The structure and evolution of the lithosphere–asthenosphere boundary beneath the Atlantic–Mediterranean Transition Region

Lithos 120 (2010) 74–95

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Lithos j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / l i t h o s

The structure and evolution of the lithosphere–asthenosphere boundary beneath the Atlantic–Mediterranean Transition Region J. Fullea a,b,⁎, M. Fernàndez b, J.C. Afonso c, J. Vergés b, H. Zeyen d a

Dublin Institute for Advanced Studies, 5 Merrion Square, Dublin 2, Ireland Group Dynamics of the Lithosphere (GDL), Institute of Earth Sciences Jaume Almera, CSIC, 08028 Barcelona, Spain ARC Key Centre for the Geochemical Evolution and Metallogeny of Continents (GEMOC), Department of Earth and Planetary Sciences, Macquarie University, North Ryde, NSW 2109, Australia d UMR8146 IDES, Université de Paris Sud XI -CNRS, Département des Sciences de la Terre, Bât. 504, F-91405 Orsay cedex, France b c

a r t i c l e

i n f o

Article history: Received 18 June 2009 Accepted 1 March 2010 Available online 19 March 2010 Keywords: Iberia North Africa Upper mantle composition Potential fields Thermal modelling Seismic velocities Sub-continental lithospheric mantle

a b s t r a c t The present-day thermal and compositional 3D structure of the lithosphere beneath the Atlantic– Mediterranean Transition Region and the lithosphere–asthenosphere interaction from Jurassic times to present has been studied. The Atlantic–Mediterranean Transition Region comprises the western segment of the Africa–Eurasia plate boundary, encompassing two main large-scale tectonic domains: the Gibraltar Arc System and the Atlas Mountains. An integrated and self-consistent geophysical–petrological methodology (LitMod3D) has been applied that combines elevation, gravity, geoid, surface heat flow, and seismic data and allows modelling of compositional heterogeneities within the lithospheric mantle. Our results reveal large variations in the depth of the Moho and the lithosphere–asthenosphere boundary (LAB) as well as a lack of spatial correlation between the thicknesses of these two boundaries. The Moho essentially mimics the topography with depths ranging from ∼ 10 km beneath the oceanic domains of the Atlantic abyssal plains and the Algerian Basin to N 34 km in the Eastern Betics and Rif, the High Atlas mountains, and the Sahara Platform. In contrast, the LAB is shallower beneath the central and eastern Alboran Basin (∼ 70 km) and all along the High, Middle and Anti Atlas (b 100 km) coinciding with the loci of Cenozoic volcanism. Deeper LAB depths are found along the central and western Betics and the Moroccan Atlantic margin (N 140 km) with values exceeding 230 km beneath the Rif and the Sahara Platform. The average bulk composition of the lithospheric mantle corresponds to that of a typical Tecton (i.e. Phanerozoic) domain, with the exceptions of the Sahara Platform, the Alboran Basin, and Atlas Mountains. Distinct mantle compositions are required in these areas to make model predictions and geophysical observables compatible. It is proposed that the highly irregular LAB topography is the result of the superposition of three different geodynamic processes: i) shortening and thickening related to NW–SE Iberia–Africa convergence lasting from Late Cretaceous to Recent, ii) impingement of a baby-like mantle plume or small-scale convection beneath the High-Middle Atlas and Anti Atlas commencing in the mid Eocene, and iii) slab roll-back or mantle delamination in the Betic–Rif–Alboran realm acting from early to late Miocene. © 2010 Elsevier B.V. All rights reserved.

1. Introduction The Atlantic–Mediterranean Transition Region comprises a narrow region along the transition between the western Mediterranean and the Atlantic Ocean along the westernmost segment of the African–Eurasian plate boundary (Fig. 1). This segment is characterized by a diffuse transpressive contact between the African and the Eurasian plates, including a wide band of active deformation (e.g. Meghraoui et al., 1996; Jiménez-Munt et al., 2001; Negredo et al., 2002). Traditionally, the Atlantic–Mediterranean Transition Region has been considered as a ⁎ Corresponding author. Dublin Institute for Advanced Studies, 5 Merrion Square, Dublin 2, Ireland. E-mail address: [email protected] (J. Fullea). 0024-4937/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2010.03.003

zone comprising the Alboran Basin, the Betic and Rif orogens, and the Gulf of Cadiz. However, the inclusion of the adjacent Atlas Mountains has been claimed by some authors (e.g. Anguita and Hernán, 2000; Duggen et al., 2005; Duggen et al., 2009) mainly on the basis of the similarities among the type and ages of the magmatism that extruded in both south-eastern Iberia and north-western Africa including the Atlas. Volcanic units are found along the Atlas in Morocco, the Alboran Basin and the south-east of Iberia (Fig. 1). Besides old Tertiary volcanism in Africa, most of the volcanic activity took place during middle and late Miocene and Pliocene to Recent. Geochemical studies suggest the presence of two clearly differentiated types of magmatic rocks: late Miocene to early Pliocene Si–K rich calc-alkaline and shoshonitic series rocks, and late Miocene to Pleistocene Si-poor Na-rich basanites and alkali basalts to hawaiites and tephrites (e.g., Wilson and Bianchini,

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Fig. 1. Elevation map of the study area showing the principal tectonic domains of the Betics, Rif and Atlas Mountain belts, together with the Cenozoic volcanic rocks (black areas) and their age. The Internal Betic and Rif zones are constituted by high pressure-low temperature metamorphic rocks whereas the External Betic and Rif zones, and the Atlas are formed by non metamorphic Mesozoic and Cenozoic cover rocks. Principal Cenozoic sedimentary basins are depicted by areas with open dots, whereas the Gibraltar Flysch Units are shown by areas of closely spaced dots. The Guadalquivir (GB) and the Rharb basins (RB) represent the foreland basins of the Betics and Rif fold belts, respectively. The Gulf of Cadiz Imbricated Wedge (GCIW) and the Horseshoe Gravitational Unit (HGU) represent the western tectonic continuation of the Betic-Gibraltar-Rif arcuate orogenic system. In the Atlantic Ocean, the Seine (SP), the Horseshoe (HP) and the Tagus (TP) abyssal plains are limited by structural highs partially formed during the Africa–Eurasia convergence.

1999; Duggen et al., 2005). The first group of magmatic rocks is likely related to subduction scenarios, while the second has a signature similar to that of lavas present in intra-plate volcanic settings such as the Atlas Mountains (i.e. mantle plumes or sub-lithospheric mantle contaminated by plumes). These mantle derived volcanic lavas have been related to regions where the lithosphere is thinner than 75-km along the eastern margin of Spain (e.g., Vergés and Fernàndez, 2006). Several competing geodynamic scenarios have been proposed to explain the tectonic and magmatic evolution of the region including Neogene subduction associated with slab roll back (Frizon de Lamotte et al., 1991; Lonergan and White, 1997), active subduction (Gutscher et al., 2002), delamination of the lithospheric mantle (e.g. Seber et al., 1996; Mezcua and Rueda, 1997; Calvert et al., 2000a; Valera et al., 2008), convective removal (Platt and Vissers, 1989; Platt et al., 2003), slab break-off (Zeck, 1996; Wortel and Spakman, 2000), and slab rollback and lithospheric tearing (Spakman and Wortel, 2004; Booth-Rea et al., 2007; Duggen et al., 2005). Most of these scenarios rely in turn on models of the present day lithospheric structure beneath the study area. However, the crustal geometry is well known only in certain areas of the study region, while mantle images rely solely on either seismic tomography studies with poor spatial resolution or 1D/2D lithospheric numerical modelling (e.g., Torne et al., 2000; Teixell et al., 2005; Zeyen et al., 2005; Missenard et al., 2006; Fullea et al., 2007). Although consistent large-scale features were obtained with 1D and 2D numerical models, significant differences persist in some areas. The main objective of this work is to obtain a regional picture of the 3D structure of the lithosphere–asthenosphere boundary beneath the Atlantic–Mediterranean Transition Region using a number of geophys-

ical and petrological data. Such a model of the present-day thermal and compositional structure represents the basis necessary to understand the evolution of the study region and test possible geodynamic scenarios. The software package LitMod3D (Fullea et al., 2009) has been used to model and interpret geophysical and petrological observables. This software is designed to perform combined geophysical– petrological modelling of the lithosphere and sub-lithospheric upper mantle within an internally consistent thermodynamic–geophysical framework, where all relevant properties are functions of temperature, pressure, and composition. By simultaneously solving the heat transfer, thermodynamic, rheological, geopotential, and isostasy (local and flexural) equations, the program outputs temperature, pressure, surface heat flow, density (bulk and single phase), seismic wave velocities, geoid and gravity anomalies, and elevation for any given lithospheric/ upper mantle domain (Fullea et al., 2009).

2. Geological setting The Atlantic–Mediterranean Transition Region consists of two main large-scale tectonic domains: the Betic–Rif arcuate orogen and the Atlas Mountains (Middle, High and Anti Atlas) (Fig. 1). The Betic– Rif orogen along the southern Iberian margin and the northern Moroccan margin is divided into Internal and External units, the Flysch Units, and the Neogene basins (Alboran back arc basin, Guadalquivir and Rharb foreland basins, and intermountain basins). Both Internal and External units are roughly continuous from the Betics to the Rif belt across the Gibraltar Strait, producing an arcuate

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shape that has been referred to as the Gibraltar Arc System (e.g. Chalouan et al., 2008). The Betic–Rif Internal Zones comprise Late Paleozoic to Triassic metamorphic rocks that were piled up during Tertiary compression. The External Zones correspond to the fold-and-thrust cover sequences of the SW-Iberian and NW-Maghrebian paleomargins. The Flysch Units along the western boundary of the Betic–Rif arcuate thrust system comprise Cretaceous to middle Miocene siliciclastic deposits folded and thrust over the External Units (Chalouan et al., 2008, and references therein). The Guadalquivir and Rharb foreland basins were formed by the flexure of the lithosphere in response to the load of the Betic and Rif thrust belts (Garcia-Castellanos et al., 2002) (Fig. 1). The infill of the distal parts of both the Guadalquivir and the Rharb basins started in middle Miocene. A comprehensive description of these basins, as well as their sediment infill through time, has been recently published by Iribarren et al (2009). Other sedimentary basins that were formed and evolved during the Neogene in the Betic–Rif orogen are represented by the so-called intermountain basins. These basins were under sea level during Tortonian times, and subsequently individualized and uplifted above sea level in different ages from late Tortonian onwards (Sanz de Galdeano and Vera, 1992). The extensional back arc Alboran basin is located along the inner side of the Betic–Rif arcuate system (Fig. 1). This basin of irregularly shape shows its maximum thickness in its western end (N8 km) (Comas et al., 1999). The basin got the most important sedimentary contribution from the basins in the Betic–Rif orogenic system (Iribarren et al., 2009). Initial sedimentary infill is late Oligocene in age overlaying metamorphic basement rocks (Comas et al., 1999). The origin of the basin is related to late Oligocene crustal thinning that ended by late Tortonian times, with sediments of this age concealing the fault tips. Most of the Alboran domain exhibits thinned continental crust, with the exception of a relatively small area of magmatic arc crust in its easternmost limit, in the transition to the Algerian Basin (Booth-Rea et al., 2007). Broadly distributed volcanism has been mapped in the Alboran basin ranging from early Oligocene (back-arc) calc-alkaline series to middle-late Miocene calc-alkaline and alkaline series related to subduction and mantle thinning (e.g., Duggen et al., 2004). The structural units forming the Betics and the Rif chains continue beneath the Gulf of Cadiz (Fig. 1). All these units are overlying a Hercynian basement that gently dips towards the East in the Gulf of Cadiz, and towards the south and north beneath the Betics and the Rif, respectively. A striking feature is the presence of large allochthonous masses with seismically chaotic reflections at the forefront of the Gibraltar Arc, named in many different ways (e.g. the “Giant Chaotic Body”, “Allochthonous Unit” or “Olistostromic Complex”). These chaotic Miocene masses can be divided into two units: an accretionary wedge formed by the imbrications of Triassic to upper Miocene sediments located in the Gulf of Cadiz continental slope (Gulf of Cadiz Imbricate Wedge, GCIW, Fig. 1), and a submarine gravitational unit in the Horseshoe abyssal plain (Horseshoe Gravitational Unit, HGU, Fig. 1) (Zeyen et al., 2005; Iribarren et al., 2007 and references therein). The Atlas orogen is an intracontinental mountain belt uplifted in the foreland of the Rif and Tell ranges, which extends for more than 2000 km through Morocco, Algeria and Tunisia (Fig. 1). This orogen is the result of the tectonic inversion of a Mesozoic extensional basin, genetically related to the opening of the Atlantic and Tethys oceans. It is composed of folded and faulted Paleozoic, Mesozoic and Cenozoic rocks, which in the western flank of the orogen are elevated more than 4 km over the sea level. From Cenozoic times to present, the chain has undergone compression associated with the NW convergence of Africa towards Europe. Shortening was achieved mainly by thick-skinned thrusting and folding, affecting the pre-Mesozoic basement and the Mesozoic–Cenozoic cover. The total shortening in the study region due to Cenozoic compression ranges from 15 to 24% according to different authors (Ayarza et al., 2005; Teixell et al., 2005, and references therein).

