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Lithosphere structure of European sedimentary basins from regional three-dimensional gravity modelling
Tectonophysics 346 (2002) 5 – 21 www.elsevier.com/locate/tecto
Lithosphere structure of European sedimentary basins from regional three-dimensional gravity modelling T.P. Yegorova*, V.I. Starostenko Institute of Geophysics, National Academy of Sciences of Ukraine, 32 Pr. Palladina, 252680 Kiev, Ukraine Received 2 October 2001
Abstract A three-dimensional (3D) density model, approximated by two regional layers—the sedimentary cover and the crystalline crust (offshore, a sea-water layer was added), has been constructed in 1 averaging for the whole European continent. The crustal model is based on simplified velocity model represented by structure maps for main seismic horizons—the ‘‘seismic’’ basement and the Moho boundary. Laterally varying average density is assumed inside the model layers. Residual gravity anomalies, obtained by subtraction of the crustal gravity effect from the observed field, characterize the density heterogeneities in the upper mantle. Mantle anomalies are shown to correlate with the upper mantle velocity inhomogeneities revealed from seismic tomography data and geothermal data. Considering the type of mantle anomaly, specific features of the evolution and type of isostatic compensation, the sedimentary basins in Europe may be related into some groups: deep sedimentary basins located in the East European Platform and its northern and eastern margins (Peri-Caspian, Dnieper – Donets, Barents Sea Basins, Fore – Ural Trough) with no significant mantle anomalies; basins located on the activated thin crust of Variscan Western Europe and Mediterranean area with negative mantle anomalies of 150 to 200 10 5 ms 2 amplitude and the basins associated with suture zones at the western and southern margins of the East European Platform (Polish Trough, South Caspian Basin) characterized by positive mantle anomalies of 50 – 150 10 5 ms 2 magnitude. An analysis of the main features of the lithosphere structure of the basins in Europe and type of the compensation has been carried out. D 2002 Elsevier Science B.V. All rights reserved. Keywords: Sedimentary basins; Europe; Earth’s crust; Upper mantle; Gravity modelling
1. Introduction Gravity modelling of a three-dimensional (3D) structure of the Earth’s crust and upper mantle is a powerful tool for studying the lithosphere in its whole volume. It is linked with the seismic method, whose data serve as a basis for constructing the density
*
Corresponding author. Fax: +380-44-450-2520. E-mail address: [email protected] (T.P. Yegorova).
models. So far, the most used method in gravity modelling is the construction of a two-dimensional (2D) gravity model on the DSS profiles. The velocity models using the correlation functions between velocity and density are converted into the density equivalents, from which the gravity effect is calculated. Using 2D modelling, velocity –density models have been obtained on most of the DSS profiles in Europe. Very often, it turns out that these models, constructed even for the same profiles, have strong discrepancies. This disagreement is caused by differences in the
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methodologies used by different authors in interpretation of data of both seismic and gravity methods. Additional factor that should be considered, especially in gravity modelling, is regarding the 3D character of the model. As a rule, main features of the observed gravity field are usually explained by distribution of density heterogeneities in the crust, generally without taking into account heterogeneous upper mantle. The 3D gravity modelling with using the unique methodology, allows to minimize the shortages of 2D modelling and to receive comparable (on a uniform ground) models of different tectonic units of the crust. An important result that might be obtained in the course of 3D modelling, is the information about density heterogeneities of the uppermost mantle received from residual gravity anomalies (obtained by subtraction of calculated crustal effect from the observed field). It is known that over young tectonic regions, such mantle density heterogeneities create gravity
anomalies reaching some hundreds 10 5 ms 2 (Artemjev et al., 1994; Yegorova et al., 1995, 1997a). Usually, such mantle gravity anomalies are not manifested in the observed field, being compensated by crustal heterogeneities. Our regional 3D gravity model for the lithosphere of Europe and part of the North Atlantic (Fig. 1) is based on generalized velocity model (Giese and Pavlenkova, 1988) and consists of two regional layers of variable thickness with laterally varying average velocity/density—the sedimentary cover and the crystalline crust. Recent publications discuss the results of 3D large-scale gravity modelling for two most interesting portions of this continent-wide model: for the Alpine – Mediterranean region (Yegorova et al., 1997b) and for the transition zone from Variscan Western Europe to East European Platform (Yegorova and Starostenko, in press). For these models, a 3D gravity influence has been calculated. Residual grav-
Fig. 1. Tectonic setting of Europe. The picture shows the region of the present large-scale 3D gravity modelling. The areas with crosses indicate Precambrian Shields on the East European Platform. The areas of Caledonides, Variscides and Alpides in Europe are shown by gray dark, medium and light shadings. A solid line indicates the position of the section of the 3D gravity model (Fig. 8).
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ity anomalies, obtained by removal of the calculated effect from the observed field, are generally caused by density heterogeneities of the uppermost mantle (Yegorova et al., 1997a). Here we present the 3D large-scale gravity model (in 1 averaging) for the lithosphere of Europe and part of the North Atlantic with the emphasis on the sedimentary basins. Although the velocity structure and geometry of most of these basins are known, new additional information might be obtained by considering them all together on the base of 3D regional model. It allows to study the type and depth of isostatic compensation, inner crustal structure and to reveal heterogeneities in the upper mantle. It has a special value since most of sedimentary basins are not manifested in the averaged observed gravity field by intensive anomalies, being
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compensated by the heterogeneities of the crust and upper mantle.
