Earth and Planetary Science Letters 199 (2002) 389^402
www.elsevier.com/locate/epsl
Moho topography in the central Andes and its
geodynamic implications
X. Yuan a; , S.V. Sobolev a , R. Kind a;b
a
GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany
b
Freie Universita«t Berlin, Berlin, Germany
Received 26 June 2001; received in revised form 12 December 2001; accepted 4 March 2002
Abstract
P-to-S converted waves at the continental Moho together with waves multiply reflected between the Earth’s surface
and the Moho have been used to estimate the Moho depth and average crustal Vp /Vs variations in the central Andes.
Our analysis confirms and significantly complements the Moho depth estimates previously obtained from wide-angle
seismic studies and receiver functions. The resulting crustal thickness varies from about 35 km in the forearc region to
more than 70 km beneath the plateau and thins (30 km) further to the east in the Chaco plains. Beneath the Andean
plateau, the Moho is deeper in the north (Altiplano) and shallower in the south (Puna), where the plateau attains its
maximum elevation. A non-linear relation exists between crustal thickness and elevation (and Bouguer gravity),
suggesting that the crust shallower than 50^55 km is predominately felsic in contrast to a predominately mafic crust
below. Such a relation also implies a 100 km thick thermal lithosphere beneath the Altiplano and with a lithospheric
thinning of a few tens of kilometers beneath the Puna. Absence of expected increase in lithospheric thickness in
regions of almost doubled crust strongly suggests partial removal of the mantle lithosphere beneath the entire plateau.
In the Subandean ranges at 19^20‡S, the relation between altitude and crustal thickness indicates a thick lithosphere
(up to 130^150 km) and lithospheric flexure. Beneath a relative topographic low at the Salar de Atacama, a thick
crust (67 km) suggests that the lithosphere in this region is abnormally cold and dynamically subsided, possibly due to
coupling with the subducting plate. This may be related to the strongest (Ms = 8.0) known intra-slab earthquake in the
central Andes that happened very close to this region in 1950. The average crustal Vp /Vs ratio is about 1.77 for the
Altiplano^Puna and it reaches the highest values (1.80^1.85) beneath the volcanic arc, indicating high ambient crustal
temperatures and wide-spread intra-crustal melting. ? 2002 Elsevier Science B.V. All rights reserved.
Keywords: Central Andes; Mohorovicic discontinuity; mantle; delamination
1. Introduction
The Andes are widely considered as a typical
* Corresponding author. Tel.: +49-331-288-1246;
Fax: +49-331-288-1277.
E-mail address: [email protected] (X. Yuan).
orogenic belt, formed due to subduction of an
oceanic plate (Nazca plate) under a continental
upper plate (South American plate) (e.g. [1]).
The central Andean plateau, bounded to the
west by the Western Cordillera volcanic arc and
to the east by the Eastern Cordillera thrust-fault
belt, comprises the Altiplano plateau in the north
and the Puna plateau in the south (Fig. 1), and is
0012-821X / 02 / $ ^ see front matter ? 2002 Elsevier Science B.V. All rights reserved.
PII: S 0 0 1 2 - 8 2 1 X ( 0 2 ) 0 0 5 8 9 - 7
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X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
next to the Tibetan plateau in size. The associated
combination of subduction and tectonic shortening processes, which resulted in an unusually large
crustal thickening at the active margin, is not well
understood. Therefore it attracted much attention
in the last decade and stimulated a number of
geophysical studies to image the deep structure
and probe into the processes of the subduction
system. Detailed information on the variations
of crustal properties and thicknesses is crucial
for understanding the mechanism of crustal thickening. Previous geophysical data have shown that
the overriding plate has a Moho at 30^80 km
depth in some places with its deepest parts beneath the Eastern and Western cordilleras. One
of the earliest comprehensive studies of crustal
thickness was that of James [2], who estimated
the maximum thickness in excess of 70 km beneath the Western Cordillera, based on surface
wave dispersion data. The wide-angle re£ection
and refraction data [3^6] indicate Moho re£ections beneath the forearc and backarc areas,
which normally could not be detected beneath
the volcanic arc. This may be due to a transitional
Moho and/or due to the highly attenuating nature
of crust there.
Seismological investigations have detected the
continental Moho across the entire plateau, using
teleseismic or local earthquake receiver function
analysis, regional waveform modeling, and underside Moho re£ection studies [7^10]. Most of these
seismic studies have concluded that the average
velocity of the crust beneath the plateau is low
(Vp = 5.9^6.2 km/s, see [4] and [9]), and its average
Poisson’s ratio is about 0.25 (Vp /Vs = 1.73)
[8,9,11,12]. This was interpreted as an evidence
for the felsic composition of the crust [7^9,11,12].
