Lithospheric transition from the Variscan Iberian Massif to the Jurassic oceanic crust of the Central Atlantic

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Tectonophysics 386

Lithospheric transition from the Variscan Iberian Massif to the

Jurassic oceanic crust of the Central Atlantic

M. Fernandez*, I. Marzan, M. Torne

Department of Geophysics and Tectonics, Institute of Earth Sciences (J. Almera), CSIC, Sole i Sabaris s/n, E-08028 Barcelona, Spain

Received 21 October 2003; accepted 25 May 2004

Available online 20 July 2004

Abstract

A 1000-km-long lithospheric transect running from the Variscan Iberian Massif (VIM) to the oceanic domain of the

Northwest African margin is investigated. The main goal of the study is to image the lateral changes in crustal and lithospheric

structure from a complete section of an old and stable orogenic belt—the Variscan Iberian Massif—to the adjacent Jurassic

passive margin of SW Iberia, and across the transpressive and seismically active Africa–Eurasia plate boundary. The modelling

approach incorporates available seismic data and integrates elevation, gravity, geoid and heat flow data under the assumptions

of thermal steady state and local isostasy. The results show that the Variscan Iberian crust has a roughly constant thickness of

~30 km, in opposition to previous works that propose a prominent thickening beneath the South Portuguese Zone (SPZ). The

three layers forming the Variscan crust show noticeable thickness variations along the profile. The upper crust thins from central

Iberia (about 20 km thick) to the Ossa Morena Zone (OMZ) and the NE region of the South Portuguese Zone where locally the

thickness of the upper crust is b8 km. Conversely, there is a clear thickening of the middle crust (up to 17 km thick) under the

Ossa Morena Zone, whereas the thickness of the lower crust remains quite constant (~6 km). Under the margin, the thinning of

the continental crust is quite gentle and occurs over distances of ~200 km, resembling the crustal attitude observed further north

along the West Iberian margins. In the oceanic domain, there is a 160-km-wide Ocean Transition Zone located between the

thinned continental crust of the continental shelf and slope and the true oceanic crust of the Seine Abyssal Plain. The total

lithospheric thickness varies from about 120 km at the ends of the model profile to less than 100 km below the Ossa Morena and

the South Portuguese zones. An outstanding result is the mass deficit at deep lithospheric mantle levels required to fit the

observed geoid, gravity and elevation over the Ossa Morena and South Portuguese zones. Such mass deficit can be interpreted

either as a lithospheric thinning of 20–25 km or as an anomalous density reduction of ~25 kg m�3 affecting the lower

lithospheric levels. Whereas the first hypothesis is consistent with a possible thermal anomaly related to recent geodynamics

affecting the nearby Betic–Rif arc, the second is consistent with mantle depletion related to ancient magmatic episodes that

occurred during the Hercynian orogeny.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Geoid; Gravity; Heat flow; Depleted mantle; SW Iberia; Plate boundary

* Corresponding author. Tel.: +34 934095410; fax: +34 934110012.

0040-1951/$ - s

doi:10.1016/j.tec

E-mail addr

(2004) 97–115

ee front matter D 2004 Elsevier B.V. All rights reserved.

to.2004.05.005

ess: [email protected] (M. Fernandez).

M. Fernandez et al. / Tectonophysics 386 (2004) 97–11598

1. Introduction

The southwestern region of the Iberian Peninsula,

which has suffered a long and complex geodynamic

evolution, is characterized by the transition from the

stable Variscan Iberian Massif to the Central Atlantic

oceanic domain. The region encompasses two late

Precambrian–early Paleozoic shear suture zones cor-

responding to continental collision and oceanic sub-

Fig. 1. Map of the study area showing the location of the modelled regiona

from North to South, CZ: Cantabrian Zone; ALZ: Astur-Leonese Zone;

Portuguese Zone. At sea from North to South, TAP: Tagus Abyssal Plain

Patch Seamount; SAP: Seine Abyssal Plain. Inset shows a global view o

(MN3) and magnetic lineations. Dashed region denotes the Africa–Eurasi

duction, respectively; a late Jurassic–early Cretaceous

passive margin; and the east end of the Azores–

Gibraltar Fracture Zone submitted to right-lateral

transpression since late Cretaceous, which limits the

present-day west termination of the Eurasian and

African plates.

The superposition of all these tectonothermal

episodes resulted in a very puzzling present-day

crustal and lithospheric structure. Up to now there

l transect (thick line). Contour bathymetry interval at 400 m. On land

CIZ: Central Iberia Zone; OMZ: Ossa Morena Zone; SPZ: South

; GB: Gorringe Bank; HAP: Horseshoe Abyssal Plain; CPS: Coral

f the Azores–Gibraltar Fracture Zone including epicenter locations

a plate boundary.

M. Fernandez et al. / Tectonophysics 386 (2004) 97–115 99

have been several attempts to study the crustal

architecture of the region although little is known

about the lithospheric mantle structure along the

transition from the Hercynian mainland to the oceanic

domain. The crustal structure of the Variscan Iberian

Massif has been investigated mainly through several

seismic wide-angle and refraction profiles (e.g.,

Mueller et al., 1973; Matias, 1996; Gonzalez et al.,

1996, 1998, among others). Crustal thickness beneath

the Iberian Massif lies around 30 km increasing

slightly towards the center of the Iberian Peninsula

(Banda, 1988; Surinach and Vegas, 1988). Major

discrepancies among authors refer to the Vp depth

distribution at intermediate/deep crustal levels, and to

the existence of noticeable crustal thickening (4–6

km) beneath the South Portuguese Zone (Fig. 1). On

the other hand, the Bouguer anomaly data show a

prominent positive anomaly affecting the South

Portuguese and the Ossa Morena zones (Fig. 1)

indicating an anomalous density distribution inside

the crust. Lateral variations in the nature/composition

of rocks at upper/middle crustal levels have been

envisaged in recent magnetotelluric and deep seismic

experiments showing the presence of high conductiv-

ity and high reflectivity bodies interpreted as sill-like

intrusions with assimilated pyritic graphite-rich rocks

(Almeida et al., 2001; Carbonell et al., 2003;

Simancas et al., 2003). Offshore the transition from

the continental crust of the shelf to the oceanic crust of

the Horseshoe Abyssal Plain have been investigated

by (Gonzalez et al., 1996) using coincident near-

vertical incidence and wide-angle seismic reflection

data. These authors conclude that the crust of the

Southwestern Iberia Margin undergoes a strong but

continuous thinning from 31 km onshore Iberia to less

than 15 km in the Horseshoe Abyssal Plain, and that

thinning occurs over an area which is about 120 km

wide. Two lithospheric transects across the western

Iberian margin were carried out by Torne et al. (1995).