Alkaline volcanism is present throughout the chain: basanites, alkali basalts and nephelinites in the Middle Atlas (Harmand and Cantagrel, 1984), phonolites, trachytes, rhyolites and comendites in the Anti Atlas (Berrahma and Delaloye, 1989), and nephelinites plus gabbro to carbonatite complex in the High Atlas (Le Bas et al., 1986). The age of the magmatism ranges from the oldest Eocene–Oligocene (45–35 Ma) nephelinites in the High Atlas to the most recent Pleistocene (1.8– 0.5 Ma) basalts, basanites and nephelinites in the Middle Atlas. It is worth noting that all volcanic rocks present in the Atlas are represented in the Canary Islands (Anguita and Hernán, 2000; Duggen et al., 2009). 3. Geophysical setting Numerous geophysical studies have been carried out over the past 20 years in the Atlantic-Mediterranean Transition Region (see Fullea et al., 2007 for a brief summary). These include wide-angle and vertical seismic surveys, body- and surface-wave tomography, and joint geophysical modelling of elevation, gravity anomalies, surface heat flow, and geoid height. In the following sections the available geophysical datasets and constraining information on the crust and LAB structures are summarized. 3.1. Regional geophysical data Regional datasets (elevation, gravity, surface heat flow, geoid height; Fig. 2) were collected from different sources. Free-air gravity anomalies were obtained from the global satellite altimetry model V16.1 (Sandwell and Smith, 1997; updated 2007). Onshore free-air anomalies in Morocco were obtained from Hildenbrand et al. (1988) and gridded to match the altimetry data prior to the final grid merging. Bouguer anomalies for the entire region were computed from the free-air grid with the software FA2BOUG (Fullea et al., 2008) (Fig. 2 A). The elevation data, i.e. topography and bathymetry, come from the new release of ETOPO2 Global Data Base (V9.1) (Smith and Sandwell, 1994; Smith and Sandwell, 1997) (Fig. 2 B). Geoid height data were taken from the recently released Earth Geopotential Model EGM2008 (Pavlis et al., 2008), which includes spherical harmonic coefficients up to degree and order 2190 (Fig. 2C). Since all masses within the Earth contribute to all harmonic degrees of the observed geoid, a high-pass filter needs to be applied to the complete geoid in order to remove the deep mantle signal, which is out the scope of this work. The wavelengths N4000 km (i.e. degrees 2–9) have therefore been removed from the complete geoid to retain the effects of density anomalies shallower than ∼ 400 km depth (Bowin, 2000). The compilation of surface heat flow data used in this study includes reports from Fernàndez et al. (1998) for the Iberian Peninsula, Polyak et al. (1996) for the Alboran Basin, Verzhbitsky and Zolotarev (1989) for the Atlantic domain, and Rimi et al. (1998) for North Africa. Fig. 2 D shows the location and values of available measurements. The data exhibit a wide scatter around a mean value of ∼ 70 mW m− 2 in the Northern Gulf of Cadiz and Rharb Basin, mainly due to groundwater flow and hydrothermal activity (Rimi et al., 1998). In the central and southern Gulf of Cadiz, the surface heat flow decreases to 40– 50 mW m− 2. The highest values are found in the central and eastern Alboran Basin (100–120 mW m− 2), most likely associated with the Neogene lithospheric extension (Torne et al., 2000). 3.2. Moho depth from previous studies An updated and thorough compilation/revision of the Moho depth in the Iberian Peninsula, Gulf of Cadiz, and Alboran Basin has recently been presented by Diaz and Gallart (2009). Although the data compiled in this work show considerable scatter and poor coverage in many areas, their map of interpolated Moho depth is an excellent starting point as input for our models.

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Fig. 2. Geophysical data. A) Bouguer anomaly map. Onshore, in the Iberian Peninsula and Morocco, the Bouguer anomaly comes from Mezcúa et al. (1996), and Hildenbrand et al. (1988), respectively. Offshore, and in the rest of the African region (shaded area), Bouguer anomaly is calculated from the free air anomaly satellite data compilation by Smith and Sandwell (1997) using FA2BOUG software (Fullea et al., 2008). Contour interval is 40 mGal. B) Elevation map from 1′ × 1′ ETOPO2 Global Data Base (V9.1) (Smith and Sandwell, 1994; Sandwell and Smith, 1997). C) Geoid anomaly map from EGM2008 Global Model (Pavlis et al., 2008). Long wavelengths (N 4000 km) have been removed. Contour interval is 1 m. D) Surface heat flow measurements. Colours of solid circles denote heat flow values (mW/m2). Data from Fernàndez et al. (1998), Polyak et al. (1996), Verzhbitsky and Zolotarev (1989), and Rimi et al. (1998).Geophysical data. A) Bouguer anomaly map. Onshore, in the Iberian Peninsula and Morocco, the Bouguer anomaly comes from Mezcúa et al. (1996), and Hildenbrand et al. (1988), respectively. Offshore, and in the rest of the African region (shaded area), Bouguer anomaly is calculated from the free air anomaly satellite data compilation by Smith and Sandwell (1997) using FA2BOUG software (Fullea et al., 2008). Contour interval is 40 mGal. B) Elevation map from 1′ × 1′ ETOPO2 Global Data Base (V9.1) (Smith and Sandwell, 1994; Sandwell and Smith, 1997). C) Geoid anomaly map from EGM2008 Global Model (Pavlis et al., 2008). Long wavelengths (N 4000 km) have been removed. Contour interval is 1 m. D) Surface heat flow measurements. Colours of solid circles denote heat flow values (mW/m2). Data from Fernàndez et al. (1998), Polyak et al. (1996), Verzhbitsky and Zolotarev (1989), and Rimi et al. (1998).

There are no source-controlled seismic studies in the Rif chain, and crustal models based on other geophysical data are scarce. The Moho beneath this orogen has been estimated to be located between 34 km and 38 km depth according to the recent 1D inversion model of Fullea et al. (2007). Torne et al. (2000) arrived to a similar result modelling gravity anomalies, surface heat flow, and elevation data. Further south, available seismic refraction data revealed a maximum crustal thickness of 38–39 km beneath the most elevated parts of the High Atlas, and a 30km-thick crust in the Anti Atlas and Meseta continental area (Makris

et al., 1985; Wigger et al., 1992). Receiver function studies show that the crustal thickness in the junction between the High and the Middle Atlas ranges between 36 km (Sandvol et al., 1998) and 39 km (van der Meijde et al., 2003). The upper mantle Pn-velocities in the High and Middle Atlas have relatively low values of 7.7–7.9 km/s according to Wigger et al. (1992). A similar average value of 7.8 km/s is given by other authors in the High Atlas, which is increased slightly to 8 km/s beneath the Anti Atlas (Makris et al., 1985). Anomalously low average Pnvelocities of 7.5–7.8 km/s are also reported in the central part of the

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Alboran Basin according to seismic refraction data (Hatzfeld et al., 1978), although other authors increase this average value to 8–8.1 km/s based on anisotropic Pn travel time inversion (Calvert et al., 2000b). Recent wide-angle reflection and refraction seismic data along a NW–SE profile from the oceanic Jurassic crust to the NW Moroccan continental margin show that the Moho deepens from 15 km beneath the oceanic crust to 35 km under the Meseta (Contrucci et al., 2004). Pn velocities found by these authors along the profile are 8–8.2 km/s. 3.3. LAB depth from previous studies The LAB beneath the Atlantic–Mediterranean Transition Region has been the subject of several studies integrating different geophysical data (e.g. Torne et al., 1995; Fernàndez et al., 2004; Teixell et al., 2005; Zeyen et al., 2005; Missenard et al., 2006). Although these studies are based on 2D lithospheric profiles, in which lateral density changes in the mantle are assumed to be consequence of lateral temperature changes only, they provide useful first-order information of the LAB structure. For instance, the works by Teixell et al. (2005), Zeyen et al. (2005), and Missenard et al. (2006) have all identified an anomalously thin lithosphere beneath the Atlas Mountains, although the magnitude, shape, and exact location of this lithospheric thinning vary considerably amongst authors, essentially due to the different crustal structure modelled in each case (i.e. density, radiogenic heat production, etc.). Missenard et al. (2006) ascribe the lithospheric thinning to a narrow band centered in the Anti Atlas with a relatively thick lithosphere in the adjacent High Atlas (∼130 km), whereas Teixell et al. (2005) propose a smoother LAB geometry with the minimum lithospheric thickness beneath the Meseta and central High Atlas. The 2D lithospheric transect of Zeyen et al. (2005) extends further north, crossing the junction of the High and the Middle Atlas, where these authors propose a relatively shallow LAB of about 70 km. Similarly, Zeyen et al. (2005) and Fullea et al. (2007) have both noted a long-wavelength lithospheric thickening (LAB at ∼160–190 km) beneath the Gulf of Cadiz and the Rharb Basin, regardless of using different methods and constraining information. In the case of SW Iberian Margin and Alboran Basin, the LAB structure is still poorly known, and different estimations exist in the literature. Torne et al. (1995) reported lithospheric thicknesses of 110 and 120 km beneath the south Iberian Massif and the Tagus abyssal plain, respectively. However, other authors suggest a lower value of ∼95 km in the SW Iberian Peninsula, invoking a deep mass deficit at the LAB in

order to make the high Bouguer anomaly values compatible with the moderate elevation and geoid anomalies observed in this area (Fernàndez et al., 2004). In the Alboran Basin, LAB depth estimations from different methods and constraining data range from b45 km (e.g. 3D model of Torne et al., 2000) to N100 km (e.g. 1D inversion of geoid anomaly and elevation, Fullea et al., 2007). 4. Method Although the term “lithosphere” originally referred to material strength (cf. Anderson, 1995), it is usually defined in different ways depending on what particular property is under study (e.g. thermal lithosphere, mechanical lithosphere, seismic lithosphere, geochemical lithosphere, electrical lithosphere, etc., Eaton et al., 2008). In this work, the term lithosphere will be used as a synonym of thermal lithosphere, which is defined as the “cold” outermost layer of the Earth in which heat transfer is dominated by conduction (e.g. Schubert et al., 2001). This definition is preferred over others for the following reasons: (a) there is a close correlation between geochemical and thermal definitions (Griffin et al., 1999), (b) there are relatively simple functional relationships between the thermal definition and other definitions based on thermophysical properties (e.g. Eaton et al., 2008), and (c) the thermal definition eliminates any ambiguity between different lithospheric domains, since it must exists everywhere (i.e. there is always an intermediate isotherm between the convective interior and the surface temperatures). The 1330 °C isotherm has been taken as the lower limit of the lithosphere (i.e. LAB) in agreement with thermo-mechanical models that use realistic rheologies (cf. Schubert et al., 2001). In the sublithospheric mantle heat transfer is dominated by convection, and thus the vertical temperature distribution is assumed here to follow an adiabatic gradient (e.g. Afonso et al., 2008; Fullea et al., 2009). The 3D finite-difference package LitMod3D (Fullea et al., 2009) was used to generate the models presented in this work. LitMod3D is a forward code intended to model the lithospheric structure down to 400 km depth by solving the heat transfer, thermodynamical, rheological, geopotential, and isostasy equations (Fig. 3). It requires as input data a lithospheric model given by a set of different crustal and lithospheric mantle layers characterized by its petrophysical and thermal properties., The output of LitMod3D comprises the 3D distributions of temperature, density and mantle seismic velocities (dark yellow boxes in Fig. 3), as well as other geophysical observables:

Fig. 3. Flow chart showing the main input and output parameters in LitMod3D. In the input lithospheric structure box, kc and km0 are the thermal conductivities (STP conditions), and Hc and Hm are the radiogenic heat productions for the crustal and mantle layers, respectively; γ is the thermodynamic Grüneisen parameter, KT, K0' are the isothermal bulk modulus and its pressure derivative, respectively, and β is the compressibility. For a list of the mantle compositions used in this work see Table 2. The dark and pale yellow boxes represent the output 3D fields and geophysical observables, respectively.

