Published as: Calvert, A., Sandvol, E., Seber, D., Barazangi, M., Roecker, S., Mourabit, T., Vidal, F., Alguacil, G., and Jabour, N., 2000. Geodynamic evolution of the lithosphere and upper mantle beneath the Alboran region of the western Mediterranean: Constraints from travel time tomography, J. Geophys. Res., 105, 10871-10898.
Geodynamic evolution of the lithosphere and upper mantle beneath the Alboran region of the western Mediterranean:
Constraints from travel time tomography
Alexander Calvert, Eric Sandvol, Dogan Seber, and Muawia Barazangi
Institute for the Study of the Continents and Department of Geological Sciences,
Cornell University, Ithaca, New York
Steven Roecker
Department of Earth and Environmental Sciences,
Rensselaer Polytechnic Institute, Troy, New York
Taoufik Mourabit
Tetuoan University, Tetuoan, Morocco
Francisco Vidal
Instituto Geografico Nacional, Madrid, Spain
and Instituto Andaluz de Geofisica, Granada, Spain
Gerardo Alguacil
Instituto Andaluz de Geofisica, Granada, Spain
Nacer Jabour
Centre National de Coordination et de Planification de la Recherche Scientifique
et Technique, Rabat, Morocco
2
Abstract. A number of different geodynamic models have been proposed to explain
the extension that occurred during the Miocene in the Alboran Sea region of the western
Mediterranean despite the continued convergence and shortening of northern Africa and
southern Iberia. In an effort to provide additional geophysical constraints on these
models, we performed a local, regional, and teleseismic tomographic travel time
inversion for the lithospheric and upper mantle velocity structure and earthquake
locations beneath the Alboran region in an area of 800 x 800 km2. We picked P and S
arrival times from digital and analog seismograms recorded by 96 seismic stations in
Morocco and Spain between 1989 and 1996 and combined them with arrivals carefully
selected from local and global catalogs (1964-1998) to generate a starting data set
containing over 100,000 arrival times. Our results indicate that a N-S line of
intermediate-depth earthquakes extending from crustal depths significantly inland from
the southern Iberian coast to depths of over 100 km beneath the center of the Alboran Sea
coincides with a W to E transition from high to low velocities imaged in the uppermost
mantle. A high velocity body, striking approximately NE-SW, is imaged to dip
southeastwards from lithospheric depths beneath the low velocity region to depths of
~350 km. Between 350 and 500 km the imaged velocity anomalies become more diffuse.
However, pronounced high velocity anomalies are again imaged at 600 km near an
isolated cluster of deep earthquakes. In addition to standard tomographic methods of
error assessment, the effects of systematic and random errors were assessed using block
shifting and bootstrap resampling techniques, respectively. We interpret the upper
mantle high velocity anomalies as regions of colder mantle that originate from
lithospheric depths. These observations, when combined with results from other studies,
3
suggest that delamination of a continental lithosphere played an important role in the
Neogene and Quaternary evolution of the region.
1. Introduction
The processes responsible for recycling of lithospheric material back into the
mantle are usually described in terms of subduction. The contribution of other processes
such as lithospheric delamination (sensu strictu, s.s.) [Bird, 1979] or convective removal
[Houseman et al., 1981] continues to be controversial. Both delamination and
subduction models have been proposed to explain the formation of many of the Neogene
extensional basins located along the Alpine convergent zone between Africa and Eurasia.
This study focuses on the Alboran Sea and the surrounding regions, located at the
westernmost limit of this zone (Figure 1a). A variety of models have been proposed to
explain the synchronous extension and contraction that occurred in the Alboran region
and its surroundings during the Early Miocene despite the continued convergence of
Africa and Iberia. Presently, delamination (s.s.) [e.g., Seber et al., 1996a; Mezcua and
Rueda, 1997; Platt et al., 1998] or convective removal models [e.g., Platt and Vissers,
1989; Vissers et al., 1995] are gaining popularity but subduction models [e.g., Lonergan
and White, 1997] are also still being proposed.
Previous tomography studies that have imaged mantle structure beneath the
western Mediterranean [Blanco and Spakman, 1993; Plomerová et al., 1993; Spakman et
al., 1993; Seber et al., 1996b; Piromallo and Morelli, 1997; Bijwaard et al., 1999] have
found a pronounced high velocity anomaly located beneath the Alboran Sea and Southern
4
Spain. However, the possible depth and lateral extent of the anomaly were not well
defined with results interpreted both in terms of a subduction [e.g., Blanco and Spakman,
1993] and delamination models [e.g., Seber et al., 1996b]. Blanco and Spakman [1993],
used both local and teleseismic events extracted from the International Seismological
Centre’s (ISC) catalog up to 1986. Seber et al. [1996b] used more recent data (1989-
1994) from the new Moroccan national network but did not use any local events or
readings from Spanish stations. In this tomographic study, we take advantage of the
expansion of seismic networks in Spain and Morocco and the large volume of new data
that have been collected since these earlier studies and use all available data (1964-1998)
to investigate and accurately map the velocity structure of the region. The resulting
velocity model is carefully evaluated using not only the standard synthetic tests and
examination of the resolution matrix but also bootstrap resampling of the travel time data
and shifting of model block boundaries to eliminate the dependence of the derived model
on a particular parameterization. The implications of our results for previously proposed
models are discussed and synthesized with observations from other recent studies into a
model for the Neogene evolution of the Alboran Sea region.
2. Tectonic Setting
The entire western Mediterranean region has undergone significant extension
despite continuous convergence of Africa and Eurasia since the Cretaceous [Dewey et al.,
1989; Srivastava et al., 1990]. The most striking evidence for this extension is the
Neogene age of the oceanic crust beneath the majority of the western Mediterranean
5
(Figure 1a). The original Mesozoic oceanic crust of the Neo-Tethys that lay between
Africa and Europe was almost completely recycled into the mantle during the Late
Oligocene and Miocene as a subduction zone rolled back to the south towards Africa and
southeast towards Calabria (See Figures 1b and 1c) [Rehault et al., 1985; Bouillin et al.,
1986; Dewey et al., 1989; Yilmaz et al., 1996; Carminati et al., 1998a; Carminati et al.,
1998b]. A narrow NE-SW trending mountain belt generated by Cretaceous-Paleogene
convergence that is believed to have existed along the present day coasts of SE France
and E Spain [Alvarez et al., 1974] collapsed behind the retreating subduction zone and
was dispersed across the western Mediterranean. Fragments of this mountain belt now
form the Internal zones of the Betics, Rif, Tell, and Calabria [Guerrera et al., 1993].
Situated at the western edge of the extended region, the Alboran region suffered
the least extension. It has also experienced the least amount of convergence relative to
the rest of the Alpine belt (~200 km since the Cretaceous [Dewey et al., 1989; Srivastava
et al., 1990]). Convergence was initially NNE-SSW to N-S before shifting to NW-SE in
the Tortonian (~10-7 Ma) [Dewey et al., 1989; Mazzoli and Helman, 1994]. The plate
boundary between Africa and Iberia is diffuse with Neogene continental deformation
distributed over a broad zone more than 500 km wide stretching from the High Atlas in
Morocco to the Betic Cordillera in Spain (Figure 2). Marine studies in the Gulf of Cadiz
and off the Atlantic coast of Morocco demonstrate that the diffuse character of this
deformation extends offshore with no clear plate boundary visible to the west until the
Gloria Fault (Not on map, ~20°W) [Hayward, 1997].
The Alboran Sea is bounded by the arcuate Betic and Rif mountain belts (See
Figures 1a and 2). These mountains are typically subdivided into Internal and External
6
zones [e.g., Choubert and Faure-Muret, 1974] (Figure 2) owing to a pronounced
difference in timing, style of deformation, and lithology. A region of highly deformed
flysch separates these zones in the western Betics and Rif. The Internal zones are
fragments of a collisional mountain belt composed of Mesozoic marine sediments
shortened during Cretaceous and Paleogene convergence. The crustal thickness beneath
this belt may have been in excess of 50 km in places [Platt and Vissers, 1989]. The
flysch was deposited on the southern and western margins of the belt on thinned
continental or possibly oceanic crust [Comas et al., 1999]. During formation of the belt,
the External zones were relatively stable continental margin areas subject to marine
deposition.
During the Late Oligocene-Early Miocene (~25 Ma), the belt collapsed along
with limited injection of basaltic dikes beneath the Betic (northern) portion of the belt
(Figure 2) [Torres-Roldan et al., 1986]. The extension seems to have occurred initially
along NNW dipping faults followed by more substantial extension along WSW dipping
low angle faults [Martínez-Martínez and Azañón, 1997]. Deep crustal rocks were rapidly
exhumed from depths of 40 km at a minimum of 4 km/Ma [Platt et al., 1998]. Peridotites
emplaced at the base of the extending crust in the Early Miocene were rapidly exhumed
beneath extensional faults to form the Spanish Ronda and Moroccan Beni Boussera
peridotite massifs [Van der Wal and Vissers, 1993] (Figure 2). Extension in this
convergent setting was accommodated by thrusting in the bounding flysch and External
zones with a primarily WNW, W, and WSW directed transport direction in the Betics,
Gibraltar, and Rif regions, respectively [Frizon de Lamotte et al., 1991; Kirker and Platt,
1998]. This coeval extension of the Internal zones and shortening in the External zones
7
was accompanied by distributed calc-alkaline volcanism predominantly in the Internal
zones but also on the perimeter of the External zones (Figure 2). By Burdigalian times
(~21-16 Ma) the western portions of the belt (now the West Alboran Basin (WAB)
(Figure 2)) had subsided below sea level and were subject to marine deposition [Comas
et al., 1992]. Active extension of the Alboran sea ceased in the Middle Miocene
(Serravallian/Tortonian, ~10 Ma) [Watts et al., 1993; Chalouan et al., 1997] but
subsidence continued in the WAB and to a limited extent in the east Alboran basin (EAB)
[Watts et al., 1993]. Extension ceased in the Betics in the Tortonian [Johnson et al.,
1997] (perhaps as early as the Serravallian [Andriessen and Zeck, 1996]). Predominantly
west directed transport continued in the Rif and Betics until the end of the Miocene
[Frizon de Lamotte et al., 1991]. The total amount of westerly directed shortening is
significant but not well determined with estimates ranging between 200 and 500 km
[Platt and Vissers, 1989; Balanyá et al., 1997; Martínez-Martínez and Azañón, 1997].