The teleseismic receiver function technique is
now routinely used to investigate the crustal and
upper mantle discontinuities beneath permanent
and mobile seismic networks. The technique enables generation of high-resolution structural images of the lithosphere and the entire upper mantle, similar to those obtained for the crust by the
near vertical re£ection method. A steeply incident
P-wave from a distant earthquake penetrates the
underlying upper mantle and the crust, before it
reaches a seismic station. At each discontinuity
the P-wave generates an S-wave that follows it
to the station. This converted S-wave is usually
very weak, but can be identi¢ed by its delay and
polarization if a signi¢cant number of records are
available.
Using a modi¢cation of the receiver function
method, Yuan et al. [13] have shown that the
continental Moho is generally seen beneath the
central Andes at depths of 40^80 km, with large
changes in the north^south direction, thinning by
10^20 km from the Altiplano to the Puna. In this
paper we will focus on mapping the Moho topography in the central Andes in greater detail.
2. Data and method
Several temporary arrays of seismic stations
were operated in northern Chile, southern Bolivia
and northwestern Argentina (Fig. 1) for time periods ranging from 2 months to more than 1 yr,
supported by the Sonderforschungsbereich (SFB)
267 of the Deutsche Forschungsgemeinschaft and
the US PASSCAL project in cooperation with
South American institutions (see [13] for details).
Most of the stations were equipped with shortperiod 1-Hz Mark-L4 seismometers (black triangles in Fig. 1) while a small number were
equipped with Guralp-3T and Streckeisen STS-2
broadband sensors (squares in Fig. 1). The experiments were designed to improve the images of
local seismicity, seismic velocity structure and
seismic attenuation structure in the central Andes.
Besides the local earthquake recordings mostly
from the subducted Nazca plate, a number of teleseismic earthquakes and deep regional events were
recorded within the operation periods and have
been used in our receiver function study.
P-waveform data with high signal/noise ratio
are selected at each station from the teleseismic
earthquake recordings with epicentral distances
between 30 and 95‡. Because of the global earthquake distribution pattern and the short observation periods, relatively few useful teleseismic
earthquakes could be recorded. Receiver functions for each earthquake^station pair were calculated in the way described by Yuan et al. [14]. The
response of the short-period instruments was
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391
Fig. 1. Topographic map of the central Andes with major tectonic features indicated. Small triangles are recent volcanoes. Seismic stations are denoted by the the large triangles (short-period stations, operated for about 3 months) and the black squares
(broadband stations, operated for more than 1 yr). The entire area is subdivided into 1U1‡ grids marked by crosses, which are
labeled in alphabetical and numerical order shown at the left and bottom.
broadened by deconvolving their instrument response. Three-component seismograms were then
rotated into a ray-oriented coordinate system in
directions of P-, SV- and SH-waves. The rotation
angles (back azimuth and incidence angle) were
determined by the eigenvalues of the covariance
matrix over a time window spanning the ¢rst few
seconds following the P-wave arrival. Receiver
functions were computed by deconvolving the
P-waveforms from the corresponding SV-wave
components. In total 642 receiver functions were
obtained for more than 170 stations used in this
study.
It is well known that a relatively large trade-o¡
exists between the velocity^depth estimation.
However, since receiver function analysis uses
the di¡erential travel times between the P-to-S
converted (Ps) waves and the incident P-wave,
they are less sensitive to the absolute P- and
S-wave velocities compared to the Vp /Vs ratio.
Although the Ps-waves at the Moho are usually
the strongest phases in receiver functions, multiply re£ected waves between the surface and the
Moho (Pps and Pss) are often observed. Zandt
et al. [15] have demonstrated that multiple reverberations within the crust, if well observed, can be
used together with directly converted waves to
accurately constrain the Vp /Vs ratio and the crustal thickness.
The Ps-waves at the continental Moho beneath
the central Andes and their multiples are well observed in the receiver function data (see also [13]
and its supplements), providing a good opportunity to determine the Moho depth and average
Vp /Vs ratio for the crust. For this purpose, the
grid-search method described by Zhu and Kanamori [16] is adopted, which relies on stacking of
receiver functions for varying Moho depth and
mean crustal Vp /Vs ratio. The best estimates of
the Moho depth and the Vp /Vs ratio are found
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X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
when the Moho conversion and their multiples
are stacked coherently. This is illustrated in
Fig. 2 for a single broadband station in the area
B4 (Fig. 1) located in the volcanic arc of the
Western Cordillera in northern Chile, which was
operated for more than 1 yr. Note that wide-angle
seismic studies [4] could not identify the Moho in
this region. In Fig. 2a receiver function data (20
traces) are displayed equally spaced and ordered
by their back azimuth. The Ps-phase at the Moho
(labeled Ps) and the two major multiple re£ections
at the Moho (labelled Pps and Pss, respectively)
can be correlated among the individual receiver
functions. We summed the individual traces after
applying moveout correction for the Ps and Pps
phases, respectively (the two traces at the top of
Fig. 2a) and found that both the Ps and the Pps
phases have been optimally coherently enhanced.
