kaban_2010_crustal model for europe

Earth and Planetary Science Letters 296 (2010) 195–209

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Earth and Planetary Science Letters

j ourna l homepage: www.e lsev ie r.com/ locate /eps l

An integrated gravity model for Europe's crust and upper mantle

M.K. Kaban a,⁎, M. Tesauro a,b, S. Cloetingh b

a Deutsches GeoForschungsZentrum Potsdam (GFZ), Germanyb Netherlands Research Centre for Integrated Solid Earth Science, Faculty of Earth and Life Sciences, VU University Amsterdam, The Netherlands

⁎ Corresponding author. Deutsches GeoForschungsZgrafenberg, 14473 Potsdam, Germany.

E-mail addresses: [email protected] (M.K. Kaba(M. Tesauro), [email protected] (S. Cloetingh).

0012-821X/$ – see front matter © 2010 Elsevier B.V. Adoi:10.1016/j.epsl.2010.04.041

a b s t r a c t
a r t i c l e i n f o
Article history:Received 13 October 2009Received in revised form 12 April 2010Accepted 20 April 2010Available online 11 June 2010

Editor: L. Stixrude

Keywords:3D gravity modelinglithospheredensity structureupper mantle

We present an integrated gravity model of the European lithosphere based on an analysis of a number ofnew data-sets, leading to a much higher resolution than provided by previous models. First of all, a recentcrustal model (EuCRUST-07) is used to quantify the crustal contribution to the observed gravity field and toidentify the effect of mantle heterogeneity. The new gravity field model is based on a combination of satellite(CHAMP and GRACE) and terrestrial data. We also use these data-sets to estimate residual mantle gravityanomalies and residual topography, reflecting the effect of mantle density variations induced by temperatureand compositional heterogeneity. The separation of these effects is vital for a proper assessment of mantlestructure and evolution. In addition, we utilize a new tomographic model for P- and S-velocity anomaliesbeneath Europe, which is a-priori corrected for crustal structure using EuCRUST-07. The seismic velocityanomalies were subsequently converted into temperature anomalies using a mineral physics approach. Weestimate the effect of temperature variations on the gravity field and subtract it from the total mantle field.The residual fields point to an important role of compositional density anomalies in the upper mantle. Anumber of key features of the compositional density distribution, so far invisible in seismic tomography data,are detected for the first time.

entrum (GFZ) Potsdam, Tele-

n), [email protected]

ll rights reserved.

© 2010 Elsevier B.V. All rights reserved.

1. Introduction

Quantification of density inhomogeneities in the crust and uppermantle is important for geodynamic studies. Thermal and compositionaldensity anomalies in the upper mantle exert a prime control on thecouplingof deepEarth and surface processes. Previous studies (e.g. Kabanet al., 2003) have demonstrated that seismic tomography can identifyinhomogeneities in mantle structure mainly related to temperatureanomalies. However, many other features, which play also a key role indynamic processes, remainhidden fromseismic studies. These anomaliescan partly be resolved by an integration with gravity modelling. Inaddition, density variations in the crust and mantle lithospheresignificantly affect stress fields within the lithosphere. Gravity modellinghas also been frequently used to investigate the lithospheric structure ofEurope and to constrain magnitudes of tectonic forces.

So far most investigations were carried out to obtain informationdirectly from the observed field (like Bouguer anomalies), to inferoverall crustal structure (e.g. Gomez-Ortiz, 2005; Rotstein et al., 2006;Pinto et al., 2005). However, the employment of only gravity data isnot sufficient to obtain a reliable model, since the solution of theinverse gravity problem is usually non-unique (e.g. Kaban et al.,2004). Therefore, other geophysical data (primarily seismic) are

required and used in many studies to investigate the lithosphere (e.g.Kaban and Mooney, 2001; Kozlovskaya et al., 2001; 2004; Ayala et al.,2003; Lyngsie et al., 2006; Pedreira et al., 2007). Furthermore, severalattempts weremade to use gravity data togetherwith topography andsurface heat flow to constrain thermal structure of the lithosphere asfor example in the Pannonian Basin (Zeyen et al., 2002) and theEastern Carpathians (Dérerová et al., 2006).

However, most of these studies cover relatively small areas ofEurope, with a number of them (e.g. Kozlovskaja et al., 2004; Pedreiraet al., 2007) based on interpretations of separate seismic sections.These studies have also been performed using different modellingapproaches, data-sets and reference models. Therefore, on a fullEuropean scale, it is difficult to compare the results obtained fordifferent subregions. The first gravity model of the Eurasianlithosphere was constructed by Artemjev et al. (1994). However,this large-scale model, based on sparse and by now obsolete data,does not resolve many important details of the European lithosphere.A number of subsequent studies focused on the European region (e.g.Yegorova and Starostenko, 1999, 2002; Artemieva et al., 2006;Tesauro et al., 2007). The initial data used in these studies are alsoincomplete and largely outdated in the light of crustal and gravity dataonly recently acquired. It should also be noted that most of thesestudies are limited to an investigation of the composite densitystructure of the mantle without a separation of temperature andcompositional effects. Up till now, this has been investigated only forvery large structures (ca. 1000 km of larger) (e.g. Kaban et al., 2003;Artemieva et al., 2006).

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196 M.K. Kaban et al. / Earth and Planetary Science Letters 296 (2010) 195–209

Reliability of the gravity modelling strongly depends on robust-ness of the crust/upper mantle models, which are mainly based onseismic data. As widely demonstrated, the observed gravity field isstrongly affected by density anomalies within the crust, being themost heterogeneous layer in the Earth (e.g. Kaban et al., 2004). Thecrustal effect can almost completely mask the effect of mantleanomalies, especially on a regional level. A common way to overcomethis difficulty invokes the a priori calculation of the crustal gravityeffect and its subsequent subtraction from the observed gravity field.The resulting residual mantle anomalies directly image densityvariations in the mantle (e.g. Kaban et al., 2003). The latter aresuitable for geodynamic reconstructions and for modelling of theprocesses responsible for the evolution of the lithosphere. Attempts tocalculate mantle gravity anomalies have been made since the '70s(e.g. Nersesov et al., 1975), when the first seismic experimentsprovided data on the crustal structure. However, a prerequisite forreliable 3D models is the availability of a sufficiently large body ofhigh-quality seismic data. In this context, it has been demonstratedthat significant errors in Moho position might lead to errors in theresidual mantle anomalies, in excess of 100 mGal (Tesauro et al.,2008). Therefore, the construction of a new mantle gravity model forEurope, using the most recent data on crustal structure and a self-consistent approach, is obviously a great challenge.