2. Initial data 2.1. Gravity data The initial observed gravity field that we used, averaged on 1 1 grid, was based on the Bouguer anomaly (BA) and free-air anomaly (FAA) data obtained by ZNIIIGAiK (the Central Research Institute of Geodesy, Airsurvey and Cartography of the former Soviet Union). We have used a combined reduction joining the BA on land and FAA offshore (on the coastal line, values of BA and FAA became
Fig. 2. Generalized map of the observed gravity anomaly field for Europe and North Atlantic (Yegorova et al., 1995). The map uses Bouguer anomalies on land and free-air anomalies offshore, averaged over 1 1 grid. Data courtesy of the Central Research Institute of Geodesy, Airsurvey and Cartography of the Former Soviet Union (ZNIIGAiK). Contour interval 10 10 5 ms 2. A solid line indicates the location of the section of the 3D gravity model. AB: Alboran Basin; AL: Alps; BB: Bay of Biscay; BS: Black Sea; BSH: Baltic Shield; BSB: Barents Sea Basin; CA: Calabrian Arc; CC: Caucasus; DDB: Dnieper – Donets Basin; FUT: Fore – Ural Trough; PGB: Polish – German Basin; NSB: North Sea Basin; PB: Pannonian Basin; PT: Polish Trough; PY: Pyrenees; PCB: Peri-Caspian Basin; SCB: South Caspian Basin; TS: Tyrrhenian Sea, TTZ: Teisseyre – Tornquist Zone; UR: Urals, USH: Ukrainian Shield. A solid line indicates the position of the section of the 3D gravity model (Fig. 8).
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Fig. 3. The 3D density model for the crust of Europe and North Atlantic. Data averaging is over 1 1 grid. The model consists of a sea-water layer with q = 1.03 103 kg m 3 and two regional layers with laterally varying average density: a sedimentary cover and a crystalline crust. (a) Sea-bottom topography, km; (b) depth to the bottom of the sedimentary layer (Giese and Pavlenkova, 1988), km; (c) depth to the bottom of the crystalline crust (the Moho), km, according to Giese and Pavlenkova (1988). A solid line indicates the position of the section of the 3D gravity model in Fig. 8. For the abbreviation of tectonic units, see Fig. 2.
equal). Accordingly, offshore, a layer of sea water was included in the model. A map of the observed gravity field using such a combined reduction for the whole European continent and the North Atlantic was published by Yegorova et al. (1995). This map is demonstrated also in Fig. 2. Quiet and non-anomalous gravity field with BA values fluctuating around zero ( ± 0 20 10 5 ms 2), observed over Western European and East-European Platforms, means that in general, the gravity field here is compensated. A specific feature of the observed field gob of Europe is the gravity low over Western Scandinavia where the lowest BA values in the north of the continent, reaching 80 10 5 ms 2, are found. The orogens of the Alps, Caucasus and Carpathians are also marked by great distinctive gravity lows. Generally anomalous gravity field pattern is seen over the Alpine – Mediterranean region and North Atlantic. As a rule, deep sedimentary basins are not manifested in the observed field by strong gravity anomalies. This
means that in the whole, the sedimentary basins are supposed to be compensated units, and the characteristic type of compensation depends on the structure and evolution of the lithosphere. 2.2. Structure of the crustal model according to seismic data Our gravity model is based on the information about the crustal structure of Europe (thickness of the sedimentary cover, crustal thickness, average crustal seismic velocity and characteristic types of the crust) generalized by Giese and Pavlenkova (1988). These data have also been used in the compilation of the Geothermal Atlas of Europe (Hurtig et al., 1992). In the Alpine – Mediterranean region, we have taken into account recent seismic data on Moho topography obtained by investigations of the EGT profile (Blundell et al., 1992). Thus, the crustal model adopted for Europe and the northern part of Northern Atlantic
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comprises two regional layers of variable thickness— the sedimentary cover and the crystalline crust. It assumes lateral variation of average density inside the layers. The model layers are bounded by reliable seismic boundaries of first order. They are (1) the depth to the ‘‘seismic basement’’ (Hsb), that is, the surface of the crustal layer specified by the compressional wave velocity Vp = 5.8 –6.2 km s 1 and (2) the total crustal thickness or depth to the Moho discontinuity. Offshore, a sea-water layer was added in the model. Averaged on 1 1 grid in a plane, these structure maps are shown in Fig. 3a – c. On the map of depth to the basement (Fig. 3b), one can clearly see sedimentary basins of different depths. Very deep basins are located in the southern and southeastern part of the East European Platform (EEP)—Dnieper – Donets Basin and Peri-Caspian Basin of more than 18 km depth—and in the vicinity of its northern (Barents Sea Basin) and southern margins (Black Sea and South Caspian Basins of more than 10 and 16 km depth, respectively). The shapes of these basins are regular with the predominance of round and oval basins (Peri-Caspian, South Caspian, Barents Sea Basins) as well as linear rift basins (Dnieper – Donets Basin). From the east, the EEP is confined by meridional Fore – Ural Trough with the thickness of sediments of more than 4 – 6 km. From NW and W, the Baltic Shield and the EEP are bounded by a series of sedimentary troughs associated with the NW trending Permo-Mezosoic European Basin (Ziegler, 1990). This series of basins begins with the North Sea (thickness > 8 km), followed by the Polish – German Basin and Polish Trough of more than 4 –6 km thickness. Nearly the same depth (4 –6 km) is typical for the Western Mediterranean (Alboran and Tyrrhenian Seas) and Pannonian Basins. The map of Moho depth in Europe (Fig. 3c) reveals the division of the continental lithosphere into two major blocks: thick (40 –44 km) consolidated crust of the East European Platform and thin activated Phanerozoic crust of West Europe. From the analysis of both maps for the basement (Fig. 3b) and for the Moho (Fig. 3c), one can see that very deep sedimentary basins in Europe are located over thick crustal block of the East European Platform. Below these basins, the thickness of the crystalline crust becomes thinner because of the basement subsidence and the Moho shallowing to a depth of 36– 38 km. However,
the Polish Trough, where the Moho boundary deepens down to a depth of 50 km, is assumed to be an exception from this regularity. Beneath the basins in Variscan Western Europe and Alpine –Mediterranean belt, the Moho shallowing is estimated in the range of 16 –28 km.