Their arrival times are 7.9 and 25.7 s, respectively,
after the moveout correction for a reference slowness of 6.4 s/‡. Using the method proposed by
Zandt et al. [15] we estimate a Vp /Vs ratio of
1.80 beneath this station.
Fig. 2b shows the results of applying the gridsearch algorithm of Zhu and Kanamori [16] to the
receiver functions for this station. For each parameter pair of the Moho depth and the crustal
average Vp /Vs ratio, travel times of the Ps and the
multiples Pps and Pss phases have been calculated. Amplitudes of each receiver function trace
corresponding to these calculated travel times of
the three phases are stacked. The optimum values
of the Moho depth and average Vp /Vs ratio correspond to the maximum energy of stacked
phases (58 km Moho depth and 1.80 average crustal Vp /Vs ratio, in agreement with the above estimated values).
C
Fig. 2. (a) Individual receiver functions of a broadband station located in grid B4. Receiver function data in a window
of 310 to 50 s are displayed equally spaced and sorted by
back azimuth (BAZ), which is indicated to the left of each
trace. Vertical straight lines mark the directly converted
phase and two multiple phases, labeled by Ps, Pps and Pss,
respectively. The top two traces are summations after moveout correction for the Ps and Pps phases, respectively, have
been applied to the individual receiver functions. (b) Gridsearch result for di¡erent Moho depths and crustal average
Vp /Vs ratios for receiver functions at the station. High energy (black and dark gray colors) corresponds to the values
of the Moho depths and the average Vp /Vs ratios in the
crust which optimally ¢t the observed P-to-S conversion at
the Moho and its multiples. The white circle indicates the
best ¢t.
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Fig. 3. Synthetic receiver functions displayed in the similar
way as in Fig. 2. Di¡erence is that the individual receiver
functions are sorted by epicentral distances in order to show
clearly the moveout for di¡erent phases of Ps, Pps and Pss.
The e¡ect of the moveout corrections for di¡erent waves can
be seen in the corresponding summation traces.
393
To understand the moveout correction of the
Ps multiples and to demonstrate the estimations
of the Vp /Vs ratio and the crustal thickness using
the grid-search algorithm, we calculated synthetic
receiver functions using re£ectivity method [17]
for a single layer crustal model with an average
Vp of 6.1 km/s, Vp /Vs ratio of 1.80 and crustal
thickness of 58 km (Fig. 3a). We presented the
data as in Fig. 2, but with a small di¡erence.
Unlike those sorted by azimuth, the individual
receiver functions in Fig. 3a are sorted by epicentral distance in order to show the di¡erence in
moveout behavior of Ps and its multiples. The
summation of moveout corrected traces enhances coherently those phases, which have been
used to calculate the moveout correction. Other
phases are suppressed. The grid-search diagram in
Fig. 3b gives exactly the same estimates of the
Vp /Vs ratio and the crustal thickness as the real
data in Fig. 2.
A set of synthetic receiver functions has been
calculated for a slowness of 6.4 s/‡ to test the
sensitivity of the receiver function data to variations in Vp /Vs ratio and crustal thickness (Fig. 4),
using the re£ectivity method [17]. For simplicity,
we used a model with a homogeneous crust. The
model of Fig. 3 was used for reference. We varied
the Vp /Vs ratio between 1.70 and 1.90 and the
crustal thickness between 48 and 68 km and calculated synthetic receiver functions for each model (Fig. 4a,b). Increase in both the Vp /Vs ratio and
in the crustal thickness delays the travel times
of all the phases in the receiver functions. The
delay of the Pps phase caused by the increase in
the Vp /Vs ratio equals the delay of the Ps phase
(Fig. 4a), while the delay of the Pps phase caused
by the increase in the crustal thickness is more
signi¢cant than that of the Ps phase (Fig. 4b).
In Fig. 4c we selected some combinations of
Vp /Vs ratios and crustal thicknesses by keeping
the time of the Ps phase ¢xed (at 7.9 s). The
time variations of the multiple phases are very
signi¢cant in response to the parameter variations. If only the Ps phase is observed, the error
in estimating the crustal thickness is V14 km for
Vp /Vs ratios in the range 1.70^1.90. However, if
the multiple phases are also clearly observed, they
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X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
can be very useful to determine the Vp /Vs ratio
and the crustal thickness more accurately.