Below we present a new gravity model of the Europeanlithosphere (35°N–71°N, 25°W–35°E). This model is based onseveral data-sets, which are greatly improved in comparison withprevious studies. To remove the crustal effect from the observedgravity field, we use a new 3D model of the crust (EuCRUST-07,Tesauro et al., 2008). The observed gravity field model is based onjoint interpretation of existing terrestrial and new satellite data,which provide for the first time a homogeneous coverage of thewhole study area (Förste et al., 2008). Temperature variations in theupper mantle are determined from an inversion of a recenttomography model (Koulakov et al, 2009). This model has been apriori corrected for the crustal effect, using EuCRUST-07 (Tesauroet al., 2008). These constraints provide the opportunity to constructfor the first time an integrated density model of Europe's uppermantle, which enables to resolve some key temperature andcompositional anomalies.

2. Calculation of the gravity effect of the crust

2.1. Gravity data

Although a large amount of surface gravity observations isavailable for Europe, it has even till recently been rather problematicto construct a homogeneous gravity map for the entire region.Existing national gravity surveys are often incomplete and thecorrections (including, for example, terrain corrections) applied tothe initial observations are often not specified. New satellite data(CHAMP and GRACE) provide for the first time a basis for constructionof such a homogeneous gravity model. We use EIGEN-GL04C, which isbased on a combination of satellite (CHAMP and GRACE) andterrestrial data (Förste et al., 2008). The model is complete to adegree/order 360 in terms of spherical harmonic coefficients andresolves features of about 55 km width in the geoid and gravityanomaly fields. This resolution is more than sufficient for the presentstudy since we focus on the mantle structure. A special band-limitedcombination method has been applied in order to preserve the highaccuracy from the satellite data in the lower frequency band of thegeopotential and to form a smooth transition to the high frequencyinformation obtained from surface data. For the specified resolution,the GL04C model is very close to the new EGM2008 model (Pavliset al., 2008). The Bouguer correction has been computed usingETOPO-2 data for elevations/bathymetry (http://dss.ucar.edu/data-sets/ds759.3/). The resulting Bouguer anomalies, together with the

topography and the map showing main tectonic features of Europeare displayed in Fig. 1(a–c).

2.2. Key features of the gravity calculations

As common, we perform gravity calculations relative to astandard reference density model, which depends on depth only.This method intrinsically reduces “edge” effects. We assume for theaverage density of the crystalline crust a value of 2.85 g/cm3, inagreement with regional and global compilations (Christensen andMooney, 1995). The anomalous effect of each crustal layer iscomputed relative to this value. Although density within the crustincreases with depth, the total gravity effect of the whole crust willbe the same. In addition, a constant reference density is also moresuitable for the calculations (Kaban et al., 2004). The Moho depth ofthe reference model corresponds to its average level (28.1 km)beneath the entire study area. It should be noted that a particularchoice of a horizontally homogeneous reference model essentiallyonly affects a constant level of the calculated fields. The latter is outof the scope of the present study, where we address spatialvariations in gravity. The most important parameter is the averagedensity of the upper mantle, which is set to 3.32 g/cm3. However, ithas been demonstrated that variations of this parameter withinreasonable limits (±0.04 g/cm3) do not change the results signifi-cantly (Kaban et al., 2004).

Within the study area the grid resolution for the calculations is15′×15′, as for EuCRUST-07. Outside this, the data from the globalcrustal model of Kaban et al. (2004) are used, which has a resolutionof 1°×1°. This global model represents a combination of the regionalmodels for Northern and Central Eurasia (Kaban, 2001), NorthAmerica (Kaban and Mooney, 2001) and of the global modelCRUST2.0 (Bassin et al., 2000) outside these areas. These data areless reliable than EuCRUST-07. However, due to their location faraway from the points used in the calculation, uncertainties in thecrustal model do not significantly affect the results (Kaban et al.,2004). Therefore, the gravity effect of each crustal layer or boundary iscalculated, considering variations of the corresponding layer/bound-ary over the whole Earth up to antipodes. This is a main differencebetween our approach and most previous analyses. The gravityanomaly of each layer within the Earth's crust andmantle is calculatedusing 3D algorithms for a spherical Earth, taking into account changesof density in the horizontal and vertical direction and the averageelevation of each cell. The sum of the gravity effect of elementaryvolumes corresponding to the initial grids is computed. The algorithmof Artemjev and Kaban (1994), based on the formulas of Strakhovet al. (1989), and improved in Kaban and Mooney (2001) and Kabanet al., (2002), is used for the calculation. The estimated accuracy of thecalculations is within 1 mGal.

2.3. Gravity effect of sediments

The thickness of sediments in the study area is specified in theEuCRUST-07 model with considerable detail (Tesauro et al., 2008).However, the density structure of the sedimentary layer is nothomogeneous in both vertical and lateral directions. Furthermore,density variations within sediments often produce even larger gravityeffects than the variations in its thickness (e.g. Kaban and Mooney,2001). Geophysical exploration methods and drilling data reveal avery complex structure of the sedimentary cover, including numerouslocal boundaries. Due to this strong heterogeneity, it is very difficult tojoint separate strata in well-logs into uniformed layers over the wholeregion. Whenever possible we use a detailed division of thesedimentary layer obtained from local studies, like in NorthernEurope (Scheck-Wenderoth et al., 2007) and in the NorthwesternMediterranean (Ayala et al., 2003). In the areas where detailedmodelsare not available, we construct a smooth density–depth relationship

http://dss.ucar.edu/datasets/ds759.3/


Fig. 1. (a) Main tectonic features of the study area modified from Jarosiński et al. (2006). Abbreviations are as follows: B-VM—Bruno–Vistulicum massif; E-H—Elba–Hamburg fault zone; MHF—mid-Hungarian fault zone; MM—Malopolskamassif; MZF—Mur–Žilina fault zone; RG—Rhine graben; STZ—Sorgenfrei–Tornquist fault zone; TB—Transylvanian basin; TESZ—Trans-European suture zone; TTZ—Teisseyre–Tornquist fault zone. (b) ETOPO-2 European topography (http://dss.ucar.edu/datasets/ds759.3/) averaged to 15′×15′ resolution (km). (c) Bouguer anomaly map (mGal).

