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.

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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