In order to apply the grid-search method to the
entire data set of the receiver functions, the study
region was subdivided into grids of 1U1‡ size.
They are labeled in alphabetical and numerical
order shown at the left and bottom of Fig. 1. In
each grid, we use directly converted and multiply
re£ected energy and estimate the average Moho
depth and the Vp /Vs ratio. Sums of all individual
receiver functions within each grid are plotted in
Fig. 5. The traces have been sorted by increasing
Fig. 4. Synthetic receiver functions to test the sensitivity to
di¡erent model parameters. The reference model has a single-layer crust with crustal thickness of 58 km, P-wave velocity of 6.1km/s, Vp /Vs ratio of 1.80. The synthetics were calculated for a slowness of 6.4 s/‡. (a) Synthetic receiver
functions for di¡erent Vp /Vs ratios between 1.70 and 1.90.
(b) Synthetic receiver functions for di¡erent crustal thicknesses between 48 km and 68 km. (c) Synthetic receiver functions for di¡erent combinations of the Vp /Vs ratios and the
crustal thicknesses. The arrival time of the Ps-wave in each
receiver function was kept unchanged.
Fig. 5. Summations of receiver functions in each 1U1‡ grid,
displayed in the order of their Ps conversion times. Di¡erent
¢lters were used to display the Moho Ps and the multiples.
While the Ps window (between 0 and 15 s) is broadband, the
multiples window (between 15 and 40 s) has been 5 s lowpass ¢ltered. Moveout corrections have been applied correspondingly in each window prior to summation. Only data
showing clear Pps multiples were used for the Vp /Vs estimates. Other multiples are not displayed. For some traces
the Pss multiples can be also seen. The corresponding grid
index is marked at the left of each trace. The numbers to the
right of each trace denote the number of receiver functions
stacked. Phase arrivals are marked with white squares.
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times of Ps conversions of the continental Moho.
Because the multiples travel three times longer in
the crust compared to the Ps-waves, they are usually weak and scattered. The data in Fig. 5 are
split into two windows. The ¢rst 15 s window
contains mainly the energy from the Moho conversions. The second window between 15 and 40 s
contains Moho multiples and has been low-pass
¢ltered with a 5 s cut-o¡ period. Moveout corrections have been applied corresponding to the individual receiver functions in each window prior
to summation. Only data showing clear Pps multiples have been used for the Vp /Vs estimates which
are displayed in the second window. In some
traces the Pss multiples can be also seen. The
large amplitude peaks on most traces directly
after the P arrival at 0 s are probably caused by
interference between converted energy at shallow
depth and some P energy remaining from imperfect rotation of the coordinate system.
3. Results
The grid-search algorithm has been applied for
receiver functions at stations within each grid in
Fig. 1. The method works generally well for most
grids, but at some grids the Moho conversions are
di⁄cult to identify. For A1, B1, B2 and C1,
strong converted energy of the oceanic Moho
dominates the receiver functions and masks the
continental Moho conversion. The data at D3
are quite noisy, preventing us from obtaining a
stable Moho stack. The two easternmost stations
in the Chaco plain have been summed together to
estimate the Moho depth (F9).
The resulting map of Moho topography is
shown in Fig. 6a. The Moho depth varies from
35 km in the forearc to more than 70 km beneath
the plateau and it becomes thinner (30 km) further to the east in the Chaco plain. Our results are
very close to the results of previous receiver function studies [7,10], and also to the results of wideangle re£ection studies [4,6]. From our additional
new data we also see a clear variation of the
Moho depth between the northern and southern
parts of the region. North of 23‡S the Moho is
generally deeper than 60 km with its deepest part
395
beneath the Eastern and Western cordilleras
(more than 70 km at E5, C5, E3 and G5), while
south of 23‡S it is generally shallower than 60 km
(in A4 it is as shallow as 42 km), although the
average altitude in the south is higher than in the
north. Note that the tendency of the Moho depth
to decrease southwards was previously reported
from the few wide-angle data and is now fully
con¢rmed with our observations using a di¡erent
method. Another interesting result is the unexpected deep Moho (67 km) beneath the Salar de
Atacama (B3), which has a much lower elevation
than the adjacent mountain regions.
The Vp /Vs ratios obtained from the grid-searching procedure are shown in Fig. 6b. Compared to
the Moho depth estimates, the Vp /Vs ratio estimates are less robust. Nevertheless a general pattern in variation of the ratio can be seen. Most of
the Vp /Vs values are rather high in accordance
with Yuan et al. [13]. The highest Vp /Vs ratios
are mostly observed beneath the volcanic arc
(Fig. 6b). Note that neither the Moho depths
nor Vp /Vs ratios could be estimated at the Altiplano^Puna volcanic complex (APVC, marked by
a red curve in Fig. 6b) and at the volcanic arc,
immediately to the north of the APVC. This is
due to the extremely energetic conversions from
the low-velocity zone in the middle crust and their
multiples [12,15] which tend to mask the Moho
conversions.
