JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 109, B02312, doi:10.1029/2002JB002345, 2004
Crustal fabric in the Tibetan Plateau based on waveform
inversions for seismic anisotropy parameters
Heather Folsom Sherrington and George Zandt
Department of Geosciences, University of Arizona, Tucson, Arizona, USA
Andrew Frederiksen
Department of Geological Sciences, University of Manitoba, Winnipeg, Manitoba, Canada
Received 12 December 2002; revised 20 October 2003; accepted 10 November 2003; published 24 February 2004.
[1] The Tibetan Plateau has the thickest continental crust on Earth, and fabrics within the
crust that are anisotropic to seismic waves may provide clues to how it reached such
extreme proportions and how it is currently deforming. Waveform modeling using a global
minimization inversion technique applied to receiver functions computed from 11 stations
spanning the north-south length of the eastern plateau has yielded a suite of crustal
models that include anisotropy. These models suggest that the Tibetan crust contains
4–14% anisotropy at different depths that is likely a result of both fossil fabrics and more
recent deformation. All models contain anisotropy in the surface layer, and for most
stations the alignment of the slow symmetry axis suggests a relationship with crustal
fabrics associated with E-W trending thrust faults or sutures. Middle to lower crustal
anisotropy is present at most stations with a fast axis trending N-S to NW-SE in the south,
nearly E-W in the central plateau, and N-S to NE-SW in the northern plateau. This pattern
appears consistent with recent ductile deformation due to both topographically induced
flow and to boundary forces from subducting lithosphere at the northern and southern
margins of the plateau. The orientations of crustal anisotropy determined for most stations
in this study are significantly different from shear wave splitting fast polarization
INDEX TERMS: 7205
directions, implying distinct deformation in the crust and mantle.
Seismology: Continental crust (1242); 7260 Seismology: Theory and modeling; 8102 Tectonophysics:
Continental contractional orogenic belts; 8159 Tectonophysics: Rheology—crust and lithosphere; 9320
Information Related to Geographic Region: Asia; KEYWORDS: Tibetan Plateau, crust, seismology, anisotropy,
modeling, inversion
Citation: Sherrington, H. F., G. Zandt, and A. Frederiksen (2004), Crustal fabric in the Tibetan Plateau based on waveform
inversions for seismic anisotropy parameters, J. Geophys. Res., 109, B02312, doi:10.1029/2002JB002345.
1. Introduction
[2] The Tibetan Plateau is composed of an amalgamation
of tectonostratigraphic terranes along roughly E-W trending
sutures (Figure 1). Collision between India and Eurasia,
beginning approximately 50 Ma, is believed to be a major
driving force behind plateau development and current deformation, as approximately 2500 km of convergence have
taken place since the collision [Molnar and Tapponnier,
1975; Patzelt et al., 1996]. The thick crust of the Tibetan
Plateau probably contains fabrics that are a result of deformation involved in uplifting the plateau to its current average
elevation of 5 km. A number of models for plateau formation
include underthrusting of Indian crust, and perhaps mantle
lithosphere, beneath the plateau or injection of some or all of
the Indian crust within the Eurasian crust [Ni and Barazangi,
1984; Zhao and Morgan, 1987; DeCelles et al., 2002].
Underthrusting Indian lithosphere may have exerted a basal
shear stress that contributed to the development of a metaCopyright 2004 by the American Geophysical Union.
0148-0227/04/2002JB002345$09.00
morphic fabric in some part of the overlying Tibetan crust.
Models involving pure shear thickening of Eurasian crust
and/or mantle lithosphere may similarly imply formation of
a large-scale crustal fabric in response to compressive
stresses generated by collision [England and Houseman,
1986; Molnar et al., 1993]. Ongoing crustal flow may be
active to maintain a uniform plateau elevation during convergence and/or in response to lateral pressure gradients
produced by the high topography of the plateau [Bird, 1991;
Clark and Royden, 2000; Vanderhaeghe and Teyssier, 2001;
Shen et al., 2001]. In addition, some geological evidence
suggests that significant shortening may have actually
occurred before India collided with Eurasia, generally
involving north or south directed subduction and collision
between microcontinents [Yin and Harrison, 2000, and
references therein], and some fabrics present in the modern
Tibetan crust could conceivably be due to Mesozoic or early
Cenozoic tectonics. Hence the regional mapping of these
crustal-scale fabrics could provide important constraints for
how the crust of the plateau has behaved throughout its
history or how and where it is deforming at present.
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Figure 1. Regional tectonic map of the Tibetan Plateau, showing major suture zones, terranes, fault
systems, and the Gangdese magmatic arc [after Yin and Harrison, 2000; Tapponnier et al., 2001].
Elevations are after Fielding et al. [1994]. Locations of 1991– 1992 PASSCAL stations and GSN station
LSA are shown with triangles. Double arrows show shear wave splitting fast polarization directions from
McNamara et al. [1994]. Major sutures are shown as thick gray lines (ISZ, Indus-Tsangpo suture zone;
BSZ, Banggong suture zone; JS, Jinsha suture; KS, Kunlun suture). Major faults are shown as thin black
lines (MFT, Main Frontal thrust; MBT, Main Boundary thrust; MCT, Main Central thrust; STDS, South
Tibetan Detachment System). Heavy dashed line encloses the central Tibet conjugate fault zone [after
Taylor et al., 2003].
[3] Portions of the Earth’s crust that contain tectonic
fabrics defined by certain seismically anisotropic minerals
can yield a bulk anisotropic response when traversed by
seismic waves [Babuska and Cara, 1991; Rabbel and
Mooney, 1996]. The receiver function method involves
using distant earthquakes as sources of seismic P-to-S
converted waves that sample the crust directly beneath a
recording station. Converted phases are sensitive to the
presence of anisotropy on vertical scales of a few hundred
meters to several kilometers and horizontal scales of several
tens of kilometers. This technique is particularly well suited
to studying crustal anisotropy, as radial and transverse
component receiver functions show systematic variations
that are a function of the back azimuth, or angle between a
recording station and seismic source, and can be used to
discern the presence and orientation of crustal-scale fabrics.
Furthermore, unlike shear wave splitting studies, the receiver
function method can provide more definitive depth
constraints on anisotropy and can be utilized for frequencies
sensitive to thin (<1 km) layers with sufficient velocity
contrasts [Zandt et al., 2003]. During the 1991 – 1992
PASSCAL broadband seismic experiment on the plateau,
a large number of teleseisms were recorded at 11 stations
that span the north-south length of the eastern plateau.
These data, along with data from Global Seismic Network
(GSN) station LSA, are used in global minimization inver-
sions to determine the details of crustal anisotropy within
the plateau.
2. Sources of Crustal Anisotropy
[4] While mantle anisotropy is widely accepted to be due
to the anisotropic properties of aligned olivine crystals in the
upper mantle, potential causes of crustal anisotropy are less
straightforward. However, a number of geologically feasible
scenarios could result in relatively uniform seismic anisotropy at large scales within the crust. Among these are aligned
microcracks, perhaps developed due to a nonhydrostatic
stress field in the shallow crust, and alignment of mineral
grains in a large body of rock [Rabbel and Mooney, 1996].
These geological situations can be reasonably approximated
with hexagonal symmetry, with a unique fast or slow
symmetry axis and uniformly slow or fast velocities, respectively, for seismic wave propagation directions in the plane
perpendicular to that axis [Weiss et al., 1999] (Figure 2).
[5] Cracks are likely to be most important in the upper
crust, where pressures are low enough to allow them to
remain open. The general importance of cracks exclusively
at low pressures is supported by experimental studies of
rock anisotropy at increasing pressure, which have found
that dry cracks close at around 100– 200 MPa and no longer
contribute to bulk rock anisotropy [Barruol and Kern, 1996;
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Figure 2. Illustration of phase velocity surfaces (oblate and prolate ellipsoids) for hexagonal symmetry
anisotropy and possible geologic explanations for unique slow and unique fast axis cases. The diagram at
the top shows the convention used to define the orientation of anisotropy symmetry axes in three
dimensions, with a trend measured clockwise from north in a horizontal plane and a plunge measured
downward from that horizontal plane.
