Download pn velocities beneath continental collision zones: the Turkish

Geophys. J . Int. (1994) 117, 273-283
pn velocities beneath continental collision zones: the Turkish-Iranian
Plateau
Thomas M. Hearn and James F. Ni
Deparlrnent of Physics, New Mexico Stale University, Las Cruces, N M 88003, USA
Accepted 1993 August 6 . Reccivcd 1993 July 30; in original form 1993 February 19
SUMMARY
Tomographic imaging of t h e uppermost mantle velocity in t h e Middle East shows
normal Pn velocities (8.0-8.2 k m s - ’ ) beneath the Black Sea and t h e southern
Caspian Sea. In contrast, low Pn velocities (<7.9 k m s-I) a r e found beneath the
Aegean Sea a n d much of t h e Turkish-Iranian Plateau. A region of exceptionally
low Pn velocities (c7.6 k m s-’) beneath t h e Lesser Caucasus Mountains and along
t h e Turkish-Iranian border coincides with a region of high Sn attenuation and
extensive Neogene volcanism. All these features suggest that near solidus conditions
exist within the uppermost mantle beneath t h e Turkish-Iranian Plateau. Such
conditions may result either from t h e decompression melting associated with
upwelling convection cells o r by t h e infiltration of water released from subducted
lithosphere into the mantle above. T h e presence of a partially melted uppermost
mantle weakens t h e lithosphere beneath the Turkish-Iranian Plateau, thus allowing
it t o become t h e locus of deformation in t h e Arabian-Eurasian collision zone.
Key words: Earth’s mantle, plate tectonics, tomography, Turkish-Iranian Plateau.
INTRODUCTION
DATA
The Turkish-Iranian Plateau extends from western Turkey
to eastern Iran and is among the world’s major continental
plateaus (Fig. 1). The Plateau was formed by late Neogenc
collision of the Arabian plate with the Eurasian continent.
Prior to collision, convergence between the Arabian and
Eurasian plates was accomplished by Jurassic to Neogene
subduction of the Neo-Tethys oceanic lithosphere northward
beneath the southern Eurasian margin. Continued convergence caused the Arabian continent to collide with the
southern Eurasian margin starting in the late Oligocene and
continuing to the present day (Sengor & Kidd 1979;
Dercourt et al. 1986; Boulin 1991). Collision uplifted the
Turkish-Iranian Plateau and the Zagros Mountains by the
end of the Pliocene (Bird 1978; Sengor & Kidd 1979; Boulin
1991). Although the mode of plate convergence changed
from subduction to continental collision, subduction-related
volcanism continues to occur throughout the Plateau
(Kazmin et al. 1986), suggesting that the tectonics of the
Turkish-Iranian Plateau continue to be driven by processes
within the upper mantle beneath the Plateau. To better
understand the effects of subduction and continental
collision on the uppermost mantle structure beneath the
Plateau, we used traveltimes from regional earthquakes to
tomographically image the lateral variations in the P-wave
velocity structure beneath the Middle East.
The data used in this study are traveltimes from regional
events that are listed by the International Seismological
Centre (ISC 1990). Regional earthquakes are defined here
as events recorded at distances ranging from approximately
200km to 1200km. Regional arrivals include true Pn
headwave arrivals at the shorter distances and waves
bottoming at depths of 50-100 km at the farthest distances.
We refer to all these arrivals as the Pn phase. Pn traveltimes
are sensitive to both variations in crustal thickness and
velocity and to the laterally varying velocity structure of the
uppermost mantle. Our tomographic method uses the
regional traveltimes to model both the lateral variations in
crustal structure and the lateral velocity variations in the
uppermost mantle. Similar studies have been applied to Pn
and Pg data from southern California (Hearn 1984; Hearn
& Clayton 1986a,b) and to regional arrivals of the western
United States (Hearn, Beghoul & Barazangi 1991).