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Fig. 2. Density–depth relationships for different classes of sedimentary basins based onwell-log and seismic data. Curves 1 and 4 are characteristic for basins with minimumand maximum density values of sediments at the same depth, respectively. The firstcurve is typical for marine deposits and the latter for basins in Iberia (the Ebro, theDuero and the Tajo basin) and basins in the Northwestern European foreland (e.g. theParis and Aquitan basins). Curves 2 and 3 are characteristic for basins with intermediatedensity values of sediments at the same depth. Curve 2 is typical for sediments locatedin some Central European basins, (e.g. the Molasse Basin) and Curve 3 is representativefor sediments of the Pannonian basin.

Table 1Thickness (km) and density (g/cm3) of the sub-layers of the sedimentary cover, usedfor the gravity calculations.

Layers Thickness(km)

Density (min–max)/mean(g/cm3)

1 0–0.5 1.88–2.67/2.102 0.5–1.5 1.9–2.57/2.203 1.5–3 1.9–2.69/2.354 3–5 2.15–2.67/2.485 5–8 2.39–2.66/2.566 8–16.262 2.47–2.65/2.60


based on averaged borehole and seismic data and onwell-determineddensity–compaction relations. Therefore, we consider only generallithology characteristics of each specific basin, while the small-scaleheterogeneity of the sedimentary cover is left for local studies. Forinvestigations of the crystalline crust and especially of themantle, thisapproach provides sufficient resolution. The validity and efficiency ofthis approach has been demonstrated in numerous previous studiesrelated to gravity modelling of sedimentary basins (e.g., Jachens andMoring 1990; Langenheim and Jachens, 1996; Artemjev and Kaban,1994; Kaban and Mooney, 2001; Kaban et al., 2004).

The density–depth relationships, which are used for the basinswhere more detailed data are not available, are shown in Fig. 2. Themost important changes occur in the upper 6–8 km, where the effectof compaction dominates over other factors. For other basins,lithology changes lead to a more rapid increase of density withdepth. These relationships are based on seismic and geological data,which are summarized in the Appendix A.

For the gravity calculations, the sedimentary cover has beendivided into 6 sub-layers (Table 1). Following this approach, we canwith sufficient accuracy incorporate sharp changes of density withdepth. The thickness of the layers is increased to account for thechanges of the vertical density gradient, which decreases with depth.

Estimation of the gravity effect of the sedimentary cover, as acontribution of 6 sub-layers instead of a single layer, lead to a markedimprovement of the results, with differences up to 70 mGal. This fieldranges between 0 and −280 mGal (Fig. 3a) and reflects thesedimentary thickness distribution. However, the influence of thelateral density variations is also clearly visible when we compare thegravity effect of basins with similar thicknesses of sediments. Forinstance, the contrast in density between the Ebro Basin (character-ized by a relatively high density) and the Pannonian Basin(characterized by a low density) gives a difference of ∼30 mGal intheir gravity effect.

2.4. Gravity effect of the crystalline crust and Moho

The gravity effect of the crystalline crust was computed in twosteps. We have estimated the impact of the Moho variations takingconstant density values for the crust and mantle from the referencemodel. Separately, the effect of density variations within thecrystalline crust (independently for the upper and lower crust) wasestimated relative to the initially used constant value. All together,these fields represent the anomalous effect of the heterogeneous crustrelative to the laterally homogeneous reference model.

The velocities in the crystalline crust have been converted todensities using non-linear relationships of Christensen and Mooney(1995), which depend on depth. The anomalous gravity effect of theconsolidated crust varies from−140 to 350 mGal (Fig. 3b). Significantdifferences are found between the areas West and East to the TransEuropean Zone (TESZ) and also within Western Europe. EasternEurope is mostly characterized by higher values thanWestern Europe(up to 350 mGal over the Baltic Shield). In Western Europe theanomalies mostly range between −100 mGal (e.g. in the Paris basin)to about +110 mGal in England. The continental part of WesternEurope is mainly affected by negative anomalies, while in someoceanic domains, especially those characterized by a thick mafic crust(e.g. in the Vøring basin), positive anomalies are observed (Fig. 3b).

The gravity effect of the Moho variations spans from −450 mGalto 450 mGal (Fig. 3c), with the lowest values localized in areascharacterized by deep Moho, such as, for example, in the Baltic Shieldand the East European Platform (EEP) and the highest valuescorresponding to areas of shallow Moho (oceanic or sub-oceaniccrust).

3. Mantle gravity anomalies and the residual topography

Residual mantle anomalies obtained after removing the crustaleffect from the observed gravity field are displayed in Fig. 4a. Theyreflect variations in both temperature and mantle composition andrange from approximately−240 mGal to +310 mGal. The procedureof removing the crustal effect is illustrated for a cross-section inFig. 4c. For example, the Moho uplift under the Balearic Sea causes thegravity maximum, which significantly exceeds the Bouguer maxi-mum. However, this difference is fully compensated by the effect oflow-density sediments, which results in near-zero mantle anomalies.Similarly, a negative effect of theMoho root under Southern Norway iscompensated by an increased density within the crystalline crust.

Potential errors of the residual field depend on the reliability of thecrustal model since the initial gravity field and topography aredetermined with high accuracy. This issue has been extensivelydiscussed by Kaban and Schwintzer (2001) and Kaban et al. (2003). Asthe density of sediments normally approaches crystalline crustaldensity at depths greater than 5 km, errors in the sedimentarythickness are not the main error source. The density variations withinsedimentary layers (relative to the normal density 2.7 g/cm3) areassumed to be known with an accuracy of about 15% (Kaban andMooney 2001). This might give a difference in the gravity effect up to

Fig. 3. (a–c) Gravity effect of the crustal layers (mGal). (a) Sedimentary layer. (b) Gravity Crystalline crust. (c) Moho variations.

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Fig. 4. (a–c) Residual mantle anomalies of the gravity field obtained after removal of the crustal effect (including topography and water) from the observed field (mGal). (a) Total anomaly (200bL). (b) Mid-wavelength component(200bLb2000 km) of residual mantle anomalies, displaying a close correlation with specific tectonic structures. The black line depicts location of cross-section displayed in Fig. 4c. (c) Cross-section showing the gravity effect of each crustalcomponent (sediments, crystalline crust and Moho), the Buguer anomalies, and the residual mantle anomalies (total and mid-term component).