The results show a rather constant lithospheric thick-

ness of about 110 km from mainland to the Iberian

Abyssal Plain, and a more variable thickness ranging

from 100 to 120 km beneath the Tagus Abyssal Plain

and the Madeira-Tore Rise.

Unlike the Variscan Iberian Massif, the southwest-

ern margin of the Iberian Peninsula and the oceanic

domain north and south of the Gorringe Bank region,

east of the Madeira Tore Rise, show active tectonic

activity characterized by a diffuse seismicity with

hypocenters down to 100–120 km (e.g., Grimison and

Chen, 1986; Mezcua and Rueda, 1997). This seis-

micity pattern has been interpreted by many authors as

indicating that the plate boundary is diffuse and that

the deformation spans a broad area (e.g., Sartori et al.,

1994; Jimenez-Munt et al., 2001a,b; Negredo et al.,

2002). Likewise, from structural geology and tomo-

graphic studies several authors have envisaged the

region as an example of westward delamination of the

lithospheric mantle (e.g., Royden, 1993; Seber et al.,

1996; Gutscher et al., 2002). However, neither the

position and dipping direction of the delaminated

mantle nor the origin of intermediate seismicity has

been yet identified.

In this paper, we model a 1000-km-long profile

that runs from central Iberia to the deep waters of the

Seine Abyssal Plain on the African plate (Fig. 1). The

profile crosses almost perpendicularly the three main

units of the Variscan Iberian Massif (Central Iberian

Zone, Ossa Morena Zone, and South Portuguese

Zone), the Southwestern Iberian Margin, and the

Africa–Eurasia plate boundary. The main objective of

this work is to image the lateral changes in crustal and

lithospheric structure when passing from a complete

section of an old and stable orogenic belt—the

Variscan Iberian Massif—to the adjacent Jurassic

passive margin of SW Iberia, and across the trans-

pressive and seismically active eastern segment of the

Azores–Gibraltar plate boundary. The results are

interpreted in terms of possible constraints on the

geodynamic evolution that affected the region. The

lithospheric modelling is based on a forward finite

element code that combines four coupled geophysical

observables: heat flow, gravity, geoid and absolute

elevation (Zeyen and Fernandez, 1994; Fernandez et

al., 2004).

2. Geological and tectonic setting

For simplicity, we can divide the study area into

two main tectonic regions, the Variscan Iberian Massif

and the Southwest Iberian Margin. The Variscan Belt

formed as a consequence of the collision between the

Balto-Laurentia and Gondwana continents, which

occurred during the late Paleozoic and is recognizable

in north and west Europe, northwest Africa and


eastern North and South America. In the Iberian

Peninsula, the Variscan Belt occupies the western half

of Iberia and in its southwestern region it encom-

passes, from NE to SW, the following tectonic units

(Julivert et al., 1974): the Central Iberian Zone (CIZ),

the Ossa Morena Zone (OMZ), and the South

Portuguese Zone (SPZ). These units together with

the Cantabrian and the Astur-Leonese Zones located

in the northern part of Iberia form the so-called

Variscan Iberian Massif (VIM) (Fig. 1).

The major tectonic episodes that affected the VIM

during the Paleozoic are (e.g., Simancas et al., 2001):

(i) the accretion of the Ossa Morena Zone to the

Centro Iberian Zone along the Badajoz-Cordoba

suture zone due to a continent–continent collision

which occurred during Mid–Late Devonian, (ii) the

accretion of the South Portuguese Zone to the former

terranes along the Beja-Acebuches suture zone as a

consequence of the closure of the Rheic ocean, which

was coeval with the continent–continent collision, (iii)

oblique extension with abundant bimodal magmatism

in the SPZ and intrusion of gabbros and diorites in the

OMZ during the Early Carboniferous, and (iv) oblique

compression that produced the left-lateral shear

deformation in the abovementioned suture zones and

diverse degrees of metamorphism.

As a result of these tectonic episodes, the Central

Iberian Zone represents the most internal part of the

Variscan paleo-cordillera and is made up of Precam-

brian and Paleozoic rocks with a variable degree of

metamorphism and large granitic intrusions. The Ossa

Morena Zone is made up of Precambrian to early

Paleozoic sediments and volcanic and plutonic rocks,

whereas the allochthonous terrane of the South

Portuguese Zone is made up of Carboniferous and

late Devonian folded and thrusted sediments with

almost no metamorphism. The Badajoz-Cordoba

shear suture zone separating the CIZ and the OMZ

is made up of Precambrian and early Paleozoic rocks

with high metamorphic grade and mafic rocks.

Whether this suture zone was created as a conse-

quence of the Variscan orogeny as proposed by

several authors (e.g., Burg et al., 1981; Matte, 1983;

Azor et al., 1994) or is the reactivation of a

Precambrian structure as argued by other authors

(e.g., Ribeiro et al., 1990; Quesada, 1991; Abalos et

al., 1991, 1993) is still a debatable question. The

contact between the OMZ and the SPZ, which is made

through the Beja-Acebuches shear zone, corresponds

to the subduction/obduction of the Rheic oceanic

lithosphere and encompasses three subunits: the Beja-

Acebuches Ophiolitic Complex made up of metaba-

salts, metagabbros and ultrabasic rocks; the Pulo do

Lobo Accretionary Terrane, underthrusting the former

subunit, which is made up of metasediments and has

been interpreted as the accretionary prism; and the

Iberian Pyrite Belt, made up of turbiditic sediments

interbedded with continental volcanic rocks (e.g.,

Moreno-Garrido and Vera, 1985; Mitjavila et al.,

1997).