First arrival times for events and receivers were selected
for rays travelling between 200 km and 1223 km in the
Middle East region shown in Fig. 2. These traveltimes were
corrected for station elevation and screened by event depth
(<30 km), event size (>25 arrivals recorded world-wide per
event), ISC listed recording precision (<0.2 s), and ISC
residual (<I2 s). An iterative selection procedure then
eliminated insufficiently recorded events and stations and
273
274
T. N . Hearn and J . F. Ni
50"N
40"N
30"N
Figure 1. A map showing locations in the Middle East. Shaded area represents elevations higher than 1oOOm. Solid triangles represent
Quaternary volcanic centres (Dercourt ef al. 1986).
Figure 2. All Pn ray paths used are plotted on a base map. The coverage is thickest in the Aegcan Sea and thinnest in the south
Pn velocities beneath the Turkish-Iranian Plateau
removed obvious outliers from the data set (Hearn et al.
1991). The final data set was fit by a straight line regression
of traveltime to distance and residuals computed relative to
that fit. In the final data set, each event was recorded by at
least 10 stations and each station recorded at least 10 events.
The maximum residual allowed is 9 s relative to the initial
straight line fit. Out of our initial set of 188 095 traveltimes,
a total of 70 755 f n first arrival times that fit our selection
criteria were finally used.
The final set of ray paths for the selected data is shown in
Fig. 2. The number of arrivals per station ranges from the
imposed minimum of 10 arrivals to a maximum of 603
arrivals. As one might expect from the distribution of
earthquakes and seismic stations, ray path density is highest
beneath the Aegean Sea.
TOMOGRAPHIC METHOD
The regional traveltimes are all modelled as refracted Pn ray
paths. The tomographic method fits traveltime perturbations
from the simple linear refraction traveltime curve found
during the data selection process. A refracted ray path can
be described in three portions: the source to mantle ray
path, the ray path through the mantle, and the ray path
from the mantle to the receiver. The source and receiver
components of the ray path are parameterized by source and
receiver static delays. The mantle velocity variations are
parameterized by dividing the mantle into a 2-D grid of
cells. Small cells, of one-quarter degree by one-quarter
degree, were used to obtain a more detailed approximation
to the continuous velocity distribution of the real Earth. A
total of 8210 cells were used.
The traveltimes for these f n ray paths are described by a
modified time term equation (Hearn 1984):
t =a
+ b + 2 disi,
where d, is the distance the ray travels in cell i, si is the
slowness (inverse velocity) for cell i, a is the event static
delay, b is the station static delay, and the sum is over all
cells through which the ray travels. Station and event delays
may be expressed in terms of crustal thickness and crustal
velocity:
station or event delay =
I
(sf - s2)"2 dz,
where s, is the crustal slowness profile as a function of
depth, and s is the average mantle slowness. The integral is
evaluated from sea level to the Moho depth or from the
hypocentre depth to the Moho depth.
The station delays represent the combined effect of crustal
thickness and crustal velocity. Crustal thickness and crustal
velocity cannot be independently determined from the delay
times, but approximate values for the relative variations in
crustal thickness and crustal velocity can be found. For a
crust with a constant velocity of 6.3 km s-', 1s of relative
delay time indicates about 10.4 km of additional crust
thickness. For a 40 km constant thickness crust 1 s of delay
corresponds to a change in crustal velocity of about
1.6 km s-'. Observed variations in mean crustal velocity of
6.2-6.6 km s-' (Braile et af. 1989) can account for a total of
275
0.5 s of relative delay time at most, although large
sedimentary basins can increase that delay substantially.
Since the computed station delays span considerably more
than 3 s, our interpretation of the station delays is primarily
in terms of crustal thickness. Another factor affecting station
delays is systematic picking errors at the station (Grand
1990). These could occur due to clock errors, timing errors
or phase identification errors that repeatedly occur at a
given station. We have no way to estimate such systematic
errors, so individual station delays must be approached with
caution.
Event delays depend on crustal velocity and crustal
thickness and also on the unknown event depth and errors
in the origin time. Because of this, the event delays are of
little use in evaluating crustal structure. They serve only to
isolate the effects of errors in event depth and origin time
from the rest of the tomographic solution.