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15 mGal. For the Moho depth, it is usually assumed that possibleerrors of a single determination might be about 5% for the recent andup to 10% for the old ones. However, for relatively wide anomalies,each anomaly is always based on many determinations, includingindependent data. In this case, an average determination error mightbe reduced to 0.5 km for well-studied areas and to 1 km for the areaswith moderate data coverage (Kaban and Mooney, 2001; Kaban andSchwintzer, 2001). This could result in an error from 10 to 20 mGal.Error standard deviations in the mean density of the crystalline crustas determined from seismic velocities could be from 0.02 g/cm3 to0.03 g/cm3 (e.g. Christensen and Mooney, 1995). Assuming athickness of 15 km for the upper and lower crustal layers, this maylead to an error of 12–18 mGal for each. Comprising all error sourcesand considering that they are independent, the estimated standarddeviations in the residual gravity anomalies amount approximately to30–50 mGal, depending on the data coverage.

The amplitudes of the most prominent anomalies significantlyexceed possible determination errors, with the signal-to-noise ratiofrom 3 to 7. A pronounced large-scale difference (∼200 mGal) isfound between Eastern andWestern Europe. Strong negative residualmantle anomalies suggest the presence of low density masses withinthe upper mantle and provide an indirect evidence for high mantletemperatures in Western Europe. The transition from positive tonegative residual gravity anomalies coincides with the TESZ, as alsoobserved in previous studies (e.g. Kaban, 2001; Yegorova andStarostenko, 2002; Tesauro et al., 2007). Distinct characteristics ofsmaller scale tectonic units obtained in the present study are resolveddue to the higher resolution of the initial crustal data.

As demonstrated in continental-wide and global studies, mantlegravity anomalies can be separated into two components dependingon wavelength (e.g. Kaban, 2001). The long-wavelength componentreflects large-scale structural heterogeneities of the lithosphere,mostly related to its thermal regime. To emphasize the differencesbetween specific tectonic units, the long-wavelength component(approx. LN2000 km) is subtracted from the total mantle gravity field.For a reliable estimation of the long-wavelength component also dataoutside the region under study are incorporated from the mostcomplete global database (Kaban et al., 2004).

The amplitude of the “regional” field (wavelengths ca.200 kmbLb2000 km) is about ±180 mGal, as shown in Fig. 4b. Theregional mantle gravity field is heterogeneously distributed inWestern Europe, showing in some areas sharp lateral changes. Forinstance, from the Tyrrhenian to the Adriatic Sea an increase of290 mGal is observed over a distance of 600 km, while from thePannonian basin to the Carpathians a similar rise (250 mGal) is foundover a distance of only 300 km (Fig. 4b). A chain of negative mantleanomalies (between −50 mGal and −150 mGal) depicts the areas ofactive neotectonics and recent back-arc extension in the WesternMediterranean (Tyrrhenian Sea and Valencia Trough–Balearic Basin),the Pannonian basin, the Massif Central and the Upper Rhine Graben(Fig. 4b). These negative anomalies are likely of thermal origin (e.g.Tesauro et al., 2009). The high temperature regime in these areas ispossibly controlled by upwelling of hot asthenosphere (e.g. in theWesternMediterranean Sea), which is confirmed by high values of thesurface heat flow and low P- and S-wave velocities (Zito et al., 2003;Koulakov et al., 2009). A very distinctive positive anomaly(N150 mGal) is located over the Dinarides, extending with loweramplitude over the Adriatic Sea (∼100 mGal) and the Ionian Sea(∼60 mGal). These positive anomalies support the notion of a coldand high density domain in the uppermost mantle, which can berelated to an increase of lithospheric thickness (e.g. beneath theAdriatic plate) or to the presence of a cold subducting slab (e.g.beneath the Dinarides). On the other hand, the Alps and the Hellenicarc, although characterized by thick lithosphere due to the subductingslab, show negligible anomalies (less than ±50 mGal), possibly onaccount of compensation of thermal and compositional effects.

Another strong positive anomaly (N150 mGal), observed over theCarpathians and extending to the Moesian Platform, might be relatedto the presence of a thick lithosphere in this area, as also imaged bytomography data (Koulakov et al., 2009). Two other smaller positiveanomalies (60–80 mGal) characterize the Aquitaine Basin, extendingover the Pyrenees and Cantabrian Mountains and over the Paris Basin,where also a thick lithosphere is expected. The positive anomalyfound over the North Sea (up to 110 mGal) is in a good correspon-dence with a low thermal anomaly (Tesauro et al., 2009).

Anomalies of small amplitude (less than ±50 mGal) characterizemost part of the Baltic Shield and the EEP (Fig. 4b). In this area thetemperature anomalies, related to the presence of a thick coldlithosphere, are possibly compensated by the compositional anoma-lies due to the iron depleted mantle, which effect will be investigatedin the next section. This compensation is not extended to the Southernpart of Norway, as revealed by the negative anomaly in this area(−60 mGal). Therefore, the high topography present there might bepartly sustained isostatically by a low density mantle, as alsohypothesized in previous studies (e.g. Ebbing, 2008). Differenceswith previous models (e.g. Kaban, 2001; Yegorova and Starostenko,1999 and 2002; Tesauro et al., 2007), are due to the updated crustalmodel used for the calculations. Our results are more robust, have ahigher resolution and show better correspondence between theanomalies and geological features. According to our study, forexample, the Tyrrhenian and Balearic Sea are characterized bynegative anomalies, with two minima located in the Southern partof the Tyrrhenian Sea and in the Valencia Trough, in agreement with ahigh thermal regime (Tesauro et al., 2009). In comparison with earlierstudies, in some European regions (e.g. Bay of Biscay and North Sea)our results show anomalies having opposite sign (positive instead ofnegative), which raises questions concerning previous interpreta-tions. In addition, the positive anomaly previously detected over thePyrenees by Tesauro et al. (2007) is in the present study extended tothe West over the Cantabrian Mountains. These findings point to thepresence of thick lithospheric roots also below this mountain chain.

In addition to the residual mantle gravity anomalies, residualtopography (Hres) is computed. Residual topography represents thatpart of the surface relief, which is under- or over compensated bydensity variations in the initial crustal model (Kaban et al., 2003):

Hres =1ρ

ρtop� �

Hobs +1ρ∫T

0Δρ zð Þ R−z

R

� �2dz ð1Þ

where ρtop is the average block density (including the effects of iceand sediments) of topography (Hobs), ρ– is the average density of theresidual topography, which is set to 2.67 g/cm3, Δρ zð Þ is the densityanomaly (including all discontinuities like ocean floor topography andMoho variations) relative to a horizontally homogeneous referencemodel with zero density above the geoid, Hobs is the topographyheight (zero for sea areas), z is the depth below geoid, R is the radius ofthe Earth. The bottom level of isostatic compensation (T) is taken hereat the base of the crust since the mantle material is homogeneous inthe initial model.

