Pelona, Rand and Orocopia schists
along the San Andreas and Garlock faults, in Proceedings of the Conference on Geologic Problems of the San Andreas Fault, edited by W. R.
Dickinson and A. Grantz, Stanford Univ. Publ. Geol. Sci., 11, 294 – 306,
1968.
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 W. G. Ernst, pp. 253 – 283, PrenticeHall, Old Tappan, N. J., 1981.
Feigl, K. L., et al., Space geodetic measurement of crustal deformation in
central and southern California, 1984 – 1992, J. Geophys. Res., 98,
21,677 – 21,712, 1993.
Fleitout, L., and C. Froidevaux, Tectonics and topography for a lithosphere
containing density heterogeneities, Tectonics, 1, 21 – 56, 1982.
Fuis, G. S., West margin of North America—A synthesis of recent seismic
transects, Tectonophysics, 288, 265 – 292, 1998.
Fuis, G. S., and W. D. Mooney, Lithospheric structure and tectonics from
seismic-refraction and other data, in The San Andreas Fault System,
California, edited by R. E. Wallace, U.S. Geol. Surv. Prof. Pap., 1515,
207 – 236, 1990.
Fuis, G. S., et al., Images of the crust beneath southern California will aid
study of earthquakes and their effects, Eos Trans. AGU, 77, 173, 176,
1996.
Fuis, G. S., T. Ryberg, N. J. Godfrey, D. A. Okaya, and J. M. Murphy,
Crustal structure and tectonics from the Los Angeles basin to the Mojave
Desert, southern California, Geology, 29, 15 – 18, 2001a.
Fuis, G. S., T. Ryberg, W. J. Lutter, and P. L. Ehlig, Seismic mapping of
shallow fault zones in the San Gabriel Mountains, from the Los Angeles
Region Seismic Experiment, southern California, J. Geophys. Res., 106,
6549 – 6568, 2001b.
Gardner, G. H., L. W. Gardner, and A. R. Gregory, Formation velocity and
density: The diagnostic basis for stratigraphic traps, Geophysics, 39,
770 – 780, 1974.
Gibson, J. D., D. G. Howell, G. S. Fuis, J. G. Vedder, Pacific Abyssal Plain
to the Rio Grande Rift, DNAG Continent/Ocean Transect, vol. 5, 233 pp.,
Geol. Soc. of Am., Boulder, Colo., 1985.
Godfrey, N. J., B. C. Beaudoin, and S. L. Klemperer, Ophiolitic basement
to the Great Valley forearc basin, northern California, from seismic and
gravity data: Implications for crustal growth at the North American continental margin, Geol. Soc. Am. Bull., 109, 1536 – 1562, 1997.
Godfrey, N. J., N. I. Christensen, and D. A. Okaya, Anisotropy of schists:
Contribution of crustal anisotropy to active-source seismic experiments
and shear wave splitting observations, J. Geophys. Res., 105, 27,991 –
28,008, 2000.
Hadley, D., and H. Kanamori, Seismic structure of the Transverse Ranges,
California, Geol. Soc. Am. Bull., 88, 1469 – 1478, 1977.
Hauksson, E., Seismotectonics of the Newport-Inglewood fault zone in the
Los Angeles basin, southern California, Bull. Seismol. Soc. Am., 77,
539 – 561, 1987.
Hauksson, E., The 1991 Sierra Madre earthquake sequence in southern
California: Seismological and tectonic analysis, Bull. Seismol. Soc.
Am., 84, 1058 – 1074, 1994.
Hauksson, E., Crustal structure and seismicity distribution adjacent to the
Pacific and North America plate boundary in southern California, J.
Geophys. Res., 105, 13,875 – 13,903, 2000.
Hauksson, E., and J. S. Haase, Three-dimensional Vp and Vp/Vs velocity
ETG
8 - 18
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
model of the Los Angeles basin and central Transverse Ranges, California, J. Geophys. Res., 102, 5423 – 5453, 1997.
Hauksson, E., et al., The 1987 Whittier Narrows earthquake in the Los
Angeles metropolitan area, California, Science, 239, 1409 – 1412, 1988.
Hearn, T. M., Pn travel times in southern California, J. Geophys. Res., 89,
1843 – 1855, 1984.