In the Pliocene and Quaternary, gross plate motions reasserted control of the
deformation in the region with primarily NNW-SSE shortening and strike-slip faulting
[Comas et al., 1999]. Uplift of the margins of the extended region differentiated the Rif
and Betic Internal zones from the Alboran Sea. Limited potassic volcanism (shoshonites
and lamprophyres) [Hernandez and Bellon, 1985] occurred in the eastern External zones
during the Pliocene and was superseded by alkaline basaltic magmatism in the
Quaternary.
8
3. Tomography Study
3.1. Data
The data set used for this study was selected from data recorded by multiple
networks in Spain and Morocco (Figure 3). Seismic stations in the region (~96) are
predominantly short-period vertical component with a limited number of 3-component
and intermediate band stations. Digital waveforms from approximately 200
local/regional and teleseismic events between 1989 and 1996 were provided by Instituto
Geografico Nacional (IGN) in Madrid (Spain); Instituto Andaluz de Geofisica (IAG) in
Granada (Spain) and Centre National de Coordination et de Planification de la Recherche
Scientifique et Technique (CNCPRST) in Rabat (Morocco). Supplementary analog
records were also provided by CNCPRST and Physique du Globe at Mohammed V
University (MOH V) in Rabat. P and S arrivals were repicked for the local events and P
arrivals from the teleseismic events located between 30 and 90 degrees from the
recording stations. These picks were associated and merged with catalog data from the
International Seismological Centre (ISC) (1964-1994), USGS National Earthquake
Information Center (NEIC) Earthquake Data Report (EDR) (1995-August 1998), and
Moroccan CNCPRST bulletins to generate comprehensive databases of local and
teleseismic arrivals recorded in Morocco, Spain, Portugal, and Algeria from 1964-August
1998. Errors in the central clock for CNCPRST network required that readings from this
network not be merged with other networks during certain periods. The combined
local/regional data set contained 7100 events with 66,200 P and 48,600 S arrivals. The S
arrivals for local/regional events were included to improve constraints on the earthquake
9
locations (particularly depth). The teleseismic data set comprised 15,700 events with
47,700 P arrivals. Teleseismic arrivals reported as emergent were not extracted from the
ISC and USGS catalogs, as these are likely to contain significant uncertainties.
The independent repicking of phases by the authors using the waveform data
allowed some quantitative assessment of the reading uncertainty (σreading) and hence
determination of reasonable data weights (1/σ2reading) for the larger catalog dataset.
Differences between the independent picks (assuming no systematic biases) made by the
authors and those reported in bulletins were calculated and the standard deviation of this
difference (SDD) was determined for the different reported qualitative picking precisions
(e.g., Impulsive as opposed to Emergent) for each network. The SDD of measurements
was calculated after removal of any arrival time differences (outliers) that exceeded a
certain threshold value. A threshold of 3 seconds was found to significantly reduce the
SDD while still allowing geologically reasonable signals in the data. The standard
deviations calculated after a 3 second cutoff are summarized in Table 1. Not
surprisingly, arrivals reported as impulsive in the ISC and USGS bulletins are of better
quality than those reported as emergent especially for the Spanish networks where picks
are made from digital records. Teleseismic residuals were also examined and found to be
consistent with clear azimuthal variations for most stations in Spain. However, some
stations in Morocco showed some scatter and were assigned an increased reading
uncertainty. An additional theoretical uncertainty of 0.1 seconds was also added to
account for uncertainties caused by inexact ray tracing and model discretization. Local
and teleseismic events recorded by 6 or more stations in the Alboran region were
extracted from the database resulting in a final selected data set of 4,500 local events with
10
55,800 P arrivals and 39,500 S arrivals (Figure 4a) and 1,600 teleseismic events with
16,400 arrivals (Figure 5).
Seismic source regions that provide local data are scattered, indicating the
distributed nature of active deformation in the region (Figure 4). However, the majority
of well constrained crustal earthquakes are concentrated in the Betic and Rif mountain
belts. Intermediate-depth earthquakes occur predominantly in two regions [Buforn et al.,
1995]: (1) A diffuse zone off the coast of Spain in the Gulf of Cadiz, and (2) relatively
concentrated N-S line spanning the Alboran Sea and dipping S from crustal depths
beneath the Betics to a depth of ~150 km beneath the center of the Alboran Sea. This
dipping band of seismicity is overlain by a relatively aseismic zone [Seber et al., 1996a].
A limited number of well-constrained intermediate-depth earthquakes have also been
located off the Atlantic coast of northern Morocco [Seber et al., 1996a]. No well-
constrained earthquakes have been located between depths of 200 km and 550 km in the
region. However, four significant and well-constrained events have been located at
depths of 550-650 km beneath southern Spain [Grimson and Chen, 1986; Buforn et al.,
1995]. One of these events (Not shown in Figure 4), that occurred in 1954 with Mw =
7.8 [Chung and Kanamori, 1976], is the largest instrumentally recorded earthquake in the
region.
The teleseismic events are located predominantly along the subduction zones of
the Americas, the Atlantic ridge system, Zagros and the India-Asia collision zone (Figure
5). Azimuthal coverage is generally good, especially from the east and west, although
there is limited illumination from the south.
11
3.2. Relocation of Local/Regional Events
An initial 1-dimensional velocity model (Figure 6 and Table 2) was chosen using a
combination of a reference model (PM2) determined for Europe by Spakman et al.
[1993] and regional refraction surveys [Carbonell et al., 1998]. Each event was relocated
using the same starting epicentral location and origin time but multiple starting
hypocentral depths of 0.5, 14, 40, 90, 300, and 600 km. The depth was fixed for the early
location iterations to lessen the influence of the starting epicentral location and origin
time on the final hypocenter location. Residual cutoffs (used to downweight outliers)
were progressively reduced from 60 seconds to 3 seconds. Despite the different starting
depths, the solutions for a single event usually converged to a single hypocentral location
and origin time. However, approximately 10% of events converged to two or sometimes
even three locations suggesting the presence of multiple residual minima. In these cases,
the location with the lowest weighted RMS residual was chosen. Events with poorly
constrained locations were removed using criteria 1-8 shown in Table 3. To ensure that
only events with relatively velocity insensitive locations were used for the inversion, the
events were relocated in 10 different velocity models generated by adding a +/- 5%
random perturbation to the reference 1-D model. Any event that moved by greater than
15 km horizontally, 20 km vertically or 25 km in total was discarded [e.g., Abers and
Roecker, 1991]. Finally, to improve homogeneity of the ray coverage in the model and
reduce the biasing of station corrections and 1-D inversions by arrivals from aftershock
sequences, the local data set was geometrically thinned by preserving the event with the
most recorded arrivals in each 5 km cube of the model [Abers, 1994]. A final local data
set of 1100 events consisting of 16,000 P and 11,200 S arrivals was used for subsequent
12
1-D inversions. The three deep earthquakes with available data were not included in this
data set owing to excessive coupling between the event depths and the velocity model
(greater than 20 km variations).
3.3. Inversion Method
The 1-D and 3-D inversions were conducted using the inversion code
SPHYPIT90 developed by Roecker [1982] and further modified by numerous researchers
[e.g., Roecker et al., 1987; Roecker et al., 1993; Abers, 1994; Jones et al., 1994]. The
algorithm follows the nonlinear maximum likelihood formalism of Tarantola and
Vallette [1982] to iterate towards the best-fit solution for the model parameters:
pk+1 = pk+(GkTCdd
-1Gk+Cpp-1)-1[Gk
TCdd-1Gk[d-g(pk)]-Cpp
-1(pk-p0)] (1)
where
p0 - a priori or starting estimate of the model parameters
pk - estimate of model parameters for the kth iteration
d - data vector of observed travel times
g - expected travel times (from ray tracing the model)
Gk - matrix of partial derivatives (dg/dp)
Cdd - diagonal a priori data covariance matrix of reading uncertainties
Cpp - a priori model covariance matrix
The elements of Cdd are modified by two dynamic weights scaled between zero and one.
Local/regional events are multiplied by a distance weight to account for a decrease in
13
picking accuracy at longer distances owing to a reduction in the amplitude and frequency
content of phases with distance. Local and teleseismic readings with relative residuals
greater than 3 seconds are also significantly down-weighted to reduce the effect of
outliers. So the elements of Cdd are:
Cdd(i) = σ2reading/ (distance_weight * residual_weight) (2)
Cpp is our estimated uncertainty in the slownesses assigned to the a priori model
parameters p0 (1-D inversion model). In this case, Cpp is a diagonal matrix with no
covariance assumed between model parameters. The elements of Cpp were chosen to
correspond to a 0.2 km/s uncertainty in the a priori model. Other values were
investigated, including allowing larger uncertainties in the crustal layers, however a
constant uncertainty of 0.2 km/s was found to be optimal with a minimum of
unconstrained assumptions.
Travel times and partial derivatives are recalculated for each iteration by ray
tracing through an average 1-D model calculated between each station-event pair and
accumulating the travel time along this path in the 3-D velocity model. Although this
approach is not true 3-D ray tracing, it was found to be a good approximation by Thurber
and Ellsworth [1980]. Teleseismic arrivals are included using a modification of the ACH
[Aki et al., 1976] method by ray tracing to the base of the model volume using the radial
IASPEI91 model [Kennett and Engdahl, 1991] before tracing onto the station through the
3-D model. Parameter separation [Pavlis and Booker, 1980] allows the new slowness
model and event relocations to be determined in separate steps. First, the slowness
14
inversion is conducted after the removal of the hypocenter effects then the local/regional
events are relocated in the new velocity model using a separate location program. After
each slowness-relocation iteration, events that fail criteria 1-10 shown in Table 3 are
removed. The matrix inversions were performed using Singular Value Decomposition.