Okaya et al., 1995; Kern and Wenk, 1990; Ji and Salisbury,
1993; Siegesmund et al., 1989; Weiss et al., 1999]. A series
of aligned microcracks can be modeled using hexagonal
symmetry anisotropy with a unique slow axis of symmetry
(Figure 2).
[6] While cracks may be important at shallow depths,
numerous studies have found that aligned minerals are the
most likely cause of anisotropy in rocks at middle to lower
crustal depths [Rabbel et al., 1998; Weiss et al., 1999;
Burlini and Fountain, 1993; Ji et al., 1993; Siegesmund et
al., 1989; Ji and Salisbury, 1993; Kern and Wenk, 1990;
Barruol and Kern, 1996]. Surprisingly, even though a
variety of minerals are common in rocks, and most minerals
are considerably anisotropic as single crystals [Babuska and
Cara, 1991], a small number of minerals seem to dominate
the bulk anisotropy of most rock types. In particular, micas,
such as biotite and muscovite, typically have cleavage
planes aligned with a foliation, while amphibole and sillimanite commonly have crystallographic axes aligned with a
lineation in strained rocks (Figure 2); these minerals are
often the primary cause of bulk rock anisotropy, even if they
are not the most abundant minerals in a rock [Barruol and
Kern, 1996; Kern and Wenk, 1990; Ji et al., 1993; Ji and
Salisbury, 1993; Siegesmund et al., 1989; Weiss et al., 1999;
Burlini and Fountain, 1993; Mainprice and Nicolas, 1989].
Quartz and feldspar, while displaying considerable anisotropy as single crystals, typically do not orient in such a way
in a rock as to cause significant net anisotropy [Barruol and
Kern, 1996; Ji et al., 1993; Kern and Wenk, 1990; Ji and
Salisbury, 1993; Weiss et al., 1999]. Thus quartzites and
high-grade metamorphic rocks such as granulites, which are
mostly devoid of hydrous minerals, i.e., micas and amphiboles, may only display a weak anisotropy [Christensen and
Mooney, 1995; Weiss et al., 1999; Kern and Wenk, 1990; Ji et
al., 1993; Barruol and Kern, 1996]. On the other hand, rocks
rich in micas and amphiboles, such as slates, schists, and
certain mylonites exhibit the highest degree of anisotropy.
3. Receiver Functions
3.1. Data and Receiver Function Computation
[7] Earthquake seismograms used in this analysis were
chosen from a data set collected as part of a Sino-American
collaborative PASSCAL broadband experiment onthe Tibetan
Plateau during 1991 –1992 and from data collected at GSN
station LSA from 1991 to present. Data from PASSCAL
station LHSA, colocated with LSA, were not used in this
study because the sampling rate of LHSA data is too low for
the frequency content of receiver functions investigated here.
The PASSCAL experiment included 11 stations that span the
north-south length of the plateau (Figure 1), and well over a
hundred teleseismic events were recorded at most stations.
The back azimuth distribution of events is somewhat biased
toward the east due to the active seismicity of the western
Pacific, but overall the back azimuth coverage is reasonably
good due to the fortuitous positioning of the Tibetan Plateau
with respect to global seismicity. Clean teleseismic events
were chosen from this data set for receiver function computation at distance ranges between 30 and 90, with a few
events slightly closer or further from a station if a clear direct
P arrival could be identified.
[8] Receiver functions were computed using an iterative,
time domain deconvolution algorithm developed by Ligorria
and Ammon [1999] that constructs a receiver function as a
sum of Gaussian pulses. A Gaussian filter value of 5 was
used, roughly corresponding to a low-pass filter with a
maximum frequency slightly greater than 2 Hz. Receiver
functions were then stacked according to back azimuth and
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ray parameter to reduce noise and to decrease the number of
waveforms being modeled in the inversion procedure.
Stacking was performed by sorting the data into back
azimuth bins with a width of 10 and further dividing these
bins into ray parameter bins with a width of 0.02 s/km. Each
waveform was sorted into only one bin, and averaging
waveforms within a bin yielded around 20 – 40 average
receiver functions for each station that sample the crust in
a circular swath beneath the station.
3.2. Anisotropy and Receiver Functions
[9] A receiver function waveform has an initial arrival
with P polarity, while the remainder of the waveform
consists of S waves that have converted from P waves at
velocity contrasts in the subsurface. In the absence of lateral
heterogeneity or anisotropy, teleseismic converted S waves
should remain in the source-receiver plane (radial-horizontal) and have exclusively SV (radial-vertical) particle motion
(Figure 3). Energy would then only be present on the radial
component. The presence of lateral heterogeneity, such as
dipping layers, or anisotropy will result in rotation of energy
out of the source-receiver plane and conversion from P to
both SV and SH particle motion at velocity contrasts in the
subsurface for most source-receiver back azimuths. This
effect appears as azimuthally varying amplitudes on both
the radial and transverse record section plots. Dipping layers
and anisotropy that does not exhibit rapid spatial variation
should both yield patterns of receiver function amplitudes
and arrival times on both the radial and transverse components that are a systematic function of source-receiver back
azimuth, whereas isotropic structures should yield energy
only on the radial component, in a pattern that does not vary
for different back azimuths (Figure 3). Laterally heterogeneous structure other than dipping layers would be expected
to produce radial and transverse component amplitudes that
vary in a highly erratic way as a function of back azimuth.
Distinguishing between dipping layers and anisotropy
as explanations for systematic back azimuth variation of
receiver function appearance is somewhat nonunique
[Savage, 1998; Levin and Park, 1997a; Peng and
Humphreys, 1997]. However, matching observed transverse
amplitudes with dipping layers can be challenging, typically
requiring unreasonably steep dips or large velocity contrasts
between layers [Leidig and Zandt, 2003; Savage, 1998]. For
simplicity in modeling and interpretation, this study deals
only with modeling of anisotropy in receiver functions.
[10] Several studies have investigated the potential presence of crustal anisotropy using receiver functions [Levin
and Park, 1997a, 1997b; Peng and Humphreys, 1997;
Savage, 1998; Leidig and Zandt, 2003; Frederiksen and
Bostock, 2000; Frederiksen et al., 2003; Vergne et al.,
2003]. Many of these studies have managed good simultaneous fits to radial and transverse component amplitudes for
multiple back azimuths using anisotropy magnitudes on the
order of 10– 20%. Unlike shear wave splitting, thick layers
of anisotropy are not required in order to produce a
noticable effect in receiver function waveforms, since the
characterization of anisotropy is not based on identification
of split S phases and measuring of split time between them
but instead depends upon the waveform shape and amplitude of converted phases. In addition, the receiver function
method allows distinction of anisotropy that is completely
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due to crustal structure and additionally provides some
resolution of the depth distribution of anisotropy within
the crust. This is in contrast to shear wave splitting, which is
a composite result of all anisotropic media between the
interface of S wave conversion or reflection, usually the
core-mantle boundary, and the surface of the Earth.
4. Inversion
4.1. Inversion Using the Neighborhood Algorithm
[11] Forward modeling of multiple back azimuths of both
radial and transverse component receiver functions is a
tedious process given the large number of variable parameters involved in characterizing anisotropic crustal structure
and is unlikely to be a thorough search of the vast parameter
space. Global minimization inversions are an efficient way
of exploring a dimensionally large model parameter space
and are thus well suited for receiver function modeling with
complex structure such as anisotropy. One particular method
of global minimization is known as the neighborhood
algorithm, developed by Sambridge [1999]. This algorithm
begins by randomly choosing some number of models from
a multidimensional model parameter space whose size is
defined by user-specified ranges in model parameters.