Henstock, T. J., A. Levander, and M. W. Group, Images of the Mendocino
and San Andreas transforms in the Mendocino triple junction region:
Continuous seismic profiling across the continental margin, Eos Trans.
AGU, 77(46), Fall Meet. Suppl., F741, 1996.
Henyey, T. L., Heat flow and tectonic patterns on the southern California
borderland, in Aspects of the Geologic History of the California Continental Borderland, edited by D. G. Howell, pp. 427 – 440, Am. Assoc. of
Pet. Geol., Pac. Sect., Bakersfield, Calif., 1976.
Holbrook, W. S., T. M. Brocher, U. S. ten Brink, and J. A. Hole, Crustal
structure of a transform plate boundary: San Francisco Bay and the
central California continental margin, J. Geophys. Res., 101, 22,311 –
22,334, 1996.
Hole, J. A., Nonlinear high-resolution three-dimensional seismic travel time
tomography, J. Geophys. Res., 97, 6553 – 6562, 1992.
Howie, J. M., K. C. Miller, and W. U. Savage, Integrated crustal structure
across the south central California margin: Santa Lucia escarpment to the
San Andreas fault, J. Geophys. Res., 98, 8173 – 8196, 1993.
Humphreys, E. D., and R. W. Clayton, Tomographic image of the southern
California mantle, J. Geophys. Res., 95, 19,725 – 19,746, 1990.
Humphreys, E. D., and B. H. Hager, A kinematic model for the recent
development of southern California crust and upper mantle, J. Geophys.
Res., 95, 19,747 – 19,762, 1990.
Humphreys, E. D., R. W. Clayton, and B. H. Hager, A tomographic image
of mantle structure beneath southern California, Geophys. Res. Lett., 11,
625 – 627, 1984.
Jackson, J., and P. Molnar, Active faulting and block rotations in the western Transverse Ranges, California, J. Geophys. Res., 95, 22,073 –
22,087, 1990.
Jennings, C. W., Geologic map of California, scale 1:750,000, Calif. Div. of
Mines and Geol., Sacramento, 1977.
Kohler, M. D., Lithospheric deformation beneath the San Gabriel Mountains in the southern California Transverse Ranges, J. Geophys. Res., 104,
15,025 – 15,041, 1999.
Kohler, M. D., and P. M. Davis, Crustal thickness variations in southern
California from Los Angeles Regional Seismic Experiment passive phase
teleseismic travel times, Bull. Seismol. Soc. Am., 87, 1330 – 1344,
1997.
Langenheim, V. E., Gravity and aeromagnetic models along the Los Angeles Region Seismic Experiment (line 1), California, report, 21 pp., U.S.
Geol. Surv., Menlo Park, Calif., 1999.
Langenheim, V. E., R. C. Jachens, Gravity data collected along the Los
Angeles Regional Seismic Experiment (LARSE) and preliminary model
of regional density variations in basement rocks, southern California,
report, U.S. Geol. Surv., Menlo Park, Calif., 1996.
Lisowski, M., J. C. Savage, and W. H. Prescott, The velocity field along the
San Andreas fault in central and southern California, J. Geophys. Res.,
96, 8369 – 8389, 1991.
Luetgert, J. H., MacRay: Interactive two-dimensional seismic raytracing
for the Macintosh, U.S. Geol. Surv. Open File Rep., 92-356, 43 pp.,
1992.
Lutter, W. L., G. S. Fuis, C. H. Thurber, and J. Murphy, Tomographic
images of the upper crust from the Los Angeles basin to the Mojave
Desert, California: Results from the Los Angeles Region Seismic Experiment, J. Geophys. Res., 104, 25,543 – 25,565, 1999.
Luyendyk, B., A model for Neogene crustal rotations, transtension, and
transpression in southern California, Geol. Soc. Am. Bull., 103, 1528 –
1536, 1991.
Luyendyk, B., M. J. Kamerling, and R. R. Terres, Geometric model for
Neogene crustal rotations in southern California, Geol. Soc. Am. Bull.,
91, 211 – 217, 1980.