Since f n ray paths can be modelled roughly as refracted
ray paths, a straight line fit determines the initial crustal
model from the slope and intercept. The slope of the
traveltime versus distance line yields an estimate of the
mean slowness for these ray paths; the intercept gives an
estimate of mean crustal delay. These initial estimates of Pn
velocity and mean crustal delay form the starting model for
the subsequent iterative tomography.
The tomographic method used here is a preconditioned
version of the LSQR algorithm of Paige & Saunders (1982).
This algorithm iteratively finds a least-squares solution to a
set of linear equations of form
Ax==,
(3)
where t is a vector of observed traveltimes; x is a vector of
station delays, event delays and slowness perturbations; and
A is a matrix corresponding to the coefficients for the delays
and slowness perturbations of eq. (1).
Preconditioning is implemented by applying LSQR to find
the least-squares solution for z in the equations
AM-'z = t,
(4)
where z = M x , and M is a preconditioning matrix.
Preconditioning does not change the least-squares norm of
the solution but has a large effect on the convergence rate
(Paige & Saunders 1982). Any matrix M that is an
approximate inverse to A works as a preconditioner. We
used a diagonal matrix that makes z dimensionless and
scales it for the ray path density of each station, event and
cell. For the diagonal elements of the preconditioning
matrix corresponding to the station and event delays, we
used the square root of the number of arrivals at that station
or event. Each diagonal element of the preconditioning
matrix corresponding to a slowness erturbation of a cell
was assigned the quantity m = C Did,, where Dlis the
total ray path length for ray i, d, is the length of ray i within
the cell, and the sum is over all ray paths crossing the cell.
This preconditioning matrix yields a fast convergence rate
for the algorithm.
The solution was regularized by requiring the Laplacian
(second derivative) of the slowness image to be small. The
Laplacian is implemented as a weighted 2-D finite difference
operator in the same manner as Lees & Crosson (1989). The
Laplacian regularization equations were applied directly to
+
276
T. N . Hearn and J . F. Ni
(a) 20"E
50"N
30"E
40"E
60"E
50"E
50"N
40"N
40"N
30"N
30"N
20"E
(b)
20"E
50"N
30"E
30"E
40"E
40"E
50"E
50"E
--z
I
1
1
1
60"E
60"E
0
+ t +
ED
W
-1.5 -1.0 -0.5 0.5 1.0 1.5
;
Station Delays
40"N
40"N
4
30"N
30"N
20"E
50"N
30"E
40"E
50"E
60"E
Figure 3. Uppermost mantle Pn velocity is plotted. High velocities are found beneath the Black Sea and the southern Caspian Sea while low
velocities are found beneath the Aegean Sea and beneath the Turkish-Iranian Plateau. (b) Crustal station delays for the Middle East. The +
symbols indicate station location only; the circles and squares have radii proportional to the size of the station delays. Negative delays (circles)
indicate thin crust or higher velocity crust. Positive delays (squares) indicate thick crust or slower velocity crust. A delay of zero represents a
crustal thickness of approximately 47 km.
Pn velocities beneath the Turkish-Iranian Plateau
the slowness components of the preconditioned solution
vector z rather than the original vector solution x. This
adjusts for variations in ray path density and gives a much
smoother image. The level of smoothing is controlled by the
weight given to the regularization equations. The weight
trades off the error size with the rcsolution width. For this
data set, setting the weight equal to 1.0 gave acceptably
small standard errors (<0.1 km SKI) yet also kept the
resolution width under 3".
RESULTS
The original straight line fit to the data set indicates a mean
Pn velocity of 8.0 k m s - ' and a mean Moho depth of
approximately 41 km with respect to sea level. This crustal
thickness was based on the intercept time of 7.06s, an
assumed crustal velocity of 6.3 km s-I, and an assumed
mean source depth of 10 km. The lateral velocity variations
and the station delays found in the tomography are
perturbations from these values. The crustal depth estimate
calculated from the intercept time is too crude to calibrate
the relative station delays to an absolute crustal thickness. A
better calibration may be made by comparing the station
delays along the northern Aegean coast (-1.0 to - 1.5 s) to
known crustal thickness of 32 to 34 km (Makris 1985). This
gives an estimate for the zero delay level that corresponds to
a Moho depth of 46 km below sea level. While the absolute
crustal thickness may be in error, relative crustal thicknesses
are more accurate.