Although we inverted for both P and S wave velocity structure in the lithospheric layers,
we will only discuss the P inversion results below. As mentioned previously, the main
reason for including local/regional S wave arrivals was to provide additional stability to
the earthquake locations.
3.4. 1-D Inversion
The local data set was used to invert for a best-fit 1-D model for the upper 150
km. Station corrections were calculated before each iteration step from the mean travel
time residual at a station. The solution converged after three iterations. The best-fit 1-D
model is shown in Table 2 and in conjunction with the station corrections results in a
residual variance improvement of 30%. Mean station delays show a strong regional
variation (Figure 7) suggesting that some crustal structure is being mapped into the
station corrections. The most significant features are the consistently late arrivals to the
stations in the Strait of Gibraltar region and the variation in the stations east of Malaga
from early in the south to marginally late in the north. The substantial station delays in
the Gibraltar region may be partly caused by the significant thickness (up to 10 km) of
intensely deformed low velocity (2.3-5.2 km/s) sediments that were observed in the
flysch regions by refraction experiments [Medialdea et al., 1986]. The significant
shortening and associated faulting that occurred in this region probably further
contributes to the low velocities in this area. The high velocity signal in the station
15
corrections on the southern coast of Spain was also identified by refraction studies
[Banda et al., 1993] and is discussed in more detail below. These station corrections
were held constant for the 3-D inversions.
3.5. 3-D Inversion
The 3-D model is parameterized by subdividing each layer into blocks using
vertical interfaces. After a number of trial inversions and examination of the ray
coverage, we chose the final model parameterization shown in Figure 8. The upper 4
layers were divided into 50 km blocks to take advantage of the dense ray coverage from
the local/regional events. The deeper layers that are controlled predominantly by
teleseismic data were divided into 100 km blocks. Figure 9 shows a hit map for a 50 km
block parameterization throughout the model volume, to provide more detail of
illumination in the deeper layers.
Four iterations were performed with the last iteration providing an improvement
in fit of less than 1%. Comparison of the diagonal elements of the partial derivative
matrix Gk for the first and last iterations indicated that the inversion was fairly linear with
elements rarely different by more than 10% and usually by less than 3%. The final 3-D
velocity model produced an overall variance reduction of 32% relative to the 1-D model.
The final variance was 0.35 sec2 whereas the a priori variance was 0.19 sec2 suggesting
that the model was not over parameterized.
As the location of block boundaries is arbitrary and ray illumination is not
homogeneous, sharp velocity contrasts that do not lie close to a block boundary may be
smoothed over a block or displaced (“leaked”) to a block boundary. Furthermore, any
16
artifacts that are introduced by a particular choice of block boundaries will probably not
be detected by the standard tests used to assess the quality and stability of models. In an
attempt to limit these problems, we used a block shift-and-average approach [Evans and
Achauer, 1993]. 100 separate inversions were conducted using the same starting 1-D
model but with the block boundaries uniformly shifted by (10n+5) km in the north-south
direction and (10m+5) km in the east-west direction for n = -5 to 4 and m = -5 to 4. The
mean RMS residual for all these different model parameterizations was 0.349 sec2 with a
2σ of 0.006 sec2 (minimum = 0.345 sec2 and maximum = 0.368 sec2 ) suggesting that no
one model parameterization provide a better (in a statistically significant way) fit than
any other. All 100 models were resampled onto the same grid with a 10 km horizontal
node spacing. The mean velocity at each grid node was calculated resulting in a smooth
model that was not a function of a particular parameterization with any significant
anomalies in the final model caused by an anomaly being imaged by many different
parameterizations. A model derived using only 16 block shifts of 25 km showed
identical features (allowing for sampling differences) so this less computationally
demanding shifting scheme was used for the synthetic tests described below. Some
measure of the parameterization dependent stability of the velocity model was provided
by the standard deviation of the velocities. Areas of high variance will not only occur in
regions of poor stability but may also indicate the presence of large but spatially limited
(on the scale of blocks) velocity anomalies.
3.6. Uncertainty Analysis
The influence of systematic and random errors on the model is often hard to
quantify owing to the non-linear nature of seismic tomography, the difficulty of
17
measuring the amount of noise in the input data, and the effects of parameterization. A
number of traditional and non-traditional methods of uncertainty estimation were used to
ensure that only robust well-constrained features were interpreted. These methods are
reviewed below using an example cross section located 25 km north and parallel to cross
section BB' (Figure 10). Figure 10a shows a cross section through the final shift-and-
average velocity model generated by stacking the results of 100 inversions using
different model parameterizations. Figure 10b shows the velocity model resulting from a
single inversion demonstrating that the first order features are the same as those found in
the shift-and-average model.
3.6.1. Bootstrap resampling of residuals. The increasing speed of modern
workstations now allows the use of non-linear uncertainty estimators to evaluate
inversion results. Hearn and Ni [1994] used a bootstrap resampling technique [Tichelaar
and Ruff, 1989] to determine uncertainties for 2-D Pn velocity models. The same
approach was used to evaluate the 3-D tomography model determined in this study. The
bootstrap method involves repeated inversions of resampled data sets and examining the
variance in the derived models. To generate a resampled data set, arrivals times are
randomly sampled from the data pool. However, after a data point is sampled it is
returned to the pool so it may be picked multiple times. The same number of data are
picked from the pool as existed in the original data set but each data set will be different
owing to the aforementioned data redundancy and resulting omission of other data points.
A velocity model is determined from these resampled data using the same inversion
parameterization. This model is stored and the process repeated approximately 100 times
to generate robust parameter distributions. The standard deviation of a particular model
18
parameter is a measure of the uncertainty in that value in the final model. An important
strength of a bootstrap method is that, in contrast to Monte-Carlo simulations, it does not
require an estimate of the noise in the data to assess the uncertainty in the resulting
model. The resampling technique allows the noise in the data to contribute to the
variance in the derived models. However, it should be noted that in this case the
bootstrap approach does not completely remove the need for noise estimates to measure
model uncertainty because they are required by the inversion method (the elements of
Cdd). A bootstrap conducted by resampling events rather than individual phases
produced virtually identical results. The bootstrap uncertainties are invariably higher
than standard errors calculated using the a posteriori covariance matrix (compare Figures
10c and 10d). The standard errors are less than 1% beneath the Betics in the lithospheric
layers and less than 1% beneath the Alboran, Betics, and Rif in the deeper layers. The
bootstrap uncertainties determined for layer depths of 100-670 km, which are controlled
predominantly by teleseismic data, were encouraging with important features imaged
with anomalies greater than twice the bootstrap error (see Figures 10b and 10d).
Although the crust and uppermost mantle layers showed significant variance (sometimes
approaching 4%), these regions also contained the largest perturbations with robust
anomalies again above twice the bootstrap uncertainty. A useful method for interpreting
our model data was to subtract between one and two times the bootstrap uncertainty from
the imaged anomalies setting an anomaly to zero if the anomaly changed sign after this
subtraction. Only the more robust anomalies remained for interpretation.
3.6.2. Resolution matrix. The resolution matrix provides an estimate of the
relative contributions of the observations and the a priori model to the final model. The
19
diagonal elements of the resolution matrix are very close to one beneath the Betics
suggesting that the ray geometry and choice of model parameterization and damping
parameters allows the imaged crust and mantle structure to be controlled by the
observations rather than the a priori model [Jackson and Matsu'ura, 1985] (Figure 10e).
Crustal structure beneath Morocco is not well controlled owing to the limited number of
local earthquakes. Off-diagonal elements of the resolution matrix were also examined to
evaluate smearing. The off-diagonal elements are usually an order of magnitude less
than the diagonal elements suggesting that at least in theory there is very limited
smearing of anomalies or generation of compensating anomalies in adjacent blocks.
3.6.3. Spike test. Spike tests are traditionally used to assess resolution in
tomographic inversions. Synthetic data are generated by ray tracing a velocity model
consisting of a regular 3 dimensional pattern of perturbations to the reference model (e.g.,
Figure 10f). A representative amount of noise is often added and the synthetic travel
times inverted following the same procedure that was used for the real data. The
recovered velocity model is compared to the velocity model used to generate the
synthetics. Regions where the recovered model closely matches the input model are
considered well resolved in the real velocity model whereas a poorly recovered region is
considered to be suspect in the final model. The addition of noise usually degrades the
recovered data but in this case it has a limited effect (compare Figures 10g and h).
The degree of recovery is also sensitive to the geometry of the synthetic test
[Lévêque et al., 1993] and in particular its relationship to the model parameterization
used to conduct the inversion. Figure 10h would seem to indicate that even with a
representative level of noise added to the data, very good model recovery is possible.
20
However, if the model parameterization is shifted laterally 25 km with respect to the
synthetic model, the amount of recovery is significantly reduced. A shift-and-average
inversion demonstrates that contrary to what would be inferred from examination of the
resolution matrix or from conducting a synthetic spike inversion with velocity
perturbations matching block boundaries (Figure 10h), discrete blocks of 100*100*50 km
are not resolvable at depths of 100-400 km (Figure 10i). A synthetic test conducted using
a single +5% spike of 150x150km in the 200-260 layer showed that this was not an
artifact caused by the choice of spike distribution. Synthetic anomalies spanning two
layers (Figures 10i & j) are imaged suggesting that depth rather than lateral resolution is
the limiting factor. The recovery in the deepest layers of even the thin anomalies is
surprising. It probably results from rays becoming more horizontal with depth and hence
traveling a greater horizontal distance before crossing a layer boundary. However, it
should be noted that these spike inversions might be giving an optimistic assessment of
resolution in the deepest layers. At these depths, the radius of the first Fresnel zone for a
1 Hz wave is ~75 km but the effects of signal frequency are not included in the synthetic
travel times generated by ray tracing. Certainly, there is little to be gained from using
blocks much smaller than 100 km at these depths.
3.7. Interpretation and Evaluation of Velocity Model
Figures 11a-m and Figures 12a-d show horizontal and vertical slices through the
block shift-and-average model. The ray segments close to the cross sections are also
shown to provide an indication of the control of the travel time observations on the
different portions of the model (please see our website at: atlas.geo.cornell.edu for depth
21
slices of these ray segments). Regions that contain many crossing rays are better
controlled than regions that contain a few sub-parallel rays.