Apart from determination errors, the medium and large-scalecomponent of the residual topography arises from two sources(Kaban et al. 2003). First, continental Hres depends on the densitydistribution in the uppermost mantle: high residuals are supported bylow density lithosphere roots, while low residuals are balanced byhigh density anchors. The second source of Hres is the normal stress atthe base of the lithosphere due to mantle flow, defined as ‘dynamic’topography. Despite the formal discrimination between the twosources, the effect of large-scale upper mantle density variations,whether from below or from inside the lithosphere, on Hres is similar.Large-scale variations of Hres may not be supported by a rigidlithosphere and consequently they should be compensated by densityinhomogeneities in the upper mantle (Kaban et al., 2003). There are


some other factors, like non compensated post-glacial deformations,but their amplitude is relatively small (b100 m). Therefore, Hres mayserve as an additional parameter to improve the knowledge aboutmantle structure. The local small-scale component is mostly due tonot compensated local density anomalies in the crust (isostaticdeviations) and errors in the initial crustal model.

It is difficult to estimate the dynamic input of deep inhomogene-ities to residual topography over continents without considering acomplete global dynamic model of the Earth. Different authors giveamplitudes of the dynamic topography ranging from 0.5 km to 1 kmand 2 km (e.g. Pari and Peltier, 2000; Steinberger et al., 2001).However, an indirect guess may be obtained by comparison with themantle gravity field. Actually, the effect of dynamic deformations ofthe Earth's surface is subtracted in themantle gravity anomalies, sincethese deformations are part of the observed topography and crustalstructure. Thus, the dynamic contribution can be considered lessimportant when mantle gravity anomalies and Hres are inverselycorrelated with an appropriate scaling factor.

The residual topography with removed small-scale features(LN200 km) varies between −3.5 km and +3.2 km and is displayedin Fig. 5. The uncertainties of the calculated variations of Hres stemfrom the same error sources as in the mantle gravity. On thecontinents, Hres error is estimated to be about 0.35 km for the areaswith a well known crustal structure (e.g. Central Europe) and to be aslarge as 0.8 km for regions having a poor seismic data coverage (e.g.the oceanic domain). The residual topography is usually notcorrelated with topography (Figs. 1b and 5). The high elevated areasmight be characterized by a positive residual topography (e.g.Western Alps), negative (e.g. Carpathians) or close to zero (e.g.Eastern Alps). On the other hand, not negligible residual topographyoften appear in flat regions, as discussed below.

The comparison with the residual gravity anomalies (Fig. 4a)shows that the regions with large negative residual gravity anomaliesare characterized by large positive Hres and vice versa. Bothamplitudes vary similarly after appropriate scaling. This goodcorrespondence indicates that the dynamic contribution is smallrelative to the total variations of Hres (Kaban et al., 2003). However,some exceptions are visible around Iceland, where the strong positiveresidual topography (∼1.5 km) corresponds to a relative smallnegative mantle anomaly (−80 mGal), supposing some dynamicsupport. These differences hold the potential, to be explored in futurework, to resolve density anomalies, which are responsible for both

Fig. 5. Residual topography (km). T

residual gravity and topography, and to determine their depth (e.g.Kaban et al., 2007).

Furthermore, it appears that most parts of Eastern Europe are overcompensated, being predominantly characterized by negative anoma-lies of Hres. By contrast, Western Europe appears mostly compensatedor for some areas even under compensated. More in detail, in EasternEurope the strongest negative anomalies in residual topography arelocated over the Carpathians (about −2.5 km). These findings couldbe due to the presence of an anomalously high density upper mantleand/or dynamic support. A more interesting result is the negativeresidual topography (up to −1.5 km) observed in some parts of theArchean–Paleoproterozoic blocks of the EEP, where the topography isalmost flat, showing undulation within 200 m (Fig.1b). Therefore,without excluding the possibility of the dynamic contribution toresidual topography, the mantle lithosphere of this region might benot as strongly iron depleted as commonly assumed, but show a morecomplex density distribution (e.g. Artemieva, 2007).

In Western Europe the only large negative anomaly is observed inthe Eastern Alps and the Dinarides (−2 km), which extends to thewest over the Adriatic plate with a lower amplitude (−1.5 km). Asmaller negative anomaly (up to −1 km) is observed over the NorthSea and the Paris basin. The former is characterized by a strong Plio-Quarternary subsidence (e.g. Kooi et al., 1998), while the latter showsalmost a flat topography.

It should be noted, that although the Western Alps are character-ized by a thick lithospheric root and subducting slabs, as imaged bytomography data (e.g. Bijwaard and Spakman, 2000; Piromallo andMorelli, 2003; Koulakov et al., 2009), they appear slightly undercompensated (0.5 km). The most significant positive anomalies areobserved for North and South Iceland (N3 km), which can be partlyexplained by large-scale and deep mantle sources (e.g. Kaban et al.,2002). Other parts of Europe, such as the European Cenozoic RiftSystem (ECRIS) and the Paris basin, are mainly characterized byanomalies of smaller amplitude (up to ±1.5 km) with a reverse signcompared to the residual mantle gravity anomalies.

4. Gravity effect of temperature variations and composition in theupper mantle

As shown in the previous section, patterns of the residual mantleanomalies correspond to tectonic division of the region. It is importantto understand, which factors (temperature or composition) control

he density is set to 2.67 g/cm3.

Fig. 6. (a) Gravity effect of temperature variations in the upper mantle (mGal), based on the inversion of a tomography model. (b) Gravity anomalies (mGal) obtained after removalof the temperature induced field from the total mantle field. These residuals provide evidence for strong compositional variations in the upper mantle, which are not reflected in theseismic model (see text for discussion). Black lines depict locations of cross-sections displayed in Fig. 8.


density distribution.We use a new thermalmodel of the uppermantleby Tesauro et al. (2009) to quantify the gravity effect of temperaturevariations. By comparing this effect with the residual mantle gravityanomalies, inferences can be made on compositional density varia-tions in the mantle.

The thermal model has been derived based on new tomographydata of Koulakov et al. (2009). This tomography model has twoprincipal differences from previous ones. First, it is a-priori correctedfor the crustal effect using EuCRUST-07 model. Second, it has beenpaid special attention to an estimate of real amplitudes of velocityanomalies. It has been demonstrated that the new model containsmuch more detail than previous models covering the same region(Koulakov et al., 2009). Therefore, using this model for temperatureestimations gives more robust results (Tesauro et al., 2009).Temperatures have been estimated from P-wave velocities usingvelocity derivatives calculated taking into account the effects ofanharmonicity and anelasticity (e.g. Sobolev et al., 1997; Goes et al.,2000).