3. Density parameterization of the model The two regional layers (the sedimentary layer and the crystalline crust) in our continent-wide 3D crustal model were assigned laterally varying average density values. Offshore, the model was supplemented by a sea-water layer with density of q = 1.04 103 kg m 3. Since the initial data on the velocity structure of the model (Giese and Pavlenkova, 1988; Hurtig et al., 1992) have no information on seismic velocity distribution in the sedimentary cover, we paid particular attention to data acquisition on the density (q) and velocity (Vp) distributions in this layer. The lateral variation of average density q of the sedimentary cover was estimated by a number of approaches. The main one deals with estimations of mean interval density according to laboratory data measurements. For that purpose, densities of dry sediments of the East European Platform (Podoba and Ozerskaya, 1975) were recalculated into the average densities, taking into account 100% water saturation. We adopted the increase of average density with depth q(h) obtained for the sediments on the EEP as a base function for evaluating the densities of sediments of Western Europe. According to Fig. 4, the average density of sediments q on the EEP (slanting shading) gradually increases with depth according to exponential function. It varies from 2.15 –2.25 103 kg m 3 on the depositional surface to 2.6 103 kg m 3 at the depth level 4– 5 km. At greater depths, the densities approach values typical for the crystalline rocks. For density parameterization of the sediments in the regions of Urals, Caspian Sea and South Caspian Basin, we have used laboratory data measurements from Dortman (1976). The second approach consists of P-wave velocity conversion into its density equivalent by using appropriate correlation density/velocity (q/V) functions. For the Black Sea Basin, density values for the sediments have been derived from P-wave velocity data using the
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Fig. 4. Dependence of sediment density with depth. (1) Distribution of average density with depth for sediments of the East European Platform (Podoba and Ozerskaya, 1975) adopted as a basic distribution for the gravity modelling; (2) the function used by Granser (1987) for the Pannonian Basin; (3 – 5) areas of density variations obtained by Kinsman (1975) for limestones (3), clays (4) and sandstones (5); (6) density distribution for oceanic sediments according to Rusakov (1990).
correlation function q =1.8822 0.0871V + 0.1104V 2 0.01132V 3 (Balavadze et al., 1975). For the Western Mediterranean, where sediments are distinguished as ‘‘uncompacted’’(V = 2.0 km s 1), ‘‘normal’’ or ‘‘semicompacted’’(V = 3.0 km s 1 ) and ‘‘compacted’’(V = 4.5 km s 1) (Moskalenko, 1975), we have estimated the average P-wave velocity for the whole sedimentary-cover thickness. These data were recalculated into density values using the mentioned above q/ V function. Assuming that sediment compaction with depth is similar in the Atlantic and Pacific Oceans (Rusakov, 1990), a generalized density/depth function for young oceanic sediments (Rusakov, 1990), shown in Fig. 4 by a solid line, was used to estimate q of sediments in the North Atlantic and North Sea. By this approach, only approximate estimates of average density of the sedimentary layer can be made. Nevertheless, its expediency is substantiated by the evaluation of the sensitivity of the method to changes in density. For example, a density decrease of 0.1 103 kg m 3 in
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the density of North Sea sediments of 6 km thickness will create a gravity effect of 25 10 5 ms 2 less than in the model ( 250 10 5 ms 2 instead of 225 10 5 ms 2) that should be compared with the residual gravity anomaly of 200 10 5 ms 2 magnitude. The resulting distribution of average density of the sedimentary cover for the study region is shown in Fig. 5a. The estimates of average density in the crystalline crust (Fig. 5b) were derived by using the conversion function of the type q = 2.7 + 0.27 (Vp 6.0) (Pavlenkova and Romanyuk, 1991) for q units of 103 kg m 3 when Vp is in km s 1. The average crustal Pwave velocities were taken from Giese and Pavlenkova (1988). The crystalline crust of very high density reaching 2.90 103 kg m 3 underlies extremely deep sedimentary basins over the EEP—the Dnieper – Donets and Peri-Caspian Basins (Fig. 5b). The same high densities are also typical for the whole crystalline crust of the Urals. The crystalline crust of the Variscan West Europe is characterized by lower average densities than that of EEP; however, under the Western Mediterranean and Pannonian Basins, a high-density crystalline crust (up to 2.83 103 kg m 3) is evident. The same high-density domain has been revealed in the crystalline crust beneath the Black Sea.