The most robust feature imaged in the upper crustal layer (Layer 1) is the
pronounced low velocity region seen in vicinity of the Strait of Gibraltar (Figure 11a)
both in Spain and Morocco. The low velocities are obtained despite the inclusion of
significant station corrections for stations in this region (see Figure 7). Inversions
conducted without station corrections yielded low velocity perturbations of ~18%
indicating that 8% is probably a minimum value and that a significant portion of the
anomaly may be being absorbed by the station corrections. These low velocity anomalies
are significantly larger than the bootstrap standard deviation of ~3%. Low velocities in
this region were imaged by other tomography studies [Blanco and Spakman, 1993;
Gurria et al., 1997; Piromallo and Morelli, 1997] and refraction studies [Medialdea et
al., 1986] and, as mentioned earlier, may be caused by a combination of low velocity
sediments and significant crustal deformation in the flysch regions. A high velocity
anomaly is imaged in the eastern Rif to the SE of the Nekor fault (NF, Figure 2) in both
crustal layers. This anomaly coincides with an aeromagnetic anomaly subparallel to the
Nekor fault that was interpreted by Michard et al. [1992] as indicating the presence of
peridotite.
The most significant feature in the lower crustal layer (Layer 2) is the high
velocity anomaly imaged off the Alboran coast of Spain in the vicinity of Malaga (Figure
11b). High velocities were also identified in this region by a refraction experiment
[Banda et al., 1993]. These authors reported that ray trace modeling was unable to
discriminate between a high velocity peridotite body and a thin crust as the source of this
22
anomaly. The absence of stations in the Alboran Sea results in very poor control of its
shallow velocity structure with only limited rays passing through this layer beneath the
Alboran Sea (See Figure 9). Gravity modeling and limited refraction profiling [Working
group for deep seismic sounding in the Alboran sea, 1978; Ben Sari, 1987] provide
evidence of thinned (~20 km) crust but the underlying mantle velocities reported were
highly variable ranging from 7.5-8.4 km/s on intersecting profiles and, therefore, may be
suspect. Low velocities are also imaged in the Rif and Betics.
Layer 3 from 30-45 km contains striking anomalies, consisting of a ‘horseshoe’ of
anomalously low velocities (relative to an already low mantle velocity of 7.9 km/s)
beneath the Betics, west Alboran Basin (WAB) and Rif surrounding a region of high
velocity beneath the Alboran Sea (Figure 11c). The low velocity region may result from
a mixture of two phenomena: (1) A crustal signature resulting from crustal thicknesses of
up to 38 km in places beneath the Betics, and (2) The presence of anomalously low
velocity mantle [e.g., Banda et al., 1993]. The high velocity region in the center of the
Alboran Sea is surprising and difficult to reconcile with our expectations of what to find
beneath a region that underwent extension during the Miocene. One possible geological
explanation for this anomaly is that a significant thickness of thermal lithosphere has
been recovered since the termination of extension in the Middle Miocene. Although,
Polyak et al. [1996], based on heatflow measurements, interpreted a lithospheric
thickness in this region of only ~40 km. The central Alboran Sea region contains few
crustal earthquakes and effectively no stations. Another possible explanation is that it is
an artifact caused by assumption inherent in our approach. For example, we assume that
a 1-D velocity model is a good approximation for the first ray tracing and inversion
23
iteration. The 1-D model places the Moho at depth of 30 km. Rays connecting
earthquakes on one coast of the Alboran with stations on the opposite coast, in reality, are
probably traveling at depths of ~15-20 km while they are traced through the model at
depths of 30 km. This extra segment of ray path before refracting along the Moho results
in the theoretical arrival time being delayed relative to the actual arrival times. However,
as the rays are not passing through the 15-30 km layer above, the velocities are increased
in the 30-45 km layer in the least constrained region in the center of the Alboran Sea. If
this is true, then one must also consider the possibility that the surrounding lows might be
enhanced by the inversion attempting to compensate for the effect of these elevated
velocities on longer ray paths. Another possibility is that significant anisotropy may
exist in the mantle and that inhomogeneous ray coverage might be causing this
anisotropy to be mapped into the isotropic velocity model. To check these hypotheses
we conducted three tests:
(1) Synthetics were generated using a mantle velocity of 7.5 km/s in the 15-30 km layer
beneath the Alboran. After adding noise, these data were inverted using the same starting
1-D velocity model. A significant portion of the anomaly in the 15-30 km layer was
mapped into the 30-45 layer below, confirming that thinned Alboran crust might be a
viable explanation for this anomaly. However, although slight compensating lows were
generated beneath the Betics and Rif, they were minor compared to the very pronounced
low velocity anomalies imaged in the actual model.
(2) As a further check on the independence of the Betic/Rif lows from this artifact,
inversions were conducted after Pn rays passing beneath the Alboran were removed. The
African and Iberian phase data were split into separate data sets and all local rays that
24
intersected an E-W line passing through the Strait of Gibraltar removed. These data sets
were inverted separately using the same model parameterization. Although these
inversions used no ray paths traveling beneath the Alboran Sea significant low velocity
anomalies were still imaged beneath the Betics and Rif.
(3) To investigate these anomalies further, as well as the possible influence of anisotropy,
the Pn arrivals extracted for various station-event separations were inverted for Pn
velocity and anisotropy using the code of Hearn and Ni [1994]. Inversions were
conducted using varying amounts of Laplacian damping on the isotropic and anisotropic
velocity components to evaluate any trade-offs. Even when anisotropy was very weakly
damped; the low velocity beneath the Betics was imaged for all distance ranges. The Rif
anomaly was also imaged but found to be less robust than the Betic anomaly perhaps
owing to the sparser ray coverage. Zero or positive velocity perturbations were imaged
beneath the Alboran Sea for all distance ranges and damping parameters. The use of
shorter ray paths did accentuate the high velocities beneath the Alboran Sea lending
support to the thin crust hypothesis; however, shorter paths might also be less affected by
a thin mantle lid. Even with strong Laplacian smoothing it was not possible to extend the
imaged low velocities beneath the Alboran Sea, suggesting that uppermost mantle
velocities beneath the central Alboran may be higher than 7.9 km/s.
The 45 - 60 km layer (Layer 4) again shows low velocity anomalies beneath the
eastern Betics and Rif (Figure 11d), but these lows are now separated by a region of
slightly elevated velocity beneath the western Betics and Rif. The low velocity beneath
the Betics conforms well to the surface outcrop of the Internal zone and its western
boundary is marked by intermediate-depth seismicity. Higher velocities are again
25
imaged beneath the Alboran but these anomalies are relatively small and within one
bootstrap error and so should not be considered robust.
The 60-100 km layer (Layer 5) is arguably the most important layer for evaluating
any geodynamic models, but also one of the most difficult to constrain as it marks the
depth where ray coverage from local/regional data is limited to only those from
intermediate-depth earthquakes and teleseismic rays are still limited to the regions close
to the stations (Figure 11e). The most robust feature is the low velocity zone along the
southern coast of Spain that extends eastwards from the N-S line of intermediate-depth
seismicity. Despite the appearance that this anomaly extends beneath the Alboran Sea to
the African coast, this feature is predominantly controlled by teleseismic rays to the
Spanish coastal stations that do not sample the central Alboran. Beneath the western
Betics and Rif a striking N-S high velocity anomaly is imaged to the west of the
intermediate-depth seismicity. The bootstrap uncertainties indicate that the portion
beneath the flysch unit of the Betics is probably robust.
In layer 6 (100-150 km), a robust high velocity anomaly is imaged beneath the
WAB and west and central Betics (Figure 11f). The SE extent of this anomaly is not well
defined owing to the absence of good ray coverage (Figure 9). Assuming that positive
velocity anomalies may be taken to indicate a region of cooler mantle, rather than a
compositional difference, the most likely interpretation is that it indicates the presence of
mantle lithosphere in view of the coherent continuation of the anomaly to depth. The
high velocity anomaly beneath central Iberia was not found to be stable using the
bootstrap error test.
26
Layers 7, 8, and 9 (150-200-260-330 km) show a continuous progression of the
high velocity anomaly to the southeast beneath the Alboran Sea (Figures 12g,h, and i).
However, it should be noted that if the entire Alboran Sea is underlain by a high velocity
region at these depths, this progression might in part be a reflection of the progressive
improvement of ray coverage beneath the Alboran Sea with depth.
The anomalies become relatively small and diffuse in layers 10 and 11 (330-410-
490 km) (Figure 11j and k). However, significant concentrated anomalies are imaged in
layers 12 and 13 (490-570-650 km). Any features imaged towards the base of the model
must be interpreted with great care, as these regions are the most likely to contain
artifacts caused by deep mantle structure outside the model volume. One of the
important assumptions of an ACH-type approach is that any deviations from a spherically
symmetric earth outside the model volume will be sampled equally by all the teleseismic
rays from a single event. Masson and Trampert [1997] showed that this assumption is
sometimes not valid for a true heterogeneous earth. Examination of a recent global
tomography model [Bijwaard et al., 1999] suggests that some heterogeneity may exist
directly beneath the model volume but the size of imaged anomalies is limited (variations
of less than 1% from mean).
Cross sections AA', BB', CC', and DD' (Figure 12) provide a good summary of the
major features seen in the velocity model and its relationship to seismicity. A low
velocity anomaly is imaged in the uppermost mantle along the SE coast of Spain. Its is
bounded on its western edge by the intermediate depth seismicity. A high velocity body
appears to dip south and east from near lithospheric depths beneath Spain and the Strait
of Gibraltar under the low velocity region. This high velocity anomaly is imaged to be
27
continuous to depths of 350-400 km where it reduces in amplitude and becomes more
diffuse. A pronounced high velocity anomaly is again imaged at depths of 500-650 km
near the base of the model in the vicinity of the deep earthquakes.