Assuming that the seismic model is well resolved and thecomposition is known, the uncertainty of the absolute temperaturesmay reach ±100 °C (Cammarano et al., 2003). However, for thepurposes of the gravity modelling only temperature variations at anyspecific depth are important. The latter are much better constrainedthan absolute temperatures. In this case the main uncertainty stemsfrom temperature derivatives of the elastic parameters. Their errorsare estimated to be between 10% and 20% (e.g. Goes et al, 2000;Cammarano et al., 2003). This may lead to an uncertainty of theinferred temperatures about ±70 °C above 300 km (Tesauro et al.,2009), which corresponds to a change of the gravity effect of about50 mGal at most.

Another important issue is the effect of compositional differences.Hyndman et al. (2009) discuss in detail this effect on temperatureestimations for various depths. The maximum temperature differenceis about 150 °C between the ‘Off-Craton’ and ‘Archean’ composition atshallow depths (b120 km). At greater depths, this difference isreduced progressively to about 65 °C, on account of the anelasticity

Fig. 7. (a–b) Total mantle gravity anomalies versus the gravity effect of temperaturevariations (mGal). (a) Mid and long-wavelength component (LN350). (b) Mid-wavelength component (350bLb2000 km).


effect. Therefore, the compositional effect is particularly important forthe Archean cratons (e.g. Hieronymus and Goes, 2010). By contrast, inthe tectonically active areas, the melting temperature is reached atshallow depths (about 100 km) and the effect of compositionaldifferences is strongly reduced. For those areas, the main uncertaintyis due to the anelasticity model used (e.g. Shapiro and Ritzwoller,2004). The area of our study mostly covers tectonically active andrelatively young structures and only a small part of the East EuropeanPlatform. The latter is not as strongly depleted as the Archean BalticShield located north of the study area, where we may assumemaximal compositional changes (e.g. Artemieva, 2009; Artemievaet al., 2006; Kaban et al., 2003). For these reasons, we assume that inour case the compositional changes are less important.

The effect of water and melt content on seismic velocities could beimportant, especially for high-temperature zones. This means that forsome deep low-velocity anomalies (e.g. in the Tyrrhenian Sea) thetemperature increase might be not as pronounced as predicted fromthe inversion. We discuss this possibility analyzing the residual

gravity field. In general there is a good agreement of the mantlegeotherms of the present thermal model and those obtained inprevious studies (e.g. Goes et al., 2000), as well as with independentdata based on the observed heat flow (Tesauro et al., 2009).

In order to eliminate small-scale artefacts, seismic velocitiesemployed in the temperature inversion have been processed by alow-pass filter retaining only the wavelengths greater than 350 km.Therefore, some local features revealed by seismic studies (e.g.receiver functions) might not be visible in the smoothed map due todifferent resolution (Tesauro et al., 2009). To be consistent, theresidual mantle field has also been filtered with the same filter toleave only the wavelengths, which are presented in the temperaturefield. These features should be sufficiently well-resolved in thetomography model, which is demonstrated by various resolutiontests (Koulakov et al., 2009). Since the gravity effect of the mantlelayers is reduced by a factor of exp(−2πZ /L) (where L is thewavelength and Z is the depth), the impact of the layers deeper than50 km to the observed field at the shorter wavelengths is negligible.

Outside the study area we used the global tomography model ofBijwaard and Spakman (2000). This is more important with respect tothe edge effects than for the crustal calculation, since the depth to theanomalous masses is much higher. The calculated temperaturevariations are converted into density applying a depth dependentthermal expansion coefficient (Tesauro et al., 2009).

The gravity anomalies induced by temperature variations rangefrom−110 mGal to 230 mGal, as shown in Fig. 6a. In contrast with theresidual mantle anomalies, they are distributed rather homogeneouslywith maximum positive values associated to the EEP, on account of itsthick cold lithospheric roots, and negative values overWestern Europe.Theminimumvalues are located over the Tyrrhenian Sea, the ECRIS andthe Pannonian Basin. Probably in the Tyrrhenian Sea even lower valuesmight be expected on account of its extremely thin lithospherepredicted from previous studies (e.g. Panza and Raikova, 2008). Thetemperature induced mantle gravity field shows only an overallsimilarity with the whole mantle gravity anomaly. Eastern Europe ismostly characterized by strong positive anomalies in both cases, incontrast with Western Europe. However, there are many pronounceddifferences even for large tectonic units within these domains, asshown in Fig. 7a–b. The relationship between the residual gravityanomalies and the temperature induced gravity variations is charac-terized by a general trend with very strong deflections from theregression line (Fig. 7a). Furthermore,whenwe compare thefields afterremoving the long-wavelength component, almost no correlation isfound (Fig. 7b). Amplitudes of the temperature-induced gravityanomalies are reduced by a factor of 2 or even more, while amplitudesof the total mantle field remain significant. At first instance it appearslogical to explain such a large difference just by a lack of resolution ofthe temperature model based on seismic tomography. However, anincrease of the amplitudes of the mid-wavelength temperature-induced gravity anomalies inevitably requires at least one of thefollowing options: either amplitudes of the tomography model shouldbe increased drastically, or the conversion factor V-to-T should beseveral times larger (in the extreme case thermal variations are notreflected in the tomography image at all). To our opinion this isunlikely. Asmentioned, corresponding features are well-resolved in thetomography model (Koulakov et al., 2009) and possible errors intemperature estimations cannot result in such large differences(Tesauro et al., 2009). Another explanation for this phenomenon istherefore required. Thermal differentiation within the lithosphere iscontinuously reducing through time (especially for mid- or small-scalefeatures). By contrast, the compositional differentiation of the uppermantle is largely controlled by tectonic evolution and remainspreserved within the lithosphere through time. This evidences thatthe density structure within the mantle under Europe largely dependson both factors (temperature and composition) and cannot beadequately described by using only seismic data.

Fig. 8. (a–c) Residual mantle anomalies of the gravity field, gravity effect of temperature variations in the upper mantle and the residual “compositional” anomalies along three cross-sections (mGal). Sections location is shown in Fig. 6b. Abbreviations are as in the text.