4. Modelling methodology and main results For the purposes of our regional gravity analysis, a two-part calculation scheme was developed. It consists of two main elements: a database and a computer program. A database contains all information about the model structure and parameterization (in 1 averaging), results of the computations for every stage of the modelling as well as the resulting maps. The kernel of the computer program block is solution of the direct gravity problem from a 3D mass distribution taking into the account the spherical configuration of the Earth (Starostenko and Manukyan, 1987). The gravity effect of the model, approximated by a set of 1 1 parallelepipeds, has been calculated by this program. For gravity effect calculations, we have used anomalous densities of the layers obtained by subtracting the upper mantle density (taken to be 3.3 g cm 3 ) from the average density of the layer: Dq = q 3.3 103 kg m 3. As a result of the mod-
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Fig. 6. Gravity inf luence of the 3D density model, obtained by summing up the effects of all three layers: the sea-water layer, the sedimentary cover and the crystalline crust. Isoline interval 50 10 5 ms 2. A solid line indicates the position of the section of the 3D gravity model in Fig. 8. For the abbreviation of tectonic units, see Fig. 2.
elling, the maps of gravity influence have been obtained for each the model layer (sea water, sediments, consolidated crust). Total gravity effect gcr, obtained by summing up the effects of the layers, is shown in Fig. 6. Since the range of anomalous density of the model layers varies from 2.27 103 kg m 3 (sea-water layer) up to 0.47 103 kg m 3 (crystalline crust), the values of the calculated gravity effect reach large negative values (up to 900 10 5 ms 2 over the Alps; see Fig. 6). Therefore, it is necessary to choose a certain background level (a ‘‘norm’’ gnor) to normalize calculated gravity effect gcr. Proceeding from our experience of 3D gravity modelling (Yegorova et al., 1995; 1997a,b), we have used the average value of gravity influence gcr for the
crust of the EEP as that value gnor. The EEP is the stable Precambrian core of the European continent, which is characterized by an observed gravity field with Bouguer anomalies up to ± 20 10 5 ms 2. For the 3D crustal model considered and the q/V function, the average value of gcr for the EEP ( gnor) is equal to 790 10 5 ms 2. The distribution of residual anomalies Dgr, obtained by subtracting the calculated effect of the model gcr (Fig. 6), normalized by constant gnor, from the observed gravity field (Fig. 2): Dgr = gob ( gcr gnor), is demonstrated in Fig. 7. After the normalizing procedure, the residual anomalies, shown in Fig. 7, are in the range + 200 to 200 10 5 ms 2. Since our model has been constructed on the base of a simplified velocity model for the crust of the
Fig. 5. Distribution of the average density in the model layers, 103 kg m 3. (a) Density in the sedimentary cover obtained with the dependence in Fig. 4; (b) in the crystalline crust, average densities were calculated from average velocities by using the conversion function q = 2.7 + 0.27 (Vp 6.0). A solid line shows the location of a section of the 3D density model (Fig. 8). For the abbreviation of tectonic units, see Fig. 2.
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Fig. 7. Residual gravity anomalies Dgr obtained by subtracting the modelled gravity effect gcr (Fig. 6), normalized by constant gnor = 790 10 5 ms 2, from the observed gravity field gob in Fig. 2: Dgr = gob ( gcr gnor) (isoline interval 50 10 5 ms 2). A solid line indicates the position of the section of the 3D gravity model in Fig. 8. For the abbreviation of tectonic units, see Fig. 2.
European continent, we are obliged to make accuracy evaluations of our modelling by choosing an optimal isoline interval for maps of gravity effects from the model layers. The sedimentary cover, with an average thickness of about 2 km (contour interval is 1 km) and with 0.05 103 kg m 3 isoline interval of average density, should be presented by a gravity effect map with a contour interval not less than 50 10 5 ms 2 . The crystalline crust of the study area, described by an average thickness of 35 –40 km (with 4 km isoline interval) and by 0.1 km s 1 isoline interval of average velocity and correspondingly by that of 0.02 103 kg m 3 of average density, is featured by a gravity effect map no less than 35 10 5 ms 2. Therefore, we have defined the main isoline interval for our gravity maps as 50 10 5 ms 2 which is in correspondence with a rms of about 20 – 25 10 5 ms 2. If the model presented reflects the real structure of the Earth’s crust, the residual anomalies should be caused by density heterogeneities in the uppermost
mantle. However, since structure of the crust is investigated incompletely because of the rare and irregular set of the DSS profiles in Europe, it is quite probable the existence of density heterogeneities in the crust (especially in the lower crust poorly investigated by seismic methods), not taken into account in our initial model. The gravity influence of the ‘‘unknown’’ heterogeneities of the upper and middle crust, not exceeding 50 10 5 ms 2, has been filtered using a 50 10 5 ms 2 contour interval in the gravity maps. We therefore presume that large residual anomalies are caused mainly by heterogeneities in the uppermost mantle, though heterogeneities of the lower crust can also contribute to the residual gravity anomalies. It is necessary to consider the origin of the residual anomalies for each specific geologic unit on the basis of comparison of configuration and magnitude of the anomalies with contours of geological units, by invoking the data from other geophysical methods, first of all, the results of seismology and geothermal data.