J. Fullea et al. / Lithos 120 (2010) 74–95 Table 1 Properties of the different crustal bodies used in the 3D model. Layer

Density (kg/m3)

Heat production (W/m3)

Thermal conductivity (W/m K)

Peridotites Miocene and Quaternary sediments External units Internal units Intermediate crust Upper-Middle Crust Lower Crust

3100 2350 + 30 ⁎ Z(km)

0 1 10− 6

3.2 2

2430 + 15 ⁎ Z(km) 2650 + 10 ⁎ Z(km) 2840 2720 2920

1 10− 6 1.5 10− 6 1.8 10− 6 1.5 10− 6 0.2 10− 6

2.3 2.5 2.3 2.5 2.1

surface heat flow, elevation, and gravity and geoid anomalies (pale yellow boxes in Fig. 3). Comparison of model outputs against observed data is used to obtain a self-consistent lithospheric/sublithospheric model that simultaneously fits all geophysical and petrological observables, and consequently reduces the uncertainties associated with the modelling of these observables alone or in pairs, as commonly done in the literature. In contrast to other available 3D codes, LitMod3D is built within a thermodynamic/geophysical self-consistent framework, where all thermophysical properties in the mantle are functions of the Gibbs free energy of the stable assemblages. The main advantage of this feature is that essential parameters describing the mantle structure (e.g. density and seismic velocities) are obtained consistently as a function of temperature, pressure and composition, rather than being imposed ad hoc. Moreover, since the observables used in the modelling are differently sensitive to shallow/deep and thermal/compositional density anomalies, this approach allows us to have a better control on thermal and compositional variations at different depths than other available methods. Details on the methodology and numerical implementations can be found in Afonso et al. (2008) and Fullea et al. (2009). In contrast to other approaches where a probabilistic inversion of geophysical data is performed to obtain the 1D thermal and compositional variations in the mantle (e.g. Khan et al., 2006;2009), LitMod3D is based on a recursive run of user-guided forward models (Fig. 3). Due to the non-linearity of the physical problems involved (e.g. Fullea et al., 2009), any inversion scheme must be based on a either systematic or stochastic exploration of the parameter space, i.e. running several successive forward models. In the 3D case, the elevated number of degrees of freedom (e.g. the petrophysical and compositional parameters of each body, and the depth of each node defining the limits of the layers), as well as the relatively long time required by each forward model to run (∼20–30 min), currently makes the inversion approach impractical. 5. Crustal model Following the main structural units described in Section 2 (Table 1), the crustal structure of the Atlantic–Mediterranean Transition Region is modelled with seven different bodies (i.e. layers) according to their physical properties. Whenever possible, the spatial distribution of the bodies was validated (within uncertainties) with data from wells and seismic experiments (e.g. Diaz and Gallart, 2009). However, in areas with no seismic coverage, the crustal structure is derived from the joint interpretation and modelling of other geophysical observables. A single layer comprising the Miocene–Quaternary sediments is defined in the Atlantic oceanic domain as well as in the Alboran, Rharb and Guadalquivir basins (Table 1; Fig. 4A). In the Alboran Basin an isochore map of the Neogene and Quaternary sediments has been constructed using a depth-conversion function that combines stack velocities with measured depths in wells (Soto et al., 1996; Torne et al., 2000). The maximum sediment thickness (N8 km) is located in

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the west Alboran Basin, near the Gibraltar Strait. The other two major depocentres (N4 km) are located in the Atlantic oceanic area near the North African continental shelf, the Gulf of Cadiz and the Rharb Basin (González et al., 1998; Litto et al., 2001; Gràcia et al., 2003). For the rest of the model, we started with a 1 km-thick layer of sediments that was modified during the modelling process as required by the fitting of the observables. The so-called External and Internal Unites are included in our model as two separate bodies with specific properties (Table 1). The former is composed of relatively unconsolidated-to-consolidated sediments and metasediments, while the latter comprises high-pressure metamorphic rocks. Fig. 4 B and C shows the isochore maps of these two units. The density of the External Units varies with depth from 2430 kg m− 3 at the surface to about 2550 kg m− 3 at the depocentres, where the units reach N8 km of thickness (e.g. Gulf of Cadiz and in the Gibraltar Strait). In the case of the Internal Units, the density is assumed to vary from a surface value of 2650 kg m− 3 to maximum values of ∼2750 kg m− 3 at the deepest levels (∼10 km depth). The presence of intra-crustal bodies of mantle peridotites has been reported in the Betic–Rif orogen on the basis of gravimetric modelling (Torne and Banda, 1992) and seismic refraction/wide-angle reflection experiments (Barranco et al., 1990; Banda et al., 1993). Similarly, dredging campaigns in the Gorringe Bank revealed that this seamount is mainly composed of serpentinized peridotites crosscut by dikes of gabbro and overlain by extrusive rocks (Cyagor II Group, 1984; GORRINGE cruise 1996, Girardeau et al., 1998). Although the volume of these peridotitic bodies is considerably small in comparison with the surrounding crustal units (Fig 4 C), their high density significantly affects the long-wavelength gravity field and therefore they cannot be neglected when fitting gravity anomalies (Torne and Banda, 1992). These bodies are accounted for by introducing a high-density crustal layer (Peridotites in Table 1), which has a non-zero thickness only in the Betics, the Rif, and the Gorringe Bank. The assumed density for this layer is comparable to that used in previous gravity models (Torne and Banda, 1992) and compatible with laboratory measurements on serpentinized peridotites at crustal P–T conditions (Delescluse and Chamot-Rooke, 2008). One of the main difficulties faced in providing a proper crustal description of such a complex area is the irregular spatial coverage of previous studies (i.e. seismic and other geophysical models) and the differences among them. Nevertheless, some features can be relatively well constrained attending to the common characteristics of previous works: an average crustal density of 2810–2840 kg/m3, and a lower crust with a radiogenic heat production and thermal conductivity of 0.2 mW/m3 and 2.1 W/m K, respectively. In the case of the uppermiddle crust there is less agreement between the different authors, with the reported values for thermal conductivity and radiogenic heat production ranging from 3.5 to 2.4 W/m K, and 2.5–0.6 mW/m3, respectively (e.g. Fernàndez et al., 2004; Teixell et al., 2005; Zeyen et al., 2005, Missenard et al. 2006, Torne et al., 1995). Bearing in mind these restrictions, our parameterisation for the remaining crystalline crustal units is performed to keep the model as general as possible. Three layers with constant density are considered: an upper-middle crust (Fig. 4 E) with a density of 2720 kg/m3 (i.e. granitic/granodioritic), a lower crust (Fig. 4 F) with a density of 2920 kg/m3 (i.e. granulitic), and an intermediate crust (Fig. 4 D) with some of its thermophysical properties computed as an average of those in the upper-middle and lower layers (Table 1). This intermediate layer represents locally the upper-middle crust of the SW Iberian Peninsula (basement of Precambrian to early Paleozoic sediments with abundant intrusions of volcanic and plutonic rocks), and the entire crust in the eastern Alboran Basin. The upper/lower crust boundary is modified through the modelling process to obtain an average density distribution for the crust that is compatible with the different observables. Moho depths, on the other hand, are required to be consistent with seismic experiments,

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J. Fullea et al. / Lithos 120 (2010) 74–95 Table 2 Bulk mantle compositions used in the 3D model. C1 to C4 denote different bulk compositions of Tecton (i.e. Phanerozoic) lithosphere. APM and PUM stands for Average Proterozoic Massif and Primitive Upper Mantle, respectively. M&S95 refers to McDonough and Sun (1995), J79 refers to Jagoutz et al. (1979).

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O NiO Mg# Cr/(Cr + Al)

C1

C2

C3

C4

APM

PUM M&S95

PUM J79

44.5 0.14 3.5 0.4 8.0 0.13 39.8 3.1 0.24 0.26 89.9 0.07

45 0.16 3.9 0.41 8.1 0.07 38.7 3.2 0.28 0.24 89.5 0.07

44 0.09 2.3 0.39 8.4 0.14 41.4 2.2 0.24 0.26 89.8 0.1

44.4 0.09 2.6 0.4 8.2 0.13 41.1 2.5 0.18 0.27 89.9 0.09

45.2 0.09 2 0.38 7.9 0.11 41.6 1.9 0.13 0.28 90.4 0.11

45 0.2 4.5 0.38 8.1 0.14 37.8 3.6 0.36 0.25 89.3 0.05

45.2 0.22 4 0.46 7.8 0.13 38.3 3.5 0.33 0.27 89.7 0.07

whenever available (e.g. González-Fernández et al., 2001; Wigger et al., 1992; Contrucci et al., 2004; Diaz and Gallart, 2009). 6. Lithospheric model 6.1. Analysis of different lithospheric mantle compositions Fullea et al. (2007) recently presented a preliminary lithospheric model of the Atlantic–Mediterranean Transition Region based on the joint inversion of elevation and geoid anomalies. These authors assumed a simple two-layer lithospheric model (i.e. crust and lithospheric mantle) in which crustal density varied linearly with depth whereas lithospheric mantle density was allowed to vary as a function of temperature only (i.e. neither compositional nor pressure effects were included). Moreover, since their inversion scheme is strictly based on a 1D formulation, 3D effects on the gravity and thermal fields caused by the finite extension of the structures was not properly accounted for. However, despite the simplicity of the method, the regional LAB structure obtained by Fullea et al. (2007) is compatible to the first order with most available geophysical information, and represents a valuable basis to generate more comprehensive models (see next section). The depth to the LAB reported in Fullea et al. (2007) has been used as the starting geometry in our calculations and modified it in successive steps to obtain a 3D best fitting model of the long-wavelength part of the observables. To study the mantle contribution to the observables four common Tecton (i.e. Phanerozoic) bulk compositions (C1, C2, C3, and C4, Table 2), have been tested in accordance to the average tectonic age of the crustal units in the Atlantic–Mediterranean Transition Region. All four compositions are only moderately depleted compared with the sublithospheric mantle, which is assumed to have the PUM (Primitive Upper Mantle) composition of McDonough and Sun (1995) throughout the modelling process (Table 2). Compositions C1 and C2 are estimates based on garnet-bearing xenoliths and garnet xenocrysts, while C2 and C3 are derived from spinel-bearing xenolith suites (Griffin et al., 2009 and references therein). From the four compositional models generated (Table 2), only those with compositions C1 and C2 are found to be compatible to the first order with the long-wavelength part of the geophysical observables. In particular, compositions C3 and C4 predict average residual elevations of N450 m, suggesting an unrealistically low density distribution (Table 3). We acknowledge, however, that a pertinent question is whether compositions C3 and C4 could be made compatible with observations by modifying the crust and lithospheric mantle structures within permissible bounds. Parallel experiments have

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Table 3 Statistics of the residuals of the different compositional models in the mantle (see Table 2). First row: average residual elevation. Rows second to fourth: standard deviations of the residual elevation, Bouguer anomaly, and geoid anomaly, respectively.

ΔE average (m) Std. ΔE (m) Std. ΔgB (mGal) Std. ΔN (m)

Preferred model

C1

C2

C3

C4

51 292 12.5 0.97

175 330 14.9 2.72

152 363 18.5 3.85

474 – – –

460 – – –

been run to address this issue and found that it is not possible to obtain a good regional fit of the geophysical observables and simultaneously satisfy seismic information on the crustal structure. It is therefore concluded that compositions C3 and C4 cannot be representative average compositions for the whole lithospheric mantle in the Atlantic– Mediterranean Transition Region. Within this context, it is worth noting that a similar observation was recently pointed out by Griffin et al. (2009) based on Re–Os analyses, which suggest that compositions C3 and C4 may represent (metasomatised?) relicts of Proterozoic lithospheric mantle preserved at shallow depths instead of “true” Tecton compositions. Compositions C1 and C2 result in a reasonable regional fit of the average residual elevation, with C1 producing the best regional fit (Table 3). Yet, high residuals still persist locally in some parts of the study region (Fig. 5). From this exercise it is concluded that C1 is a representative average for the whole mantle in the Atlantic– Mediterranean Transition Region. Our model with composition C1 still displays localized areas of poor fit. Foremost amongst these are the West African Craton/Mobile Zone, the Alboran basin, and the Atlas Mountains (Fig. 5). These three areas are characterized by distinct tectonic ages and histories, and therefore it is likely that they retain different compositional signatures in their lithospheric mantles (Griffin et al., 2009). Consequently, we have attempted to improve the fit by introducing lateral compositional heterogeneities in the lithospheric mantle. The West African Craton/Mobile Belt contains remnants of Proterozoic and Archean crust that have been affected mostly by accretionary tectonics during the Proterozoic and Pan-African orogeny (Begg et al., 2009). In a recent continental-scale analysis of the lithospheric architecture of Africa, Begg et al. (2009) interpreted this zone as a collage of Tecton, Proton, and Archean lithospheric domains, although the smallscale distribution of these domains could not be assessed. In our model, the magnitude and spatial distribution of the misfits suggest a possible compositional contribution to the observables. Both geoid and elevation residuals (positive and negative, respectively), which are sensitive to deep (i.e. mantle) density anomalies, indicate a mass excess in this area (lower right corner in Fig 5 A and D). Trying to fit these two observables by modifying either the crustal structure or the depth to the LAB results in a strong deviation of the long-wavelength signal of the gravity anomalies. Alternatively, a combination of both effects could be proposed to improve the fit. However, complementary tests reveal that the 3D effects of these modifications considerably affect the fit in the surrounding areas, which arguably are better constrained in terms of their crustal structure. Unfortunately, at present a completely combined crustal/LAB cause for these anomalies cannot be ruled out due to lack of constraining data. Nevertheless, given the difference in tectonic ages between this region and the surrounding areas, we favour a deep compositional (i.e. more depleted mantle) origin for the obtained residuals. Therefore, a bulk lithospheric mantle composition corresponding to an Average Proterozoic Massif (APM in Table 2) has been used. This choice not only makes it possible to fit surface observables, but it also brings our synthetic velocities closer to the high seismic velocities imaged in seismic tomography studies (e.g., Priestley et al., 2008). This is due to the

Fig. 4. Thickness maps of the different crustal layers used in the model (see Table 1). A) Miocene and Quaternary sediments. B) External Units. C) Internal Units and peridotitic bodies (in black). D) Intermediate Crust. E) Upper-Middle Crust. F) Lower Crust.