The tomographic inversion used 25 LSQR iterations
which reduced the root-mean-square error of the residuals
to 1.8s from 2.4s. Further iterations had little additional
effect on either the standard error or the velocity image.
The tomographically imaged Pn velocity variations range
from about 7 . 6 k m s - ' to more than 8.2kms-l (Fig. 3a).
The highest Pn velocities are found beneath the Black Sea
and the southern Caspian Sea. These high velocities are
probably related to the Mesozoic oceanic mantle beneath
these seas (Zonenshain & Le Pichon 1986). High Pn
velocities also are indicated beneath the north-west margin
of the Arabian Plate in northern Iraq, but the Syrian portion
of the Arabian Plate shows low velocities. Low Pn velocities
are found beneath much of the Turkish-Iranian Plateau and
beneath the adjoining Aegean Sea, with some exceptionally
low velocities (7.6 km s-I) beneath the Lesser Caucasus
region. Average Pn velocities (8.0-8.1 km s-I) are found
beneath the southwestern Mediterranean.
The station static delays shown in Fig. 3(b) span over 2 s
and represent predominantly the effect of crustal thickness
variations. Unfortunately, the station delays also may
represent systematic clock or picking errors and this may
account for some of the scatter between neighbouring
stations. Early delays around the edge of the Aegean Sea
represent the thinner crust there. Delays of a similar
magnitude found in Armenia, indicate that the crust is also
approximately 34 km there. Delays within the TurkishIranian Plateau imply crustal thicknesses greater than
50km. The large delays along the western Caspian coast
may be due t o the thick sedimentary section in the Caucasus
Valley and the thick crust of the Turan Platform t o the
north.
277
ERRORS A N D R E S O L U T I O N
The solution variances are calculated by the bootstrap
method (Efron 1979; Koch 1992). The bootstrap works by
using repeated inversions of bootstrap data sets picked at
random from the original data set. By using many different
bootstrap inversions, the resulting set of bootstrap solutions
is then used to estimate the solution variances (see Koch
1992, for details). W e used 100 bootstrap inversion
iterations. The standard deviations for the velocity cells of
the tomographic model are shown in Fig. 4(a) and the
standard deviations of the station delays are shown in Fig.
4(b).
Standard deviations of the station delays (Fig. 4b) range
from nearly 0 s to just over 1 s a t the more poorly recorded
stations. The average station delay standard deviation is
0.25 s, implying an estimated error of approximately 3 km in
estimating crustal thickness variations. Neighbouring stations in Fig. 3(b) often show differences much greater than
0.25s. We believe that this indicates the presence of
station-dependent systematic errors in the original traveltime picks.
Station delays in the Lesser Caucasus region of Armenia
compare in magnitude t o those found along the Aegean
coast, indicating a crustal thickness of nearly 34 km. The
errors in the station delays there are as high as 0.9s, thus
the probable error is approximately 9 km. Furthermore, n o
trade-off between the station delays of that region and the
velocity image is indicated by the checkerboard resolution
tests. Thus, we conclude that crustal thickness beneath
Armenia is less than 45 km.
The average standard deviation for the estimated velocity
perturbations is 0 . 1 4 k m s - ' but some of the cells at the
edge of the image have errors well above 1 km s-'. The
central portion of the image has velocity errors below
0.1 kms-'; this is well below the size of the velocity
anomalies we image. Thus, errors are only a serious
problem for cells at the edge of the image and for the more
poorly recorded stations.