Synthetic seismograms are computed for each of these
models, and cross correlation based misfits between these
synthetics and input data are determined. Progressively
smaller regions of model parameter space containing low
misfit models are iteratively searched in more detail to find
a best fitting model. The implementation of this algorithm
for modeling of receiver functions using anisotropic crustal
structure with dipping layer interfaces was developed by
Frederiksen et al. [2003].
[12] Computation of synthetic seismograms as part of the
inversion process involves a ray-based approach that is
described in detail by Frederiksen and Bostock [2000]. This
method of synthetic seismogram computation is efficient
because only specified phases are produced, as opposed to
reflectivity methods, which reproduce the entire seismic
response of a model. In particular, computation of multiples
can be avoided if desired, and deconvolution is not necessary,
as phases with P polarity are not present on the synthetic
seismograms when multiples are not computed; the synthetic
seismograms themselves are thus approximately receiver
functions. Multiples were not calculated for any of the
inversions, primarily to minimize computation time. For
example, calculation of synthetic seismograms for a model
with eight layers can take hours if multiples are included in
the calculations, versus a computation time of less than one
minute when multiples are not included. Cross correlation
based misfits between synthetics produced during the inversion runs and synthetics generated using a reflectivity code
(anirec by J. Park, personal communication, 2002) range
from 0.04 to 0.07 on a scale from 0 (perfect correlation) to 2
(perfect anticorrelation), implying that multiples have very
little impact on the overall waveshapes. While shallow
impedance contrasts in the models may result in multiples
that arrive within the first 10 s after the direct P arrival, the
combined effects of attenuation, scattering, and anisotropy
can often be expected to yield multiples from upper crustal
layers with negligible amplitudes in real data, especially at
the higher frequencies used in this study.
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Figure 3. Illustration of the effect of anisotropy on receiver function waveforms. The receiver functions
are plotted as a function of the station-to-source back azimuth, with time increasing downward to
emphasize the correlation with depth. Dark gray is used for positive polarities, and light gray is used for
negative polarities. The top plot shows a record section generated using a model consisting of an isotropic
layer over a half-space. The right plot is a perspective diagram showing an example P-to-S conversion
and the event-station coordinate system, where N is north, R is radial, and T is transverse. Radial receiver
functions do not show any variation for different back azimuths, and no energy is present on the
transverse component. The addition of anisotropy to the layer results in energy on both the radial and
transverse components, shown in the bottom plot, due to conversion from P to both SV and SH particle
motions. Radial component waveforms show a systematic variation as a function of back azimuth and
include some negative polarities, even though there are no back azimuths for which the velocity
relationship between the layer and the half-space is a decrease with increasing depth. The transverse
component similarly shows a variation with back azimuth and contains no energy for propagation
directions parallel to the trend of the anisotropy symmetry axis. P and S wave velocities as a function of
depth are shown for each model, and the extra lines on the anisotropic velocity plot show the minimum,
maximum, and average velocities for the anisotropic portion of the model.
[13] Anisotropy is incorporated in the code through
specification of a magnitude, which is essentially the
percent difference between maximum and minimum velocities, positive for unique fast axis symmetry or negative for
unique slow axis symmetry, and an anisotropy symmetry
axis trend and plunge. Trend is measured clockwise from
north, while plunge is measured downward from horizontal
in the direction of the trend (Figure 2). The percentages of P
and S anisotropy and the average P and S velocities in a
layer determine four elastic constants. A fifth parameter is
required to uniquely define the five elastic constants necessary for hexagonal symmetry; in this study, a parameter
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called c, based on an elastic tensor parameterization used by
Levin and Park [1998], determines the deviation of a phase
velocity surface from ellipsoidal (see Appendix A). The
parameter c was set to zero in all runs, such that phase
velocity surfaces for anisotropic layers are ellipsoidal, based
on the observation from seismic refraction studies that c is
typically small [Anderson, 1989].
[14] The anisotropy axis was constrained to be uniquely
fast or slow for a given anisotropic layer in a particular run,
since the type of hexagonal symmetry cannot be uniquely
resolved in receiver function data. Numerical experiments
have shown that fast (positive) and slow (negative) axis
anisotropy can generate very similar effects in the data when
the two have opposite (180) axis trends and supplementary
(90 d) plunges [Erickson, 2002]. Thus an a priori
assumption is made on the sign of the anisotropy within
the Tibetan crust. The shallow crust is constrained to have
unique slow axis symmetry, as crack systems are thought to
dominate uppermost crustal anisotropy, while unique fast
axis symmetry is assumed for the middle to lower crust for
reasons discussed below. A number of geophysical techniques have provided evidence for weak middle to lower
crustal material within part or all of the Tibetan Plateau
[Alsdorf and Nelson, 1999; Nelson et al., 1996; Brown et
al., 1996; Makovsky et al., 1996; Kola-Ojo and Meissner,
2001; Chen et al., 1996; Wei et al., 2001]. Weak middle to
lower crust is further implied by the general lack of earthquakes below 10– 15 km depth [Molnar and Lyon-Caen,
1989; Chen and Molnar, 1983] and by flexural studies of
rifts in the southern to central plateau [Masek et al., 1994].
Active, long-distance crustal flow has been proposed as a
mechanism to maintain a uniform plateau elevation during
convergence and/or in response to lateral pressure gradients
produced by the high topography of the plateau [Bird, 1991;
Clark and Royden, 2000; Vanderhaeghe and Teyssier, 2001;
Shen et al., 2001]. Experimental and natural studies of rock
deformation have shown that foliation and lineation fabrics
in rocks are often statistically coincident with flow planes
and flow directions, respectively [Mainprice and Nicolas,
1989]. Significant crustal flow is likely to impose a regionalscale mineral lineation that will exhibit unique fast axis
anisotropy. Therefore the inversions in this study allowed
only unique fast axis symmetry in the deeper crust.
[15] As many as 30– 40 inversion runs were performed
using data from each of the Tibetan Plateau stations,
including preliminary runs to determine appropriate bounds
for model parameter variation in order to fit targeted phases
on waveforms; results discussed here are based on about 5 –
20 runs for each station. The number of layers in models for
a given station was fixed based on the number of clear
arrivals before the P-to-S conversion from the Moho in the
data. For all stations, only the first 10 s following the direct
P arrival were modeled. In the inversions, the ratio of P to S
velocity (Vp/Vs) was fixed to 1.73 in all layers to decrease
the number of variable parameters. Percent anisotropy,
anisotropy symmetry axis trend, and anisotropy symmetry
axis plunge were allowed to vary in layers in which
anisotropy was permitted. Standard deviations for inverted
parameters were estimated based on the range of values
determined from multiple runs. Although these error estimates are meaningful within a particular set of parameterizations, they do not fully account for potential errors from
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the full range of nonuniqueness of the problem. The
nonuniqueness problems are discussed further in later parts
of this paper. A variety of runs were performed for each
station, allowing anisotropy in different combinations of
layers; anisotropy was generally not allowed in more than
two consecutive layers, based on experiments with synthetic
data, which found that the true depth distribution and
orientation of anisotropy can be hard to resolve if anisotropy
is allowed to be present in too many consecutive layers in
an inversion run [Erickson, 2002].
4.2. Example Inversion: Station BUDO
[16] BUDO is located in the Songpan-Ganzi terrane, near
the Kunlun fault (Figure 1), and has been modeled in
several previous studies [Zhu et al., 1995; Frederiksen et
al., 2003; Erickson, 2002; Vergne et al., 2003]. BUDO
receiver functions contain four clear arrivals, including a
distinct Moho P-to-S conversion at around 8.5 s, and
models with four layers yield good fits to these waveforms
(Figure 4). Note that the inclusion of multiples in synthetic
waveform computation for the best BUDO model has very
little effect on their appearance (Figure 4a); the crosscorrelation misfit value between synthetics with and without
multiples is 0.05. On the basis of numerous inversions of
BUDO data with anisotropy permitted in layers 1, 2, and 4,
or in layers 1, 3, and 4, anisotropy is well defined in the
first, second, and third layers but never in the fourth layer.