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Fig. 5. Residual values (calculated minus measured) for a model with a homogeneous mantle layer with composition C1 (see Table 2). A) Geoid anomaly. B) Elevation under flexural isostasy (assuming a uniform elastic thickness of 30 km). C) Bouguer anomaly. D) Elevation under local isostasy.

fact that depleted peridotites typically exhibit faster velocities than their less depleted counterparts at equal P–T conditions (e.g. Griffin et al., 2009; Afonso et al., 2010). Mantle xenoliths recovered from alkaline basalts in the Middle Atlas reveal a widespread metasomatism of the mantle in this area since Late Cretaceous-? or Eocene times (Raffone et al., 2009). The Mg# (i.e. MgO/(MgO + FeO) ⁎ 100) of olivine in xenolith suites that have not been subject to metasomatism suggests that the bulk composition of the unmetasomatised part of the mantle is similar to our C1–C2 compositions (see Table 2 in Raffone et al., 2009), supporting our assumption of considering C1 as the representative average composition for the whole mantle in the Atlantic–Mediterranean Transition Region. These “pristine” mantle sectors were re-fertilized later by the passage of migrating fluids and alkaline melts originating in the sublithospheric mantle. It is therefore likely that the lithospheric thinning predicted by our models in this region is associated with, or the reflection of, the thermo-chemical erosion related to the Late Cretaceous-? or Eocene metasomatic event. A more-than-average fertile composition for the lithospheric mantle in this region (e.g. PUM J79, see Table 2) would also make our model more compatible with the low Pn velocities obtained in seismic transects (Makris et al., 1985; Wigger et al., 1992).

There is geophysical and petrological evidence that supports the presence of an “anomalous” mantle beneath the Alboran Basin. From the geophysical point of view, there are the elevated SHF measurements (Fernàndez et al., 1998), the negative velocity anomalies imaged by seismic tomography from 40 km to 100 km depth (Calvert et al., 2000a; Gurria and Mezcua, 2000), and the extremely low Pn velocities reported at the Moho by seismic refraction studies (Hatzfeld et al., 1978). From a petrological point of view, we observe the presence of a Si–K rich group (calc-alkaline and shoshonitic series rocks), and an Upper Miocene to Pleistocene Si-poor Na-rich group (basanites and alkali basalts to hawaiites and tephrites) magmas in the Alboran Basin. The first group of magmas is related to partial melting of a lithospheric mantle enriched by fluids or melts associated with a previous subduction, whereas the second one points to intraplate melting of asthenospheric material enriched by plume material (Duggen et al., 2005). Other authors suggest that the early decompression melting event was accompanied by an increase in the mantle temperatures of 50 °C–100 °C (Soto and Platt, 1999) produced by a complete convective removal of the lithosphere (Platt et al., 1998; Turner et al., 1999). The presence of a re-fertilized mantle layer in the Alboran Basin with a composition closer to that of the PUM sublithospheric upper mantle (PUM M&S95, see Table 2) would have two positive effects in our model. Firstly, it would help us to reduce

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the prominent negative and positive residual values of both Bouguer anomaly and elevation data, respectively, which point to a lack of mass in the lithosphere (Fig. 5), and, secondly it would reduce our predicted seismic velocities allowing a better match of the low Pn velocities (Hatzfeld et al., 1978) and negative velocity anomalies (Calvert et al., 2000a; Gurria and Mezcua, 2000).

6.2. Preferred model for the Atlantic–Mediterranean Transition Region Maps of the LAB and crust-mantle boundary depths of our final best fitting model are shown in Fig. 6. Figs. 7 and 8 show representative cross-sections with the mantle density. Thickness maps of the anomalous lithospheric mantle bodies used in the Alboran Basin, Atlas Mountains, and West African Craton/Mobile Belt are shown in Fig. 9 We

Fig. 6. A) Moho depth map corresponding to our preferred model superimposed to the structural map with the main tectonic units and volcanism (see Fig. 1). Isolines every 2 km. B) Lithosphere–asthenosphere boundary depth map corresponding to our preferred model superimposed to the structural map with the main tectonic units and volcanism (see Fig. 1). Isolines every 20 km.

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define the upper boundary of the anomalous mantle bodies at the Moho, and extend them downwards as suggested by independent geophysical and petrological evidence (see Section 6.1), and as required to fit the geophysical observables. The depth of the crust-mantle boundary (Fig. 6 A) is less than 12 km in the easternmost part of the Alboran Basin, towards the Algerian Basin. Beneath the central Alboran Basin, the Moho is relatively flat (around 16 km), while near the Gibraltar Arc is about 30 km. The Eastern Betics are characterized by a crustal thickness of 34–36 km, shallowing abruptly southwards, towards the Alboran Basin. The crustal thickness in the SW Iberian Peninsula is about 30–32 km. In northern Africa, the Rif and Atlas Mountains are underlain by a thick crust, with Moho depths of N34 km and N36 km, respectively. The crust-mantle boundary shallows towards the Moroccan margin (∼28 km) and steadily deepens beneath the Saharan Craton to N38 km. In the Atlantic domain, the Moho lies at a depth of 12–14 km in the abyssal plains, and 25–30 km in the Gulf of Cadiz region. The resulting LAB topography (Fig. 6 B) is characterized by strong lateral gradients in the northern, southern and eastern limits of the thick lithosphere imaged beneath the Gulf of Cadiz, the Betics and the Rif (170–240 km). These regions coincide with the contact between the Iberian Variscan Massif and the Betic chain in the north, the contact between the Middle Atlas and the external Rif domain to the south, and the contact between the Betic–Rif orogen and the Alboran Basin to the east. This lithospheric thickening, with a reduced magnitude, continues to the southwest, encompassing the NW Moroccan margin (140– 170 km). The thinnest lithosphere obtained in this model is in the eastern Alboran Basin, where the LAB depth is about 70 km. Moderate lithospheric thicknesses are obtained in the SW Iberian Variscan Massif (90 km) and the Atlas Mountains, particularly in the central High Atlas and eastern Anti Atlas (90 km), and the Middle Atlas (b80 km). The eastern branch of the Atlas does not seem to be affected by a lithospheric thinning. The LAB depth increases to N230 km in the SE corner of our model, beneath the eastern ends of the Sahara Platform (West African Craton/Mobile Zone). As put forward by the residuals shown in Fig. 10, this preferred model clearly represents a significant improvement over our initial model with homogeneous mantle composition (Table 3). In addition, predictions of Pn velocities are in better agreement with observed Pn velocities at the Moho (Fig. 11), bearing in mind the uncertainties commonly associated with seismic experiments (±0.1 km/s). In the Iberian Massif and Betics the sub-Moho P-velocities of our model (8– 8.1 km/s) are in agreement with the reported values, which range between 8 and 8.2 km/s (Banda and Ansorge, 1980; Banda et al., 1993; Barranco et al., 1990; González et al., 1998; Palomeras et al., 2009; Simancas et al., 2003). Similar values of 8 km/s and 8–8.2 km/s are observed in the Gulf of Cadiz and NW Moroccan Margin, respectively (González-Fernández et al., 2001; Contrucci et al., 2004) which are consistent with the average value of 8.1 km/s obtained in our model. Anomalous low velocities of 7.8 km/s are found in the High and Middle Atlas (Makris et al., 1985; Wigger et al., 1992). Our results suggest slightly higher values (7.94–7.96 km/s) which, however, are close to the upper bounds in these areas if we assume standard uncertainties (±0.1 km/s). Furthermore, in view of the relatively shallow LAB found in these areas, the presence of small amounts of partial melting and/or hydrous phases in the uppermost mantle, which is not considered in our model, cannot be ruled out. In the central and eastern Alboran Basin the reported Pn velocities exhibit a wide range of variation, i.e. 7.5–7.9 km/s (Hatzfeld et al., 1978) and 8– 8.1 km/s (Calvert et al., 2000b). Two factors complicate the determination of a well constrained seismic velocity structure in this case: i) the lack of stations and significant seismicity (Barranco et al., 1990), and ii) the pronounced anisotropy present in this region (Calvert et al., 2000b; Díaz et al., 1998). Anisotropic effects can account for velocity variations of 0.2–0.3 km/s in the Alboran Basin, with the fast axis showing a predominantly N–S orientation (Calvert et al.,

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Fig. 7. Selected lithospheric cross-sections from our preferred model (thick solid lines) and superimposed crust-mantle boundary and LAB geometries from previous works: stippled line (Fullea et al., 2007), dotted line (Teixell et al., 2005), dashed line (Zeyen et al., 2005), dash-dotted line (Missenard et al., 2006). The vertical arrows show the crossover between the different profiles. Lateral inset shows the location of profiles.

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Fig. 8. Selected lithospheric cross-sections of the mantle density obtained from our preferred model (thick solid lines). Isolines of density each 20 kg/m3. The vertical arrows show the crossover between the different profiles. Lateral inset shows the location of profiles.

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2000b). Keeping in mind these restrictions, our calculated velocities (7.82–7.88 km/s) lay well within the average values of the observed range (7.5–8.1 km/s).

One of the main advantages of using LitMod3D in comparison with other approaches is that the main thermophysical properties of the mantle are derived self-consistently from a free energy minimization formalism (Stixrude and Lithgow-Bertelloni, 2005; Afonso et al., 2008; Fullea et al., 2009). Therefore, seismic velocities derived from our model can be compared with tomography data obtained independently from travel time inversion. Since most of the available tomography studies in the Atlantic–Mediterranean Transition Region are based on body wave inversion using AK135 (Kennett et al., 1995) as a 1D reference model, the AK135 model has been subtracted from our calculated mantle Pwave velocities. Fig. 12 compares the Vp anomalies at 145 km depth obtained from global travel-time tomography with those calculated from our model. In order to make a representative and meaningful comparison, given the inherent vertical smearing of body-wave tomography, we show an average of our model anomalies at depths of 100 and 150 km. Global tomography data have been provided by A. Villaseñor (pers. comm.) using the same methodology than Bijwaard et al. (1998) with additional travel-time data at regional distances (see also Villaseñor et al., 2003 and Valera et al., 2008). To the first order, our model satisfactorily reproduces the main Vp anomalies (both location and magnitude) from 40 km depth down to 200 km: low velocities beneath the Middle Atlas, SW Iberian Peninsula and eastern Alboran Basin, and high velocities in the Betics, Rif, Rharb Basin and NW Moroccan margin (Bijwaard and Spakman, 2000; Calvert et al., 2000a; Gurria and Mezcua, 2000; Piromallo and Morelli, 2003; Spakman and Wortel, 2004; Schmid et al., 2008) (Fig. 12). At depths greater than 200 km most tomography images suggest that the positive Vp anomaly migrates north-eastwards, focused mainly in the Betic domain (Blanco and Spakman, 1993; Bijwaard and Spakman, 2000; Calvert et al., 2000a; Spakman and Wortel, 2004). This NE displacement in depth of the positive Vp is likely related to a deep feature, i.e. a cold drip originated by any of the proposed mechanisms (e.g. subduction, delamination, convective removal, etc., see discussion section). In its present version, LitMod3D does not allow to model dynamic sublithospheric anomalies which are decoupled from the lithospheric structure (see Appendix C in Fullea et al., 2009) for a brief discussion on this restriction). Therefore, the velocity anomalies of our model are only dependent on the temperature distribution and the compositional variations within the lithospheric mantle, and that is the reason why deep sub-lithospheric mantle velocity anomalies are absent in our model. A recent surface wave tomography study of the European upper mantle using fundamental-mode Rayleigh and Love waves (Boschi et al., 2009) shows interesting correlations with the results derived in this work (Fig. 13). More specifically, these authors image a low vertical shear wave velocity beneath the NW Morocco at 100 km depth. This negative anomaly, which is limited to the east by a positive anomaly, persists with decreased amplitude down to 150 km depth (Boschi et al., 2009). However, the positive anomaly below Rif, Rharb Basin and NW Moroccan margin imaged in our model (Fig. 13) and in most body-waves tomography works (e.g. Bijwaard and Spakman, 2000; Calvert et al., 2000a; Piromallo and Morelli, 2003; Spakman and Wortel, 2004; Schmid et al., 2008) is absent from the surface-wave tomography studies, either those using ray-theory (Boschi et al., 2009) or finite frequency theory (Peter et al., 2008). Unlike in bodywave tomography, the short-wavelengths structures tend to be absent in surface-wave tomography due to the low frequency seismic signals used to obtain the anomalies. Differences between our model and surface-wave derived anomalies could be, to some extent, associated with this inherent lack of short wavelength features. In the Sahara

Fig. 9. Thickness map of the mantle layers of our preferred model. All layers are defined from the Moho depth (Fig. 6 A) downwards (see Fig. 7). A) West African Craton/Mobile Belt (composition APM, see Table 2). Isolines every 10 km. B) Atlas Mountains (composition PUM J79, see Table 2). Isolines every 5 km. C) Alboran Basin (composition PUM M&S95, see Table 2). Isolines every 10 km.