The regions with the highest errors in velocity
(20.2 km s-') are beneath the northern and south-eastern
edges of the image. These regions (the northern Black Sea,
the northern Greater Caucasus, and northern Iraq) have the
thin ray path coverage that causes the large errors in
estimated velocity. Thus, the low velocities imaged beneath
the Black Sea and the Greater Caucasus and the high
velocities of northern Iraq d o not have reliable magnitudes.
To show the resolution of the inversion, we used two test
models. Both models consisted of a checkerboard velocity
pattern, with the first model having a square size of 3" by 3"
(Figs 5a and b) and the second model having a square size of
2" by 2" (Figs 6a and b). The squares were assigned
velocities that alternate between +0.3 km s-' and -0.3
kms-'. Station and event delays were set to zero and no
noise was included. Traveltimes computed from these models
were inverted in the same manner as the original data. For the
first model, the 3" by 3" squares are resolvable for most of the
imaged region in Fig. 5(a) with distortion occurring mostly
around the edges. In the second model, 2" by 2" squares are not
well resolved in central Turkey, the southern Mediterranean
Sea, and around the edge of the image (Fig. 6a). The 2" by 2"
squares are resolvable over eastern Europe, the Black Sea to
278
T. N . Hearn and J . F. Ni
50"N
40"N
4O"N
30"h
30"N
I
L
(b) 20"E
50"N
30"E
40"E
50"E
60"E
50"N
40"N
40"N
30"N
30"N
20"E
30"E
40"E
50"E
60"E
Figure 4. (a) The standard deviation of the cell velocity is plotted for each cell. These were computed by the bootstrap method. (b) The
standard deviation of the station delays is plotted. These were computed by the bootstrap method. The + symbols indicate station location
only; the circles have radii proportional to the size of the station delay error.
Pn velocities beneath the Turkish-tranian Plateau
(b) 2(
50"N
50"N
40"N
40"N
30"N
30"N
279
2
Figure 5. (a) Output Pn velocity model created from a checkerboard test model with cell sizes of 3" by 3". Most o f the squares are resolved. (b)
Output station delays created from a checkerboard test model with cell sizes of 3" by 3". Station delays of zero were used for the input
checkerboard model. The small value for the delays shown here demonstrates that there is little trade-off between the computed station delays
and the velocity image. The + symbols indicate station location only; the circles and squares have radii proportional to the size of the station
delays.
280
7'. N . Hearn and J . F.
(4 20"E
50 "N
Ni
30"E
40"E
50"E
60"E
50"N
40 "N
4O"N
30
30"N
(b)
20"E
30"E
40"E
50"E
60"E
2O"E
30"E
40"E
50"E
60"E
50"N
50"N
40"'v
40"N
30"N
30"N
2
Figure 6. (a) Output Pn velocity model created from a checkerboard test model with cell sizes of 2" by 2". Slightly more than half the cells are
resolved indicating that the resolution width is between two and three degrees. (b) Output station delays created from a checkerboard test
model with cell sizes of 2" by 2". Station delays of zero were used for the input checkerboard model. The small value for the delays shown here
demonstrates that there is little trade-off between the computed station delays and the velocity image. The + symbols indicate station location
only; the circles and squares have radii proportional to the size of the station delays.
Pn velocities beneath the Turkish-Iranian Plateau
the Caspian Sea, and northern Iran. A model with a square size
of 1" by 1" also was used and none of the cells were resolved.
These resolution tests show the velocities imaged represent the
true Pn velocities averaged over a 2" to 3" diameter region.
The checkerboard models produced significant station
delays only for the larger square size, and those stations were at
the edge of the imaged region (Figs 5b and 6b). Even when a
checkerboard with a square size of 5" by 5" was tested, there was
little effect on the station delays. This demonstrates the lack of
trade-off between the velocity image and the station delays.