At this point, we discuss the errors associated
with the Moho depth and Vp /Vs determinations
(shown in Fig. 6) obtained by the method of Zhu
and Kanamori [16]. We compare these results
with travel time observations of the Ps and Pps
phases in Fig. 5. From these phases alone (see
Fig. 3) we can uniquely determine the average
Moho depth and Vp /Vs ratio for a homogeneous
crust. The errors depend on the reading accuracy,
provided the phases are correctly identi¢ed and
the signal^noise ratio is good enough. It can be
seen in Fig. 5 that the reading error of Ps (marked
by squares) is a small fraction of a second (0.1^0.2
s), whereas the errors of Pps (upright triangles)
are about 0.5 s, although we have selected only
the traces with the best available multiples. Sorting the traces (in Fig. 5) with increasing Moho
conversion time makes it also easier to identify
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X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
the multiples by considering the whole suite of
seismograms. We disregard the error of Ps in
comparison with the larger errors of Pps and estimate from Fig. 4c the errors of the Moho depth
and Vp /Vs due to the estimated U 0.5 s error of
Pps. It follows that the errors of the data in Fig. 6
resulting from the inaccurate readings of the multiples are about 0.02 for Vp /Vs and U 1.5 km for
the Moho depth, with Moho conversion times
kept ¢xed. These errors are about three times
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397
Fig. 6. (a) Map of the Moho depth estimated from the P-to-S conversion at the Moho and its multiples by the grid-search method. The Moho depths are color-scaled and also indicated in each grid with numbers (in km). Shaded grids indicate where the
Moho depth is estimated by primary and multiple phases and therefore an average crustal Vp /Vs ratio is available. For the unshaded grids the Vp /Vs ratio is set to 1.73 for estimating the Moho depth by the primary converted phase alone. White grids
with numbers show Moho depth estimated from near-critical re£ections. We note good consistency between the receiver function
and wide-angle re£ection results. (b) Map of the average Vp /Vs ratio in the crust estimated from the P-to-S conversion at the
Moho and its multiples by the grid-search method. The highest Vp /Vs values are observed close to the volcanic arc. Data at grids
D3, D4 and especially C3 and C4 (indicated by a red contour) are strongly dominated by the negative conversions at middle
crustal depths (see [13] and [18]) which do not allow a reliable estimate of either the Moho depth or the average Vp /Vs ratio but
most likely indicate partially molten crust.
6
smaller than the increments in the color scales
chosen in Fig. 6. In the following we discuss
some extreme values in Fig. 6.
The Moho depth at the neighboring grids E5
(80 km) and F4 (57 km) is quite di¡erent. The
Ps conversion in F4 is good (Fig. 5). It is not
that clear in E5, because there are several more
oscillations following the identi¢ed Moho phase,
and E5 has only four traces. The Pps multiple in
E5 looks well, whereas Pps in F4 has relatively
large noise preceeding it. The other Moho conversions in the vicinity of E5 (F3 and G5) are also
good, con¢rming a region of deep Moho there,
with large depth variation over short distances.
F4 is the only grid in the Altiplano region with a
Moho depth less than 60 km, the largest Moho
depths being 70^80 km. Swenson et al. [11,12] reported Moho depths of 60^65 km under the Altiplano using regional earthquake records. The lateral resolution of the receiver function method is
certainly better than that of any wide-angle method. The Moho depths at grids A2^A6 seems to
oscillate. The suite of traces sampling A2, A4,
F8 and A6 (Fig. 5) varies homogeneously, leading
to a steady increase of Moho depth from A2 over
A4 to A6. The Moho conversion times of A3 and
A5 are later. Most Moho signals are clear in the A
sequence, except that A2 and A3 have double signals, making these signals less certain. However,
we believe that the change in Moho depth at 24‡S
by up to 15 km over short distances is a robust
result. Another clear result is that all the grids in
the Puna have a Moho depth less than 60 km.
Vp /Vs of F8 is questionable since it does not ¢t
in the homogeneous sequence of Ps conversions
and Pps multiples in A2, A4, F8 and A6. It
should rather be between 1.75 and 1.80 according
to the travel time observations alone. Such a value would also agree better with neighboring grids
F7 and F9. The reason for this problem is probably that the optimum amplitude summation
method of Zhu and Kanamori [16] considers a
very late Pss of F8, due to noise. A Vp /Vs ratio
of 1.73 has been reported by Swenson et al.