3.8. Synthetic Smearing Test and Comparison to Previous Velocity Models
The depth extent and volume of the high velocity body in the upper mantle are
important constraints for any geodynamic model for the region. Despite no explicit
regularization between blocks, the smoothing inherent in least squares and the shift-and-
average method will result in smearing of the imaged anomalies. To examine the effect
of this smearing, synthetic travel times generated using a representative velocity model of
the region (Figure 13) were inverted. Gaps were placed in key regions of the model
(Layers 5, 10, & 11 – Figures 13e,j & k) where smearing may have been occurring.
Only limited smearing occurred into the 60-100 km layer (Figure 13e). A small
positive anomaly was mapped into the region just to the west of the intermediate-depth
seismicity, in the vicinity of the imaged high velocity in the real model, suggesting that
smearing may have played some role in generating this anomaly. However, low
velocities are not smeared into this layer from above suggesting that the low velocities
beneath the southern coast of Spain in this layer are real. The existence of anomalous
low velocity mantle beneath the coastal Betics is further supported by the low Q values
found for this region by Canas et al. [1991]. Smearing clearly occurs into layers 10 and
11, indicating that the inversion is probably biased towards filling in any gap if it exists.
Most previous studies have reported a relatively continuous high velocity body
from depths of 200 to 650 km [Blanco and Spakman, 1993; Spakman et al., 1993;
Piromallo and Morelli, 1997; Bijwaard et al., 1999]. Examination of Blanco and
28
Spakman's depth slices shows that similar to our model, a high velocity is also imaged
beneath the western Betics, Gibraltar, and Rif in their 70-120 km layer before moving
east in deeper layers to join with the deeper anomaly at 200 km. The studies by
Plomerová et al. [1993] and Seber et al. [1996b] image a high velocity body beneath the
Alboran and Southern Spain but owing to bottom layers beginning at 200 and 150 km,
respectively, they provide limited information on the depth extent of this body. Bijwaard
et al. [1999] show a section cut from their global model (BSE) in the Alboran region that
also shows a pronounced reduction in the size of the anomaly beneath 400 km (although
this reduction in amplitude in the transition zone may be a global feature of the BSE
model [Personal communication, Spakman, 1999]). All these inversions used similar
data sets and so they possess the same potential for systematic smearing errors.
Establishing continuity of this body from 350-500 km is probably beyond the resolution
of studies using existing station distributions. However, existence of a single coherent
body extending from < 200 km to 650 km is difficult to explain given the constraints on
the relatively small amount of Tertiary convergence between Africa and Iberia (~200
km). The hypothesis that the deep earthquakes are occurring in lithospheric bodies of
limited size is supported by the waveforms recorded in Spain. Buforn et al. [1997]
suggest that one possible explanation for puzzling arrivals from these deep earthquakes
may be caused by S to P conversions near the source and propose that they are occurring
in a body of limited size.
29
4. Discussion
4.1. Previously Proposed Models
Recent publications have converged to two competing but not necessarily mutually
exclusive hypotheses, one proposing a progressive rolling or peeling back of lithosphere,
the other proposing detachment of lithosphere. Within these broad categories, there is
also debate about the involvement of crust in any recycling of lithosphere. Figure 14
shows a simple categorization and schematic cartoons of the most current models taken
from their original publications.
4.1.1. Alboran block and retreating subduction. The predominant westerly
transport direction in the Betics and Rif lead to the proposal of the Alboran block model
[Andrieux et al., 1971]. The Internal zones were considered to have behaved as rigid
block that escaped to the west during N-S convergence. Subsequent observations that the
Internal zones underwent extension during westerly migration of the Alboran block
necessitated modification of the model. Migration of a collapsing Alboran block were
proposed to result from roll-back to the west of an east-dipping subducting slab, with the
Alboran block extending in a back arc setting [Royden, 1993]. Royden suggested that arc
migration may have proceeded past Gibraltar. No evidence of such subduction zone
migration through the Gulf of Cadiz was found by recent marine surveys [Hayward,
1997]. More recent models propose that subduction stopped retreating in the Gibraltar
region [Sengör, 1993; Lonergan and White, 1997].
30
4.1.2. Slab break-off. An alternative subduction related model was initially
proposed by Blanco and Spakman [1993] based tomography results and further modified
by Zeck [1996]. Zeck suggested that a NW dipping slab was subducted beneath SE Iberia
during the eastward movement of Africa and Iberia relative to Eurasia during the
Cretaceous and Paleogene. This subducted slab was proposed to have broken off during
the earliest Miocene initiating uplift, extension, and peripheral thrusting, and now is
oriented vertically beneath southern Spain. This hypothesis is attractive as it provides a
mechanism for considerable lithospheric thickening and/or subduction of significant
quantities of lithosphere beneath Spain, as the significant length of the imaged high
velocity anomaly (~450-500 km) is hard to explain given the limited convergence
between Africa and Iberia. However, it should be noted that most of the eastward motion
of Africa and Iberia probably occurred before the Mid Cretaceous [Srivastava et al.,
1990], requiring any subducted lithosphere to remain hanging in the mantle for 60
million years before finally detaching and descending into the mantle.
4.1.3. Convective removal. P-T paths determined for rocks in the Betics and
Alboran as they were exhumed to the surface lead Platt and Vissers [1989] to propose
convective removal of mantle lithosphere as a possible mechanism. They proposed that
lithospheric thickening as a result of convergence had generated a collisional ridge with a
substantial upper mantle root by Oligocene time. In their model, this gravitationally
unstable lithospheric root was convectively removed in the Late Oligocene-Early
Miocene and replaced by hot low-density asthenospheric material, resulting in uplift and
extension of the overlying crust. Thermal modeling of additional P-T paths derived from
31
data from ODP drill holes in the Alboran Sea have recently lead Platt et al. [1998] to
propose that lithospheric material was removed at approximately 60 km depth, very close
to the base of the over-thickened crust. Platt et al. [1998] comment that removal of
lithosphere at such a shallow depth is more in keeping with a delamination (s.s.) model,
but may still be modeled by a convective removal process given specific mantle rheology
and temperature boundary conditions.
4.1.4. Delamination. The apparent inability of a convective removal model to
explain the predominant westward migration of deformation lead García-Dueñas et al.
[1992] to hypothesize that a delamination model [Bird, 1979] might be applicable to the
Alboran. Docherty and Banda [1995] and Tandon et al. [1998] proposed a SE migrating
delamination model based on marine seismic reflection profiles and wells. Observations
of the 3-D pattern of earthquake hypocenters, gravity modeling and regional wave
propagation characteristics also lead Seber et al. [1996a] to propose a delamination
model. The N-S line of intermediate-depth seismicity apparently overlain by an
aseismic, attenuating, low velocity mantle and underlain by a region of high velocity
mantle [Seber et al., 1996b] was considered evidence that Iberian and African lithosphere
peeled back from the east to west [Seber et al., 1996a] perhaps initiated by detachment of
a gravitationally unstable thickened lithospheric root beneath the Internal zones. Mezcua
and Rueda [1997] and Morales et al. [1997] supported this hypothesis based on their
earthquake relocations. Seber et al. [1996a] further suggested that the delamination
process is still active today in the western Alboran.
32
4.2. Implications of Tomography Results
The most striking feature of the velocity model we obtained is the significant
positive anomaly that appears to extend from lithospheric depths beneath the Strait of
Gibraltar and southern Spain to depths of about 350-400 km beneath the Alboran Sea.
We interpret this anomaly as a lithospheric body that has descended into the upper
mantle. Although a low velocity anomaly is present above this body beneath the
southeastern coast of Spain, the body appears to be attached to Iberia to the west and
possibly north (Figures 11 and 12). The continuation of intermediate depth seismicity
from the top of this body into the Iberian crust may be a further indication of this
attachment (Figure 4). Slab break off or convective removal models proposing complete
detachment of a lone lithospheric body in the Early Miocene and no further mantle
processes are probably not sufficient to explain such a geometry. The tomography results
alone do not allow discrimination between retreating subduction and delamination
processes, as both would have very similar velocity signatures. However, these
processes would have produced a different thermal structure in the Alboran lithosphere
during extension in the Late Oligocene and Early Miocene [Platt et al., 1998]. A
retreating subduction model would predict stretching of the overriding plate in a back arc
setting whereas delamination (or detachment) models would predict a sudden heating at
the base of a thinned lithosphere associated with upwelling of asthenosphere to replace
the sinking body. Platt et al. [1998] present strong evidence that lithosphere was
removed from near the base of an over thickened crust and explicitly state that a
retreating subduction mechanism cannot explain the P-T path followed by rocks
33
presently beneath the western Alboran Sea. The above process of elimination leaves a
delamination based model as the simplest way to explain these observations.
The above argument is probably an over simplification, as it does not allow more
than one active process during the Neogene. The peridotites in the western Betics and
Rif (Figure 2) show evidence of cooling in the mantle at 70 km depth before
emplacement at the base of the crust [Van der Wal and Vissers, 1993]. This cooling
signature is interpreted by Vissers et al. [1995] to result from the location of these
peridotites in the hanging wall of a subduction zone some time before the Neogene,
suggesting that subduction probably played a role in the formation of the collisional belt.
The existence of significant high velocity anomalies at depths of 550-650 km, coincident
with the few deep earthquakes, and perhaps detached from the shallower body raises the
possibility that some lithosphere may have become detached completely during the
Neogene. It is also possible that the pull of this hanging lithosphere may be allowing
some subduction of Iberian continental crust as Africa and Iberia continue to converge
[Morales et al., 1999].
4.3. Synthesized model
Figure 15 shows our modified version of the working hypothesis model proposed
by Platt and Vissers [1989] (compare with Figure 14). Lithospheric thickening occurred
during Paleogene convergence and probably involved limited subduction. At the end of
Oligocene and Early Miocene the gravitationally unstable upper mantle root partially
detached beneath the thickened Alboran crust and peeled back to the west with a portion
of the lithosphere perhaps detaching completely and descending into the upper mantle.
34
Crust above the inflowing asthenospheric material was heated, uplifted, and extended.