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The general decrease of average residual anoma-lies from 0 50 10 5 ms 2 on the east to 100 150 10 5 ms 2 on the west can be explained by a decrease in the upper mantle density below the continent. The division of the continent in two main blocks of different structure and activity of the crust and upper mantle—activated Phanerozoic Western Europe and stable Precambrian East European Platform—has been revealed from numerous geophysical investigations: seismic data (Giese et al., 1976; Fuchs, 1979; Guterch et al., 1984; Blundell, 1990), seismotomographic investigations with using P- and S-waves (Panza et al., 1980; Babuska et al., 1984; Spakman et al., 1993; Zielhuis and Nolet, 1994; Marquering and Snieder, 1996) and geothermal data (Cermak, 1982, 1993). In the area of the Teisseyre –Tornquist Zone (TTZ), a positive anomaly of 50 100 10 5 ms 2 has been revealed against the background of the gradient zone of residual anomalies. The results on this part of regional 3D gravity model in the transition zone from Western Europe to East European Platform are presented by Yegorova and Starostenko (in press). Against the regional background of general decrease of the residual gravity field from east to the west, positive and negative anomalies have been distinguished over specific tectonic units. The negative Dgr anomalies with maximal amplitude (more than 250 10 5 ms 2) are found in the North Atlantic. Negative mantle anomalies of 150 200 10 5 ms 2 are typical for the Western Mediterranean Basin, Pannonian Basin, Polish –German Basin, North Sea and the Iberian Peninsula. Local mantle gravity highs are found mainly over the Alpine orogeny. Alps and Adriatic plate/Dinarides are marked by relative Dgr highs with more than + 50 10 5 ms 2 amplitude. Local highs are also observed above the Calabrian Arc and the Pyrenees. The strongest mantle anomaly with more than 150 10 5 ms 2 amplitude is found in the Caucasus area. The regional character of this anomaly is confirmed by its extension over the Kura and South Caspian Basins (Yegorova et al., 1995). The orogens are considered as structures formed as the result of the collision of the continental plates and microplates, which led to the conservation of a rigid, cool, highvelocity/density lithospheric block in the upper mantle. Interpretation of the 3D gravity modelling results for the Alpine –Mediterranean area has been done earlier by Yegorova et al. (1997b).
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5. Imaging the sedimentary basins by 3d largescale gravity analysis The main results of the 3D gravity analysis showing the contribution of each of the layer into the model are demonstrated in the section (Fig. 8) crossing the 3D model along the Tyrrhenian Sea-Pannonian BasinTTZ-Dnieper – Donets Basin-Peri-Caspian Basin. Sedimentary basins in Europe regarding their manifestation in the calculated gravity fields, type of isostatic equilibrium state and geotectonic evolution might be related to three major groups. (1) Deep sedimentary basins located in the East European Platform and in its northern and eastern margins—the Peri-Caspian, Dnieper – Donets, Barents Sea Basins and Fore –Urals Trough. These basins are characterized by transitional type (from continental to oceanic) of the crust (Belousov and Pavlenkova, 1984, 1989). According to our gravity modelling, mantle gravity anomalies over these basins are absent or slightly positive (reaching 50 10 5 ms 2 amplitude). These deep basins have originated on the thick continental crust of the EEP. Rapid subsidence of the crust and sediment accumulation occurred in the basins during Late Devonian – Carboniferous. The common characteristics of these basins are the great thickness of the sedimentary cover, the uplift of the Moho boundary (up to the 36-km depth) below the most subsided part of the basins, strong increase of average velocity/ density in the consolidated crust and normal values of the terrestrial heat flow. The increase of average velocity in the crystalline crust is the result of the reduction of thickness of ‘‘granitic’’ layer (with Vp velocity up to 6.5 km s 1) or its total elimination. It is much more probable that the destruction of the continental crust is affected by the action of hot melted upper mantle material on the crust. As a driving process resulting in a rapid increase of velocity (density) in the crystalline crust below the basins, an ‘‘oceanisation’’ or ‘‘basification’’ of the crust—intrusion of mantle magmas of basic and ultrabasic composition into the crust is suggested by Belousov (1990). As a result of density increase of the crust due to ‘‘basification’’ or metamorphism, the crustal blocks sink into the upper mantle; they may experience eclogitization (Belousov, 1990; Artyushkov, 1992, 1993). These authors believe that the destruction of continental crust occurs in situ, without extension
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Fig. 8. A density section of the 3D density model of the European crust. A section crosses the continent through Tyrrhenian Sea, Pannonian Basin, Dnieper – Donets Basin and Peri-Caspian Basin (for location, see Figs. 2 and 3, 5 – 7). The numbers in the crustal section are density values (in 103 kg m 3) derived from the dependence in Fig. 4 (for sediments) and from the conversion function q = 2.7 + 0.27 (Vp 6.0) (for the crystalline crust), where Vp was taken from Giese and Pavlenkova (1988). The gravity effects (middle part of the figure) were derived for each the model layer (sea water, sediments, crystalline crust). Upper part of the figure characterizes the thermal regime of the region: terrestrial heat f low density (simplified version of Hurtig et al., 1992) and calculated temperatures at 40 km depth (Cermak, 1993). The crustal section was complemented by the thickness of the lithosphere taken from Panza et al. (1980) and Babuska and Plomerova (1992), numbers in the brackets are the shear-wave velocity (in km s 1) at 80 km depth according to seismotomography study of Marquering and Snieder (1996).