[11,12] under the entire Altiplano. This more or
less agrees with most of our results from the
BANJO and SEDA stations, but clearly disagrees
with our results at grids F3 and F4. The large Vp /
Vs in F3 and F4 depends on the early Pps arrival
(Fig. 5). In order to be consistent with a Vp /Vs of
1.73, the Pps phase must arrive 2 and 3 s later in
F3 and F4, respectively, than indicated in Fig. 5.
This seems to be unlikely.
4. Discussion
Here we analyze the estimated crustal thickness
in conjunction with the surface topography and
gravity and discuss the possible geodynamic consequences. Fig. 7a shows the average topography
vs. Moho depth in each 1U1‡ grid. Blue diamonds correspond to the grids north of the 22‡S
and red diamonds to the grids south of the 22‡S.
The most interesting feature in the Fig. 7a is a
poor correlation between surface topography
and Moho depth below the Altiplano and Puna.
Although the Moho depth varies by more than
20 km, there is no considerable variation of the
surface topography. Fig. 7b shows that the same
is true for the Bouguer gravity; variable crustal
thickness does not result in signi¢cant variations
in Bouguer gravity under the high plateaus. This
means that either (i) mass de¢cits (excesses) due to
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Fig. 7. (a) Average altitudes versus depths to the Moho in the 1U1‡ grids. Blue diamonds correspond to the grids north of 22‡S
and red diamonds to the grids south of 22‡S. Dotted line shows calculated isostatic altitude^Moho depth relations for the crust
composed from felsic rocks. Solid line, labelled Hl = 100 km, shows the best ¢t of the data using a two layer lower crust composed of felsic and ma¢c rocks. Curve shape is mostly controlled by the crustal composition. The curve is shifted parallel to the
vertical (altitude) axis with changes in thermal lithospheric thickness. Solid curves show expected altitude^Moho depth relations
for the mixed felsic^ma¢c lower crust for 50, 100 and 150 km lithospheric thickness. Most of the data are consistent with mixed
felsic^ma¢c composition of the lower crust and lithospheric thickness about 100 km. However, there are four grids with large deviations from the average curve. Corresponding points are outlined by the circles. (b) Bouguer gravity anomalies [33] versus
depths to the Moho in 1U1‡ grids.
variable Moho depth are compensated by excesses
(de¢cits) of lithospheric density or (ii) the lowermost crust beneath the plateaus has density close
to the mantle and thus changes in thickness do
not produce signi¢cant gravity or topographic
anomalies. Let us consider ¢rst case (i) and assume that the entire crust beneath the plateaus
is felsic [8]. We use the thermodynamic modelling
approach of Sobolev and Babeyko [19] to calculate mineral equilibrium, densities and seismic velocities for the central Andean rocks with bulk
chemical compositions taken from Lucassen et
al. [20]. The thin dotted line in Fig. 7 shows the
calculated isostatic altitude^Moho depth relation
for the crust composed from average central Andean felsic rocks after [20] and an Altiplano geotherm suggested in [21] (the results are the same
for another Altiplano geotherm [22]). This curve
is consistent with our data for crustal thickness
smaller than 50^55 km, but completely inconsistent with a thicker crust. Topography of the
lithosphere^asthenosphere boundary could compensate this mis¢t. However, to do so the lithospheric thickness must vary by more than 200 km,
which is unlikely. The observations, however, can
be readily explained if the lowermost crust below
the plateau has predominantly ma¢c composition
(some 80 vol.% of ma¢c rock and 20 vol.% of
felsic rock) with average density about 3.2 g/cm3
(solid curve in Fig. 7a). Average P-wave velocity
of such lower crustal layer is estimated to be
about 7.3 km/s at temperatures above 1000‡C,
expected at a depth larger than 50^55 km below
Altiplano [21,22]. For comparison, pure felsic
lowermost crust at these conditions would have
a P-wave velocity of about 6.6 km/s, owing to
beta quartz stability. Average velocities of the
pure felsic crust and the felsic crust with a 10
km thick layer of predominantly ma¢c rocks at
the crustal base would di¡er by only 0.09 km/s, a
di¡erence which is hardly detectable by relatively
low-resolution seismological methods.
Summarizing this part of the discussion we
emphasize that our data suggest that the crust
below Altiplano and Puna is felsic to the depth
of 50^55 km but has more ma¢c composition below this depth. This may suggest that the thick
crust in the high-plateau region has achieved its
critical thickness. Further thickening must have
generated high-grade metamorphic rocks with a
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X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
density higher than the upper mantle, leading to
their delamination and detachment [23,24]. If correct, this means that a part of the high-plateau
crust could have been lost during shortening. It
is interesting that ma¢c lowermost crust is present
mostly beneath the Altiplano and is almost absent
below the Puna (with the exception of the grid
C5) in accordance with mantle delamination hypotheses suggested for the Puna [24].