Today, although this peeling back has significantly reduced in speed or perhaps
terminated, a large lithospheric body remains beneath the Alboran Sea attached to Iberia
beneath the western Betics and possibly the Rif. The intermediate depth earthquakes
appear to be occurring near the top surface of this hanging body. Whereas the deep
earthquakes may be located in detached pieces of lithosphere perhaps as they undergo
phase transitions at the base of the upper mantle.
5. Conclusions
A tomographic inversion of local, regional, and teleseismic phase arrivals
recorded by Spanish and Moroccan regional networks images significant P velocity
anomalies beneath the Betics, Rif, and Alboran Sea. Evaluation of the resulting velocity
model using bootstrap resampling, block shifting, and synthetic tests suggest that a robust
high velocity anomaly is located beneath southern Spain steeply dipping to the SE from
depths of about 60 km to 400 km. Below this depth, the high velocity anomaly becomes
diffuse. A concentrated high velocity anomaly in the 570-650 km layer, although
probably poorly constrained, correlates well with the location of the few but significant
deep earthquakes beneath southern Iberia. We interpret these high velocity bodies to
represent cold lithospheric material recycled into the upper mantle.
Previous tomographic studies have imaged a single high velocity body beneath
the Betics and Alboran from about 200 – 650 km depth, but detailed synthetic and error
tests that we performed suggest that the apparent continuity of this high velocity anomaly
may be partially caused by smearing from shallower and deeper depths owing to poor
station coverage within the Alboran Sea. The high velocity upper mantle body that we
35
mapped is overlain by a region of low velocity uppermost mantle beneath the Betics.
This velocity model together with recent results from the latest Alboran ODP leg provide
strong evidence that lithospheric delamination (s.s.) played an important role in the
Neogene evolution of the Alboran region.
36
Acknowledgments. This work was supported by National Science Foundation
Grant EAR-9627855. We would like to thank the members of Instituto Andaluz de
Geofisica (Spain), Instituto Geografico Nacional (Spain), Centre National de
Coordination et de Planification de la Recherche Scientifique et Technique (Morocco);
and the Physique du Globe at Mohammed V University (Morocco) for their hospitality
and for providing the waveform data used in this study; Claudia Piromallo for
enlightening discussions about the different tomographic studies in the Alboran region;
Harmen Bijwaard and Wim Spakman for providing the Alboran portion of their global
model; Tony Watts for insights into the tectonics of the region and for providing
preprints of upcoming papers; and Rob Van der Hilst for valuable discussion and review.
The authors would also like to thank Andrea Morelli, Wim Spakman, and Rinus Wortel
for their careful reviews and suggestions, and Graham Brew, Francisco Gomez, and José
Morales for their reviews of early drafts. Institute for the Study of the Continents
contribution number 255.
37
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48
Table 1. P and S reading uncertainties assessed by comparing independent picks of the same
arrival
Uncertainties Data Type Impulsive
(P phase /S phase) sec
Unreported+ (P phase /S phase)
sec
Emergent (P phase /S phase)
sec Local, IGNa and IAGb 0.2/1.0* 0.4*/- 0.4/0.8 Local, CNCPRSTc 0.5/0.8 0.8*/0.7 0.8/0.7 Local, Mohammed Vd 0.4/0.8 0.4/0.8 0.4/0.8 Teleseism, IGN and IAG 0.3 0.3 - Teleseism, CNCPRST‡ 0.3 0.6 - +Phase not reported as either impulsive or emergent
*Limited data aInstituto Geografico Nacional, Madrid, Spain bInstituto Andaluz de Geofisica, Granada, Spain cCentre National de Coordination et de Planification de la Recherche Scientifique et Technique,
Rabat, Morocco dPhysique du Globe, Mohammed V University, Rabat, Morocco
‡ Teleseismic residuals for IFR, TAF, & RBA were found to be particularly scattered hence
arrivals reported to these stations were assigned an increased reading uncertainty (1 second).
Table 1
Table 2. Initial and final 1-D models
Interface Depth (km)
Initial P vel (km/s)
Final 1-D P vel (km/s)
Initial S vel (km/sec)
Final 1-D S vel (km/s)
-5 - 15 5.80 5.75 3.40 3.32 15 - 30 6.50 6.53 3.75 3.80 30 - 45 7.90 7.89 4.50 4.54 45 - 60 8.00 8.09 4.50 4.64
60 - 100 8.10 8.12 4.50 4.67 100 - 150 8.10 8.12 4.50 4.67 150 - 200 8.15 8.15 4.52 4.67 200 - 260 8.20 8.20 4.55 4.67 260 - 330 8.40 8.40 4.65 4.67 330 - 410 8.80 8.80 4.80 4.80 410 - 490 9.45 9.45 5.15 5.15 490 - 570 9.80 9.80 5.30 5.30 570 - 650 10.15 10.15 5.50 5.50
The final best-fit 1-D model was obtained from inversion of local data. The average velocity in
each layer of final 3-D model was insignificantly different from 1-D inversion results.
Table 2
Table 3. Criteria used to remove poorly constrained events
Criteria No. Criteria Value 1 Condition Number < 100 2 Weighted RMS residual < 2.0 sec 3 Closest Station Distance < 150 km 4 Azimuth Gap < 200 degrees 5 Average Residual < 1.0 sec 6 Theoretical Location Error < 10 km 7 Final Iteration step < 3 km 8 Depth > 0.1 km 9 Horizontal Shift < 10 km
10 Vertical Shift < 10 km The criteria were applied after relocation. Criteria 1,3, and 4 check that a good station geometry
is used to determine the location. Criteria 2 and 5 check that the location provides a reasonable
fit to the data. Criterion 7 checks that the solution converges. Criterion 8 was required to remove
a significant number of events that attempted to converge to a location above ground. Criteria 9
and 10 were used during the 1-D and 3-D inversion iterations to remove any events that were
shifted a significant amount when relocated in the new velocity models.
Table 3
49
Figure Captions
Figure 1. Maps showing regional setting of the Alboran Sea and schematic
paleogeographic reconstructions of western Mediterranean [Modified from Alvarez et al.,
1974; Rehault et al., 1985; Dewey et al., 1989; Yilmaz et al., 1996; Lonergan and White,
1997].
(a) The Alboran Sea is one of a number of Neogene extensional basins that formed
along the Africa-Eurasia plate boundary zone despite continued convergence between
the two plates since the Cretaceous. MA = Middle Atlas and SA = Saharan Atlas.
Box shows Alboran region and location of Figure 2.
(b) Extension behind a subduction zone retreating to the E and S in the Late-
Oligocene/Early Miocene resulted in the collapse of a thickened mountain belt that
used to extend along the coast of Spain and France. The Paleogene location of the
thickened crust that later formed the Betic and Rif Internal zones and the westerly
extent of the subduction zone are not well constrained. Arrows show the transport
direction.
(c) Extension persisted until the Early/Middle Miocene with continued westward
transport of the Internal zones in Alboran region.
Figure 2. Simplified tectonic map of Alboran region. AH = Al-Hoceima, AR = Alboran
Ridge, EAB = East Alboran Basin, NF = Nekor fault, SAB = South Alboran Basin, SG =
Strait of Gibraltar, WAB = West Alboran Basin, p = location of Rhonda (Spain) and
50
Beni-Bousera (Morocco) peridotites, b = Early Miocene basaltic dykes, c = Mid Miocene
calc-alkaline volcanism, s = Late Miocene potassic volcanism, and b = Pliocene and
Quaternary alkali basalt flows [Bellon and Brousse, 1977]. Note absence of volcanism
west of Malaga and concentration of calc-alkaline volcanism in the center of the Alboran
Sea. [Map modified after Lonergan and White, 1997; Platt et al., 1998; Comas et al.,
1999].
Figure 3. Seismic networks in the region used for this study. CNCPRST = Centre
National de Coordination et de Planification de la Recherche Scientifique et Technique
(Morocco); MOH V = Mohammed V University (Morocco); IAG = Instituto Andaluz de
Geofisica (Spain), IGN = Instituto Geografico Nacional (Spain). Bracketed term in
legend describes the type of data used in this study. D = Digital, A = Analog, and B =
Bulletin.
Figure 4. Block diagrams showing seismicity in the Alboran region (same area as Figure
2). (a) Data compiled from local (CNCPRST) and global (ISC & NEIC) catalogs and
relocations made by authors. NS and EW projections of seismicity located between the
dashed lines are shown on block sides with x6 vertical exaggeration. Note the N-S line
of intermediate depth seismicity that dips beneath the Alboran Sea from beneath city of
Malaga (see Figure 2 for location) and the three deep earthquakes located beneath the
central Betics. Scattered intermediate depth earthquakes are also located beneath the
Gulf of Cadiz. The horizontal lines of seismicity at depths of 10 and 33 km are result of
ISC and NEIC operators fixing the depth to obtain an epicentral location with limited
51
data. Data shown formed the initial pool of local/regional data used for the velocity
inversions.
(b) Events relocated in derived 3D velocity model after removal of poorly constrained
events (failed criteria 1-8 shown in Table 3).
Figure 5. Map showing the 1600 teleseismic earthquakes used for the velocity inversion.
The size of the symbol marking each event indicates the number of stations in the
Alboran region that recorded an impulsive arrival from this event. Note good E-W
azimuth coverage. Illumination is limited from N and S especially from the south.
Circles mark 30 and 90 degrees from the center of the Alboran Sea.
Figure 6. 1-D Velocity model used for earthquake relocation and other representative
velocity models.
(a) Velocities for the uppermost layers were chosen based on local refraction
experiments. Refraction velocities are from a compilation presented by [Carbonell et
al., 1998] and references therein.
(b) Velocities at greater depths were chosen to lie close to PM2 [Spakman et al.,
1993]. Jefferys-Bullen and IASPEI91 models are shown for reference.
Figure 7. Map showing station delays obtained from 1-D inversion. Note significant
delays for stations in the western Betics and Rif, and early arrivals near the Spanish
Alboran coast.
52
Figure 8. Block diagram showing model parameterization and locations of cross-
sections (Figure 12). 50 x 50 x 15 km blocks were used in upper 60 km. 100 x 100 km
blocks increasing from 40 to 70 km in thickness with depth were used from 60 to 650 km.