being essential. As it is shown by Artyushkov and Beer (1983), to make the crust subside for several kilometers, it should be extended to three or more times its size. Belousov (1990) considers that there is just no
room for that on the continents, where basins are surrounded by stable structures. Now, there are new data available on a particular role of extension in formation of deep basins on the
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EEP. These data have been obtained in the framework of EUROPROBE project ‘‘GeoRift’’ focused on the well-documented Dnieper– Donets Basin (DDB), the largest and deepest Devonian – Carboniferous rift in Europe (Stephenson et al., 1996). It was shown by Nikishin et al. (1996) that origin and evolution of the sedimentary basins on the EEP was governed by repeatedly changing regional stress field. Intraplate magmatism was controlled by changes in stress field and by mantle hot-spot activity. Based on geochemical study of magmatic rocks of the DDB, combined with estimations of the amount of extension (b = 1.1– 1.3), Wilson and Lyashkevich (1996) suggest that magmatism was trigged by the upwelling of a thermally and geochemically anomalous mantle plume from the deep mantle. They concluded that magmatism, rifting and domal basement uplift were contemporaneous at several localities within the EEP, suggesting that the thermal and geodynamic evolution of the platform could have been influenced by a cluster of mantle plumes during Late Devonian. The mass compensation of these very deep basins (up to 20 km) is reached on the depth level of he bottom of the crust (the Moho), generally due to rapid increase of average density (up to 2.9 103 kg m 3) in the crystalline crust (see Fig. 5b) and a slight Moho uplift (up to 36 km). Fig. 8 shows that the gravity effect of the sediments of Peri-Caspian Basin is practically compensated by the effect of the highdensity thin crystalline crust. Slight residual anomaly (up to 50 10 5 ms 2) may be indicative of the higher densities of the sedimentary cover (especially in its lower part) than those taken in the model (2.52 103 kg m 3), though an additional consolidation of the subcrustal lithosphere below the basin also can take place as a result of phase transition in the lower crust –upper mantle. (2) Basins of active rift and taphrogenic regime mainly in Western Europe, manifested by negative mantle gravity anomalies of 150 200 10 5 ms 2 amplitude. As a rule, these basins (Pannonian Basin, Tyrrhenian and Alboran Basins, North Sea, Rhine Graben and Polish – German Basin) are characterized by crust of transitional and oceanic types with thinner (15 – 28 km) crust and lower average velocities/densities in the consolidated crust in comparison with the EEP. The crust below the Western Mediterranean Basin is of oceanic type due to its considerable
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thinning. The lithosphere beneath these basins, as it is seen from Fig. 8, is characterized by low-velocity heterogeneities in the upper mantle (Spakman et al., 1993; Marquering and Snieder, 1996) and high-temperature regime caused by the upwelling of the heated substance of the asthenosphere up to the depth from 30 (Western Mediterranean Basin) to 50 – 60 km (North Sea, Rhine Graben, Pannonian Basin) (Panza et al, 1980; Babuska and Plomerova, 1992). It is generally accepted that the subsidence in these young basins has happened in the result of stretching the lithosphere, during which the continental crust becomes thinner and finally ruptures [shown by Sclater and Christie (1980), Wood and Barton (1983) and Ziegler (1983) on the example of the North Sea Basin]. It is considered also that a fast extension, which has taken place during some epochs, originated an initial pulse for basin subsidence (McKenzie, 1978). Subsequent slow cooling of the crust and upper mantle caused an additional subsidence. As factors leading to the subsidence of the crust, there are also considered the density increase of the lower crust and its fracture in the result of: phase transitions of gabbro – eclogite, processes of ‘‘basification’’, subcrustal erosion, thermoelastic compression, etc. Negative mantle anomalies, revealed over these basins, are caused by thermal density decrease of the asthenosphere. If the density decrease due to mantle heating is assumed to occur below the Moho to a depth of 200 km, the anomalies with 150 to 200 10 5 ms 2 amplitude [as it has been shown on the example of Western Mediterranean and Pannonian Basins by Yegorova et al. (1997b)] are caused by a density decrease of 0.045 103 kg m 3. This corresponds with the calculated thermal density decrease values of 0.01 103 kg m 3 for each 100 C. The maximal depth of this layer is evaluated as 200 km. The choice of the lower depth level is based on the results of seismotomographic study (Spakman et al., 1993), showing that the strongest anomalies are located above this level. The compensation of these basins is realized as follows. Mass deficit of sediments is overcompensated by strong Moho uplift (Fig. 8), the full mass compensation is reached on the depth level in the upper mantle, more probably related to the bottom of the asthenosphere, due to thermal expansion of the asthenosphere substance.
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(3) Basins associated with the suture zones and located at the western and southern margins of the East European Platform. In the transition zone from Western Europe to the EEP, marked by the Polish Trough on the surface, a positive mantle anomaly of 50 10 5 ms 2 amplitude has been revealed. This anomaly coincides with the high-velocity inhomogeneity in the upper mantle recognised from seismic data (Marquering and Snieder, 1996). It can be explained by a cold lithosphere block sunk and frozen in the upper mantle due to the complex evolution of tectonic regimes in the suture zone between Variscan West Europe and Precambrian East European Platform (Yegorova and Starostenko, in press). However, recent interpretation of seismic data on the LT-7 profile has the Moho (M1) below the TTZ at 35 km, whereas a second reflector at the depth 50 –55 km may represent a ‘‘second Moho’’ (M2) (Guterch et al., 1994). These authors suggest the existence of a crust –mantle layer in the depth interval between M1 and M2. Basin modelling of sediment accumulation due to rifting in the Polish Trough (Dadles et al., 1995) indicates that in the central part of the TTZ, thickness of the crust has to be about 40 km. These authors suggest that Permo-Mesozoic basin development may be related partly to the intrusion of mantle material into and densification of the lower crust (Dadles et al., 1995). That might be another explanation of positive residual gravity anomaly revealed below the TTZ due to our regional gravity analysis. Strong positive mantle anomaly of 150 10 5 ms 2 amplitude has been also evident over the South Caspian Basin. The regional character of this anomaly is confirmed by its extension over the Kura Depression and Caucasus orogen. In contrast to the PeriCaspian Depression, the evolution of the South Caspian Depression is well under way. It is considered (Zonenshain and Le Pichon, 1986; Nikishin et al., in press) that the formation of the South Caspian Basin has happened in the Upper Jurassic – Lower Cretaceous as result of back-arc extension. Rapid basement subsidence has happened in the Middle Pliocene – Pleistocene. At present, the basin belongs to the region of compressional regime (Pristley et al., 1994; Kopp and Shcherba, 1998). The isostatic equilibrium in the South Caspian Basin is disturbed, since the sharp subsidence of the sedimentary basin bottom is accompanied only by a