McCaffree-Pellerin, C. L., and N. I. Christensen, Interpretation of crustal
seismic velocities in the San Gabriel-Mojave region, southern California,
Tectonophysics, 286, 253 – 271, 1998.
McKenzie, D. P., and W. J. Morgan, Evolution of triple junctions, Nature,
224, 125 – 133, 1969.
Miller, K. C., J. M. Howie, and S. D. Ruppert, Shortening within underplated oceanic crust beneath the central California margin, J. Geophys.
Res., 97, 19,961 – 19,980, 1992.
Morton, D. M., R. F. Yerkes, Introduction, in Recent Reverse Faulting in
the Transverse Ranges, California, edited by D. M. Morton and R. F.
Yerkes, U.S. Geol. Surv. Prof. Pap., 1339, 1 – 5, 1987.
Murphy, J. M., et al., Report for explosion data acquired in the 1994 Los
Angeles Region Seismic Experiment (LARSE94), Los Angeles, California, U.S. Geol. Surv. Open File Rep., 96-536, 120 pp., 1996.
Nafe, J. E., and C. L. Drake, Variation with depth in shallow and deep water
marine sediments of porosity, density and the velocities of compressional
and shear waves, Geophysics, 22, 523 – 552, 1957.
Nicholson, C., C. C. Sorlien, T. Atwater, J. C. Crowell, and B. P. Luyendyk,
Microplate capture, rotation of the western Transverse Ranges, and initiation of the San Andreas Transform as a low-angle fault system, Geology, 22, 491 – 495, 1994.
Okaya, D. A., J. Bhowmik, G. S. Fuis, J. M. Murphy, M. C. Robertson, A.
Chakraborty, M. L. Benthien, K. Hafner, and J. J. Norris, Report for
airgun data acquired at onshore stations during the 1994 Los Angeles
Region Seismic Experiment (LARSE), California, U.S. Geol. Surv. Open
File Rep., 96-297, 224 pp., 1995.
Oliver, H. W., S. L. Robbins, W. L. Rambo, R. F. Sikora, Bouguer gravity
map of California, Chico sheet, Calif. Div. of Mines and Geol., Sacramento, 1980.
Page, B. M., and T. M. Brocher, Thrusting of the central California margin
over the edge of the Pacific plate during the transform regime, Geology,
21, 635 – 638, 1993.
Powell, R. E., Balanced palinspastic reconstruction of pre-late Cenozoic
paleogeology, southern California: Geologic and kinematic constraints
on evolution of the San Andreas fault system, in The San Andreas Fault
System: Displacement, Palinspastic Reconstruction, and Geologic Evolution, edited by R. R. Powell, R. J. Weldon, and J. C. Matti, Mem. Geol.
Soc. Am., 178, 1 – 106, 1993.
Raikes, S. A., Regional variations in upper mantle structure beneath southern California, Geophys. J. R. Astron. Soc., 63, 187 – 216, 1980.
Richter, C. F., Elementary Seismology, 768 pp., W. H. Freeman, New York,
1958.
Ryberg, T., and G. S. Fuis, The San Gabriel Mountains bright reflective
zone: Possible evidence of young mid-crustal thrust faulting in southern
California, Tectonophysics, 286, 31 – 46, 1998.
Shearer, P. M., and K. B. Richards-Dinger, Estimating crustal thickness in
southern California by stacking PmP arrivals, J. Geophys. Res., 102,
15,211 – 15,224, 1997.
Sheffels, B., and M. McNutt, Role of subsurface loads and regional compensation in the isostatic balance of the Transverse Ranges, California:
Evidence for intracontinental subduction, J. Geophys. Res., 91, 6419 –
6431, 1986.
Shen, Z., Regional tectonic deformation in southern California, inferred
from terrestrial geodesy and the Global Positioning System, Ph.D. thesis,
175 pp., Univ. of South., Calif., Los Angeles, 1991.
Shen, Z.-K., D. D. Jackson, and B. X. Ge, Crustal deformation across and
beyond the Los Angeles basin from geodetic measurements, J. Geophys.
Res., 101, 27,957 – 27,980, 1996.
Shor, G. G., R. W. Raitt, Seismic studies in the southern Continental Borderland, paper presented at the 20th International Geological Congress,
Section IX (Geophysical Applications), Mexico, 1958.