C O M P A R I S O N WITH O T H E R D A T A
In eastern Europe there have been many previous studies of
the crustal and uppermost mantle structure. Our results
indicate that Pn velocities as low as 7.6 km s-' exist beneath
the Aegean Sea, a modern backarc basin. Pn velocities
increase to 7.9-8.0 km 5 - l beneath the Hellenides in
Greece, the Rhodope Massif in Bulgaria, and in southern
Yugoslavia. Traveltimes from local events show that the
average Pn velocity for the entire region is approximately
7.9 km s - ' (Panagiotopoulos & Papazachos 1985). The
tomographic studies of Ligdas, Main & Adams (1990) and
of Spakman & Nolet (1088) indicate seismic uppermost
mantle P-wave velocities of 7.6-7.8 kms-l beneath the
southern Aegean Sea. Spakman (1991) found high
(8.1 km s-I) uppermost mantle velocities beneath the
northern Aegean coast where we find a lower velocity of
about 7.9 km s I . This difference in the uppermost mantle
P -wave velocities may arise because, in our model, the thin
crust of the Aegean coast is directly accounted for by the
station delays.
Crustal thicknesses in the Aegean region are 20 km
beneath the southern Aegean Sea and 32-34 km beneath
Crete and the Greek coast surrounding the sea (Makris
1985). This represents the thinnest crust in eastern Europe.
Our early station delays around the northern Aegean coast
agree with thin crust. We cannot explain the +0.5 s delay at
Crete other than to suggest either that the crust of Crete is
extremely slow or that there may be a systematic picking
error at that station.
Beneath Turkey, Pn velocities from regional arrivals were
reported as high as 8.3 km
(Kadinsky-Cade et al. 1981)
and as low as 7.7 km s-' (Chen, Chen & Molnar 1980). Our
tomography results indicate average Pn velocities there of
approximately 7.9 km s-'. For the Iranian Plateau, Pn
velocities of 7.6 km s- ' to 8.2 km s-' (Kadinsky-Cade et al.
1981). 8.0 km s-' (Chen eb al. 1980), and 7.9-8.2 km s-'
(Asudeh 1982) have been reported. A low Pn velocity of
7.6 km s-' was reported for the seismic station at Tabriz,
Iran, using events to the north in the Caucasus region
(Kadinsky-Cade et al. 1981). Our study shows that the
lowest Pn velocities are restricted to the Lesser Caucasus
and the region along the Turkey-Iran border. South of the
Zagros Main Thrust, data generally yield higher Pn
velocities of 8.2 km s-' (Kadinsky-Cade 1981; Asudeh
1983). Our tomography results also suggest high velocities
there, but have too high an error to make a quantitative
estimate of Pn velocity.
Low mantle velocities beneath western Eurasia have been
tomographically mapped by Spakman (1991) using both
teleseismic and regional arrivals. In general, in Spakman's
s
C
'
281
maps for the Turkish-Iranian Plateau, the layer 2
(33-60 km) velocities agree with the Pn velocities that we
found. For example, his maps show low uppermost mantle
velocities (approximately 7.6 km SKI) beneath the Lesser
Caucasus and high uppermost mantle velocities (approximately 8.1 km s-I) beneath the Black Sea, southern Caspian
Sea, and parts of the Arabian Shield. Some of the details of
his study differ with what we found. We do not image the
region of relatively high uppermost mantle velocity beneath
Bulgaria that he does, and our region of low Pn velocities
spatially corresponds better with the Turkish-Iranian
Plateau. Because we used mainly horizontal ray paths and
corrected for crustal delays, we feel our results obtained
better lateral resolution than previous studies.
One important conclusion inferred from Spakman's
results is that the low velocities beneath the Lesser Caucasus
extend from the base of the crust at least 100 km into the
mantle. Thus, the lateral zone of low Pn velocities we
mapped beneath the Lesser Caucasus is indeed a feature of
the uppermost mantle and not a crustal effect.
Zones of high Sn attenuation are found in the uppermost
mantle beneath the southern margins of the Black Sea and
the Caspian Sea, the region between the two seas, and the
region from north-eastern Iran to Afghanistan (Molnar &
Oliver 1969; Toksoz & Bird 1977; Kadinsky-Cade et al.
1981). A similar zone of high S-wave attenuation exists
beneath the Aegean Sea (Kovachev et al. 1991). These
regions roughly coincide with our zones of low Pn velocity.