slight uplift of the Moho. Revealed strong positive mantle anomaly assumes a subcrustal layer with increased density which underlies both the South Caspian Depression and the Caucasus mountain ranges. Therefore, we suggest that the upper-mantle structures beneath the depression and adjacent orogenic structures of the Caucasus have been inherited from a common pre-alpine source. These relatively heavy portions in the upper mantle under the South Caspian Depression and Caucasus may be interpreted as a result of collision between the Eurasian and Arabian plates. Thus, from the west and south, the EEP is framed by series of high-density domains, located in the uppermost mantle and lower crust, which have been formed probably as a result of collision and accretion on the junction zone of Precambrian EEP with young plates. The location, structure and role of the southern margin of the EEP in Late Paleozoic rifting and in governing the tectonic style and distribution of subsequent (primarily compressional) reactivation are the targets of ongoing national and international programmes of geophysical data acquisition on the EEP (Stephenson, 1997). Vast positive residual anomaly has been also revealed over the Bay of Biscay and adjoining part of the North Atlantic. The value of this anomaly is 50 10 5 ms 2 amplitude (Fig. 7) against the background of mantle anomalies of 100 150 10 5 ms 2. This residual anomaly may be caused by nonbalanced state of the crust of oceanic type. Namely, strong mass deficiency in the Bay of Biscay due to large thickness of both sea-water layer (>4 km) and sedimentary cover (up to 6 km) is not balanced by the Moho uplift up to the depth of 16 km. For reaching the total compensation of the crust, it is necessary to introduce additional high-density domains not only in the crystalline crust, but also in the upper mantle. Mantle origin of this anomaly is confirmed by the highvelocity domain found out in the upper mantle to the west from Iberian peninsula and Bay of Biscay (Marquering and Snieder, 1996). The presence of heavy block in the upper mantle below the Bay of Biscay corresponds with nonactivated regime of the upper mantle of this part of Northern Atlantic. According to Cermak (1982, 1993), the calculated temperatures in the upper mantle at 40 km depth are extremely low ( < 500 C), comparable with that below the Adriatic
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plate and the EEP. To a certain extent, it might be related with the most ancient lithospheric block that occurred in this part of Northern Atlantic (Marquering and Snieder, 1996). Thus, the positive residual anomaly, revealed in the Northern Atlantic to the west from Iberian peninsula and Bay of Biskay, confirms the presence of a cold, high-density/velocity lithospheric block of ancient consolidation traced in the upper mantle down to the depth of approximately 140 km.
6. Conclusions A large-scale 3D density model for the whole European continent has been constructed. The model, approximated by two regional layers—sedimentary cover and the consolidated crust (an offshore seawater layer was added), is based on a simplified velocity model (Giese and Pavlenkova, 1988; Hurtig et al., 1992). The gravity effects of this model were calculated with the program developed by Starostenko and Manukyan (1987), designed to solve the 3D direct problem, taking into account the spherical form of the Earth. Residual gravity anomalies, obtained by subtracting the gravity influence of the crustal model from the observed field, reach amplitudes of a few hundreds 10 5 ms 2. Assumed mantle origin of residual anomalies is supported by their clear correlation with both velocity heterogeneities of the subcrustal upper mantle obtained by seismic tomography and geothermal data. The sedimentary basins in Europe are manifested differently on the map of mantle gravity anomalies (Fig. 7). Taking into account its character, specific features of the evolution of the basin and type of the isostatic compensation, the sedimentary basins in Europe might be classified into three groups. (1) Deep sedimentary basins located in the Precambrian East European Platform and its northern and eastern margins—the Peri-Caspian, Dnieper– Donets, Barents Sea Basins and Fore – Ural Trough. These basins are characterized by great thickness of sediments, slight uplift of the Moho and by an increase of average velocity/density in the consolidated crust (as a result of the reduction of the ‘‘granitic’’ layer). No strong mantle gravity anomalies have been revealed here. The mass compensation of this very deep basins (up to 20 km) is reached on the Moho boundary due to
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strong ‘‘basification’’ of the crust, resulting in a rapid increase of velocity and density up to 6.8 km s 1 and 2.9 103 kg m 3, respectively, and a slight uplift of the Moho up to the depth of 36 km. (2) Basins of active rift and taphrogenic regime located mainly in Western Europe and Western Mediterranean. Negative mantle anomalies of 150 200 10 5 ms 2 have been obtained in the Pannonian Basin, the Tyrrhenian Sea, the Balearic Basin, the Polish – German Basin, the North Sea and Rhine Graben, which are characterized by thin (15 –28 km) crust of transitional and oceanic type. The lithosphere beneath these basins is characterized by lowvelocity domains and high-temperature regime caused by the rise of the asthenosphere up to the subcrustal depths. Negative mantle anomalies obtained over these basins are caused by thermal expansion of the material in the shallow asthenosphere. The compensation of these basins is reached in the upper mantle (probably on the bottom of the asthenospheric layer) due to thermal expansion of the asthenosphere substance. (3) Sedimentary basins located over suture zones at the western and southern margins of the EEP and marked by positive mantle anomalies of 50 – 150 10 5 ms 2 amplitude. Such anomaly has been revealed in the transition zone from Phanerozoic Western Europe to Precambrian East European Platform over the Polish Trough and correlates well with highvelocity domain in the upper mantle. Over the South Caspian Basin, the strongest positive anomaly (up to 200 10 5 ms 2), with its extension to the Caucasus, has been also obtained. Thus, the lithosphere of the East European Platform is surrounded from west to south by a series of upper mantle domains of high density formed as a result of collision and accretion of different lithospheric plates. Thus, the continent-wide 3D gravity modelling, based on the results of generalization of seismic data, may serve as unique structure basis for getting additional important information (revealing the main features of similarity and distinction) on the structure and evolution of main tectonic elements of the study region.