Anisotropy is not required in both the second and third
layers, and better fits are generally obtained for anisotropy
in the second layer. The first layer, with a thickness of
around 10 km, has very strong anisotropy, with a magnitude
of 14 ± 1% and a SW to SSW trending symmetry axis
with a shallow plunge of 15– 20. Anisotropy in the second
layer has a magnitude of 10 ± 1%, with a trend to the SW to
SSW and a plunge of 40– 55. Trade-offs between symmetry axis orientation and percent anisotropy are illustrated for
each of the two anisotropic layers in the gray scale plots in
Figure 4b. The darker pixels represent models with low
misfits, and the lighter pixels represent models that do not
fit the data well. Note that the anisotropy axis trends are
generally well resolved in both layers. The goodness of fit is
shown in a different format in Figure 4c, which illustrates
the fit for individual phase amplitude versus back azimuth
for the major P-to-S conversions in the BUDO data.
[17] The BUDO model presented above differs from that
of Frederiksen et al. [2003] in a number of important ways
that are dependent largely upon modeling assumptions. In
particular, constraints placed on layer thickness, layer
velocity, V p/V s ratio, anisotropy symmetry type, and
amounts of P and S anisotropy are quite different. Inversions were performed for models with three, rather than
four, layers, and receiver functions were of a lower frequency content than in the present study. Vp/Vs ratio was
fixed for all inversion runs in this study, while P velocity
was fixed, and S velocity was allowed to vary, in the study
by Frederiksen et al. [2003], yielding somewhat extreme
Vp/Vs ratios for some layers. P and S anisotropy were
allowed to vary separately by Frederiksen et al. [2003],
while P and S anisotropy were held fixed to each other for
BUDO inversion runs here. Experiments with synthetic data
have shown that the appearance of receiver functions is
generally more sensitive to P anisotropy than to S [Levin
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Figure 4. (a) BUDO receiver function record section data and corresponding synthetic record sections
for the average model estimated from numerous inversions. The adjacent plot shows receiver functions
computed for the average model including multiples in the calculations; note the negligible difference
between these waveforms and the synthetic waveforms computed without including multiples. Dark
gray is used for positive polarities, and light gray is used for negative polarities. Extra lines on the
plots of P and S velocities as a function of depth indicate minimum, maximum, and average velocities
for anisotropic layers. Horizontal lines on plots of BUDO data (left) and model (right) mark direct P
arrivals and P-to-S conversions that were targeted in modeling, and back azimuth versus amplitude
plots for each of these arrivals are shown in Figure 4c. (b) Trade-off plots give an indication of the
resolution of symmetry axis trends. Dark shadings indicate low misfits, while light shadings indicate
high misfits. These trade-off plots are from a single inversion run, in which about 16,000 models were
generated, but are representative of similar plots for other inversion runs on BUDO data. (c) Back
azimuth versus amplitude plots for select arrivals. The solid lines represent synthetic data for a single
ray parameter of 0.06 s/km. Triangles, circles, and squares indicate data points for ray parameters of 0.04,
0.06, and 0.08 s/km, respectively; error bars around these points show the variation in values of waveforms that
were stacked together to produce the average waveforms that were modeled in the inversion procedure.
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Figure 4. (continued)
and Park, 1998], and P and S anisotropy are difficult to
resolve individually when both are present. Layer interfaces
were also allowed to dip in the Frederiksen et al. [2003]
model, resulting in a possible modification of anisotropy
orientations. A final important difference between the two
models is the definition and constraint of the elastic
tensor parameterization used for anisotropic layers (see
Appendix A). In the present study, the phase velocity surfaces
were constrained to be ellipsoids, as discussed previously,
while Frederiksen et al. [2003], an anisotropy parameter, h,
was fixed to a value appropriate for upper mantle fabrics
defined by olivine alignment [Farra et al., 1991].
[18] Vergne et al. [2003] recently published an anisotropic
model for BUDO using slightly lower frequency receiver
functions than in the present study. Their model has two layers
of anisotropy: The surface layer is 13 km thick and contains
15% anisotropy oriented with a trend of 22 and zero
plunge; the second layer is also 13 km thick and contains
15% anisotropy oriented with a trend of 30 and a plunge
of 48. The best fitting BUDO model in the present study
includes a surface layer with a thickness of 11 km and 14%
anisotropy oriented with a trend of 207 and a plunge of 18.
The parameters for the surface layer are essentially the same
in both studies because the zero plunge in the Vergne et al.
[2003] model makes the trend indeterminate within a factor of
180. The second layer of the preferred BUDO model in the
present study is 14 km thick and has 10% anisotropy with a
199 trend and 48 plunge. These parameters can be reconciled with the Vergne et al. [2003] model by taking the sign of
the anisotropy into account. Fast (positive) and slow (negative) axis anisotropy can generate very similar effects in the
data when the two have opposite (180) trends and supplementary (90 d) plunges [Erickson, 2002]. In the case of
BUDO, the two models have opposite anisotropy signs, are
169 different in trends, and are within 3 of supplementary
plunge angles.
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[19] Despite these differences, the models discussed
above fit the data nearly equally well, providing an important example of the nonuniqueness inherent in inverting for
crustal anisotropy using receiver functions, even with simplified symmetry, due to the large number of variable
parameters. The most robust parameter is the orientation
of the unique symmetry axis, either fast or slow. Geological
and geophysical constraints can help to decrease the inherent nonuniqueness in modeling. For example, detailed
knowledge of the composition of the crust would allow a
more robust determination of whether the anisotropy should
be modeled as fast or slow. Measurements of the plunge
directions of dominant structures in the crust would allow
constraint of that parameter and inversion for either fast or
slow anisotropy, perhaps helping to constrain the composition of the crust. Clearly, further studies of crustal anisotropy, both in the field and laboratory, are needed to address
some of these issues.
4.3. Inversion Results
[20] The results of multiple inversion runs have been
compiled for each station, yielding average values of model
parameters that were allowed to vary in each inversion run,
and these average values were used to create average
models with estimated standard errors for each station
(Table 1 and Figure 5). Anisotropy symmetry axis trend
and layer thickness are nearly always well resolved for
certain layers in models for each station. P wave velocity,
percent anisotropy, and anisotropy symmetry axis plunge
are not as well resolved.
[21] Receiver function record sections plotted according to
station-to-source back azimuth and synthetics corresponding
to the models in Figure 5 are shown in Figure 6 (with the
exception of the plots for station BUDO, which are shown
in Figure 4). Amplitude variations in the data are generally
well matched with synthetics, supporting crustal anisotropy
as a reasonable explanation for observed polarity variations
as a function of back azimuth. In Figure 6a, the plots for all
the PASSCAL stations located in the Lhasa terrane are
shown, while the plots for PASSCAL stations in northern
Tibet are shown in Figure 6b.
[22] Receiver function data are somewhat insensitive to
absolute crustal velocity but are fairly sensitive to the
velocity contrast between two layers; the crustal thicknesses
and absolute magnitudes of velocities determined through
inversions in this study should therefore not be interpreted
too strictly. Furthermore, the P-to-S conversion from the
Moho, marking the base of the crust, is not clear at all
stations for all back azimuths, an observation supported by
previous authors who have examined these PASSCAL data
[Zhu et al., 1995; Zhao et al., 1996]. Receiver functions
examined here are higher in frequency than some of the
published migrated receiver function images for the plateau
that display more prominent Moho arrivals [e.g., Kosarev et
al., 1999]. The lack of a clear, high-frequency Moho P-to-S
conversion in data from many of the Tibetan PASSCAL
stations may be the result of gradational boundaries at the
base of the crust in some parts of the plateau. Alternatively,
anisotropy within the crust could yield a very complex
converted phase from the Moho that is difficult to distinguish in the higher-frequency data. Approximate Moho
P-to-S converted phase arrival time picks are based on
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several studies using a number of techniques, including
seismic reflection [Zhao et al., 2001], receiver functions
[Zhao et al., 1996], and S-to-P conversions [Owens and
Zandt, 1997].