An interesting question is whether a signi¢cant
amount of partial melt is present in the Altiplano^
Puna crust or not. Swenson et al. [12] have concluded that there is no evidence for the partial
melting in the Altiplano crust after comparing
their estimated average Poisson’s ratio (0.25)
and Vp estimates (6.0 km/s) in the crust, with
those expected for a felsic crust. However, possible beta quartz stability in the hot crust of the
Altiplano is not accounted for in their calculation.
We have calculated the expected average Vp and
Vp /Vs ratios (Poisson’s ratio) for a 60 km thick
melt-free crust of felsic composition for convective and conductive temperature distributions in
the crust of Altiplano. The ‘conductive’ model
[22] does not take into account possible convective heat transfer by the partially molten or close
to solidus felsic rocks and therefore generates very
high temperatures in the crust in an attempt to ¢t
high surface heat £ow observed in Altiplano. The
‘convective’ model [21] is based on thermo-mechanical modelling of the tectonically shortened
felsic crust and it predicts temperatures in the
lower crust lower than the ‘conductive’ model.
In our calculation we take into account alpha^
beta quartz transition and associated non-linear
changes of elastic moduli close to the transition
[19]. The estimated average Vp appears to be 6.22
km/s for the ‘conductive’ and 6.17 km/s for the
‘convective’ temperature models. The corresponding Vp /Vs ratios are close to 1.68 (Poisson’s ratio
0.23). From these calculations it is clear that in
order to achieve an average Vp of 6.00 km/s [12] a
small degree of partial melt is required even for
the purely felsic crust. The same conclusion holds
for the crust with a thin ma¢c layer at its base,
because the resulting average velocity and Vp /Vs
ratio is almost the same (6.31 Rm/s, 1.68 for the
conductive and 6.26 Rm/s, 1.68 for the convective
399
model). Depending on melt geometry, presence of
a small degree of partial melt decreases Vp and
increases Vp /Vs ratio. Using the model by Watanabe [25], we obtain the amount of partial melt
required to decrease Vp by 4^5% (from 6.25^6.3
to 6.0 km/s) as 3^4 vol.% with a simultaneous
increase of Vp /Vs ratio by some 4%, reaching
1.75 (0.26). This is just in-between our estimate
of the average Vp /Vs ratio for the crust in Altiplano and Puna (1.77) and that in [11,12] (1.73).
Note, that high Vp /Vs ratios obtained by us correlate well with the pronounced low-velocity zone
in the middle crust (20^40 km depth), detected by
both receiver functions [13] and wide-angle re£ections [4] in Altiplano, suggesting that partial melt
is mostly concentrated in the middle crust, rather
than in the lower crust. This result is in good
agreement with the results of thermo-mechanical
modelling of the Altiplano crust [21].
Thickness of the lithosphere (crustal plus
mantle parts) is another factor, which a¡ects the
isostatic equilibrium altitude of the crust. If only
the lithospheric thickness is varied keeping crustal
densities ¢xed, then the shape of the altitude^
Moho depth curve (solid line in Fig. 7a) does
not change, but this curve shifts parallel to the
vertical (altitude) axis. To a ¢rst approximation, we can quantify this e¡ect using data from
other, better studied regions of the world. Lithospheric thickness in the volcanic regions of the
French Massif Central is estimated to be about
50^60 km [26] from interpretation of the high-resolution tomographic models. Average altitude
(slightly above 1 km) and crustal thickness (close
to 30 km) [27] ¢t the curve labeled Hl = 50 km in
Fig. 7a. Lithospheric thickness is around 100 km
away from the volcanic ¢elds of the French Massif Central [26], which might be typical for tectonically active western Europe [27]. Taking
28^30 km as typical crustal thickness and 0^0.2
km as the altitude for active western Europe we
¢x the curve corresponding to a lithospheric
thickness of 100 km, which is labeled Hl = 100
km in Fig. 7a. Continuing this procedure linearly
we obtain the curve labeled Hl = 150 km which
corresponds to the 150 km thick lithosphere.
These estimates are rather crude, but nevertheless
suggest that the data for the high-plateau are con-
EPSL 6204 21-5-02 Cyaan Magenta Geel Zwart
400
X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
sistent with a lithospheric thickness less than
100 km rather than 150 km. This conclusion is
important because due to strong tectonic shortening, the lithosphere must almost be doubled in the
high-plateau region. This means that either the
lithosphere had to be very thin (some 50 km)
before shortening or a signi¢cant part of the lithospheric mantle had to be removed during shortening. As there is no evidence for a very thin
lithosphere in the entire high-plateau region before shortening, we consider lithospheric mantle
removal to be the most plausible explanation.