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Fig. 10. Upper panel, residual values (calculated minus measured) for our preferred model (Figs. 6, 7, 8 and 9, see text for further details): A) Geoid anomaly. B) Elevation under flexural isostasy (assuming a uniform elastic thickness of 30 km). C) Bouguer anomaly. D) Elevation under local isostasy. In the lower panel: statistics of the residual E) elevation, F) geoid anomaly and G) Bouguer anomaly for our preferred model (see Table 3 for comparison with other compositional models).

Platform and West African Craton positive velocity anomalies have been obtained (Fig. 13) as expected from the deep LAB of our model (Fig. 6 B) and the compositional variation in the mantle (Fig. 9 A). This outcome is consistent with the results obtained in the West African Craton by a recent work of SV-wave tomography at an African scale (Priestley et al., 2008). Similarly to body-wave tomography models, surface wave studies depict positive anomalies at depths greater than 200 km,

extended in this case, due the lateral smearing, to the entire Atlantic and Mediterranean Moroccan margin, and southern Iberian peninsula (Boschi et al., 2009). The temperature distribution within the lithosphere is essentially controlled by two factors: the crustal heat production, and the crustal and lithospheric thickness. The temperature at the Moho tends to be high both when the crust is thick (as the amount of heat producing

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Fig. 11. Calculated Pn mantle velocities of our preferred model. The white velocity values inside the green boxes indicate existing data from seismic refraction studies (see text for references).

elements increases and the temperature augments with depth) and when the LAB depth is shallow. In Fig. 14 A, a map of the calculated temperature at the crust-mantle boundary is shown for our preferred model. The continental areas of Iberia and Africa are characterized by Moho temperatures above 450 °C. The temperatures at the Moho are lower in marine areas: b∼200 °C in the oceanic crust of the Atlantic abyssal plains, and b250 °C in the thinned continental crust of the easternmost Alboran Basin. In Africa, the highest temperatures are located beneath the relatively thick crust of the central High Atlas (N700 °C), the Middle Atlas (550–600 °C), and the Anti Atlas (500– 600 °C). In the Rif, where the crustal thickness is similar to that of the Atlas, the temperatures are lower (300–500 °C) owing to the deep LAB depth in the area (Fig. 6). In Iberia, two maxima for the Moho temperature in the SW Variscan Massif and the Internal Betics (N550 °C) have been found. The first one is related to the lithospheric thinning, while the second one is likely due to the thick crust present beneath the Internal Betics (Figs. 4 and 6 B). In the Alboran Basin and surroundings, the obtained Moho temperature follows a similar pattern to that proposed by Soto et al. (2008), with minimum values in the eastern and western areas of the Alboran Basin and maximum values in the eastern Betics, central Alboran and north Morocco. However, our estimates are about 100– 150 °C lower than those calculated in Soto et al. (2008). These differences are mainly related to the different numerical approaches used in both studies and, to a lesser extent, to the different radiogenic heat production distributions assumed. To calculate the temperature at the Moho, Soto et al. (2008) used a 1D steady-state approach with fixed surface temperature and surface heat flow as boundary conditions, whereas our model is fully 3D and uses fixed temperatures at the surface and LAB as boundary conditions. As a result, Soto et al. (2008) obtained a temperature of 1350 °C at 45 km depth for most of the central and western Alboran Basin resulting in a considerably shallower LAB than that obtained in our study (around 70 km depth).

Fig. 12. A) Synthetic P-wave velocity anomalies predicted from our preferred model with respect to the reference model AK135 (Kennett et al., 1995) averaged at depths of 100 and 150 km. Isolines every 0.4 %. B) Horizontal slice of P-wave velocity anomaly at 145 km depth obtained from a seismic tomography study (Villaseñor et al., 2003, see text for details). C) Hit count with the number of rays crossing the model at 145 km depth (Villaseñor et al., 2003). Anelasticity effects are included in our synthetic anomalies as explained in Afonso et al. (2008), considering the dominant frequencies of the seismic tomography study (∼1–0.5 Hz) and an average grain size of 1 cm.

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Fig. 13. Synthetic S-wave velocity anomalies, with respect to the reference model PREM, predicted from our preferred model at depths of: A) 100 km. B) 150 km. Isolines every 0.5 %.

Although both approaches neglect transient and advective processes in their heat transfer formalisms, in our model we simulate them by locally increasing the thermal conductivity of the lithospheric mantle to an unrealistic value of 4.5 W/m K. This modification results in an increase of heat transfer in localized areas without significantly perturbing the surrounding thermal field, therefore effectively mimicking the thermal effect of considering transient thermal regimes and/or vertical mantle advection. The latter are undoubtedly important in the Alboran Basin, which has been affected by recent tectonism and volcanism. Indeed, the inclusion of these effects is required to make SHF observations compatible with the rest of geophysical observables. In contrast, a lithospheric mantle density configuration like that inferred from Soto et al. (2008) is not able to fit simultaneously the observed elevation and potential field data (Bouguer and geoid anomalies). Likewise, imposing a 45–50 km thick lithosphere to our self-consistent 3D geophysical– petrological approach results in large residuals between measured and calculated observables. The calculated surface heat flow (SHF) in the study area (Fig. 14 B) shows a general trend compatible (bearing in mind the wide scatter and high uncertainties inherent to SHF measurements) with the observed values (Fig. 2 D). The lowest SHF values (b40 mW/m2) are

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Fig. 14. A) Temperatures at the crust-mantle boundary predicted by our preferred model superimposed to the structural map with the main tectonic units and volcanism (see Fig. 1). Isolines every 50 °C. B) Surface heat flow predicted by our preferred model superimposed to the structural map with the main tectonic units and volcanism (see Fig. 1).

located in the abyssal plains of the Moroccan Atlantic margin, where a thin crust coexists with a deep LAB (Fig. 6 B). Low SHF values (b50 mW/m2) extend also to the relatively thick continental crust of the western Betics–Rif chain, the westernmost Alboran Basin and the Gulf of Cadiz accretionary wedge, as well as to the Sahara Platform (Fig. 6 A), where the maximum lithospheric thickness values are found by our 3D model (Fig. 6 B). On the contrary, SHF values are larger in the Atlas (60–65 mW/m2), in line with the moderated lithospheric thinning obtained in this work (Fig. 6 B). The highest SHF values are associated with the thinned continental crust of the central and eastern Alboran Basin (N90 and N105 mW/m2, respectively), where the shallowest LAB is present (Fig. 6 B). The Iberian mainland is characterized by SHF values around 55–60 mW/m2. A remarkable misfit is present in the SW Iberian Massif, where the relative high SHF measurements of 80 mW/m2 are likely related to the Pyrite Belt, which is a local feature out of the scope of the present regional scale study. This narrow structure is characterized by a large thermal conductivity and radiogenic heat production causing an unusually high surface heat flow (Fernàndez et al., 1998). Other authors have

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obtained regional values up to 60–65 mW/m2 in SW Iberia using larger radiogenic heat production values in the upper crust (2–2.5 μWm− 3) in a 2D modelling profile (Fernàndez et al., 2004). 7. Discussion 7.1. Modelling results The primary aim of this study was to image the LAB topography in the Atlantic–Mediterranean Transition Region and to understand how the lithosphere and the underlying asthenosphere interacted in a region dominated by a transpressive and diffuse plate boundary. The methodology used is based on the 3D finite-difference code LitMod3D (Fullea et al., 2009), which solves the lithostatic, thermodynamic, gravity and steady-state thermal equations to calculate the temperature, compositional, density and seismological structure of the mantle down to 400 km depth. This methodology has been successfully applied to study the 2D structure of the oceanic and Archean lithosphere (Afonso et al., 2008), as well as the deep lithospheric structure of the Namibian volcanic margin (Fernàndez et al., 2010). The first caveat regarding the present study is that the steady-state thermal condition is likely not satisfied in several subdomains of the study region, where much of the deformation occurred during the Cenozoic. However, quantifying these transient effects in the temperature distribution requires a much better knowledge of the geodynamic evolution and tectonic mechanisms acting in the region than presently available. Steady-state thermal modelling tends to underestimate/overestimate the actual lithospheric thickness when the modelled thickening/thinning processes are still under transient conditions, which means that the actual LAB topography variations have probably larger amplitudes. Local departures from the actual thermal state would affect the calculated density and seismic velocity distributions, and this is something that needs to be considered in the interpretation of the results. Nevertheless, we stress that although there may be local errors in the temperature structure, the average physical properties derived in this study are compatible to a good approximation with all constraining observables. The second caveat comes from the fact that we neglect mantle flow related to lateral variations of pressure (Navier–Stokes term), which prevents the possibility of modelling sublithospheric dynamic processes (see Fullea et al., 2009 for a discussion on this effect). Despite this limitation, the obtained present-day geometry of the crust and LAB, together with surface geology data, gives relevant constraints to analyze the interaction between the lithospheric mantle and the underlying asthenosphere. Bearing in mind these considerations, the results presented in the previous section provide evidence of large variations in the crustal and lithospheric mantle thickness, which amounts to more than 30 and 150 km, respectively (Fig. 6). Although the general trends of Moho and LAB depth maps are similar to those previously published for the study region (e.g. Fullea et al., 2007), the crustal and lithospheric thickness values differ noticeably. The first improvement with respect to previous works is the incorporation of as many crustal bodies as required by a full 3D geometry. This allows a better quantification of the crustal-derived signals in the geophysical observables and, hence, a better quantification of the LAB geometry. The Moho depth obtained in this work (Fig. 6 A) shows some differences relative to that presented in Fullea et al. (2007), particularly in the Eastern Alboran basin, the Gulf of Cadiz, the Moroccan Meseta and the towards the West African Craton. The second improvement is the integration of lateral compositional heterogeneities in the lithospheric mantle, which allows modelling of both major depletion processes (e.g. Poudjom Djomani et al., 2001; Griffin et al., 2009) in the West African Craton, and recent mantle metasomatism (e.g. Raffone et al., 2009) in the Alboran Basin and Atlas Mountains. These effects together with the above mentioned crustal thickness variations result in a LAB topography (Fig. 6 B) that shows

higher lateral variations than previously proposed (e.g., Torne et al., 2000; Teixell et al., 2005; Zeyen et al., 2005; Missenard et al., 2006; Fullea et al., 2007). The predicted lithospheric thickness in our new model ranges now between 70 km in the Eastern Alboran basin to more than 230 km in the Rif and Rharb Basin, reflecting the complex tectonic evolution of the region. Fig. 7 shows four representative lithospheric cross-sections of the modelled region, where we have superimposed the crust-mantle boundary and the LAB geometries obtained from Fullea et al. (2007), Missenard et al. (2006), Teixell et al. (2005) and Zeyen et al. (2005). The largest differences with respect to the 1D model of Fullea et al. (2007) are found along profiles P1 and P3, which are very sensitive to 3D geometries. Profiles P2 and P4 show also noticeable differences relative to previous works: our 3D model generally yields a deeper LAB. The lack of a spatial correlation between the thicknesses of the crust and lithospheric mantle (Fig. 6) in large parts of the model suggest a decoupling of these two layers. This can be interpreted as a result of vertical strain partitioning that occurred during the Cenozoic evolution of the region (see e.g., Torne et al., 2000; Fullea et al., 2007). Therefore, whereas the Moho depth map essentially mimics topography with its maximum values beneath the Eastern Betics, Rif and Atlas mountains, the LAB map shows a remarkably different pattern characterized by NE–SW thickening along the western Betics, Rif, Rharb Basin, and Moroccan Atlantic margin, and thinning along the Eastern Alboran Basin and Middle, High and Anti Atlas. It seems clear that the lack of thickness correlation between crust and mantle results from the interaction of different tectonic processes, which incorporate a sub-lithospheric component. 7.2. Geodynamic interpretation Briefly, the present day thickened lithosphere beneath the western Betics, Rif, Rharb Basin, and Moroccan Atlantic margin is the result of the action of three main tectonic processes of different origin and duration (e.g., Iribarren et al., 2007; Zitellini et al., 2009). The first process is related to the slow and protracted convergence between Africa and Eurasia acting since Late Cretaceous, with a present velocity of 3–4 mm/yr (Fig. 15). The trend of this long term convergence changes from NW–SE transpression in the Alboran Basin to E–W strike-slip west of the Gorringe Bank. Superimposed to this far-field mechanism there is a second tectonic process that acted since late Oligocene–early Miocene to late Tortonian with a predominant E–W trend, which was responsible for a fast tectonic restructuration of the Gibraltar Arc System (Fig. 15). Extension in an ENE–WSW direction took place in the internal domains of the Betic–Rif System as a consequence of this second process, which is also responsible for the emplacement of west-verging Flysch Units along the Gibraltar Arc. Thrusting in the External domains ceased along the Betics, Rif and Gulf of Cadiz Imbricated Wedge at 7–8 Ma, as evidenced from seismic lines (e.g., Berástegui et al., 1998; Iribarren et al., 2007; Zitellini et al., 2009). The third tectonic process is related to the volcanic activity along the Atlas Mountains (Fig. 15). Within this tectonic framework, it is proposed that the predominant processes shaping the present crustal and lithospheric structure followed a sequence in which the locus of deformation changed laterally, in depth and through time. Since Late Cretaceous to late Oligocene, convergence between Africa and Eurasia was responsible for the closure of the Alpine Tethys and the stacking of the Internal Units of the Betics and Rif producing high-pressure/low-temperature metamorphism. Further south, inversion of the Mesozoic intra-continental basins took place in the Middle and High Atlas (e.g., Laville et al., 2004; Missenard et al., 2006; Frizon de Lamotte et al., 2008). At subcrustal levels, however, most of this shortening would have been accommodated along the Moroccan Atlantic Margin rather than in the intra-continental basins (present Atlas), partly accounting for the inferred lithospheric thickening along the margin in our model

J. Fullea et al. / Lithos 120 (2010) 74–95 Fig. 15. Map and two lithospheric transects to illustrate the three different mechanisms that shaped the present lithospheric structure of the Atlantic Mediterranean transition region. E–W transect along profile P1 is partially based, at crustal level, on Comas et al. (1999) and Iribarren et al. (2007). Back-arc thinning occurred in the Alboran domain whereas thickening beneath the External Betics is the product of concurrent NNW–SSE shortening and mantle delamination (or subduction roll-back). The deep geometry of the lithospheric mantle is inferred from seismic tomographic studies (Blanco and Spakman, 1993; Bijwaard and Spakman, 2000; Calvert et al., 2000a; Spakman and Wortel, 2004). NW–SE transect across the NW Moroccan margin is subparallel to profile P2 in its southern segment. The transect shows mantle thinning beneath the Atlas Mountains produced by either small convection cells or by a baby mantle plume. To the NW mantle thickening is mostly the product of protracted NNW African motion against Eurasia.