Silver, L. T., H. P. Taylor, B. Chappell, Some petrological, geochemical and
geochronological observations of the Peninsular Ranges batholith near
the international border of the U.S.A. and Mexico, in Mesozoic Crystalline Rocks: Peninsular Range Batholith and Pegmatites, Point Sal
Ophiolite, edited by P. L. Abbott and V. R. Todd, pp. 83 – 110, San Diego
State Univ., Dep. of Earth Sci., San Diego, Calif., 1979.
Sorensen, S. S., Petrology of basement rocks of the California Continental
Borderland and the Los Angeles basin, Ph.D. thesis, 423 pp., Univ. of
Calif., Los Angeles, 1984.
Stierman, D. J., and R. L. Kovach, An in situ velocity study: The Stone
Canyon well, J. Geophys. Res., 84, 672 – 678, 1979.
ten Brink, U. S., R. M. Drury, G. K. Miller, T. M. Brocher, D. Okaya,
Los Angeles Region Seismic Experiment (LARSE)—Offshore seismic
refraction data, U.S. Geol. Surv. Open File Rep., 96-27, 29 pp.,
1996.
ten Brink, U. S., J. Zhang, T. M. Brocher, D. A. Okaya, K. D. Klitgord, and
G. S. Fuis, Geophysical evidence for the evolution of the California inner
continental borderland as a metamorphic core complex, J. Geophys. Res.,
105, 5835 – 5857, 2000.
Tennyson, M. E., Pre-transform early Miocene extension in western California, Geology, 17, 792 – 796, 1989.
Vidale, J., Finite difference calculation of travel times, Bull. Seismol. Soc.
Am., 78, 2062 – 2076, 1988.
Weldon, R., and E. Humphreys, A kinematic model of southern California,
Tectonics, 5, 33 – 48, 1986.
Wessel, P., and W. H. F. Smith, Free software helps map and display data,
Eos Trans. AGU, 72, 441, 445 – 446, 1991.
Woodford, A. O., J. E. Schoellhamer, J. G. Vedder, R. F. Yerkes, Geology
of the Los Angeles basin, in Geology of Southern California, edited by
R. H. Jahns, Bull. Calif. Div. Mines, 170, 65 – 81, 1954.
Wright, T. L., Structural Geology and tectonic evolution of the Los Angeles
basin, California, in Active Margin Basins, edited by K. T. Biddle, AAPG
Mem., 52, 35 – 134, 1991.
GODFREY ET AL.: LOWER CRUST BENEATH CENTRAL TRANSVERSE RANGES
Yeats, R. S., Quaternary flake tectonics of the California Transverse
Ranges, Geology, 9, 16 – 20, 1981.
Yerkes, R. F., T. H. McCulloh, J. E. Schoellhamer, and J. G. Vedder,
Geology of the Los Angeles basin, California—An introduction, U.S.
Geol. Surv. Prof. Pap., 420-A, 57 pp., 1965.
Zhao, D., H. Kanamori, and E. Humphreys, Simultaneous inversion of local
and teleseismic data for the crust and mantle structure of southern California, Phys. Earth Planet. Inter., 93, 191 – 214, 1996.
Zhu, L., and H. Kanamori, Moho depth variation in southern California
from teleseismic receiver functions, J. Geophys. Res., 105, 2969 – 2980,
2000.
Ziony, J. I., R. F. Yerkes, Evaluating earthquake and surface-faulting potential, in Evaluating Earthquake Hazards in the Los Angeles Region: An
ETG
8 - 19
Earth Science Perspective, edited by J. I. Ziony, U.S. Geol. Surv. Prof.
Pap., 1360, 43 – 91, 1985.
T. M. Brocher, G. S. Fuis, and V. Langenheim, U.S. Geological Survey,
345 Middlefield Road, MS 977, Menlo Park, CA 94025. (brocher@
usgs.gov; [email protected]; [email protected])
N. J. Godfrey, Landmark EAME, Hill Park South, Springfield Drive,
Leatherhead, Surrey KT22 7NL, UK. ([email protected])
D. A. Okaya, Department of Geological Sciences, University of Southern
California, Los Angeles, CA 90089-0740, USA. ([email protected])