Low Pn velocity is thought to be primarily related to mantle
temperature (Black & Braile 1982) and Sn attenuation is
affected by partial melt within the uppermost mantle. Thus,
both high Sn attenuation and low Pn velocity may indicate
an uppermost mantle near solidus.
Partial melt begins to occur in the upper mantle when the
normal lithospheric velocity is lowered by five to six per cent
from cool mantle velocities (Sato, Sacks & Murase 1989). If
a cool lithosphere has a Pn velocity of 8.2 km s-', then
velocities of less than 7.7-7.8 km s-I must represent
partially melted mantle. At 50 km depth, this corresponds to
a solidus temperature of 1530 K for a dry peridotite mantle.
However, this temperature is drastically lower if water from
subducted lithosphere is present within the mantle wedge
above the subduction zone (Sato et al. 1989). The effects of
both increased temperature and added water on mantle
velocity are similar and cannot be separated by seismic
velocity measurements alone.
TECTONIC I M P L I C A T I O N S
The entire Turkish-Iranian Plateau has Pn velocities
substantially beneath the average continental Pn velocity of
8.1 km s-' (Mooney & Braile 1989) and we interpret these
low velocities as indicating near solidus conditions in the
uppermost mantle. The presence of small amounts ovpartial
melt also is required by the high S-wave attenuation and the
extensive volcanism that occurs within the Plateau (Molnar
& Oliver 1969; Toksoz & Bird 1977; Kadinsky-Cade et af.
1981). If the uppermost mantle is near solidus, then its
buoyancy contributes to the observed low Bouguer gravity
values in the Plateau (Bird, Toksoz & Sleep 1975) and to
the isostatic support of the high elevations, although a major
282
T. N . Hearn and J . F. Ni
part of the isostatic compensation of this plateau comes
from the thickened crust (Bird 1978). Moreover, the
presence of any partial melt weakens the lithosphere
beneath the Plateau thus allowing it to become the locus of
deformation.
While most of the Turkish-Iranian Plateau shows low to
very low Pn velocities, there are some significant regions of
more normal Pn velocity contained within it. For example,
Pn velocities of 8.0-8.1 km s-' are imaged beneath the
Taurides and in the eastern Kirsehir Massif of central
Turkey. These regions of normal Pn velocity may be stable
continental blocks trapped during the closing of the Tethys
Ocean. Another interesting feature is the apparently sharp
division along the 52"E meridian in Iran that separates a
narrow zone of high velocities on the west from low
velocities on the east. This meridian also roughly separates a
zone of low Miocene-Quaternary volcanic activity on the
west from more vigorous activity on the east. These regions
of normal to high Pn velocity, however. are small compared
to the overall region of low Pn velocity imaged beneath the
Turkish-Iranian Plateau. W e thus focus our attention on the
tectonic processes that create the large regions of low
velocity beneath the Turkish-Iranian Plateau.
Advective upwelling in extensional backarc regions results
in decompression melting of mantle material. Backarc
extension dominated the Turkish-Iranian region from the
Jurassic t o the Neogene (Dercourt et al. 1986) and, thus, a
significant amount of mantle upwelling and mixing occurred
during the past and continues to occur. If upwelling were
the only intluence on the observed pattern of Pn velocity
variations beneath the Plateau, then our results would
suggest that currently the upwelling is most vigorous
beneath the Lesser Caucasus and along the Turkish-Iranian
border where we image the lowest velocities (<7.6 km s-I).
The early station delays beneath Armenia suggest that this
region also may have a thin crust. Both low uppermost
mantle velocities and thin crust could be explained by
delaminating the lower part of the crust into the uppermost
mantle. However, the few station delays beneath the rest of
the Turkish-Iranian Plateau indicate thick crust (>50 km)
so any delamination would be limited to the Lesser
Caucasus and along the Turkish-Iranian border. Delamination has been proposed by Pearce et al. (1990) to explain the
. geochemistry of the volcanics of western Turkey. Since the
Lesser Caucasus initially formed as a volcanic arc by the
southward subduction of the Black Sea lithosphere beneath
them during the late Miocene and early Quaternary
(Kazmin et al. 1986; Boulin 1991), any delamination could
be due t o the foundering of the underthrusted Black Sea
slab.