Dashed line marks the area shown in Figures 10-13.
Figure 9. Maps showing hit count for 50 km width blocks throughout the model volume.
Note that for the inversion 100 km blocks were used below 60 km depth.
Figure 10. Cross sections showing an example section through the final 3-D P velocity
model and the different attributes used to assess it. The cross sections are all located in
the same position parallel and 25 km north of BB' along the southern coast of Spain (See
Figure 8 for location of BB’). A topographic profile along the line of the cross section is
shown for reference.
(a) Final velocity model calculated using block shift-and-average approach (average
of 100 separate inversions conducted with 10 km block shifts, see text for details).
Anomalies shown are relative to the 1-D model shown in Table 2. Earthquakes
(circles) are projected in from a maximum distance of 50 km perpendicular to the
section. The most striking feature is the high velocity anomaly in the upper mantle.
This anomaly reaches its shallowest point in the west where it is separated from the
low velocity region by the intermediate depth seismicity.
(b) Result of a single inversion using the model parameterization shown in Figure 8
(no block shifting). The first order features shown in Figure l0a, although more
difficult to interpret, are visible.
53
(c) Standard error in block velocities calculated using a posteriori covariance matrix.
(d) Uncertainty in block velocities calculated using non-linear bootstrap resampling.
The high velocity mantle anomalies are found to be of the order or greater than two
times the bootstrap error (See Figure 10b) suggesting that they are robust features for
this model parameterization.
(e) Diagonal elements of the resolution matrix. High diagonal elements indicate
strong control of observations on inversion result. Diagonal elements are higher in
the west owing to the larger number of illuminating rays.
(f) Spike test input model. 100 km blocks with +/- 0.5 km/s perturbations relative to
reference model were used, separated by a block with zero perturbation. A depth
slice would also show a similar pattern of perturbations.
(g) Result of inversion using synthetic arrival times calculated using spike model
shown in Figure 10f. No noise was added to data and block boundaries used for
inversion were coincident with the boundaries of the perturbations. Recovery is
almost perfect in areas of illumination.
(h) Identical to 10g except representative Gaussian distributed noise was added to the
data. The addition of noise degrades the recovery but not substantially.
(i) Result of shift-and-average inversion with noise added to arrival times. Recovery
is very poor throughout the upper mantle layers showing that degree of recovery of
anomalies the size of a model block is strongly dependent on the model
parameterization. Note the surprising recovery of velocity structure at the base of
model (See text for discussion).
54
(j) Modified synthetic model with spikes the same width as in Figure 10f but now
extending over two layers with a two layer vertical gap. Horizontal size and spacing
of perturbations was not changed.
(k) Shift-and-average recovery of spikes spanning two layers. Limited structure is
now imaged in the mid-upper mantle but smearing is significant, indicating that depth
resolution may be limited.
Figure 11. Horizontal slices through the center of each layer of the final P velocity
model. Bracketed term in the title is the reference velocity. Darker and lighter gray
shades indicate regions of higher and lower velocities relative to reference model,
respectively. Earthquakes were relocated from the starting data set using the 3-D
velocity model (same locations as Figure 4b) with poorly constrained events removed.
Main features (see text for detail discussion): (a) Low velocities in western Betics and
Rif; (b) High velocity anomaly off southern coast of Spain east of Malaga; (c)
‘Horseshoe’ of low velocities surrounding a high velocity region in the central Alboran;
(d) Low velocities beneath Betics and Rif; (e) High velocity beneath western Betics and
Rif separated from a low velocity along southern coast of Spain by intermediate depth
earthquakes; (f-i) Progression of high velocity anomaly to south and east with depth
beneath the Alboran Sea; (j and k) High velocity anomaly becomes more diffuse; (l and
m) more concentrated anomalies imaged in the vicinity of deep earthquakes.
Figure 12. Cross sections: (a) AA', (b) BB', (c) CC', and (d) DD' of P velocity model
(left) determined using block shift-and-average and ray coverage (right). See Figure 8
55
for locations. Darker and lighter gray shades indicate regions of higher and lower
velocities relative to reference model (Table 2), respectively. Gray rays are associated
with local events and black rays with teleseismic events. Earthquakes shown with ray
plots are events actually used for the inversion. Earthquakes shown on the velocity
model plots were relocated from the starting data set in the 3-D velocity model with
poorly constrained events removed. Triangles are stations used for the inversion. Only
earthquakes and ray segments located within 50 km of the cross sections are shown. A
significant high velocity anomaly is imaged in the upper mantle that appears to dip to the
SE from lithospheric depths to approximately 350 km before becoming more diffuse.
The intermediate depth earthquakes mark a west to east transition from high to low
velocity in the uppermost mantle. Significant anomalies are also imaged at the base of
the model near the location of the deep earthquakes (Marked by arrows on AA’ and DD').
Only one of the three deep earthquake locations in the data set was sufficiently well
constrained by arrivals in the region to be plotted.
Figure 13. Horizontal and vertical sections through the synthetic (SYN) and recovered
(REC) models used to investigate smearing of upper mantle anomalies. Synthetic travel
times were calculated using model SYN designed to include the principle features of the
derived velocity model (Compare with Figures 11 and 12). However, Layers 5, 10, and
11 (e, j, and k) were left with no perturbation in the synthetic model to determine whether
anomalies would be smeared into these layers from other layers. Perturbations are +/-
5% in the synthetic model. Inversion of the synthetic travel times after the addition of
suitable noise produced model REC. Limited smearing occurs into layer 5. However,
56
significant smearing occurs into layers 10 and 11 indicating that determining slab
continuity at these depths may be beyond the resolution of the data set used for this
study.
Figure 14. Models previously proposed to explain the geodynamic evolution of the
Alboran region. These models may be categorized in terms of roll-back or detachment of
lithosphere and whether or not crust is involved in recycling of the lithosphere into the
mantle. (a) Retreat from east to west of an east dipping subducting slab during the Early
Miocene [Lonergan and White, 1997]. (b) Break-off of a NW dipping slab at the
Oligocene-Miocene boundary perhaps subducted during earlier (mainly Cretaceous)
eastward movement of Iberia relative to Eurasia [Blanco and Spakman, 1993; Zeck,
1996; Zeck, 1997]. (c) Delamination (s.s.) of lithosphere thickened during Paleogene
convergence initiated during Early Miocene but still active today [Seber et al., 1996a].
(d) Convective removal during the Early Miocene of mantle lithosphere thickened during
Paleogene convergence [Platt and Vissers, 1989].
Figure 15. Evolutionary model showing map view and two cross sections modified from
working model proposed by Platt and Vissers [1989] (compare with convective removal
model shown in Figure 14) to explain results from this tomography study and Platt et
al.'s [1998] more recent results. Thick dashed lines show suggested position for
connection of delaminated lithosphere to surface lithosphere. G.B. = Guadalquivir Basin.
Circles show locations of present subcrustal seismicity. Note the body shown in the
upper mantle in the "Present" panel of BB' is attached to Iberia to the west of the cross
57
section. Patterns are the same as those shown in Figures 2 and 14. The crooked arrows
shown in the Strait of Gibraltar marks the relative motion of Africa to Iberia remaining
before present.
Calabrian A
rcIberia
Africa
Alboran Sea
TyrrenianSea
Vale
ncia
Tro
ughPyrenees
Tell Atlas
Rif
Betics
Apennines
Atlantic
Ocean
Bay of Biscay
High Atlas
MASA
5°W 0° 5°E 10°E 15°E10°W
40°N
35°N
0 500km
Figure 1
Cretaceous-PaleogeneThrust Belts
NeogeneThrust Belts
NeogeneOceanic Crust
Neogene ExtendedContinental Crust
NeogeneBasin
Tethyan Oceanic Crust
AtlasMountains
Stable Block
Thrust Subduction
Algerian-Provençal Basin
?
?MioceneOceanic Crust
?
Middle Miocene
kmcb
a
MesozoicOceanic
Crust
?
Sub
duct
ion
?
?
?
Flysch
?
km
Late Oligocene -Early Miocene
Pre
betic
Sub
betic
200
Mo
rocc
an M
eset
a
Gu
lf o
fC
adiz
Atl
anti
c O
cean
Med
iter
ran
ean
Sea
Bal
eari
c Is
land
s
AR
Mal
aga
0
WA
BS
G
EA
B
SA
B
Fly
sch
Nap
pes
Bas
inE
xter
nal Z
one
Inte
rnal
Zon
e
Atla
s M
ount
ains
Sta
ble
Blo
ck
Iber
ia
Gua
dalq
uivi
r B
asin
Gha
rb
Bas
in Hig
h A
tlas
Mid
dle
Atlas
Rif
Tel
l Atl
as
km
c
s as
sa
a
aa
as
ss
cc
cc
cc
cc
c
c
cc
c
p p
bb
b
a =
alka
li b
asal
t fl
ow
sb
= b
asal
tic
dyk
esc
= ca
lc-a
lkal
ine
s =
sho
shin
ite/
lam
pro
ph
yre
(po
tass
ium
en
rich
ed)
p =
per
ido
tite
Figure 2
5°W
0°10
°W
0°
35°N
NF
AH
5°W
0 km 200
Network Administrator (Data type)
Figure 3
Atlantic
Mediterranean
CNCPRST (DAB)
MOH V (AB)
IAG & IGN (DB)
Other (B)
35°N
40°N
10°W 5°W 0°
1501209060300
Dep
th (
km)
North
100 km
0
Figure 4
EQ Depth< 50 km
50-500 km
> 500 kmA
B
6-89-1213-2021-3031-50
# of Stations
Figure 5
30 d
egre
es
3 4 5 6 7 8 9 10 11700
600
500
400
300
200
100
0
Velocity (km/s)
Dep
th (
km)
3 4 5 6 7 8 9100
80
60
40
20
0
Velocity (km/s)
Dep
th (
km)
S P
Figure 6
Jefferys-Bullen (ISC)IASPEI91PM2
1D RelocationRefraction
(a)
(b)
> 10.5 -1.00.2 - 0.5-0.2 - 0.2-0.5 - -0.2-1.0 - -0.5< -1.0
P delay (sec)
Figure 7
35°N
40°N
10°W 5°W 0°
0 km 200
Atlantic
Mediterranean
800
600
400
200
0
200
400
600
800
600
400
200
0
200
400
600
600
500
400
300
200
1000
Figure 8
Depth (km)
E-W
Mod
el C
oord
inat
e (k
m)
N-S M
odel
Coo
rdin
ate
(km
)B
B'
A
A'
C
C'
D
D'
a) L
ayer
1:
-5-1
5 km
b) L
ayer
2:
15-3
0 km
c) L
ayer
3:
30-4
5 km
d) L
ayer
4:
45-6
0 km
e) L
ayer
5:
60-1
00 k
mf)
Lay
er 6
: 100
-150
km
g) L
ayer
7: 1
50-2
00 k
mh)
Lay
er 8
: 200
-260
km
i) La
yer
9: 2
60-3
30 k
mj)
Laye
r 10
: 330
-410
km
k) L
ayer
11:
410
-490
km
l) La
yer
12: 4
90-5
70 k
m
m)
Laye
r 13
: 570
-650
km
110
100
1000
50 k
m b
lock
hit
coun
t
Figure 9
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
W EAtlantic Iberia/Alboran Med.