A regional extensional tectonic episode which was

initiated during the Permian, resulted in the dismem-

berment of Pangea, the opening of the Atlantic Ocean

and the initiation of extension in the Southwest

Iberian Margin. The analysis of seafloor spreading

magnetic anomalies has permitted the kinematics of

the Iberian plate since Mesozoic times to be recon-

structed (e.g., Klitgord and Schouten, 1986; Srivas-

tava et al., 1990; Malod and Mauffret, 1990).

Continental breakup took place firstly along the

Central African margin (Triassic) and propagated to

the north until middle–late Jurassic times producing

seafloor spreading off the North African margin and

forming the dprotoT Azores–Gibraltar plate boundary

between Africa and Iberia/Newfoundland. Magnetic

anomaly M0 (120 Ma, Palmer and Geissman, 1999)

marks the separation between the Iberian and the

North American plates which coincides with the

counterclockwise rotation of Iberia relative to Eurasia

and the initiation of seafloor spreading in the Bay of

Biscay, which continued until chron 33 (80 Ma). After

that, at chron 31 (70 Ma) convergence between

Eurasia and Africa would have started. Iberia moved

together with the African plate from the latest Creta-

ceous until the mid-Eocene (chron 19, 42 Ma) when

they started to move independently each other and the

plate boundary changed from the Bay of Biscay to the

Azores–Gibraltar Fracture Zone. Essentially, the sense

of motion between Africa and Iberia/Eurasia has

remained the same since that time to the present.

Over the Southwest Iberian Margin, Triassic and

Jurassic sediments represent the pre-rift and syn-rift

deposits on top of a Variscan basement of Precam-

brian and Paleozoic age. Continental conditions are

evidenced by lower to middle Triassic sediments that

exhibit a shift towards a marine environment from


middle Triassic onwards (Tortella et al., 1997), giving

rise to a well-developed carbonate platform which

was largely fragmented during the mid Jurassic–early

Cretaceous rifting episode. Widespread deformation

associated with tectonic convergence is recorded by

Tertiary and Quaternary sediments over the region.

Major WSW–ENE trending structural highs (Gorringe

Bank and Ampere and Coral Patch seamounts)

separated by basement lows (Tagus, Horseshoe and

Seine abyssal plains) were produced after the early

Eocene (Sartori et al., 1994) (Fig. 1). According to

these authors, deformation is accommodated over a

wide region by discrete thrust faults and buckling of

varying age rather than by a single major plate

boundary. Very large chaotic sedimentary bodies, that

probably formed during the early–middle Miocene,

are recorded over the whole region resulting from

tectonic frontal wedges and/or gravitational sliding

(e.g., Torelli et al., 1997; Tortella et al., 1997).

The possible influence of the evolution of the

Betic–Gibraltar–Rif system on the Southwest Iberian

Margin is poorly known and rather unclear. Several

authors have proposed that the origin of the Betic–

Gibraltar–Rif system is related to the westward roll-

back of an east-dipping mantle slab initiated during

the Miocene (e.g., Royden, 1993; Seber et al., 1996;

Lonergan and White, 1997; Gutscher et al., 2002).

Although most of these authors imply that deforma-

tion in the Alboran and the Gorringe/Coral Patch

regions was not directly linked, major discrepancies

arise between authors suggesting that compressional

structures in the SIM are related to active subduction

(e.g., Purdy, 1975; Royden, 1993) and those who

propose that convergence does not imply subduction

at all (e.g., Sartori et al., 1994; Torelli et al., 1997).

3. Methodology

Commonly, the base of the lithosphere is consid-

ered either as an isotherm (1300–1350 8C) or as a

surface where a fraction (0.85–0.9) of the melting

temperature of dry peridotites is accomplished.

Modelling the present-day lithospheric structure

requires the combination of various independent

observables (seismics, elevation, Bouguer anomaly,

geoid height and surface heat flow) that give

information on the density and temperature distribu-

tion at different depth ranges. Therefore, the integra-

tion of these observables in a unique model reduces

substantially the range of possible solutions. We have

used a finite element code that solves simultaneously

the geopotential, lithostatic, and heat transport equa-

tions (Zeyen and Fernandez, 1994; Fernandez et al.,

2004). The modelled lithospheric section is meshed

with triangular elements and divided into a number of

bodies to which, according to an assumed lithology,

different material properties are assigned: thermal

conductivity, heat production and its variation with

depth, density and its variation with pressure and

temperature. The top of the model section is the

Earth’s surface and the bottom corresponds to the

lithosphere–asthenosphere boundary. The density of

the mantle lithosphere is considered to be temperature

dependent (Parsons and Sclater, 1977) such that

qm=qa(1+a(Ta�T(z)), where qa is the density of the

asthenosphere, a is the thermal expansion coefficient,

and Ta is the temperature of the asthenosphere.

Temperature distribution and surface heat flow are

calculated by solving the steady-state heat transport

equation with the following boundary conditions:

constant temperatures of 15 and 1350 8C at the top

and bottom of the model, respectively, and no

horizontal heat flow across the lateral boundaries.

The Bouguer gravity anomaly is calculated using

Talwani’s 2D algorithm (Talwani et al., 1959) for each

triangular element of the mesh, thus allowing for

density variations depending on pressure, temperature

and lithology. The geoid anomaly is calculated using a

3D algorithm based on the resolution of the gravity

potential for a regular prism and Brun’s formula

(Ayala et al., 2003), modified for a 2D finite element

grid (H. Zeyen, personal communication). Elevation is

calculated for every column of the mesh under the

assumption of local isostasy and following (Lachen-

bruch and Morgan, 1990). The depth to the isostatic

compensation level to calculate elevation and gravity

and geoid anomalies is taken at the maximum depth

reached by the lithospheric mantle, and the space

between this depth and the base of the model is filled

with asthenospheric material with a constant density.