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(Fig. 7). The reason is that, although Jurassic rifting affected the two regions, continental break-up took only place along the margin. Therefore, mantle stiffening related to the thermal recovery of the previously rifted lithosphere might have progressed faster in the intra-cratonic basins than in the continental margin, preventing the formation of a significant lithospheric root beneath the Atlas. This hypothesis is also supported by the fact that sublithospheric magmatic activity in the present Atlas did not commence until mid Eocene (Raffone et al., 2009). Inversion along the Atlas and the Alpine Tethyan margins (Betic-Rif) ceased at late Eocene–early Oligocene times, whereas accommodation of shortening continued northwards in the Pyrenees (e.g., Vergés et al., 2002). In the Atlas Mountains, this quiescent period lasted for most of the Oligocene, but a new phase of deformation resumed in Pliocene times (e.g., Gomez et al., 2000; Frizon de Lamotte et al., 2008) or even later during the Quaternary (Missenard et al., 2006 and references therein). Interestingly, the resumption of activity in the Alpine Tethys margins roughly coincides with the start of the second main tectonic process, which was responsible for the formation of the Gibraltar Arc System. Several mechanisms have been proposed to explain the evolution of this segment of the Africa–Eurasia plate boundary where coeval extension and compression took place in a reduced area. These mechanisms include convective removal and orogenic collapse (e.g. Dewey, 1988; Platt and Vissers, 1989; Platt et al., 2003), mantle delamination (e.g. Seber et al., 1996; Mezcua and Rueda, 1997; Calvert et al., 2000a; Valera et al., 2008), slab roll back (Frizon de Lamotte et al., 1991; Lonergan and White, 1997; Faccenna et al., 2004), active subduction (Gutscher et al., 2002), slab break-off (Zeck, 1996; Wortel and Spakman, 2000), and slab rollback and lithospheric tearing (Spakman and Wortel, 2004). Regardless of the specific mechanism, the general consensus is that it was responsible for both the lithospheric thickening beneath the Gibraltar Arc (well imaged by seismic tomography; e.g., Bijwaard and Spakman, 2000; Calvert et al., 2000a) and the adjacent thinning in the Alboran Basin (e.g., Torne et al., 2000). According to Missenard et al. (2006), the Moroccan Atlas underwent uplift and erosion at about 17 Ma coinciding with the onset of alkaline volcanic activity. Both volcanic activity and uplift migrated laterally towards the Anti-Atlas (11.6 Ma) and the Guercif Basin (6.5 Ma). This thermal uplift together with the presence of mafic lavas is related by these authors to the impingement of baby plume-like structures at the base of the lithosphere arising from a common magma reservoir. This view has been proposed by other authors as well (e.g., Hoernle et al., 1995; Zeyen et al., 2005; Teixell et al., 2005). However, in light of our results we cannot rule out thermal erosion by small-scale convection as a plausible mechanism producing lithospheric mantle thinning beneath the Atlas and Anti-Atlas. In this scenario, the Atlas and Anti-Atlas represent a relative thinned lithospheric region in comparison to the surrounding thick lithospheric domains of the West African Craton and Moroccan Atlantic Margin (previously thickened by the AfricanEurasian convergence). This undulation of the LAB topography would have promoted or enhanced small-scale convection of the sublithospheric mantle with a localized upward flow beneath the thin region (Atlas and Anti-Atlas). We deem that partial melting of this sublithospheric material can be, at least partially, the source of the Eocene volcanism (e.g. Harmand and Cantagrel, 1984; Raffone et al., 2009). The imaged lithospheric thinning along the High, Middle and AntiAtlas seems to extend to the NE beneath the Alboran Basin and the Western Mediterranean oceanic domain, and to the West below the Atlantic Ocean just north of the Canary Archipelago. The existence of such a sub-lithospheric corridor, however, is a consequence of independent mechanisms acting respectively in different areas: the westernmost Mediterranean (slab roll-back, delamination, etc.), the Moroccan Atlas System (baby plume or small scale convection) and the Canary Islands (mantle plume). We do not support a mantle delamination mechanism (e.g., Ramdani, 1998; Duggen et al., 2009) as

responsible for partial removal of the lithospheric mantle beneath the Atlas. Estimated shortening across the High and Middle Atlas amounts to less than 20% (Teixell et al., 2005; Frizon de Lamotte et al., 2008), which seems insufficient to generate conditions for mantle delamination or gravitational removal, especially if post-Jurassic thermal recovery of the original lithospheric extension had not been completed (e.g. Valera et al., 2008). Most of the Miocene mafic lavas seem to be spatially related to regions where the lithospheric mantle has been thinned (Fig. 6 A), suggesting a causal relationship between mantle upwelling and alkaline magmatism. These magmas would have a common mantle source, as indicated by geochemical analyses (e.g. Hoernle et al., 1995; Duggen et al., 2009), forming a deep reservoir extending from the Canary Islands to the Western Mediterranean (Goes et al., 1999). Therefore mantle melting and alkaline volcanism would be either related to small-scale convection involving the deep reservoir or to small plumes acting as ‘escape valves’ (Hoernle et al., 1995; Zeyen et al., 2005). This hypothesis does not require a pre-existing lithospheric corridor beneath the Atlas to enable large lateral transport of the Canary mantle plume material as recently proposed by Duggen et al. (2009). Westwards of the Gibraltar Arc System the shortening related to Miocene Africa–Eurasia convergence took place along the NW Moroccan margin and moved progressively to the Gorringe Bank and the Ampere–Coral Patch at a crustal scale. At deeper levels, this shortening continued to be accommodated along the Moroccan Atlantic margin and the western Betics. The apparent lithospheric thinning beneath the Gorringe Bank is related to the flexural support of this narrow structure which presently corresponds to a large anticline of serpentinized mantle rocks above a crustal-scale north-west directed thrust. 8. Conclusions In summary, we have modelled the 3D structure of the lithosphere– asthenosphere boundary beneath the Atlantic–Mediterranean Transition Region, the westernmost segment of the African–Eurasian plate boundary. Our preferred model of the present-day thermal and compositional structure was obtained using the software package LitMod3D, and constrained by different data sets (elevation, potential fields SHF, seismic tomography). The highly irregular topography of the LAB depicted by our results is suggestive of an active interaction between the lithosphere and the underlying sub-lithospheric mantle. The following major features can be distinguished: i. The presence of compositional lithospheric mantle heterogeneities is mandatory to reduce the residuals between measured and calculated regional observables such as gravity, geoid and elevation. In addition, these heterogeneities, related to either secular compositional variations (Sahara Platform) or recent mantle metasomatism (Atlas Mountains and Alboran Basin), allow a good reproduction of the measured Pn-velocities of the uppermost mantle and the regional travel-time tomography models. ii. The NE–SW lithospheric thickening beneath the western Betics, Rif and Atlantic Moroccan Margin, reaching a maximum thickness of 230 km, is related to the superposition of two different mechanisms: the NW–SE Africa–Eurasia convergence acting since Late Cretaceous; and the early Miocene to late Tortonian mechanism responsible for the formation of the Alboran Basin (slab roll-back, mantle delamination, subduction, etc.). iii. The lithospheric mantle thinning along the High-Middle Atlas and the Anti-Atlas can be related either to small-scale convection or to the activation of a small mantle plume. The connection with the Eastern Alboran Basin to the NE, and with the Atlantic Ocean to the W, although visible in our model, is not straightforward and may not share a common origin.

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iv. The pronounced SW–NE lithospheric thinning beneath some parts of the High Atlas, Middle Atlas and Anti-Atlas is probably responsible for, or related to, the observed alkaline volcanism rather than acting as an effective corridor through which magma displaces laterally from a unique reservoir located in the Canary Islands. v. The large lithospheric thickness imaged to the south of the Atlas System delineates the boundary of the Western African Craton and is therefore compatible with the existence of an old and cold lithospheric domain.

Acknowledgments We thank A. Villaseñor for providing us with his Vp tomography model and fruitful discussions. Z. Alasonati Tasárová and L. Boschi for helpful and detailed reviews. This work was supported by the Spanish research projects SAGAS CTM2005-08071-C03-03/MAR, Consolider TopoIberia CSD2006-00041, TopoMed CGL2008-03474-E/BTE, ESFEurocores 07-TOPOEUROPE-FP006, DIAS/IRCSET fellowship (J.F.), and Macquarie University grants (J.C.A.). This is contribution 627 from the Australian Research Council National Key Centre for the Geochemical Evolution and Metallogeny of Continents (http://www.gemoc.mq. edu.au). References Afonso, J.C., Fernàndez, M., Ranalli, G., Griffin, W.L., Connolly, J.A.D., 2008. Integrated geophysical–petrological modeling of the lithosphere and sublithospheric upper mantle: Methodology and applications. Geochemistry Geophysics Geosystems 9 (Q05008). doi:10.1029/2007GC001834. Afonso, J.C., Ranalli, G., Fernàndez, M., Griffin, W.L., O'Reilly, S.Y., Faul, U., 2010. On the Vp/Vs–Mg# correlation in mantle peridotites: Implications for the identification of thermal and compositional anomalies in the upper mantle. Earth and Planetary Science Letters 289, 606–618. Anderson, D.L., 1995. Lithosphere, asthenosphere and perisphere. Reviews of Geophysics 33, 125–149. Anguita, F., Hernán, F., 2000. The Canary Islands origin: a unifying model. Journal of Volcanology and Geothermal Research 103, 1–26. Ayarza, P., Alvarez-Lobato, F., Teixell, A., Arboleya, M.L., Teson, E., Julivert, M., Charroud, M., 2005. Crustal structure under the central High Atlas Mountains (Morocco) from geological and gravity data. Tectonophysics 400 (1–4), 67–84. Banda, E., Ansorge, J., 1980. Crustal structure under the central and eastern part of the Betic Cordillera. Geophysical Journal of the Royal Astronomical Society 63, 515–532. Banda, E., Gallart, J., García-Dueñas, V., Dañobeitia, J., Makris, J., 1993. Lateral variation of the crust in the Iberian Peninsula New evidence from the Betic Cordillera. Tectonophysics 221, 51–66. Barranco, L.M., Ansorge, J., Banda, E., 1990. Seismic refraction constraints on the geometry of the Ronda peridotitic massif (Betic Cordillera, Spain). Tectonophysics 184, 379–392. Begg, G.C., Griffin, W.L., Natapov, L.M., O'Reilly, Suzanne Y., Grand, S.P., O'Neill, C.J., Hronsky, J.M.A., Djomani, Y. Poudjom, Swain, C.J., Deen, T., Bowden, P., 2009. The lithospheric architecture of Africa: seismic tomography, mantle petrology, and tectonic evolution. Geosphere 5, 23–50. Berástegui, X., Banks, C., Puig, C., Taberner, C., Waltham, D., Fernàndez, M., 1998. Lateral diapiric emplacement of Triassic evaporites at the southern margin of the Guadalquivir Basin, Spain. In: Mascle, A., Puigdefàbregas, C., Luterbacher, H.P., Fernàndez, M. (Eds.), Cenozoic Foreland Basins of Western Europe. Geological Society Special Publications, London, pp. 49–68. Berrahma, M., Delaloye, M., 1989. Données géochronologiques nouvelles sur le massif volcanique du Siroua (Haut-Atlas, Maroc). Journal of African Earth Sciences 9, 651–656. Bijwaard, H., Spakman, W., 2000. Non-linear global p-wave tomography by iterated linearized inversion. Geophysical Journal International 141, 71–82. Bijwaard, H., Spakman, W., Engdahl, E.R., 1998. Closing the gap between regional and global travel time tomography. Journal of Geophysical Research 103 (B12), 30055–30078. Blanco, M.J., Spakman, W., 1993. The P-wave velocity structure of the mantle below the Iberian Peninsula: Evidence for a subducted lithosphere beneath southern Spain. Tectonophysics 221, 13–34. Booth-Rea, G., Ranero, C., Martínez-Martínez, J.M., Grevemeyer, I., 2007. Crustal types and Tertiary tectonic evolution of the Alborán Sea, western Mediterranean. Geochemistry, Geophysics and Geosystems 8. doi:10.1029/2007GC001639. Boschi, L., Fry, B., Ekström, G., Giardini, D., 2009. The European upper mantle as seen by surface waves. Surveys in Geophysics 30. doi:10.1007/s10712-009-9066-2. Bowin, C., 2000. Mass anomalies and the structure of the Earth. Physics and Chemistry of the Earth 25 (4), 343–353.