5. Discussion
5.1. Anisotropy Overview
[23] Average crustal velocity models determined for
Tibetan stations show that significant anisotropy is present
within the crust of the Tibetan Plateau. Anisotropy with
magnitudes of 4 – 14% is present in layers 2 – 25 km thick
that collectively compose over half of the crust beneath some
recording stations (Figure 5). The station coverage here is
much too sparse and our global understanding of crustal
anisotropy is too incomplete to make any robust general
statements about the nature of anisotropy within the plateau
crust except that it has quite variable magnitude (compare
LSA and GANZ) and that the northern plateau stations
appear to have shallower levels of anisotropy. Importantly,
the orientation of the unique anisotropy axis is the most
robust parameter modeled, and changes in orientation are
emphasized in interpretations of modeling results.
[24] Considered as a whole, this crustal anisotropy can be
divided into three depth zones. The first is a zone of shallow
upper crustal anisotropy, or the uppermost layer in the
model for each station. This layer of anisotropy could be
due to the presence of aligned cracks in the shallow
subsurface but could also indicate alignments of minerals
in the shallow crust. The second region of anisotropy is
within the upper crust, at depths between about 5 and 25 km
that are probably too deep for significant crack systems to
remain open. Anisotropy at these depths is thus probably
due to alignment of mineral grains, yet this part of the crust
is not likely at appropriate temperature and pressure conditions to have experienced recent ductile deformation and/
or recent widespread development of metamorphic fabrics.
The third depth zone of anisotropy is in the middle to lower
crust, and anisotropy at these depths may be a result of
recent or ongoing ductile deformation, metamorphism, or
crustal flow.
5.2. Shallow Anisotropy
[25] Unique slow anisotropy symmetry axis orientations in
the upper crust imply planar structures (perpendicular to the
symmetry axis) that generally show a correlation with surface
features, particularly the orientations of faults and sutures
separating terranes (Figure 7). In the southern portion of the
plateau, the upper crust beneath stations LSA and XIGA
contains planar features with an E-W strike, similar to the
strike of the Indus-Tsangpo suture in this area. These planar
orientations may also be related to the south dipping Great
Counter thrust system and the north dipping Gangdese thrust
system in the southern plateau [Yin et al., 1994; P. Kapp,
personal communication, 2002]. Nearby, anisotropy at
GANZ implies planar structures with a NE-SW strike,
consistent with a bend in the Indus-Tsangpo suture to a
slightly more NE orientation in the vicinity of this station.
Anisotropy in the upper crust beneath SANG indicates E-W
striking planar structures, consistent with E-W terrane
boundary orientations, though SANG is located essentially
in the center of the Lhasa terrane.
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Table 1. Average Anisotropic Models From Neighborhood
Inversiona
H, km
r, g/cm3
Percent
Anisotropy
Trend,
deg
Plunge,
deg
Vp, km/s
Vp/Vs
4
8(2)
0
4(2)
0
0
0
0
351(20)
250(17)
30
64(6)
330(17)
52(14)
184(7)
41(8)
24(10)
151(10)
48(9)
39(11)
6(0)
17(0)
6(0)
10(1)
9
10(1)
22(1)
0
2.7
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5.9(0.0)
5.8(0.0)
6.2(0.0)
6.0(0.0)
6.4
6.1(0.1)
6.7(0.1)
8.0
LSA
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
5(0)
12(1)
6(1)
10(0)
17(1)
24(1)
0
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5.7(0.1)
5.9(0.0)
6.1(0.0)
6.3(0.1)
6.4(0.1)
6.5(0.1)
8.0
XIGA
1.73
1.73
1.73
1.73
1.73
1.73
1.73
14(1)
0
0
11(2)
10(2)
0
0
5
8(0)
6(0)
5
5(1)
13(1)
6(1)
15
0
2.7
2.7
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5.7
6.0(0.1)
5.9(0.1)
6.2
6.1(0.1)
6.1(0.1)
6.4(0.1)
6.3
8.0
GANZ
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
10
0
10(2)
0
14(1)
0
11(2)
6(3)
0
5
4(0)
5
4(1)
7
7(1)
5
8(1)
22
0
2.7
2.7
2.7
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5.8
5.7(0.1)
6.0
6.0(0.1)
6.5
6.4(0.1)
6.1
6.6(0.1)
6.8
8.0
SANG
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
5
6(0)
5(1)
6(1)
10
11(1)
11(1)
20
0
2.7
2.7
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5.6
5.9(0.1)
5.9(0.1)
5.7(0.1)
6.4
6.3(0.1)
6.4(0.1)
6.4
8.0
4(0)
5(0)
5(1)
7(1)
16(2)
33(2)
0
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5(0)
8(0)
11(2)
16(1)
13(2)
24(2)
0
2.7
2.7
2.7
2.7
2.7
2.7
3.3
344(9)
40
51(8)
13(9)
59(7)
22(6)
63(16)
300(33)
51(8)
26(14)
4
8(2)
0
14(1)
0
11(2)
10(2)
0
0
0
0(13)
320(20)
30
61(7)
290(11)
56(8)
94(10)
298(15)
14(9)
55(5)
AMDO
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
1.73
9
0
11(2)
0
13(2)
11(2)
4(3)
0
0
40(15)
40
26(13)
40(11)
71(9)
282(26)
70(56)
11(5)
62(6)
22(15)
5.9(0.0)
6.2(0.0)
6.1(0.1)
6.2(0.1)
6.6(0.1)
6.4(0.1)
8.0
WNDO
1.73
1.73
1.73
1.73
1.73
1.73
1.73
9(1)
9(1)
0
5(1)
5(1)
0
0
78(5)
133(7)
22(4)
31(3)
64(7)
237(7)
53(5)
11(5)
5.9(0.1)
5.9(0.1)
6.3(0.1)
6.4(0.0)
6.3(0.1)
6.8(0.1)
8.0
USHU
1.73
1.73
1.73
1.73
1.73
1.73
1.73
10(1)
0
6(1)
0
4(1)
6(1)
0
278(8)
65(3)
22(6)
38(4)
158(11)
305(7)
45(6)
3(3)
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Table 1. (continued)
Percent
Anisotropy
Trend,
deg
Plunge,
deg
ERDO
1.73
1.73
1.73
1.73
1.73
1.73
14(1)
10(1)
0
11(2)
0
0
22(10)
19(9)
53(5)
48(6)
0(7)
31(8)
6.1(0.0)
6.2(0.0)
6.1(0.0)
6.4(0.1)
8.0
BUDO
1.73
1.73
1.73
1.73
1.73
14(1)
10(1)
0
0
0
207(5)
199(7)
18(3)
48(5)
2.7
2.7
2.7
2.7
2.7
3.3
5.5(0.0)
6.4(0.0)
6.3(0.1)
6.2(0.1)
6.6(0.1)
8.0
TUNL
1.73
1.73
1.73
1.73
1.73
1.73
10(1)
0
5(2)
4(1)
0
0
43(7)
28(6)
183(10)
46(17)
35(12)
47(13)
2.7
2.7
2.7
2.7
2.7
2.7
3.3
5.6(0.1)
6.2(0.1)
6.1(0.1)
6.2(0.1)
6.6(0.1)
6.7(0.1)
8.0
MAQI
1.73
1.73
1.73
1.73
1.73
1.73
1.73
6(3)
0
13(1)
0
0
0
0
345(24)
34(13)
31(10)
54(7)
H, km
r, g/cm3
Vp, km/s
Vp/Vs
7(0)
13(1)
11(1)
17(1)
13(1)
0
2.7
2.7
2.7
2.7
2.7
3.3
5.6(0.0)
6.0(0.1)
6.2(0.1)
6.3(0.1)
6.7(0.1)
8.0
11(0)
14(0)
10(0)
34(0)
0
2.7
2.7
2.7
2.7
3.3
2(0)
11(1)
8(1)
27(0)
19(1)
0
5(2)
8(1)
11(1)
7(1)
12(1)
8(1)
0
a
H, layer thickness; r, layer density; Vp, P wave velocity; Vp/Vs, ratio of
P wave and S wave velocities. Trend and plunge refer to orientation of the
anisotropy symmetry axis. Values in parentheses are estimated 1s standard
deviation for model parameters that were allowed to vary in inversion runs.