The resulting elevation, Bouguer anomaly, geoid

height variation, and surface heat flow are compared

with the measured values and the crustal and litho-

spheric mantle geometry is modified until the best fit

is obtained.


4. Regional data

The data used in this study were compiled from

different sources. Elevation including both bathy-

metry and topography was taken from ETOPO2

Global Data Base. Bouguer gravity data were taken

from Torne et al. (1995) offshore, and from the

Spanish dInstituto Geografico NacionalT (Mezcua et

al., 1996) onshore. Geoid heights were obtained

from the EGM96 Global Model (Lemoine et al.,

Fig. 2. Regional data used in this study: (a) bathymetry and topography ma

gravity anomaly map. Onshore from IGN (Mezcua et al., 1996), offshor

anomaly map from EGM96 Global Model (Lemoine et al., 1998). Harm

Contour interval is 1 m; (d) heat flow density map in mW m�2. Data fro

1998). To subtract the regional component of the

geoid, we removed harmonic coefficients up to

degree and order eight, leaving a geoid anomaly

with wavelength components shorter than 5000 km.

Thermal data were obtained from the compilation of

heat flow data in the Iberian Peninsula and its

margins by Fernandez et al. (1998) and from

Verzhbitsky and Zolotarev (1989) in the Horseshoe

Abyssal Plain. Maps of these input parameters are

shown in Fig. 2.

p from ETOPO2 Global Data Base, contours at 400 m; (b) Bouguer

e from Torne et al. (1995). Contour interval is 25 mgal; (c) geoid

onic coefficients up to degree and order eight have been removed.

m Fernandez et al. (1998) and Verzhbitsky and Zolotarev (1989).


4.1. Elevation

Fig. 2a shows the elevation data used in this study.

Onshore, maximum topographic heights of 700 m are

recorded around the northeastern end of the model

profile (Tagus Basin, central Iberia). From there, the

topography smoothly descends throughout the Varis-

can peneplain to reach the coastline at Cape San

Vicente. The transition from the shallow waters of the

continental shelf to the deep waters of the Seine

Abyssal Plain (~4400 m) is marked by a narrow zone

with steep slopes, particularly along the model profile.

To the southeast, in the Gulf of Cadiz area, this

transition is much gentler and occurs along a wider

zone. In contrast, to the northwest of the offshore

section of the profile, the bathymetry is very rough

with a series of deep basins (Tagus, Horseshoe, and

Seine abyssal plains) bounded by structural highs

(Gorringe Bank and Coral Patch Seamount) reaching

depths of few hundred meters.

4.2. Gravity

The Bouguer gravity field over the Variscan Iberian

Massif (Fig. 2b) shows an NW–SE regional trend

parallel to the main geological structures decreasing

very gently from values of ~100 mGal at Cape San

Vicente (shoreline) to about�90mGal at the NE end of

the profile (Tagus basin). The most outstanding feature

is that the 0 mGal isoline roughly follows the suture

between the Central Iberian and Ossa Morena zones

crossing the profile more than 300 km from the coast.

The significant displacement of the 0mGal isoline from

the coast indicates, as will be discussed below, the

presence of a high density crust in the Ossa Morena

Zone. Offshore, the Bouguer anomaly increases from

100mGal on the continental shelf up to 300mGal in the

deep waters of the abyssal plains reflecting the subsea-

floor contributions of the gravity field. The bathymetry

highs of the Gorringe Bank and Coral Patch Seamount

are bounded by relative gravity highs with values of

up to 400mGal, which have been related to lithospheric

mantle uplifting (e.g., Purdy, 1975; Souriau, 1984).

4.3. Geoid

The geoid anomaly map of Fig. 2c shows that

onshore the geoid variations are quite smooth with a

relative NE–SW trending geoid height centered on the

Ossa Morena Zone where the geoid reaches a

maximum value of more than 12 m. Offshore the

map shows that the TAP and SAP are associated with

local blowsQ of about �1 and �4 m, respectively,

whereas the continental shelf is associated with geoid

highs of 8–10 m. The transition from the continental

crust of Iberia to the oceanic domain is characterized

by a decrease of about 10 m in the geoid anomalies.

The two domains are separated by a steep gradient

which correlates with the increase in water depth from

the continental shelf towards the deep waters of both

the TAP and SAP.

4.4. Heat flow

Heat flow measurements in the study area are

irregularly distributed (Fig. 2d). Onshore, measure-

ments are concentrated in the southwestern Ossa

Morena and South Portuguese zones, whereas almost

no heat flow data are available in the Central Iberian

Zone. The available values, which come from oil,

water and mining exploration wells, show average

heat flow values between 60 and 70 mW m�2 with a

well defined thermal anomaly centered in the Pyrite

belt with values up to 80–85 mW m�2. Offshore, the

few available heat flow measurements in the region of

the Horseshoe Abyssal Plain indicate an average value

of 57F15 mW m�2.

5. Crustal structure from seismic data

Refraction/wide-angle seismic data acquired dur-

ing the last decades onshore and offshore have shed

light on the crustal architecture, and on the velocity–

depth distribution at crustal and uppermost mantle

levels, particularly in the SW corner of Iberia and in

the Gulf of Cadiz areas. Fig. 3 shows a compilation of

the seismic data used to constrain the crustal structure

of our lithospheric model.

In the Tajo Basin and the Central Iberian Zone, we

have used the results from Surinach and Vegas (1988)

and ILIHA DSS Group (1993) to which for compar-

ison we have added those from Pulgar et al. (1996)

corresponding to the Duero Basin. For the SW Ossa

Morena Zone and the South Portuguese Zone, we

have mainly used the compilation and reinterpretation

Fig. 3. Top map: location of seismic profiles used to constrain the crustal structure. Dotted line on lithospheric section denotes location of crustal cross section by Matte (1983) (Fig.