The gravity effect of compositional variations in the lithosphere iscalculated by subtracting “temperature” gravity anomalies from thetotal mantle field, processed like the temperature, by a low-pass filter,retaining the wavelengths greater than 350 km (Fig. 6b). Further-more, the mantle gravity field also contains the signal from depthsbelow 300 km that was the limit of the temperature calculations.However, this signal is weaker than the upper mantle signal, andamounts to about 10–15% of the total mantle anomaly (Kaban et al.,2003). This finding is consistent with results derived from jointinversion of global seismic tomography and gravity data (Ricard et al.,1993; Currieu et al., 1995; Forte and Perry, 2000).We also remove thiscomponent from the “whole mantle” anomaly using the results ofKaban and Schwintzer (2001). Although the effect of the deep mantlestrongly depends on the employed model, this uncertainty does notappear to exert a major influence on the results.

The remaining gravity anomalies predominantly reflect composi-tional variations in the mantle and span from −200 to +170 mGal.Compared to the thermal gravity field (very smooth), the composi-tional anomalies are changing more locally. The EEP is characterizedby a pronounced gravity low, which is typically within the range−150 to−200 mGal, implying corresponding compositional changes(i.e. iron depletion). In Western Europe the strongest negativeanomaly (up to −120 mGal) is observed over the Hellenic arc.Large positive compositional gravity anomalies are found over theDinarides (N120 mGal) and over a part of the North German Basin(NGB) (∼100 mGal). The latter anomaly does not exactly correspondto the tectonic division of the region. This might be partially related touncertainties in the tomography model, which is not well definedhere. The positive anomalies are located over the Caledonian province,between the Thor and the Rheic Suture. Therefore, the closure of twooceans in the Paleozoic and the consequent amalgamation of terrainsin this area (Ziegler et al., 2001) could result in a compositionaldifferentiation within the upper mantle with respect to the Sveco-Norvegian province in the North and the Variscides in the South.

The relationship of the temperature and compositional gravityanomalies is strikingly different for different tectonic units, asdisplayed in three cross-sections (Fig. 8a–c). Both components havesimilar amplitudes over the Balearic Sea and the Massif Central(Fig. 8a). In most areas characterized by a high thermal regime(Tesauro et al., 2009), such as the Pannonian basin and the TyrrhenianSea, the compositional anomalies are small, suggesting that thetemperature model satisfactorily describes the mantle densitystructure (Fig. 8b). On the other hand, the compositional anomaliesare predominant over areas characterized by subducting slabs, such asthe Adriatic plate and the Dinarides and over the Hellenic arc (Fig. 8c).However, the sign of the compositional anomalies is opposite, whichmight be attributed to a different nature of the slabs. Petrological datahave also revealed complex spatial and temporal variations ofcomposition of the volcanic rocks from the circum-Aegean region tothe central part of Anatolia (e.g. Doglioni et al., 2002). These resultslikely point to small-scale variations in the composition and physicalconditions of the subducted African plate. However, due to a relativelylarge depth, the gravity effect at the surface of these variations withinthe mantle is expected to be small and not reflected in the residualmantle fields. Therefore, the broad positive anomaly over the Hellenicarc more likely reflects the signature of the mantle wedge (lighterthan the subducted slab), instead of a strong compositional change. Bycontrast, the positive anomaly over the Dinarides and the Adriaticplate might be due to a denser subducting slab. In this context, thedifference in the anomaly sign is probably closer related to thesubduction mode than to the nature of the slabs themselves.Furthermore, it should be noted that the positive anomaly sharplydecreases south of the Adriatic Sea, in the same locations whereseismic and GPS studies (e.g. Oldow and Ferranti, 2004; Grenerczyand Kenyeres, 2004) point to the presence of the boundary betweenthe Adriatic and African plates. Although many questions remain on

the fine structure and causes of the residual mantle anomalies overthe Dinarides and Hellenic arc, the present results sustain thehypothesis that the Adriatic and African plates are independent (e.g.Mantovani et al., 2004).

5. Conclusions

In this paper we estimated the mantle field using a density model,based on a new digital crustal model (EuCRUST-07) and newcompilations of the sedimentary cover of Western and Eastern Europe.The gravity effect of the crust was estimated and subtracted from thetotal gravity field. The residual mantle anomalies, which reflectcompositional and temperature variations in the mantle, demonstratea close correlation with tectonic features. In particular, strong negativeanomalies (up to−150 mGal) areobservedover theTyrrhenianSea, theECRIS and the Pannonian basin, while positive anomalies (N150 mGal)are found over the Dinarides and the Adriatic Sea. Very small anomalies(±50 mGal) are found over the EEP, probably on account ofcompensation of the temperature (low) and compositional effects(mantle iron depletion). The positive anomaly previously observed overthe Pyrenees, is extended to the west over the Cantabrian Mountains,pointing to the presence of thick lithospheric roots beneath this area.The residual topography, estimated using the new crustal densitymodel, shows largely a negative correlation with the residual mantleanomalies. This suggests isostatic compensation in most parts of thestudy area, by additional mass anomalies located in the mantle. Thesefindings pertain to relatively large structures of about several hundredkm ormore. An exception is given by Iceland, where the relative strongresidual topography (+1.5 km) is not compensated by a relative weaknegative mantle anomaly (−80 mGal), pointing to a strong dynamicsupport of the topography in this area.

A new thermal model has been used to estimate the gravity effectinduced by temperature variations. Subtracting this field from thetotal mantle anomalies, gave the opportunity to determine compo-sitional anomalies. The anomalies related to temperature are fairlyhomogeneously distributed, showing a strong difference betweenEastern (cold) and Western (hot) Europe, with negative values(∼100 mGal) located over Neogene extensional basins (e.g. theTyrrhenian basin). By contrast, compositional anomalies demonstratea much more heterogeneous distribution. The field east of the TESZ issmoother, and characterized by negative values over the EEP (about−150 mGal), likely due to iron depletion. Both positive and negativecompositional anomalies are found over Western Europe. TheDinarides and Hellenic arc, both underlain by subducted slabs, arecharacterized by an opposite sign of the compositional anomalies withpositive values over the Adriatic plate and negative values over theAegean Sea. These findings support the existence of independentAdriatic and African plates with possibly different modes ofsubduction.

Acknowledgements

Valuable comments from the Editor Prof. Tilman Spohn, JörgEbbing and an anonymous reviewer greatly improved the originalmanuscript. Funding was kindly provided by NWO (NetherlandsOrganization for Scientific Research), SRON (Space Research Organi-zation Netherlands) and DFG (German Research Foundation) RO-2330/4-II.