93

Calvert, A., Sandvol, E., Seber, D., Barazangi, M., Roecker, S., Mourabit, T., Vidal, F., Alguacil, G., Jabour, N., 2000a. Geodynamic evolution of the lithosphere and upper mantle beneath the Alboran Region of the Western Mediterranean: Constraints from travel time tomography. Journal of Geophysical Research 105, 10871–10898. Calvert, A., Sandvol, E., Seber, D., Barazangi, M., Roecker, S., Mourabit, T., Vidal, F., Alguacil, G., Jabour, N., 2000b. Propagation of regional seismic phases (Lg and Sn) and Pn velocity structure along the Africa–Iberia plate boundary zone. Geophysical Journal International 142, 384–408. Chalouan, A., Michard, A., El Kadiri, Kh., Negro, F., Frizon de Lamotte, D., Soto, J.I., Saddiqi, O., 2008. The Atlas System. In: Michard, A., Saddiqi, O., Chalouan, A., Frizonde Lamotte, D. (Eds.), Continental Evolution: The Geology of Morocco, pp. 203–302. Comas, M.C., Platt, J.P., Soto, J.I., Watts, A.B., 1999. In: Zahn, R., Comas, M.C., Klaus, A. (Eds.), Proceedings ODP, Scientific Results: Ocean Drilling Program, 161, pp. 555–579. Contrucci, I., Klingelhoefer, F., Perrot, J., Bartolomé, R., Gutscher, M.-A., Sahabi, M., Malod, J., Rehault, J.-P., 2004. The crustal structure of the NW Moroccan continental margin from wide-angle and reflection seismic data. Geophysical Journal International 159, 117–128. Cyagor (II Group), 1984. Intraoceanic tectonism of the Gorringe Bank: observation by submersible. In: Gass, J.G., et al. (Ed.), Ophiolite and Oceanic Lithosphere, pp. 113–120. Delescluse, M., Chamot-Rooke, N., 2008. Serpntinization pulse in the actively deforming Central Indian Basin. Earth and Planetary Science Letters 276, 140–151. Dewey, J.F., 1988. Extensional collapse of orogens. Tectonics 7, 1123–1139. Diaz, J., Gallart, J., 2009. Crustal structure beneath the Iberian Peninsula and surrounding waters: A new compilation of deep seismic. Physics of the Earth and Planetary Interiors 173, 181–190. Díaz, J., Gallart, J., Hirn, A., Paulssen, H., 1998. Anisotropy beneath the Iberian Peninsula: the contribution of the ILIHA- NARS broad-band experiment. Pure and Applied Geophysics 151, 395–405. Duggen, S., Hoernle, K., van den Bogaard, P., Harris, C., 2004. Magmatic evolution of the Alboran region: The role of subduction in forming the western Mediterranean and causing the Messinian Salinity Crisis. Earth and Planetary Science Letters 218, 91–108. Duggen, S., Hoernle, K., van den Bogaard, P., Garbe-Schoenberg, D., 2005. Postcollisional transition from subduction to intraplate-type magmatism in the westernmost Mediterranean; evidence for continental-edge delamination of subcontinental lithosphere. Journal of Petrology 46 (6), 1155–1201. Duggen, S., Hoernle, K.A., Hauff, F., Klügel, A., Bouabdellah, M., Thirlwall, M.F., 2009. Flow of Canary mantle plume material through a subcontinental lithospheric corridor beneath Africa to the Mediterranean. Geology 37 (3), 283–286. Eaton, D.W., Darbyshire, F., Evans, R.L., Gruetter, H., Jones, A.G., Yuan, X., 2008. The elusive lithosphere–asthenosphere boundary (LAB) beneath cratons. Lithos 109, 1–22. Faccenna, C., Piromallo, C., Crespo-Blanc, A., Jolivet, L., Rossetti, F., 2004. Lateral slab deformation and the origin of the western Mediterranean arcs. Tectonics 23, TC1012. doi:10.1029/2002TC001488. Fernàndez, M., Marzán, I., Correia, A., Ramalho, E., 1998. Heat flow, heat production, and lithospheric thermal regime in the Iberian Peninsula. Tectonophysics 291, 29–53. Fernàndez, M., Marzán, I., Torne, M., 2004. Lithospheric transition from the Variscan Iberian Massif to the Jurassic oceanic crust of the Central Atlantic. Tectonophysics 386, 97–115. Fernàndez, M., Afonso, J.C., Ranalli, G., 2010. The deep lithospheric structure of the Namibian volcanic margin. Tectonophysics 481 (1–4). doi:10.1016/j.tecto.2009.02.036. Frizon de Lamotte, D., Andrieux, J., Guézou, J.C., 1991. Cinématique des chevauchements néogènes dans l'Arc bético-rifain: discussion sur les modèles géodynamiques. Bulletin de la Société Géologique de France 162 (4), 611–626. Frizon de Lamotte, D., Zizi, M., Missenard, Y., Hafid, M., Azzouzi, M., Maury, R.C., Charrière, A., Taki, Z., Benammi, M., Michard, A., 2008. The Atlas System. In: Michard, A., Saddiqi, O., Chalouan, A., Frizon de Lamotte, D. (Eds.), Continental Evolution: The Geology of Morocco, pp. 133–202. Fullea, J., Fernàndez, M., Zeyen, H., Vergés, J., 2007. A rapid method to map the crustal and lithospheric thickness using elevation, geoid anomaly and thermal analysis. Application to the Gibraltar Arc System and adjacent zones. Tectonophysics 430, 97–117. Fullea, J., Fernàndez, M., Zeyen, H., 2008. FA2BOUG—a FORTRAN 90 code to compute Bouguer gravity anomalies from gridded free air anomalies: application to the Atlantic-Mediterranean transition zone. Computers & Geosciences 34, 1665–1681. Fullea, J., Afonso, J.C., Connolly, J.A.D., Fernàndez, M., García-Castellanos, D., Zeyen, H., 2009. LitMod3D: an interactive 3D software to model the thermal, compositional, density, seismological and rheological structure of the lithosphere and sublithospheric upper mantle. Geochemistry Geophysics Geosystems 10 (Q08019). doi:10.1029/2009GC002391. Garcia-Castellanos, D., Fernàndez, M., Torne, M., 2002. Modeling the evolution of the Guadalquivir foreland basin (southern Spain). Tectonics 21 (3). doi:10.1029/ 2001TC001339. Girardeau, J., Cornen, G., Beslier, M.O., Le Gall, B., Monnier, C., Agrinier, P., Dubuisson, G., Pinheiro, L., Ribeiro, A., Whitechurch, H., 1998. Extensional tectonics in the Gorringe Bank rocks, Eastern Atlantic ocean: evidence of an oceanic ultra-slow mantellic accreting centre. Terra Nova 10, 330–336. Goes, S., Spakman, W., Bijwaard, H., 1999. A lower mantle source for central European volcanism. Science 286, 1928–1931. Gomez, F., Beauchamp, W., Barazangi, M., 2000. Role of Atlas Mountains (northwest Africa) within the African–Eurasian plate-boundary zone. Geology 28, 775–778. González, A., Córdoba, D., Vegas, R., Matias, L.M., 1998. Seismic crustal structure in the southwest of the Iberian Peninsula and the Gulf of Cadiz. Tectonophysics 296 (3–4), 317–331.

94

J. Fullea et al. / Lithos 120 (2010) 74–95

González-Fernández, A., Córdoba, D., Matias, L.M., Torne, M., 2001. Seismic crustal structure in the Gulf of Cadiz (SW Iberian Peninsula). Marine Geophysical Researches 22, 207–223. Gràcia, E., Dañobeitia, J., Vergés, J., Bartolomé, R., Córdoba, D., 2003. Crustal architecture and tectonic evolution of the Gulf of Cadiz (SW Iberian margin) at the convergence of the Eurasian and African plates. Tectonics 22 (4), 1033. doi:10.1029/ 2001TC901045. Griffin, W.L., O'Reilly, S.Y., Ryan, C.G., 1999. The composition and origin of subcontinental lithospheric mantle. In: Fei, Y., Berkta, C.M., Mysen, B.O. (Eds.), Mantle Petrology: Field Observations and High-Pressure Experimentation: A Tribute to Francis R. (Joe) Boyd: Special Publication Geochemical Society, 6, pp. 13–45. Griffin, W.L., O'Reilly, S., Afonso, J.C., Begg, G.C., 2009. The composition and evolution of lithospheric mantle: a re-evaluation and its tectonic implications. Journal of Petrology 50, 1185–1204. doi:10.1093/petrology/egn033. Gurria, E., Mezcua, J., 2000. Seismic tomography of the crust and lithospheric mantle in the Betic Cordillera and Alboran Sea. Tectonophysics 329 (1–4), 99–119. Gutscher, M.-A., Malod, J., Rehault, J.-P., Contrucci, I., Klingelhoefer, F., Mendes-Victor, L., Spakman, W., 2002. Evidence for active subduction beneath Gibraltar. Geology 30, 1071–1074. Harmand, C., Cantagrel, J.M., 1984. Le volcanisme alcalin tertiaire et quarternaire du Moyen Atlas (Maroc): Chronologie K/Ar et cadre géodynamique. Journal of African Earth Sciences 2, 51–55. Hatzfeld, D., and the Working Group for Deep Seismic Sounding, 1978. Crustal seismic profiles in the Alboran Sea — preliminary results. Pure and Applied Geophysics 116, 167–180. Hildenbrand, T., Kucks, R., Hamouda, M., Bellot, A., 1988. Bouguer gravity map and related filtered anomaly maps of Morocco. U. S. Geological Survey Open-File Report 88–517, 15. Hoernle, K., Zhang, Y.-S., Graham, D., 1995. Seismic and geochemical evidence for largescale mantle upwelling beneath the eastern Atlantic and western and central Europe. Nature 374 (2), 34–39. Iribarren, L., Vergés, J., Camurri, F., Fullea, J., Fernàndez, M., 2007. The structure of the Atlantic–Mediterranean transition zone from the Alboran Sea to the Horseshoe Abyssal Plain (Iberia–Africa plate boundary). Marine Geology 243 (1–4), 97–119. Iribarren, L., Vergés, J., Fernàndez, M., 2009. Sediment supply from the Betic–Rif orogen to basins through Neogene. Tectonophysics 475 (1), 68–84. doi:10.1016/j. tecto.2008.11.029. Jagoutz, E., Palme, H., Baddenhausen, H., Blum, K., Cendales, M., Dreibus, G., Spettel, B., Lorenz, V., Wanke, H., 1979. The abundance of major, minor and trace elements in the Earth's mantle as derived from primitive ultramafic nodules. Proceedings of the 10th Lunar and Planetary Science Conference: Geochimica et Cosmochimica Acta Supplement, pp. 2031–2050. Jiménez-Munt, I., Fernàndez, M., Torne, M., Bird, P., 2001. The transition from linear to diffuse plate boundary in the Azores–Gibraltar region: results from a thin sheet model. Earth and Planetary Science Letters 192, 175–189. Kennett, B.L.N., Engdahl, E.R., Buland, R., 1995. Constraint on seismic velocities in the earth from travel-times. Geophysical Journal International 122, 108–124. Khan, A., Connolly, J.A.D., Olsen, N., 2006. Constraining the composition and thermal state of the mantle beneath Europe from inversion of long-period electromagnetic sounding data. Journal of Geophysical Research 111 (B10102). doi:10.1029/ 2006JB004270. Khan, A., Boschi, L., Connolly, J.A.D., 2009. On mantle chemical and thermal heterogeneities and anisotropy as mapped by inversion of global surface wave data. Journal of Geophysical Research 114 (B09305). doi:10.1029/2009JB006399. Laville, E., Piqué, A., Amrhar, M., Charroud, M., 2004. A restatement of the Mesozoic Atlasic rifting (Morocco). Journal of African Earth Sciences 38, 145–153. Le Bas, M.J., Le Maitre, R.W., Streckeisen, A., Zanettin, B., 1986. A chemical classification of volcanic rocks based on the total alkali–silica diagram. Journal of Petrology 27, 745–750. Litto, W., Jaaidi, E.B., Medina, F., Dakki, M., 2001. Etude sismo-estructurale de la marge nord du bassin du Gharb (avant-pays rifain, Maroc): Mise en évidence d'une distension d'âge miocène tardif. Eclogea Geologicae Helvetiae 94, 63–73. Lonergan, L., White, N., 1997. Origin of the Betic–Rif mountain belt. Tectonics 16 (3), 504–522. Makris, J., Demnati, A., Klussmann, J., 1985. Deep seismic soundings in Morocco and a crust and upper mantle model deduced from seismic and gravity data. Annales Geophysicae 3, 369–380. McDonough, W.F., Sun, S.-S., 1995. The composition of the Earth. Chemical Geology 120, 223–253. Meghraoui, M., Morel, J.L., Andrieux, J., Dahmani, M., 1996. Tectonique plio-quaternaire de la chaîne tello-rifaine et de la mer d'Alboran. Bulletin de la Société Géologique de France 167, 143–159. Mezcúa, J., Gil, A., Benarroch, R., 1996. Estudio Gravimétrico de la Península Ibérica y Baleares. Instituto Geográfico Nacional, Publicaciones, 22 pp. and 2 maps. Mezcua, J., Rueda, J., 1997. Seismological evidence for a delamination process in the lithosphere under the Alboran Sea. Geophysical Journal International 129, F1–F8. Missenard, Y., Zeyen, H., Frizon de Lamotte, D., Leturmy, P., Petit, C., Sébrier, M., Saddiqi, O., 2006. Crustal versus asthenospheric origin of the relief of the Atlas Mountains of Morocco. Journal of Geophysical Research 111 (B03401). doi:10.1029/ 2005JB003708. Negredo, A.M., Bird, P., Sanz de Galdeano, C., Buforn, E., 2002. Neotectonic modeling of the Ibero-Maghrebian region. Journal of Geophysical Research 107 (B11), 2292. doi:10.1029/2001JB000743. Palomeras, I., Carbonell, R., Flecha, I., Simancas, F., Ayarza, P., Matas, J., Martínez-Poyatos, D., Azor, A., González Lodeiro, F., Pérez-Estaún, A., 2009. Nature of the lithosphere across the Variscan orogen of SW Iberia: dense wide-angle seismic reflection data. Journal of Geophysical Research 114 (B02302). doi:10.1029/2007JB005050.