5). Bottom panel: geometry of the crustal layers as inferred from seismic data. P-wave velocities are shown in km s�1.

M.Fern

andez

etal./Tecto

nophysics

386(2004)97–115

104


of seismic wide-angle and refraction data by Gonzalez

et al. (1998) and the work by Mueller et al. (1973).

The available seismic wide-angle and refraction

results show that the crust underneath the Tertiary

Tajo Basin is about 32 km thick, slightly thinning

towards the OMZ, where the crust is less than 30 km

thick. The results in the SPZ indicate a noticeable

crustal thickening with thickness values around 34

km. Nevertheless, it is worth noting that Matias

(1996), based on a combined velocity–amplitude

analysis of available refraction and wide-angle pro-

files, proposes a nearly flat Moho at the SPZ and

OMZ lying at 31 km depth and rising towards the

continental margin. An outstanding result of the

seismic experiments is that the average velocity of

the mid and lower crust remains almost constant

onshore along the studied profile, whereas significant

lateral changes are observed at upper crustal levels. In

Fig. 3, we observe how upper crust P-wave velocities

vary from minimum values of 5.6–5.7 km s�1 to

maximum values of 6.1 km s�1. These changes can be

even more dramatic near the contact between OMZ

and SPZ with velocities of 6.35 and 6.7 km s�1 at

shallow levels probably related to exhumation of deep

crustal levels and/or basic intrusions (Matias, 1996).

According to Gonzalez et al. (1998) and ILIHA DSS

Group (1993), the middle crust beneath the SPZ and

partly beneath the OMZ is thickened relative to the

CIZ and this thickening is taken up by the upper crust.

The mid and lower crust are characterized by average

P-wave values of 6.4 and 6.8 km s�1, respectively,

whereas the Pn velocity falls within the range of 8.0–

8.1 km s�1 (e.g., Surinach and Vegas, 1988; ILIHA

DSS Group, 1993; and Gonzalez et al., 1998).

Offshore, results of the coincident near-vertical

incidence and wide-angle reflection data along an NE–

SW oriented seismic profile (Gonzalez et al., 1996)

have allowed us to constrain the crustal structure at the

transition from continent to ocean (Fig. 3). These

results show that the crust undergoes a strong but

continuous thinning from less than 30 km under the

continental shelf to less than 15 km at the eastern end

of the Horseshoe Abyssal Plain (see Figs. 1 and 3), and

that thinning is mainly taken up by the lower crust.

Along the transition zone, the crust is characterized by

average velocities that range from 2.2 to 3.7 km s�1 in

the sedimentary layers to 6.0 and 6.4 km s�1 at upper

and mid crustal levels, respectively. The Pn velocity is

poorly constrained, and there is weak evidence under-

neath the continental shelf that it falls within the range

of 7.8–7.9 km s�1. The geometry of the sedimentary

layers and the water depth were taken directly from

IAM-3 MCS profile (Banda et al., 1995; Tortella et al.,

1997). Hence, the thickness of the sedimentary cover

varies from 0.5 s (twtt) on the continental shelf and

slope to about 2 s (twtt) in the central part of the profile

(around �150 km distance in Fig. 3).

The crustal structure of the oceanic domain of the

Seine Abyssal Plain was derived from results of

Profile C of Purdy (1975) and Sartori et al. (1994).

These results show that the crust is anomalously thick

in the vicinity of the Coral Patch Seamount (about 11

km, including the sedimentary column) thinning

towards the central region of the Seine Abyssal Plain

to values of less than 7 km. Results of Purdy (1975)

show that the oceanic crust in the central region of the

Seine Abyssal Plain has P-wave velocities that range

from 4.1 to 5.6 km s�1 at upper crustal levels to 6.3

km s�1 at lower crustal levels, and that the uppermost

mantle is characterized by an anomalous Pn velocity

of 7.6 km s�1 (Fig. 3).

6. Lithospheric modelling

6.1. Modelling inputs

The model transect has been divided into a mesh of

211�59 nodes with a horizontal grid spacing of 5 km

and a variable vertical grid spacing depending on the

geometry of the defined bodies. Initially, we consid-

ered the crustal structure depicted in Fig. 3 and a flat

lithosphere–asthenosphere boundary at 110 km depth.

This initial geometry was successively modified until a

best fitting model was reached (Figs. 4 and 5).

According to the considered petrophysical properties

(density, thermal conductivity and heat production),

the lithospheric section has been divided into 12

different materials (Table 1). Most of the rock

parameters have been taken from direct measurements,

empirical relationships, and from previous modelling.

The density of the sedimentary cover varies from

2400 kg m�3 for the Tagus Basin and deeply buried

marine sediments in the continental margin to 2000 kg

m�3 for upper sediments in deep oceanic basins, in

agreement with gravity studies performed in the

Fig. 4. Results along model transect. (a) Heat flow; (b) geoid anomaly; (c) Bouguer gravity anomaly; (d) elevation. Dots with bars indicate measured data and grey thick lines indicate

calculated values. (e) Model geometry with body numbers as in Table 1. Hypocenters (MN3) have been projected onto the profile from a strip of 40 km half-width. Same legend as in

Fig. 1.

M.Fern

andez

etal./Tecto

nophysics

386(2004)97–115

106

Fig. 5. (a) Blow up of the obtained crustal structure. Body numbers as in Table 1. (b) Comparison between the obtained crustal structure and that

deduced from compilation of available seismic data (Fig. 3). (c) Comparison between the crustal structure obtained in this study and the

geological cross section proposed by Matte (1983).