a) Shift P velocity perturbation (%)
Dep
th (
km)
W-E Model Coordinate (km) -4
-3
-2
-1
0
1
2
3
4
Percent velocityperturbation
relative to 1-D model
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
-4 -3 -2 -1 0 1 2 3 4
b) No shift P velocity perturbation (%)
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
0.5 0.6 0.7 0.8 0.9 1
e) P resolution diagonalsd) Bootstrap Error (%)
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
0 0.5 1 1.5 2 2.5 3 3.5 4
Figures 10a-e
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
c) Standard Error (%)
0 0.5 1 1.5 2 2.5 3 3.5 4
Figure 10f-k
% P velocity Perturbation
-4 -3 -2 -1 0 1 2 3 4
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
k) Double Layer spike: Noise & Shift
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
j) Double layer input spike model
i) Single layer spike: Noise & Shift
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
h) Single layer spike: No shift & with noise
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
g) Single layer spike: No shift and no noise
500 400 300 200 100 0 100 200 300 400 500
600
500
400
300
200
100
0
f) Single layer input spike model
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
a) L
ayer
1, -
5-15
km
< 5
.75
km/s
>
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
KM
KM
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
b) L
ayer
2,
15-3
0 km
< 6
.53
km/s
>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
c) L
ayer
3,
30-4
5 km
< 7
.89
km/s
>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
d) L
ayer
4,
45 -
60 k
m <
8.0
9 km
/s>
Figure 11a-d
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
g) L
ayer
7, 1
50-2
00 k
m <
8.1
5 km
/s>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
h) L
ayer
8, 2
00-2
60 k
m <
8.2
0 km
/s>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
f) L
ayer
6, 1
00 -
150
km <
8.1
2 km
/s>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
e) L
ayer
5,
60-1
00 k
m <
8.1
2 km
/s>
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
Figure 11e-h
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
j) La
yer
10, 3
30-4
10 k
m <
8.8
0 km
/s>
Figure 11i -l
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
i) La
yer
9, 2
60-3
30 k
m <
8.4
0 km
/s>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
k) L
ayer
11,
410
-490
km
< 9
.45
km/s
>
500
400
300
200
100
010
020
030
040
050
040
0
300
200
100
0
100
200
300
400
l) La
yer
12, 4
90-5
70 k
m <
9.8
0 km
/s>
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
500 400 300 200 100 0 100 200 300 400 500400
300
200
100
0
100
200
300
400m) Layer 13, 570 -650 km <10.15km/s>
-4 -3 -2 -1 0 1 2 3 4
% P velocity Perturbation
Figure 11m
a) R
ays,
A-A
', W
-E C
ross
sec
tion
500
400
300
200
100
010
020
030
040
050
0
600
500
400
300
200
1000
Depth (km)
Figure 12a
500
400
300
200
100
010
020
030
040
050
0
600
500
400
300
200
1000
% P
vel
ocity
Per
turb
atio
n-4
-3-2
-10
12
34
a) A
-A',
W-E
Cro
ss s
ectio
n
W-E
Mod
el C
oord
inat
e (k
m)
W-E
Mod
el C
oord
inat
e (k
m)
Depth (km)A
tlant
icM
edite
rran
ean
Iber
ia
b) R
ays,
B-B
', W
-E C
ross
sec
tion
Figure 12b
0 500
400
300
200
100
010
020
030
040
050
0
600
500
400
300
200
100
W-E
Mod
el C
oord
inat
e (k
m)
Depth (km)
500
400
300
200
100
010
020
030
040
050
0
600
500
400
300
200
1000
b) B
-B',
W-E
Cro
ss s
ectio
n
% P
vel
ocity
Per
turb
atio
n-4
-3-2
-10
12
34
W-E
Mod
el C
oord
inat
e (k
m)
Depth (km)
Atla
ntic
Med
iterr
anea
nG
ibra
ltar
Alb
oran
400
300
200
100
010
020
030
040
0
600
500
400
300
200
1000
c) R
ays,
C-C
', S
-N C
ross
sec
tion
Figure 12c
S-N
Mod
el C
oord
inat
e (k
m)
Depth (km)
400
300
200
100
010
020
030
040
0
600
500
400
300
200
1000
c) C
-C',
S-N
Cro
ss s
ectio
n
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
S-N
Mod
el C
oord
inat
e (k
m)
Depth (km)
Mor
occo
Iber
iaA
lbor
anS
ea
400
300
200
100
010
020
030
040
0
600
500
400
300
200
1000
d) R
ays,
D-D
', S
-N C
ross
sec
tion
S-N
Mod
el C
oord
inat
e (k
m)
Depth (km)
Figure 12d
600
500
400
300
200
1000
Depth (km)
400
300
200
100
010
020
030
040
0
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
S-N
Mod
el C
oord
inat
e (k
m)
Mor
occo
Iber
iaA
lbor
anS
ea
d) D
-D',
S-N
Cro
ss s
ectio
n
d)S
YN
: Lay
er 4
: 45
-60
km
d)R
EC
: Lay
er 4
: 45
-60
km
h)S
YN
: Lay
er 8
: 200
-260
km
h)R
EC
: Lay
er 8
: 200
-260
km
Figure 13a-h
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
a)S
YN
: Lay
er 1
: -5
-15
km
a)R
EC
: Lay
er 1
: -5
-15
km
e)S
YN
: Lay
er 5
: 60
-100
km
e)R
EC
: Lay
er 5
: 60
-100
km
b)S
YN
: Lay
er 2
: 15
-30
km
b)R
EC
: Lay
er 2
: 15
-30
km
f)S
YN
: Lay
er 6
: 100
-150
km
f)R
EC
: Lay
er 6
: 100
-150
km
c)S
YN
: Lay
er 3
: 30
-45
km
c)R
EC
: Lay
er 3
: 30
-45
km
g)S
YN
: Lay
er 7
: 150
-200
km
g)R
EC
: Lay
er 7
: 150
-200
km
l)SY
N: L
ayer
12:
490
-570
km
l)RE
C: L
ayer
12:
490
-570
km
s)S
YN
: Sou
th-N
orth
: CC
’
s)R
EC
: Sou
th-N
orth
: CC
’
-4-3
-2-1
01
23
4
% P
vel
ocity
Per
turb
atio
n
Figure13i-s
i)SY
N: L
ayer
9: 2
60-3
30 k
m
i)RE
C: L
ayer
9: 2
60-3
30 k
m
m)S
YN
: Lay
er 1
3: 5
70-6
50 k
m
m)R
EC
: Lay
er 1
3: 5
70-6
50 k
m
j)SY
N: L
ayer
10:
330
-410
km
j)RE
C: L
ayer
10:
330
-410
km
q)S
YN
: Wes
t-E
ast:
AA
’
q)R
EC
: Wes
t-E
ast:
AA
’
k)S
YN
: Lay
er 1
1: 4
10-4
90 k
m
k)R
EC
: Lay
er 1
1: 4
10-4
90 k
m
r)S
YN
: Wes
t-E
ast:
BB
’
r)R
EC
: Wes
t-E
ast:
BB
’
Cro
ss s
ectio
ns
Rap
id R
oll B
ack
Ear
ly-M
id M
ioce
ne
Sho
rten
ing
& R
otat
ions
Wes
tE
ast
Ext
ensi
on
Lone
rgan
& W
hite
(19
97)
Ret
reat
ing
Subd
ucti
on
?
100
km Olig
ocen
e E
arly
Mio
cene
Pre
sent
Alb
oran
Dom
ain
Pla
tt &
Vis
sers
(19
89)
NN
WS
SE
Con
vect
ive
Rem
oval
Del
amin
atio
n
Seb
er e
t al.
(199
6a)
Afr
ica
Iber
ia
Pal
eoce
neE
ocen
e-O
ligoc
ene
E. M
ioce
ne-P
rese
nt
Zec
k (1
997)
200
km
NW
SE
Ext
ensi
onIb
eria
n Li
thos
pher
eB
etic
Li
thos
.
200
km20
0 km
SinkingSlab
Pre
sent
Bre
akof
f of s
ubdu
cted
sla
bL.
Olig
ocen
e/E
.Mio
cene
Slab
B
reak
off
Pro
gres
sive
Rol
ling
or P
eelin
g B
ack
Det
achm
ent
Crust & mantle recycling Mantle recycling
Figure 14
NS
W
Lith
osph
ere
Ast
heno
sphe
re
Moh
o
colli
sion
al r
idge
exte
ndin
g pl
atea
u
deta
chm
ent f
aults
?
100 km
Paleogene L. Oligocene/E. Miocene
Present
A A'
B B'
G.B.
G.B.
?
??
100 km
Figure 15
A
A'
B
B'
Inflow
?
??
Mantle Lid
Asthenosphere
Crust
collisional ridge
extending plateau
detachment faultsAlboran Domain