The past tectonics of the Turkish-Iranian Plateau also
may still affect uppermost mantle velocities beneath the
region. Continental collision of Arabia and Eurasia during
the Neogene shut off active subduction and left the
subducted slab to founder into the mantle under its own
weight. A high velocity region imaged at depths of greater
than 320 km by Spakman (1991) could be interpreted as the
remnant slab and the foundering of this remnant slab still
may affect the convection patterns within the uppermost
mantle.
Another explanation for the low velocity patterns found
beneath the Turkish-Iranian Plateau is the addition of water
to the mantle from the subducted oceanic lithosphere.
Dehydration of subducted slabs is responsible for the
creation of subduction-related volcanism and may cause a
significant amount of water to infiltrate the mantle wedge
(Peacock 1990; Thompson 1992). Beneath the TurkishIranian Plateau, water may have accumulated in the
uppermost mantle since the early Jurassic, and regional
variations in the water content of the wedge should then
reflect changes in the subduction history. The locus of
subduction shifted southward from the current position of
the Lesser Caucasus to the southern margin of the
Turkish-Iranian Plateau during the Cretaceous (Kazmin ei
al. 1986; Dercourt et al. 1986) and some southward dipping
subduction also occurred there during the late Miocene and
Quaternary (Kazmin et al. 1986). We speculate that the
Lesser Caucasus region has had a longer history of
subduction than the remainder of the Plateau. T h e earlier
subduction episodes allowed additional water to accumulate
within the upper mantle and, hence, brought that portion of
the mantle closer to solidus than the rest of the
Turkish-Iranian Plateau.
We suggest that both the upwelling of hot mantle and the
fluxing of the mantle wedge with water from the dehydrating
slab occurred within the mantle beneath the TurkishIranian Plateau. Both of these subduction-related processes
can lower the seismic velocity within the mantle wedge.
Seismologically, however, the observed low velocities tell us
only that the uppermost mantle is near solidus; they d o not
tell us whether this is accomplished through convective
upwelling or through the hydration of the uppermost
mantle. If velocities are lowered solely through decompression melting, the velocity anomaly remains only as long as
convection continues and the mantle remains hot; whereas,
if mantle hydration is involved, the low velocities may
become a permanent feature and the magnitude of the low
velocities directly relates to the cumulative amount of
material subducted.
CONCLUSION
High Pn velocities (>8.0 km s-') underlie the oceanic crust
of the Black Sea and the southern Caspian Sea, as well as
the continental shields of Eurasia and Arabia. In contrast,
low Pn velocities ( c 7 . 9 km s ') are clearly found beneath
most of the Turkish-Iranian Plateau and the Aegean Sea,
with some exceptionally low velocities of less that 7.6 km s-'
occurring beneath the Lesser Caucasus Mountains.
The spatial correlation of the region of low Pn velocities
with the Turkish-Iranian Plateau suggests that the low
velocities intimately relate to the subduction and subsequent
continental collision processes. Subduction creates an
upwelling of material in the mantle wedge and this may
cause possible decompression melting of the uppermost
mantle. This upwelling may be enhanced after continental
collision by the foundering of a remnant slab or by
delamination of the slab from the crust. Subduction also
releases substantial amounts of water from the subducted
oceanic crust. This released water can flux the mantle wedge
and bring it nearer to solidus.
Pn velocities beneath the Turkish-Iranian Plateau
ACKNOWLEDGMENTS
This research was supported by National Science Foundation Grant EAR-9105575 and Air Force Office of Scientific
Research Grant F49620-93-1-0429. Map figures used the
GMT-System software (Wessel & Smith 1991). The
manuscript was improved by comments from Muawia
Barazangi, Noureddine Beghoul, Katherine Kadinsky-Cade,
Peter Molnar, Wim Spakman and Steve Ward.
283
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