Duero Basin (Fernandez-Viejo et al., 2000) and along

the IAM-3 MCS profile (Gonzalez et al., 1996),

respectively. At crustal levels, densities were obtained

from the empirical relationships between P-wave

velocity and density of Ludwig et al. (1970), Zelt

and Ellis (1989), Barton (1986), and Christensen and

Shaw (1970). In the upper continental crust, densities

vary from 2680 kg m�3 underneath the CIZ and OMZ

to 2740 kg m�3 below the SPZ and the continental

shelf of the Gulf of Cadiz. A constant density of 2800

kg m�3 was assigned to the middle crust and to the

Beja-Acebuches Ophiolitic Complex, whereas for the

lower continental crust we used a constant value of

2950 kg m�3. The petrophysical properties of the Pulo

do Lobo Accretionary Terrane and the Iberian Pyrite

Belt were obtained from the best fitting model, which

required relatively high values of density, thermal

conductivity and heat production (Table 1, Fig. 5).

We have defined a transition zone between the

continental domain of the margin and the oceanic

Table 1

Rock parameters used in the model transect

Body

number

Heat

production

(AW m�3)

Thermal

conductivity

(W m�1 K�1)

Density

(kg m�3)

Neogene marine

sediments

1 1.2 c 2.0 a 2000 d

Paleogene

sediments

2 1.2 b 2.3 a 2400 d

Upper crust SPZ 3 2 a 3 a 2740 d

Upper crust

OMZ–CIZ

4 2.5 a 3 a 2680 d

Pyrite belt and

accretionary

Terrane

5 3.5 e 3.5 e 2770 e

Transitional crust 6 0.3 c 2.1 c 2800 c

Oceanic crust 7 0.3 c 2.1 c 2840 c

Middle crust

and Beja-

Acebuches

8 2 c 2.4 c 2800 d

Lower crust 9 0.3 c 2.1 c 2950 d

Anomalous

upper mantle

10 0 c 3.2 c 3160 d

Lithospheric

mantle

11 0 c 3.2 c T-dependent

Lithospheric mantle density is temperature dependent being:

3200(1+3.5�10�5 8C�1(Ta�T(z))); Ta=1350 8C. Other data come

from: (a) Fernandez et al. (1998); (b) from Rybach and Cermak

(1982); (c) from Torne et al. (1995); (d) densities deduced from the

empirical relationship between P-wave velocities and density of

Ludwig et al. (1970), Zelt and Ellis (1989) and Barton (1986); (e)

from the best fitting model.


domain of the Seine Abyssal Plain. This transitional

crust has been proposed also by Gonzalez et al. (1996)

along the IAM3 MCS profile and by Torne et al.

(1995) along lithospheric profiles crossing the Atlantic

Iberian margin at the Tagus and the Iberia Abyssal

Plain. Hence, we have considered a density of 2800 kg

m�3 for the transition crust, and a density of 2840 kg

m�3 for the oceanic crust of the SAP, which following

Purdy (1975) is underlain by an anomalous mantle to

which we have assigned a density of 3150 kg m�3

(Fig. 5 and Table 1). For the lithospheric mantle, we

considered a temperature dependent density referred to

the density of the asthenosphere, which is assumed to

be constant (Lachenbruch and Morgan, 1990).

6.2. Modelling results

Figs. 4 and 5 present the results of the lithospheric

modelling along the regional model profile. Panels

(a), (b), (c), and (d) in Fig. 4 show the observed and

calculated heat flow, geoid and Bouguer anomalies,

and elevation, while panel (e) shows the lithospheric

structure. The observed heat flow values correspond

to the isolines depicted in Fig. 2d to which we have

associated an uncertainty of F10% (Fernandez

et al., 1998; Verzhbitsky and Zolotarev, 1989). The

observed Bouguer and geoid anomalies and elevation

data have been projected onto the profile from a strip

of 40 km half-width, the error bars indicating the

standard deviation associated with the lateral varia-

bility of each observable. The model results are in

good agreement with the data observed with the

exception of some local misfits. Short wavelength

misfits between the observed and calculated elevation

and Bouguer anomaly are found along the continental

slope and the structural highs bounding the abyssal

plains (Fig. 5c, d), which are attributed to local

features that are out of the scope of our regional

modelling, e.g., related to the gross geometry adopted

for the sediments and/or the top of the basement. The

geoid anomaly is fitted very well, whereas the surface

heat flow follows the observed general trend although

the large uncertainties of measured data make it more

difficult to evaluate the apparent match. The predicted

values in the vicinity of the Horseshoe Abyssal Plain

are considerably lower than those reported by

Verzhbitsky and Zolotarev (1989). These authors

propose that the Iberian–African plate boundary

behaves as a hot belt due to the active tectonics of

the region although the scarcity and variability of data

makes it difficult to confirm this hypothesis.

Our results (Figs. 4 and 5) show that across the

Variscan Iberian Massif the thickness of the crust

keeps almost constant at about 30 km depth. The

continental crust is characterized by three layers with

noticeable lateral thickness variations. The upper

crust, which is divided into different density

domains roughly coinciding with the main geo-

logical features observed in the area (Matte, 1983),

thins from 20 km below the CIZ to less than 7 km

underneath the OMZ. Towards the continental shelf,

in the South Portuguese Zone, the upper crust

thickens to values of 8–9 km. The upper crustal

thinning occurs at the expense of the middle crust

which below the OMZ and the northern half of the

SPZ reaches maximum thickness values of 18 km.

The thickness of the lower crust remains almost


constant (6 km) all along the Iberian Massif,

thinning out towards the Gulf of Cadiz. Our crustal

results confirm some of the gross crustal features

deduced from wide-angle seismic data. Major dis-

crepancies, however, are related to the upper/middle

crustal structure of the OMZ and the CIZ, which

according to seismic data should be much more

homogeneous; and to the total crustal thickness

beneath the SPZ, where Gonzalez et al. (1998) pro-

pose a crustal thickening of 3–4 km (Figs. 3 and 5b).