Fig. 8 shows deviations of the observed altitudes
from the theoretical curve for the crust with a
ma¢c layer. With a few exceptions, which will be
discussed below, the deviations are small, signi¢cantly less than 0.5 km. There is a small systematic
di¡erence between the Puna (south of 22‡S) and
Altiplano (north of 22‡S), also clearly visible in
Fig. 7a,b, which can be explained either by some
20 km thinner lithosphere beneath the Puna or by
a less dense Puna crust. Analysis of Cenozoic magmatism [24], as well as evidence from lower seismic
velocities and higher attenuation in the uppermost
Fig. 8. Di¡erence between observed altitude and theoretically expected altitude. The largest negative deviations (blue ^ theoretical
altitude is too high) are located in the Subandean ranges (F6, F7) and in the Salar de Atacama block (B3). The largest positive
is in the Puna (A4). Also shown are the grid-search results for four grids with largest di¡erences between observed and expected
altitude.
EPSL 6204 21-5-02 Cyaan Magenta Geel Zwart
X. Yuan et al. / Earth and Planetary Science Letters 199 (2002) 389^402
mantle beneath the Puna than beneath the Altiplano [28], support the idea of a relatively thinner
lithosphere beneath the Puna.
Although most of the data ¢t well with the
theoretical curve shown in Fig. 7a, there are several points which do not. They are outlined by the
circles in Fig. 7. The corresponding 1U1‡ grids
show up as strong blue or red colors in Fig. 8.
As also demonstrated in Fig. 8, the grid-search
method provides good estimations for the crustal
thickness for all these anomalous crustal blocks.
Therefore, these deviations likely re£ect anomalous deep structure and dynamic state rather
than artifacts of our analysis technique. The elevation anomaly in the Subandes (grids F7^F8,
east of 63‡W) can only be partially caused by
£exure of the Brazilian lithosphere, which is
underthrusting the Andean crust [29]. The rest
of the anomaly could be explained by some
30^50 km thicker lithosphere in the Subandes
than beneath the high-plateau. This is consistent
with the higher seismic velocity and lower attenuation in the Subandean upper mantle than in the
adjacent Eastern Cordillera at 20‡S [30].
Of particular interest is the thick (67 km) crust
observed in the Salar de Atacama block (B3),
whose altitude is about 2.5 km. This is more
than 1 km below the value expected from a crustal thickness of 67 km if isostatic equilibrium is
achieved, and it has anomalously high Bouguer
gravity. Elevation of this block would be consistent with the 150 km thick lithosphere (see Fig.
7a), but this could not be the case because the
subducting Nazca plate is located at only 100 km
depth below this block. From this discrepancy
and also from seismic tomographic data showing
high seismic velocities and very low attenuation in
the mantle below this block [28], we infer that the
Salar de Atacama block has an abnormally cold
lithospheric mantle which is mechanically coupled
with the top of the subducting Nazca plate. The
coupling between this cold block (acting as a
giant asperity) and the slab might be the reason
that the strongest (Ms = 8) known intra-slab
earthquake of 1950 in the central Andes [31] occurred close to the Salar de Atacama block.
Another anomalous block (A4) is located close
to the Salar de Atacama block to the other side of
401
the volcanic arc, and it stands higher than would
be expected from its crustal thickness of 42 km if
isostatic equilibrium is achieved. As seen from
Fig. 7a, the observed elevation of this block
(about 4 km) can not be attained even if the ambient crust is purely felsic and the lithosphere is
very thin (50 km). Therefore, we expect that this
block may be in non-isostatic equilibrium. It is
interesting that the average elevation and Bouguer
gravity of the two adjacent anomalous blocks, B3
and A4, treated as one joint block, is very close to
that expected for their average crustal thickness.
If those two blocks were mechanically coupled,
the equilibrium for such a combined heterogeneous block would require block B3 to be dynamically subsided and block A4 to be uplifted [32], in
accordance with our observations.
Acknowledgements
We thank R. Kumar, B. Schurr and D. Hindle
for carefully reading the manuscript and H.J.
Go«tze for providing Bouguer gravity data. The
¢eld experiments have been supported by the
SFB 267 of the Deutsche Forschungsgemeinschaft, the GeoForschungsZentrum Potsdam, the
Freie Universita«t Berlin, the US National Science
Foundation, the BANJO and SEDA projects, the
Universidad de Chile (Santiago), the Universidad
Catolica del Norte (Antofagasta) and the Universidad National Salta.[RV]
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