A value of 0 standard deviation for thickness or P wave velocity means that
the standard deviation is less than 0.5 km or 0.05 km/s, respectively.
[26] Further north, uppermost crustal anisotropy beneath
AMDO does not appear to bear any obvious correlation to
surface features. Shallow anisotropy beneath WNDO corresponds to NNW-SSE striking planes. WNDO is located in
the central Qiangtang terrane and is on strike with the
plunging axis of the Qiangtang anticlinorium that is overprinted by generally N-S trending normal faults [Kapp et
al., 2000, 2003]. The shallow anisotropy determined at
WNDO may thus be an eastern expression of this young
E-W extensional system. Approaching the Jinsha suture, the
anisotropy orientation at ERDO seems to correspond with
the NW-SE orientation of mapped thrust faults in the
northern Qiangtang terrane. Anisotropy within the upper
crust beneath BUDO and TUNL suggests planar structures
with an orientation similar to NW-SE trending thrust faults
in the Songpan Ganzi terrane. In particular, the south
plunging axis at BUDO is consistent with north dipping
thrust faults north of the Jinsha suture (P. Kapp, personal
communication, 2002). Toward the east, the uppermost
crustal anisotropy beneath MAQI is consistent with planar
structures parallel to the local strike of the Kunlun suture.
South of MAQI, shallow anisotropy at USHU is not
obviously correlated with known structures.
[27] Near-surface crustal anisotropy is usually thought to
be dominantly the result of aligned cracks in the shallow
10 of 20
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SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
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Figure 5. Summary plot of anisotropic crustal velocity models determined in this study by receiver
function inversions for the 10 Sino-American 1991 – 1992 Tibetan Plateau PASSCAL stations and GSN
station LSA. Extra lines on the plots of P and S velocities as a function of depth indicate minimum,
maximum, and average velocities for anisotropic layers. The adjacent stereo plots illustrate the lower
hemisphere projection of the anisotropic symmetry axis for each anisotropic layer. The open circles show
the slow axis trend and plunge for the upper crust; the closed circles represent the fast axis directions for
the upper to upper middle crust; and the stars show the fast axis directions for the middle to lower crust.
subsurface, but the orientations determined in this study
(with the possible exception of WNDO) do not appear to
reflect crack systems that have formed in response to the
present-day E-W extensional stresses within the upper crust
of the Tibetan Plateau. Upper crustal anisotropy could
alternatively be controlled by dipping metamorphic foliations formed in the vicinity of fault zones that have been
exhumed to shallow depths. Additional surface structural
11 of 20
Figure 6a. Comparison of data and synthetics for each of five stations in the Lhasa terrane of the southern Tibetan Plateau.
Inset map shows the locations of the five stations. Each panel of four record sections compares the receiver functions for
data (left) and corresponding synthetics (right) for both the radial (top) and transverse (bottom) components. Red is used for
positive polarities, and blue is used for negative polarities. See color version of this figure at back of this issue.
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SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
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Figure 6b. Comparison of data and synthetics for each of five stations in the Qiangtang and Songpan Ganzi terranes of the
northern Tibetan plateau. Inset map shows the locations of the five stations. The comparison for station BUDO is shown in
Figure 4a. Each panel of four record sections compares the receiver functions for data (left) and corresponding synthetics
(right) for both the radial (top) and transverse (bottom) components. Red is used for positive polarities, and blue is used for
negative polarities. See color version of this figure at back of this issue.
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SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
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SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
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Figure 7. Orientations of shallow upper crustal anisotropy in the Tibetan Plateau. Thick black lines
show the orientations of unique slow symmetry axes. Ellipses perpendicular to the lines represent trends
of planar structures, such as aligned microcracks or metamorphic foliations, which might correspond to
the slow symmetry axis orientations. The orientations of these planes are roughly consistent with mapped
faults or terrane boundaries in the vicinity of several stations. The first letter of each station name is used
for station identification. The full station name and explanation for the base map are given in Figure 1.
studies in Tibet, especially on exposures of metamorphic
rocks, would be useful for interpreting seismic anisotropy
results for the uppermost crust.
5.3. Upper to Upper Middle Crustal Anisotropy
[28] Upper to upper middle crustal anisotropy, generally
above 20– 25 km and below the surface layer, has a
rather consistent trend of unique fast symmetry axes in
the northern plateau but displays an erratic spatial variation in the south (Figure 8). Anisotropy at the southern
stations LSA, GANZ, and AMDO has a NE or SW
symmetry axis trend, while anisotropy in the upper
middle crust of SANG has a NW or SE orientation. At
similar depths, anisotropy has a SE orientation at WNDO.
Beneath XIGA, anisotropy has a NE axis trend in a layer
that extends slightly deeper than 25 km, although the
orientation of anisotropy in this layer greatly contrasts
with the orientation in a deeper layer of anisotropy in
models for this station and is thus considered as upper to
upper middle crustal anisotropy. In the north, ERDO,
BUDO, TUNL, USHU, and MAQI all have a NE-SW
symmetry axis orientation in the upper middle crust,
consistent with the results of Vergne et al. [2003].
[29] The relatively shallow depths of upper to upper
middle crustal anisotropy generally preclude the development of significant recent metamorphic fabrics, but this
portion of the crust is also probably too deep for significant
crack systems to be present. Thus anisotropy at these depths
is probably due to alignment of mineral grains in fabrics that
are not related to recent deformation. The great spatial
variation of this anisotropy in the southern portion of the
plateau suggests that it could be due to localized deformation manifested in fossil fabrics, perhaps associated with the
amalgamation of terranes to form present-day Tibet. Although they are more spatially consistent, the orientations
determined for the northern stations can similarly be
explained by the presence of fossil fabrics related to past
deformation, although these orientations are not particularly
different from orientations of middle to lower crustal
anisotropy in the northern plateau, described in the following section. The deeper portions of upper middle crustal
anisotropy at the northern stations could therefore be a
result of more recent deformation, as discussed below.
The higher level of crust affected by ongoing deformation
in the northern plateau could be explained by the higher
mantle temperatures in the north as inferred from numerous
seismic studies [e.g., Owens and Zandt, 1997].
5.4. Middle to Lower Crustal Anisotropy
[30] The orientations of unique fast symmetry axes in the
middle to lower crust (>25 km) show a distinct spatial
pattern (Figure 9). In the south, the middle to lower crust
beneath XIGA, LSA, and GANZ has NNW to NW symmetry axis orientations. This orientation rotates to nearly
E-W at SANG and AMDO. Slightly further north, anisotropy
at WNDO has an ENE to NE symmetry axis trend. To the
east of WNDO, USHU has a NW orientation of middle to
lower crustal anisotropy. The pattern changes abruptly at
ERDO, where middle crustal anisotropy has a trend toward
the north. Anisotropy beneath 25 km is apparently absent at
14 of 20
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SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
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Figure 8. Orientations of upper to upper middle crustal anisotropy in the Tibetan Plateau. Thick black
lines show the orientations of unique fast symmetry axes determined at each station. More than one line
at a given station indicates multiple layers of anisotropy at upper to upper middle crustal depths. This
anisotropy is most likely due to deformation fabrics that are a result of past crustal deformation. The first
letter of each station name is used for station identification. The full station name and explanation for the
base map are given in Figure 1.
BUDO and MAQI, and weak middle to lower crustal
anisotropy has a NE symmetry axis trend at station TUNL.