Appendix A. Density structure of sediments

The sedimentary cover in the study area is very heterogeneous.Very deep basins (N10 km), located in different parts of Europe, wereformed both in compressional (e.g. the Carpathians foredeep) andextensional regimes (e.g. North Atlantic rifted continental marginsand Neogene Mediterranean back-arc basins). At the same time,


sedimentary thicknesses are minor in many other areas, such as theEEP (Tesauro et al., 2008). Data available from previous studies havebeen analysed to determine the density of the sediments in the studyarea. Thickness, seismic velocities, age and, whenever possible,densities, of sediments for the European basins are displayed inTable A1.

From this table it can be observed that the density of thesedimentary layer strongly varies with depth and from basin tobasin. Previous sedimentary models are characterized by differentresolutions, which make the construction of a homogeneous modelfor the whole region rather difficult. For instance, in recent work ofScheck-Wenderoth et al. (2007) the sedimentary cover in the Møreand Vøring basins has been divided in 6 sub-layers of different age,thickness and average P-wave velocity. By contrast, for the thicksedimentary package of the Aquitaine Basin (up to 10 km) only thesurface and bottom compressional velocities are available (Pedreiraet al., 2003), without any information on possible internal boundaries.As pointed out earlier, for the areas where detailed models are notavailable, we draw smooth density–depth relationships (Fig. 2). Thesecurves reflect a general trend of density increasing with depth inaccordance with compaction laws in the upper part and lithologychanges in the lower part of the section (Kaban andMooney, 2001). Atthe same time, these relationships are consistent with averagedborehole and seismic data (Table A1). When only seismic data areavailable, velocities are converted to densities using the relationship

Table 2Age, weighted average density (g/cm3) and maximum thickness (km) of sediments fordifferent European geological settings. The values of thickness refer to the part of thegeological structures displayed in the map.

Geological feature Age Meandensity(g/cm3)

aMaxthickness(km)

Central European BasinSystem

Paleozoic–Tertiary 2.40–2.45 10

Møre and Vøring Margin Pre-Cretaceous–Quaternary 2.45–2.50 15Lofoten Margin Pre-Cretaceous–Quaternary 2.40–2.45 8Upper Rhine Graben Late Paleozoic–Quaternary 2.45 4Cantabrian–Pyrenees Mesozoic–Tertiary 2.50–2.55 5Aquitaine Basin Mesozoic–Tertiary 2.45–2.50 10Duero Basin Mesozoic–Tertiary 2.50–2.55 6Ebro and Tajo Basin Mesozoic–Tertiary 2.50–2.55 6Balearic Sea Tertiary–Quaternary 2.35 5Gulf of Cadiz Tertiary–Quaternary 2.30–2.35 4Bay of Biscay Mesozoic–Quaternary 2.25–2.35 2Iberian Abyssal Plain Tertiary–Quaternary 2.10–2.20 2Provençal–Corsica margin Tertiary–Quaternary 2.30 5Gulf of Lyon Tertiary–Quaternary 2.20–2.25 2Tyrrhenian Sea Tertiary–Quaternary 2.25–2.30 5Ionian Sea Mesozoic–Quaternary 2.30–2.35 6Eastern Mediterranean Sea Tertiary–Quaternary 2.35 8Aegean Sea Tertiary–Quaternary 2.05–2.15 2Po Plain Tertiary–Quaternary 2.45–2.50 15Molasse Basin Tertiary–Quaternary 2.45 6Adriatic Sea Mesozoic–Quaternary 2.25–2.30 4Ligurian Sea Tertiary–Quaternary 2.30–2.35 5South Rockall Basin Late Paleozoic–Quaternary 2.30–2.35 5Lousy and Rosemery Bank Mesozoic–Quaternary 2.25–2.30 2North Rockall Basin Mesozoic–Quaternary 2.30–2.35 6Iceland–Faeroe Ridge Mesozoic–Quaternary 2.20–2.25 2Porcupine Basin Mesozoic–Quaternary 2.45–2.50 12Edoras Bank Mesozoic–Quaternary 2.20–2.25 2Hatton Basin Mesozoic–Quaternary 2.30 4Focşani Basin Late Mesozoic–Quaternary 2.45–2.50 16Transylvania Basin Paleozoic–Quaternary 2.25 5Foredeep Carpathians Tertiary–Quaternary 2.45–2.50 7Pannonian Basin Mesozoic–Quaternary 2.35–2.40 5Black Sea Mesozoic–Quaternary 2.45–2.50 12Dnieper–Donets Rift Paleozoic–Quaternary 2.40–2.50 10Russian Rift Paleozoic–Quaternary 2.40–2.45 3

a The values of thickness are referred to the part of the geological features displayedin Fig. 1a.

of Nafe and Drake, (1961). Features like density inversion with depthand abrupt lateral changes within a basin are not taken into accountand left for future local studies.

The resulting distribution of the weighted average density of thesedimentary layer for the study area is summarised in Table 2. Recentto Miocene age sediments of the Tyrrhenian, Adriatic and Balearic Seahave a low density value (2.2–2.3 g/cm3), even if they reach athickness over 7 km. By contrast, along the North Atlantic riftedmargins relatively high densities (up to 2.5 g/cm3) are observed in thedeep basins (e.g. the Porcupine and the Vøring basin). Continentalsediments have a low average density (b2.1 g/cm3) only when theirthickness is low (b2 km). Increased density values (N2.4 g/cm3) aretypical for pre-Tertiary sediments, filling 4–6 km deep (and evenmore) basins. These basins are found in a continental (e.g. the PolishThrough and the Carpathian foredeep), as well as in a marine (e.g.Black Sea basin) setting. Young basins (e.g. the Pannonian Basin) areusually characterized by relatively soft sediments (∼2.35 g/cm3). Oneexception is given by the Iberian basins (the Ebro, the Duero and theTajo basin), where, despite their young age (Tertiary) and theirmoderate thickness (∼5 km), a very high density is estimated (N2.5 g/cm3) from seismic data (Vidal, 1998). These basins underwent amajorphase of Miocene till recent uplift, leading to the erosion of asignificant part of their initial basin fill (Gaspar Escribano et al., 2001).The present model has much higher resolution than previous models(e.g. Yegorova and Starostenko, 2002), showing spatial densityvariations between tectonic units, and is more robust on account ofthe recent data employed.

Appendix B. Supplementary data

Supplementary data associated with this article can be found, inthe online version, at doi:10.1016/j.epsl.2010.04.041.

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