Across the continental margin, the crust thins from

30 km to less than 10 km (excluding water layer) over

a horizontal distance of 200 km. This crustal attitude

is also observed further north in the Tagus and the

Iberian abyssal plains where Torne et al. (1995) have

modelled two lithospheric profiles showing that the

continental crust thins across the passive margin over

horizontal distances of about 125 km. Similar results

have been obtained more recently from combined

seismic refraction and gravity modelling in the

Iberian Abyssal Plain (Dean et al., 2000). There is

no evidence of a well defined ocean–continent

boundary in the region and therefore we assumed

that the crust occupying the easternmost Horseshoe

Abyssal Plain and Coral Patch Seamount, whose

nature is not well determined (body #6 in Figs. 4 and

5), corresponds to a single transitional layer with an

average density of 2800 kg m�3 and variable thick-

ness. Such transitional crust has also been proposed

by Torne et al. (1995) and Dean et al. (2000) in the

western Iberian Margin. For simplicity, the oceanic

crust (without sediments) was defined as a single

average layer with a thickness of 5–6 km and density

of 2840 kg m�3. Although our results do not differ

very much from those obtained by Gonzalez et al.

(1996) from a combined wide-angle seismic and

gravity study, we find a necking-like structure

beneath the Horseshoe Abyssal Plain where the

minimum continental crustal thickness is attained

(Fig. 5). Differences also extend to the Coral Patch

Seamount that we characterize as a transitional crust,

and where we infer a Moho depth of ~14 km in

agreement with Sartori et al. (1994). At the end of the

profile, in the Seine Abyssal Plain, we have consid-

ered an anomalous mantle about 6 km thick consistent

with measured Vp values of 7.6 km s�1 (Purdy, 1975)

and in a way similar to that proposed by Torne et al.

(1995) beneath the Tagus Abyssal Plain.

The lower panel in Fig. 4 shows a noticeable

lithospheric thinning affecting the central part of the

model profile where the lithosphere–asthenosphere

boundary (LAB) rises from about 120 km depth

beneath central Iberia and its continental margin to

less than 96 km beneath the Ossa Morena and South

Portuguese zones. Lithospheric thicknesses of around

110–120 km have been proposed by Torne et al.

(1995) in two profiles crossing the Iberian Atlantic

margin, and by Fernandez et al. (1998) in the stable

Iberian Peninsula. The lithospheric mantle thinning is

mainly enforced by the simultaneous fitting of gravity,

elevation and geoid signatures. Thereby, the high

Bouguer anomaly values measured at the OMZ and

the SPZ demand a relative high-density crust that

must be compensated by a deeper density deficit to be

compatible with moderate topography and geoid

height values (Fig. 2). Further to the SW and close

to the Seine Abyssal Plain, the lithosphere thickens

again to values of 125 km coinciding with both the

continent–ocean transition and the Iberian–African

plate boundary.

7. Discussion

The most outstanding results obtained from the

present study are: (i) the thickening of the mid-crustal

layer affecting the Ossa Morena and South Portuguese

domains; (ii) the thinning of the mantle lithosphere

beneath these regions; and (iii) the lithospheric

structure across the Iberian–African plate boundary.

In the following, we will discuss the possible origin

and significance of these findings.

7.1. Mid-crust thickening in the Ossa Morena and

South Portuguese zones

Thickening of the mid-crust layer have been

reported by Gonzalez et al. (1998) from wide-angle

seismic data in the SW corner of the Iberian

Peninsula. According to these authors, there is a high

apparent velocity (6.4 km s�1) layer at shallow depths

(7–10 km) affecting the South Portuguese Zone,

which could be related to mafic and ultramafic rocks

of the Beja-Acebuches Ophiolitic Complex. Our

results, however, show that the thick mid-crust layer

extends further landwards into the Ossa Morena Zone


and reaches a dnormalT thickness beneath the suture

between the OMZ and the CIZ. The coincidence

between the limit of the mid-crustal thickening and

the limit of the regional positive Bouguer anomaly

(Fig. 2b) suggests that thickening and hence mass

excess are related to the lithologies forming the South

Portuguese and Ossa Morena Zones. To test this

hypothesis, we calculated the gravity effect associated

with anomalous bodies representing the mid-crustal

thickening of the SPZ and the OMZ and the high

density rocks of the suture zone. The geometry and

density contrast of these bodies (Fig. 6) have been

estimated from seismic data (Gonzalez et al., 1998)

and from our model results (Fig. 5 and Table 1). Fig.

7 shows that the resulting Bouguer anomaly, after

removing the gravity effect of the abovementioned

bodies, lacks most of the anomalous positive signa-

ture and confirms that both the SPZ and the OMZ

have a distinct nature characterized by a denser upper/

middle crust. At this point, it is worth noting that our

analysis does not allow us to distinguish between an

effective mid-crustal thickening and the presence of a

denser body placed at upper/mid-crustal levels.

Evidence for processes that could have produced a

density increase at upper/mid-crustal levels is pro-

vided by recent magnetotelluric and deep seismic

reflection experiments. Magnetotelluric data show a

highly conductive layer at about 10 km depth

occupying most of the OMZ and part of the SPZ

which is attributed to pyritic graphite-rich lithologies

(Almeida et al., 2001). Similarly, deep seismic

Fig. 6. Sketch showing the 3D geometry used to calculate the gravity con

(light grey), and the Beja-Acebuches Ophiolitic Complex (dark grey). Li

seismic profiles by Gonzalez et al. (1998).

reflection data reveal a 1–2 s twtt thick highly

reflective layer located at 4–6 s twtt depth affecting

the Ossa Morena and part of the Centro Iberian zones

which is interpreted as a sill-like structure generated

by intrusion of mafic and ultramafic magmas

(Simancas et al., 2003). On top of this structure and

between 2 and 4 s twtt depth, a series of high-

amplitude east dipping reflections is observed (Car-

bonell et al., 2003). Nevertheless, the relationship

between high amplitude reflections, high reflectivity,

high conductivity and high density is still unclear.