[31] Velocity models for each station (Figure 5) show that
the majority of middle to lower crustal anisotropy is actually
in the middle portion of the thick Tibetan crust. The
dominance of anisotropy in the middle crust, rather than
the lower crust, may be a real feature, or it may be an
artifact of both the receiver function method and the
behavior of minerals under certain pressure, temperature,
fluid, and stress conditions. When anisotropy is present, the
appearance of a P-to-S conversion from a deep crustal
interface can be modified by passage across shallower
interfaces or anisotropy at shallower depths, and the anisotropy associated with the deeper layers may be harder to
resolve from the waveforms. The absence of measurable
tectonic fabrics can be due to a lack of deformation, or to
the presence of sufficient heat or fluid transport that allows
deformation mechanisms that tend to produce isotropic
fabrics. An apparent lack of anisotropy could also be due
to a deficiency in minerals that form seismically detectable
fabrics. In short, the absence of crustal anisotropy at middle
to lower crustal depths at some stations may not be an
indication of undeformed crust at those depths.
5.5. Tectonic Significance of Middle to
Lower Crustal Anisotropy
[32] The orientations of middle to lower crustal anisotropy
determined in this study correlate well with directions of
deviatoric stress calculated by Flesch et al. [2001] using both
gravitational potential energy variations and the northward
motion of India, though not considering the effect of any
lithospheric subduction beneath the northern portion of the
plateau. Fast symmetry axis trends align with large magnitude compressive stress directions near the southern and
northern margins of the plateau and with extensional stresses
in the central plateau. The N to NW orientation of fast axes in
the southern portion of the plateau may reflect a combined
influence from the northward motion of India and lateral
topographically driven flow to align minerals in a ductile
middle to lower crust. The rotation of symmetry axes to a
nearly E-W trend at SANG and AMDO, progressively closer
to the Banggong suture, implies a dominant effect from
topographically induced flow and extrusion of material
along E-W trending strike slip faults. Recent studies have
revealed a 200 – 300 km wide zone of active conjugate
strike-slip faulting across central Tibet that is accommodating coeval east-west extension and north-south contraction
[Taylor et al., 2003]. The rotation of the mid-to-lower crustal
anisotropy axes to E-W at stations SANG and AMDO,
located within this central Tibet conjugate fault zone
(Figure 9), suggests that crustal extrusion in this zone is
occurring to at least midcrustal levels. Diminishing influence
of underthrusting Indian lithosphere on the central part of the
plateau is supported by seismic evidence suggesting that
Indian lithosphere only extends as far north as the Banggong
suture [Beghoul et al., 1993; Owens and Zandt, 1997].
[33] Toward the east, NW trending fast axes in the deep
crust beneath USHU seem consistent with anisotropy resulting from alignment of material in a flow field that is
influenced by both lateral motion from pressure gradients
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Figure 9. Orientations of middle to lower crustal anisotropy in the Tibetan Plateau. Thick black lines
show the orientations of unique fast symmetry axes determined at each station. More than one line at a
given station indicates multiple layers of anisotropy at middle to lower crustal depths. No middle to lower
crustal anisotropy was well resolved at stations BUDO and MAQI. This anisotropy is most likely due to
deformation fabrics that are a result of active crustal deformation. The first letter of each station name is
used for station identification. The full station name and explanation for the base map are given in Figure 1.
and rotation around the eastern Himalayan syntaxis. The
ENE to NE orientation of middle to lower crustal anisotropy
at WNDO implies influence from a north or south directed
force, although the nearly E-W symmetry axis orientations
at SANG and AMDO would seem to suggest that flow in
the crust beneath WNDO is not significantly affected by
underthrusting Indian lithosphere. Some authors have found
evidence that lithospheric subduction may be occurring in
the vicinity of the Jinsha suture [Wittlinger et al., 1996].
Northeast trending anisotropy in the middle crust at WNDO
could thus be influenced by both lithospheric subduction in
the northern plateau and eastward or westward extrusion or
flow of material. The north trend at ERDO could similarly
be related to lithospheric subduction beneath the northern
portion of the plateau. The absence of any apparent significant influence of an east or west directed force on deeper
crustal anisotropy at BUDO, TUNL, MAQI, and ERDO is
consistent with the general lack of evidence for E-W
extension (i.e., N-S or NE-SW trending rifts) in the northern
plateau. Deeper anisotropy at TUNL has a NE trend,
possibly a result of lithospheric subduction of the Qaidam
basin beneath Eurasia [Tapponnier et al., 2001; Kind et al.,
2002] or of northward subduction of some lithospheric
fragment beneath Qaidam [Wittlinger et al., 1996].
5.6. Relationship Between Recent Crust
and Mantle Deformation
[34] Significant shear wave splitting has been measured
using ScS, SKS, and S phases in data from a number of
stations on the Tibetan Plateau, presumably indicating the
presence of upper mantle fabrics [McNamara et al., 1994;
Hirn et al., 1995; Guilbert et al., 1996; Sandvol et al., 1997;
Huang et al., 2000]. These studies have generally observed
large split times in the central to northern plateau of nearly
two seconds, values that are difficult to reconcile with
crustal anisotropy alone. Horizontal, fast polarization directions determined with split ScS, SKS, and S phases typically
have N to NNW orientations just south of the IndusTsangpo suture [Hirn et al., 1995] and NE to ENE orientations in the southern portion of the plateau (Figure 1).
Anisotropy orientations rotate to a more E-W trend toward
the north. Stations to the east, particularly USHU and
MAQI, were found to have anisotropy with a NW orientation. Importantly, the measured orientations of mantle
anisotropy from the studies cited here involve the assumption of a horizontal symmetry axis, which may yield
erroneous or biased results, in the case of multiple layers
or dipping symmetry axes of anisotropy.
[35] A previous study of anisotropy in the Tibetan region
involved the examination of split Moho P-to-S conversions
to determine details of crustal anisotropy within the plateau
[Herquel et al., 1995]. While this study found orientations
of horizontal, fast symmetry axis anisotropy similar to those
determined in the aforementioned studies of mantle anisotropy, only a very small number of events from a limited
range of back azimuths were used to determine symmetry
axis orientations. The study by Herquel et al. [1995] also
assumes that anisotropy has a horizontal symmetry axis
trend, which is almost never the case for anisotropy determined in the present study. Split phases can be hard to
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SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
identify in converted waveforms [Savage, 1998], and any
shear wave splitting from crustal structure modeled in the
present study would also most likely be minimal. If all of
the anisotropic layers for a given station had the same
orientation, a maximum split time of about 0.5– 1.0 s for
the P-to-S conversion from the Moho would result for the
various stations. However, the symmetry axis trends are not
parallel in multiple layers in most of the models, such that
any observed split time would probably be significantly less
than these maximum values. McNamara et al. [1994]
examined Moho P-to-S conversions in receiver functions
computed from the PASSCAL data set and saw very little
measurable splitting.
[36] Fast symmetry axis orientations determined from
shear wave splitting studies are somewhat different from
trends of middle to lower crustal anisotropy, especially for
stations near the northern and southern margins of the
plateau (Figures 1 and Figure 9). In particular, roughly
NE or ENE orientations of mantle anisotropy in the southern plateau are quite distinct from N to NW crustal trends at
XIGA, LSA, and GANZ. Similarly, E-W trends of mantle
anisotropy in the northern plateau are very different from N
to NE trends of middle to lower crustal anisotropy at TUNL
and ERDO. As previously discussed, the orientation of
middle to lower crustal anisotropy near the northern and
southern margins of the plateau appears to bear a resemblance to potential directions of lithospheric subduction
beneath the southern, and perhaps northern, plateau. If these
fast mantle orientations are due to recent deformation, the
nearly perpendicular, or slab-parallel, trends of mantle
anisotropy at both margins suggest that some portion of
lithospheric mantle material, or perhaps the underlying
asthenosphere, is deforming in a manner that is distinct
from the crust. This observation is consistent with several
geodynamic models for high plateau development that
involve motions of crustal material that are largely independent from flow in the mantle [Shen et al., 2001; Royden,
1996]. However, in the central portion of the plateau, ENEWSW to E-W mantle anisotropy is consistent with roughly
E-W trends of middle to lower crustal anisotropy at SANG,
AMDO, and WNDO, and NW trends of mantle anisotropy
at USHU are similarly consistent with middle to lower
crustal anisotropy at this station. Lithospheric and/or asthenospheric material and the middle to lower crust in the
central and eastern plateau may thus be deforming coherently; alternatively, they may be decoupled but responding
in a similar fashion to the same driving force.