Pavlis N.K., Holmes S.A., Kenyon S.C., Factor J.K, 2008. An Earth Gravitational Model to Degree 2160: EGM 2008, GRACE Science Applications, session G3, EGU 2008, Vienna. Peter, D., Boschi, L., Deschamps, F., Fry, B., Ekström, G., Giardini, D., 2008. Surface wave tomography: finite frequency shear velocity inversions for the EuropeanMediterranean region. Geophysical Research Letters 35 (L16315). doi:10.1029/ 2008GL034769. Piromallo, C., Morelli, A., 2003. P wave tomography of the mantle under the AlpineMediterranean area. Journal of Geophysical Research 108. doi:10.1029/ 2002JB001757. Platt, J.P., Vissers, R.L.M., 1989. Extensional collapse of thickened continental lithosphere: a working hypothesis for the Alboran Sea and Gibraltar arc. Geology 29, 299–302. Platt, J.P., Soto, J.-I., Whitehouse, M.J., Hurford, A.J., Kelley, S.P., 1998. Thermal evolution, rate of exhumation, and tectonic significance of metamorphic rocks from the floor of the Alboran extensional basin, western Mediterranean. Tectonics 17, 671–689. Platt, J.P., Whitehouse, M.J., Kelley, S.P., Carter, A., Hollick, L., 2003. Simultaneous extensional exhumation across the Alboran Basin: implications for the causes of late orogenic extension. Geology 31, 251–254. Polyak, B.G., Fernàndez, M., Khutorskoy, M.D., Soto, J.I., Basov, I.A., Comas, M.C., Khain, V.Y., Alonso, B., Agapova, G.V., Mazurova, I.S., Negredo, A., Tochitsky, V.O., Linde, J.d.l., Bogdanov, N.A., Banda, E., 1996. Heat flow in the Alboran Sea, western Mediterranean. Tectonophysics 263, 191–218. Poudjom Djomani, Y.H., O'Reilly, S.Y., Griffin, W.L., Morgan, P., 2001. The density structure of subcontinental lithosphere through time. Earth and Planetary Science Letters 184 (3–4), 605–621. Priestley, K., McKenzie, D., Debayle, E., Pilidou, S., 2008. The African upper mantle and its relationship to tectonics and surface geology. Geophysical Journal International 175, 1108–1126. Raffone, N., Chazot, G., Pin, C., Vannucci1, R., Zanetti, A., 2009. Metasomatism in the Lithospheric Mantle beneath Middle Atlas (Morocco) and the Origin of Fe- and Mgrich Wehrlites. Journal of Petrology 50 (2), 197–249. Ramdani, F., 1998. Geodynamic implications of the intermediate-depth earthquakes and volcanism in the intraplate Atlas Mountains. Physics of Earth and Planetary Interiors 108, 245–260. Rimi, A., Chalouan, A., Bahi, L., 1998. Heat flow in the westernmost part of the Alpine Mediterranean system (the Rif, Morocco). Tectonophysics 285, 135–146. Sandvol, E., Seber, E., Calvert, A., Barazangi, M., 1998. Grid search modeling of receiver functions: Implications for crustal structure in the Middle East and North Africa. Journal of Geophysical Research 103 (B11), 26899–26917. Sandwell, D.T., Smith, W.H.F., 1997. Marine gravity anomalies from GEOSAT and ERS-1 satellite altimetry. Journal of Geophysical Research 102, 10039–10054. Sanz de Galdeano, C., Vera, J.A., 1992. Stratigraphic record and palaeogeographical context of the Neogene basins in the Betic Cordillera, Spain. Basin Research 4, 21–36. Schmid, C., van der Lee, S., VanDecar, J.C., Engdahl, E.R., Giardini, D., 2008. Threedimensional S velocity of the mantle in the Africa–Eurasia plate Boundary region from phase arrival times and regional waveforms. Journal of Geophysical Research 113 (B03306). doi:10.1029/2005JB004193. Schubert, G., Turcotte, D.L., Olson, P., 2001. Mantle convection in the Earth and planets, 940 pp. Cambridge University Press, U.K. Seber, D., Barazangi, M., Ibenbrahim, A., Demnati, A., 1996. Geophysical evidence for lithospheric delamination beneath the Alboran Sea and Rif–Betic Mountains. Nature 379, 785–790. Simancas, J.F., Carbonell, R., González Lodeiro, F., Pérez Estaún, A., Juhlin, C., Ayarza, P., Kashubin, A., Azor, A., Martínez Poyatos, D., Almodóvar, G.R., Pascual, E., Sáez, R., Expósito, I., 2003. Crustal structure of the transpressional Variscan Orogen of SW Iberia: SW Iberia deep seismic reflection profile (IBERSEIS). Tectonics 22 (6), 1062. Smith, W.H.F., Sandwell, D.T., 1994. Bathymetry prediction from dense satellite altimetry and sparse shipborne bathymetry. Journal of Geophysical Research 99, 21803–21824. Smith, W.H.F., Sandwell, D.T., 1997. Global sea floor topography from satellite altimetry and ship depth soundings. Science 227, 1957–1962. Soto, J.I., Platt, J.P., 1999. Petrological and structural evolution of high-grade metamorphic rocks from the floor of the Alboran sea Basin, western Mediterranean. Journal of Petrology 40 (1), 21–60. Soto, J.I., Comas, M.C., de la Linde, J., 1996. Espesor de sedimentos en la cuenca de Alborán mediante una conversión sísmica corregida. Geogaceta 20 (2), 382–385. Soto, J.I., Fernandez-Ibañez, F., Fernandez, M., Casco, A.G., 2008. Thermal structure of the crust in the Gibraltar Arc: Influence on active tectonics in the western Mediterranean. Geochemistry Geophysics Geosystems 9 (Q10011). doi:10.1029/2008GC002061. Spakman, W., Wortel, M.J.R., 2004. A tomographic view on western Mediterranean geodynamics. In: Cavazza, W., Roure, F., Spakman, W., Stampfli, G.M., Ziegler, P. (Eds.), The TRANSMED Atlas—The Mediterranean region from crust to mantle. Springer, Berlin Heidelberg, pp. 31–52. Stixrude, L., Lithgow-Bertelloni, C., 2005. Mineralogy and elasticity of oceanic upper mantle: origin of the low velocity zone. Journal of Geophysical Research 110. doi:10.1029/2004JB002965. Teixell, A., Ayarza, P., Zeyen, H., Fernàndez, M., Arboleya, M.-L., 2005. Effects of mantle upwelling in a compressional setting: the Atlas Mountains of Morocco. Terra Nova 17, 456–461. Torne, M., Banda, E., 1992. Crustal thinning from the Betic Cordillera to the Alboran Sea. Geo-Marine Letters 12, 76–81. Torne, M., Fernàndez, M., Carbonell, J., Banda, E., 1995. Lithospheric transition from continental to oceanic in the West Iberia Atlantic Margin. In: Banda, E., Torne, M., Talwani, M. (Eds.), Rifted Ocean-Continent Boundaries. Kluwer Academic Publications, Dordrecht, pp. 247–263.

J. Fullea et al. / Lithos 120 (2010) 74–95 Torne, M., Fernàndez, M., Comas, M.C., Soto, J.I., 2000. Lithospheric structure beneath the Alboran Basin: Results from 3D gravity modeling and tectonic relevance. Journal of Geophysical Research 105 (B2), 3209–3228. Turner, S.P., Platt, J.P., George, R.M.M., Kelley, S.P., Pearson, D.G., Nowell, G.M., 1999. Magmatism associated with orogenic collapse of the Betic–Alboran domain, SE Spain. Journal of Petrology 40 (6), 1011–1036. Valera, J.-L., Negredo, A., Villaseñor, A., 2008. Asymmetric delamination and convective removal numerical modeling: comparison with evolutionary models for the Alboran Sea region. Pure and Applied Geophysics 165, 1683–1706. van der Meijde, M., van der Lee, S., Giardini, D., 2003. Crustal structure beneath broadband seismic stations in the Mediterranean region. Geophysical Journal International 152 (3), 729–739. Vergés, J., Fernàndez, M., 2006. Mountain Ranges and Basins in the Iberian Peninsula. In: Gee, D., Stephenson, R. (Eds.), European Lithosphere Dynamics: Geological Society, London, Memoirs, 32, pp. 223–234. Vergés, J., Fernàndez, M., 2002. The Pyrenean orogen: pre-, syn-, and post-collisional evolution. In: Rosenbaum, J.G., Lister, G.S. (Eds.), Reconstruction of the evolution of the Alpine–Himalayan Orogen: Journal of Virtual Explorer, 8, pp. 55–84. Verzhbitsky, E.V., Zolotarev, V.G., 1989. Heat flow and the Eurasian–African plate boundary in the eastern part of the Azores–Gibraltar fracture zone. Journal of Geodynamics 11 (3), 267–273. Villaseñor, A., Spakman, W., Engdahl, E.R., 2003. Influence of regional travel times in global tomographic models. Geophysical Research Abstracts 5, EAE03-A-08614, EGS-AGU-EUG Joint Assembly, April 6-11, 2003, Nice, France.

95

Wigger, P., Asch, G., Giese, P., Heinsohn, W.-D., El Alami, S.O., Ramdani, F., 1992. Crustal structure along a transverse across the Middle and High Atlas Mountains derived from seismic refraction studies. Geologische Rundschau 81, 237–248. Wilson, M., Bianchini, G., 1999. Tertiary–Ouaternary magmatism within the Mediterranean and surrounding regions. In: Durand, B., Jolivet, L., Horvath, F., Séranne, M. (Eds.), The Mediterranean Basins: Tertiary Extension within the Alpine Orogen: Geological Society, London, Special Publications, 156, pp. 141–168. Wortel, M.J.R., Spakman, W., 2000. Subduction and slab detachment in the Mediterranean–Carpathian Region. Science 290, 1910–1917. Zeck, H.P., 1996. Betic–Rif orogeny: subduction of Mesozoic Tethys lithosphere under eastward drifting Iberia, slab detachment shortly before 22 Ma, and subsequent uplift and extension tectonics. Tectonophysics 254 (1–2), 1–16. Zeyen, H., Ayarza, P., Fernàndez, M., Rimi, A., 2005. Lithospheric structure under the western African–European plate boundary: a transect across the Atlas Mountains and the Gulf of Cadiz. Tectonics 24 (TC2001). doi:10.1029/2004TC001639. Zitellini, N., Gràcia, E., Matias, L., Terrinha, P., Abreu, M.A., DeAlteriis, G., Henriet, J.P., Dañobeitia, J.J., Masson, D.G., Mulder, T., Ramella, R., Somoza, L., Diez, S., 2009. The quest for the Africa–Eurasia plate boundary west of the Strait of Gibraltar. Earth and Planetary Science Letters 280 (1–4), 13–50.