7.2. The mantle lithosphere in the Variscan domain

As mentioned, to make the high gravity values in

the SPZ and OMZ compatible with a moderate

topography and a relative maximum in geoid height

requires a density excess at mid-crustal levels that

must be compensated by a deep-seated density

deficit. Therefore, the lithospheric mantle has been

thinned by raising the LAB between 20 and 25 km

(Fig. 4e). This result is a direct consequence of the

assumptions made in the numerical modelling. In

particular, we assumed that the lithospheric mantle is

homogeneous and that its density is only temper-

ature-dependent. If this assumption is correct, the

lithospheric mantle thinning is real and must be

interpreted as a transient thermal anomaly related to

a recent, probably post-Mesozoic, tectonothermal

event. However, the Variscan Iberian Massif has

remained stable since late Paleozoic times, which

tributions of the mid-crustal thickening affecting the SPZ and OMZ

ne #1 denotes the modelled lithospheric profile, and #2 and 3, the

Fig. 7. Bouguer gravity anomaly corrected by the 3D gravity contribution of the crustal bodies shown in Fig. 6. Note that the 0 mgal isoline is

now displaced towards the shoreline. Contour interval is 25 mgal.


suggest that lithospheric thinning is more likely

related to lateral effects of late Oligocene–middle

Miocene tectonothermal processes that occurred in

the Alboran-Gulf of Cadiz (e.g., Gutscher et al.,

2002; Calvert et al., 2000; Bijwaard and Spakman,

2000). The main problem with this interpretation is

the long distance to which the thermal effects related

to the geodynamics of the Iberian–African plate

boundary must propagate and the apparent correla-

tion with the OMZ–CIZ suture zone. An alternative

explanation is to consider that the lithospheric mantle

can contain lateral heterogeneities capable of pro-

ducing noticeable density variations. Such mantle

heterogeneity must be placed near the base of the

lithosphere to keep the density moment associated

with the geoid anomaly, and must account for a

density contrast of about 25 kg m�3, which is the

density change associated with a temperature varia-

tion of about 200 8C. With these constraints, a

plausible interpretation is to regard the LAB as a

roughly flat surface where the space associated with

the uplifted asthenosphere would now be occupied

by depleted lithospheric mantle with a density

similar to that of the asthenosphere (Fig. 8). The

subduction/obduction of the Rheic oceanic litho-

sphere and the further continental collision that

configured the SPZ and the OMZ could have led

to partial melting of the lithospheric mantle as

evidenced by the presence of mafic Carboniferous

magmatism. Melt extraction produces a long lived

density decrease in the residual mantle due to

depletion in basaltic components, which can amount

40 kg m�3 depending on the melt fraction (e.g.,

White and McKenzie, 1989; Oxburgh and Parment-

ier, 1977; Doin et al., 1996). This alternative

interpretation also supports the mafic/ultramafic

nature of the sill-like structure corresponding to the

observed highly reflective mid-crust band at the

OMZ, as proposed by Simancas et al. (2003).

7.3. Lithospheric structure across the Iberian–African

plate boundary

Plate reconstruction and magnetic anomaly studies

reveal that continental break-up of the African plate

occurred during the Late Jurassic (~156 Ma) whereas

Fig. 8. Comparison between two different lithospheric thermal structures along the modelled profile which assume that the SPZ and OMZ are

affected either by lithospheric thinning (upper panel) or by mantle depletion at the base of the lithosphere with a density contrast of �25 kg m�3

(lower panel).


in the Iberian plate it occurred in Early Cretaceous

times (~118 Ma) (Srivastava et al., 1990). Today the

contact between the plates changes progressively

from a pure right-lateral strike-slip plate boundary

along the Gloria fault to a diffuse transpressive plate

boundary in the Gorringe Bank–Gibraltar region.

Our profile crosses the continental domain of the

southwestern Iberian margin and enters into the

northern Seine Abyssal Plane, which according to

Purdy (1975) corresponds to the oceanic domain of

the African plate. The obtained crustal structure

shows a major thinning beneath the easternmost

Horseshoe Abyssal Plane, a wide transition zone

around the Coral Patch Seamount, and a dnormalToceanic crust underlain by anomalous mantle in the

Seine Abyssal Plane. At deeper levels, the results

show a noticeable lithospheric thickening when the

profile passes from the complex and diffuse Gorringe

Bank–Gibraltar plate boundary region to the African

plate. A possible explanation for this thickening is

the age difference between the oceanic domains of

the African and the Iberian plates which would result

in a thermally thicker African lithospheric mantle.

This hypothesis implies that deformation associated

with transpression, which is evident at crustal levels

(e.g., Torelli et al., 1997; Sartori et al., 1994), did not

substantially affect deep lithospheric levels. Con-

versely, lithospheric thickening could be entirely

attributed to tectonic deformation if we consider that

oceanic lithosphere older than 120 My is very close

to thermal equilibrium as proposed by lithospheric

plate models (e.g., Parsons and Sclater, 1977). Two

reasons make us give more support to the second

hypothesis. Firstly, the lack of noticeable differences

in bathymetry and geoid height across the Gloria

transform plate boundary, where differences in plate

thickness should be even more conspicuous. Sec-

ondly, the occurrence of intermediate depth seismic-

ity in the transpressive Gorringe Bank–Gibraltar

region (Fig. 4e), which strongly suggests that


tectonic deformation is affecting lithospheric levels

to depths of 80–100 km.

Acknowledgements

We are indebted to A. Perez-Estaun and R.

Carbonell for their helpful comments and suggestions

on the structure and geological evolution of the

Iberian Variscan Belt. A. Gonzalez and an anonymous

reviewer are also thanked for their constructive

comments. Data acquisition and processing of the

IAM project were funded by EU project JOU2-CT92-

0177. This work has been supported by Spanish

research agencies through projects PB94-0013,

MAR98-0962, and REN2001-3868-C03-02/MAR,

and NATO Grant EST.CLG.978922.

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Documents

Lithospheric transition from the Variscan Iberian Massif to the Jurassic oceanic crust of the Central Atlantic