6. Conclusions
[37] Global minimization inversions of receiver functions
from data recorded on the Tibetan Plateau have yielded
average models for each of 11 stations that include 4 – 14%
anisotropy in layers 2 – 25 km thick at three different depth
zones within the crust. The presence of significant anisotropy within the crust of the Tibetan Plateau strongly
supports the idea that large-scale crustal fabrics are widespread and formed during part or all of the plateau’s crustal
deformation history and present-day tectonic activity. Upper
crustal anisotropy may be due to cracks or fossil fabrics,
while deeper, middle to lower crustal anisotropy is most
likely due to aligned mineral grains, possibly amphiboles or
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sillimanite. The orientation of uppermost crustal anisotropy
seems to be related to surface features, some of which are
not presently active structures. Upper to upper middle
crustal anisotropy shows a considerably erratic spatial
variation in the southern plateau and is most likely not a
result of recent deformation, but is fairly consistent in the
north and may be related to young deformation. The
orientation of middle to lower crustal anisotropy seems
consistent with a Tibetan crust that is experiencing
ongoing deformation due to both topographically induced
pressure gradients and boundary forces related to subduction of lithosphere at its southern, and perhaps northern,
margins. The relationship of these orientations to computed deviatoric stress directions and present-day plateau
tectonics implies that geodynamic models seeking to explain these features via either movement of large rigid
blocks or by homogeneous thickening are too simplified
[Avouac and Tapponnier, 1993; England and Houseman,
1986]. Geodetic and geologic studies imply that significant
deformation is probably occurring within the plateau crust
on a large scale [Wang et al., 2001; Taylor et al., 2003].
Orientations of crustal anisotropy do not correlate with fast
polarization directions determined from shear wave splitting
studies in the northern and southern plateau, suggesting that
crust and mantle motions may be distinct in these regions,
although lithospheric motions seem more uniform in the
central and eastern plateau. Geodynamic models for plateau
development that include a weak middle or lower crustal
layer suggest that mantle motion may be decoupled from
that of the crust, such that flow directions may not be
uniform in all parts of the lithosphere in which ductile
deformation is occurring.
[38] The data examined here come from a very small
number of stations compared with the size of the Tibetan
Plateau, and many larger data sets must be collected and
analyzed to corroborate the results of this study. In addition,
further studies are required to more uniquely relate seismically observed crustal anisotropy to actual deformation
mechanisms in rocks and crustal flow. Nevertheless, this
study has shown that significant seismic anisotropy is
present in the crust of the Tibetan Plateau that provides
some indication of how the crust may have behaved through
time and is deforming at present.
Appendix A: Comparison of Alternative
Parameterizations of Hexagonal Symmetry
Anisotropy
[39] For a hexagonally symmetric anisotropic medium,
with one unique seismically fast or slow symmetry axis and
uniformly slow or fast velocities perpendicular to the axis,
five elastic constants are required to fully determine the P
and S wave velocities as a function of propagation direction
through the medium, assuming density perturbations
are ignored. One parameterization is the following, from
Anderson [1989]:
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2
A ¼ rVP?
ðA1Þ
2
C ¼ rVPk
ðA2Þ
SHERRINGTON ET AL.: TIBETAN PLATEAU CRUSTAL ANISOTROPY
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2
N ¼ rVS?
ðA3Þ
2
L ¼ rVSk
ðA4Þ
F ¼ hð A 2LÞ
ðA5Þ
The elastic constants above depend upon velocities of P or S
waves for different propagation directions through the
anisotropic medium. In particular, VP? is the velocity of P
waves propagating perpendicular to the axis of symmetry,
and VPk is the P wave velocity for propagation parallel to
the symmetry axis. Similarly, VS? is the S wave velocity for
propagation perpendicular to the symmetry axis and
polarization parallel to the symmetry axis, while VSk is the
S wave velocity for propagation parallel to the symmetry
axis and polarization perpendicular to the symmetry axis.
The density of the medium is denoted by r, and h is an
anisotropic parameter that controls the variation of velocity
as a function of propagation direction for directions other
than parallel or perpendicular to the symmetry axis. An
elastic tensor constructed from the above constants has the
following form:
0
A
B A 2N
B
B F
Cij ¼ B
B 0
B
@ 0
0
A 2N
A
F
0
0
0
F
F
C
0
0
0
1
0 0 0
0 0 0C
C
0 0 0C
C
L 0 0C
C
0 L 0A
0 0 N
ðA6Þ
An alternative parameterization for the elastic constants,
involving a, b, c, d, and e, is used by Levin and Park [1997a,
1998] and follows a derivation by Backus [1965]. In this
parameterization, b and e are the percent P and S anisotropies,
expressed as fractions, where percent anisotropy refers to the
difference between the maximum and minimum P or S
velocity divided by the average P or S velocity, respectively,
for the medium. A negative value of b or e corresponds to a
unique slow axis of symmetry, while positive values indicate
a fast axis of symmetry. For c = 0, the phase velocity surfaces,
or three-dimensional curves that represent the variation of
velocity as a function of propagation direction, are ellipsoids.
A nonzero value of c corresponds to a distortion of the
ellipsoids. These constants can be expressed in terms of A, F,
C, L, and N as follows:
A¼abþc
ðA7Þ
F ¼ a 3c 2ðd þ eÞ
ðA8Þ
C ¼aþbþc
ðA9Þ
L¼dþe
ðA10Þ
N ¼dc
ðA11Þ
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The anisotropic parameter h is then
h¼
a 3c 2ðd þ eÞ
a b þ c 2ðd þ eÞ
ðA12Þ
When c = 0, equation (A12) becomes
h¼
a 2ðd þ eÞ
a b 2ðd þ eÞ
ðA13Þ
Equation (A13) is used in this study in a slightly modified
version of the code by Frederiksen et al. [2003] to correspond
to the Levin and Park [1997a, 1998] parameterization for the
case (c = 0) where the anisotropy is assumed to be purely
ellipsoidal.
[40] Acknowledgments. Funding for this project was provided in part
by NSF grant EAR-0125121. Data of teleseismic events recorded during
the 1991 – 1992 PASSCAL experiment were provided by Tom Owens.
Many thanks to Steve Sorenson for modifying the raysum inversion code to
run simultaneously on multiple processors. We thank Norm Meader for his
assistance in manuscript preparation. Additional thanks to Susan Beck,
Clem Chase, Paul Kapp, and Jerome Guynn for their reviews of early drafts
of this paper. We thank an anonymous reviewer and the Associate Editor for
constructive reviews.
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A. Frederiksen, Department of Geological Sciences, 341 Wallace
Building, University of Manitoba, Winnipeg, Manitoba, Canada R3T
2N2. (frederik@cc.umanitoba.ca)
H. F. Sherrington and G. Zandt, Department of Geosciences, University
of Arizona, Tucson, AZ 85721, USA. (folsom@geo.arizona.edu; zandt@
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Figure 6a. Comparison of data and synthetics for each of five stations in the Lhasa terrane of the southern Tibetan Plateau.
Inset map shows the locations of the five stations. Each panel of four record sections compares the receiver functions for
data (left) and corresponding synthetics (right) for both the radial (top) and transverse (bottom) components. Red is used for
positive polarities, and blue is used for negative polarities.
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Figure 6b. Comparison of data and synthetics for each of five stations in the Qiangtang and Songpan Ganzi terranes of the
northern Tibetan plateau. Inset map shows the locations of the five stations. The comparison for station BUDO is shown in
Figure 4a. Each panel of four record sections compares the receiver functions for data (left) and corresponding synthetics
(right) for both the radial (top) and transverse (bottom) components. Red is used for positive polarities, and blue is used for
negative polarities.
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