Acknowledgements This paper is a result of collaboration within the framework of the EUROPROBE project GEORIFT.
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The scientific investigations leading to this paper were undertaken as part of INTAS Grants 93-3346 and 970743, and the financial support of INTAS is gratefully acknowledged. We extend our appreciation to Dr. R.A. Stephenson (Free University, Amsterdam, The Netherlands), Dr. P. Szafian (Geomega, Budapest, Hungary) and an anonymous reviewer for their constructive comments and useful suggestions on improving the English of the manuscript. The authors thank Dr. V.G. Kozlenko (Institute of Geophysics, Kiev, Ukraine) and Prof. N.I. Pavlenkova (Institute of Physics of the Earth, Moscow, Russia) for precious help and fruitful discussions. We thank also the participants of the 9th Workshop of the ILP Task Force ‘‘Origin of Sedimentary Basins’’ and EUROCONFERENCE-EUROPROBE held in Oliana (Spain), September 1998, for the interest showed to the topic and stimulating discussions. References Artemjev, M.E., Kaban, M.K., Kucherinenko, V.A., Demyanov, G.V., Taranov, V.A., 1994. Subcrustal density inhomogeneities of Northern Eurasia as derived from the gravity data and isostatic models of the lithosphere. Tectonophysics 240, 249 – 280. Artyushkov, E.V., 1992. Role of crustal stretching in subsidence of the continental crust. Tectonophysics 215, 187 – 215. Artyushkov, E.V., 1993. Physical Tectonics. Nauka, Moscow (in Russian). Artyushkov, E.V., Beer, M.A., 1983. About role of vertical movements in formation the troughs on the continental crust in fold belts. Izv. Akad. Nauk SSSR, Ser. Geol. 9, 25 – 52. Babuska, V., Plomerova, J., 1992. The lithosphere in Central Europe—seismological and petrological aspects. Tectonophysics 207, 141 – 163. Babuska, V., Plomerova, J., Sileny, J., 1984. Large-scale oriented structures in the subcrustal lithosphere of Central Europe. Ann. Geophys. 2 (6), 649 – 662. Balavadze, B.K., Belokurov, V.S., Garkalenko, I.A., Malovitski, Ya.P., Mindeli, P.Sh., Moskalenko, V.N., Neprochnov, Yu.P., Neprochnova, E.A., Starshinova, E.A., Chekunov, A.V., 1975. Physical properties of the rocks. Black Sea Basin. In: Bulanze, Yu.D., Muratov, M.V., Subbotin, S.I., Balavadze, B.K. (Eds.), Earth’s Crust and Evolution of Black Sea Basin. Nauka, Moscow, pp. 254 – 269 (in Russian). Belousov, V.V., 1990. Tectonosphere of the Earth: upper mantle and crust interaction. Tectonophysics 180, 139 – 183. Belousov, V.V., Pavlenkova, N.I., 1984. The types of the Earth’s crust. J. Geodyn. 1, 167 – 183. Belousov, V.V., Pavlenkova, N.I., 1989. Types of the earth’s crust in Europe and the North Atlantic. Geotectonika 3, 3 – 34 (in Russian).
Blundell, D.G., 1990. Seismic images of continental lithosphere. J. Geol. Soc. (London) 147, 895 – 913. Blundell, D., Freeman, R., Mueller, St. (Eds.), 1992. A Continent Revealed. Cambridge Univ. Press, Cambridge. Cermak, V., 1982. Crustal temperature and mantle heat flow in Europe. Tectonophysics 83, 123 – 142. Cermak, V., 1993. Lithospheric thermal regime in Europe. Phys. Earth Planet. Inter. 79, 179 – 193. Dadles, R., Narkiewicz, M., Stephenson, R.A., Visser, M.T.M., van Wees, J.-D., 1995. Tectonic evolution of the Mid-Polish Trough: modelling implications and significance for central European geology. Tectonophysics 252, 179 – 195. Dortman, N.B. (Ed.), 1976. Physical Properties of the Rocks and Mineral Resources (Petrophysics). Nedra, Moscow (in Russian). Fuchs, K., 1979. Structure, physical properties and lateral heterogeneities of the subcrustal lithosphere from long-range deep seismic sounding observations on continents. Tectonophysics 56, 1 – 15. Giese, P., Pavlenkova, N.I., 1988. Structural maps of the earth’s crust in Europe. Izv. Akad. Nauk SSSR, Fiz. Zemli 10, 3 – 14 (in Russian). Giese, P., Prodehle, G., Stein, A. (Eds.), 1976. Explosion Seismology in Central Europe, Methods and Results, Berlin. Granser, H., 1987. Three-dimensional interpretation of gravity data from sedimentary basins using an exponential density – depth function. Geophys. Prospect. 35, 1030 – 1041. Guterch, A., Grad, M., Materzok, R., Rajduel, J., Perchuc, E., Toporkiewicz, S., 1984. Structure of the Earth’s crust in the contact zone of the Paleozoic and Precambrian Platforms and the Carpathian Mts in Poland. Acta Geophys. Pol. XXXII, 25 – 41. Guterch, A., Grad, M., Janik, T., Materzok, R., Luosto, U., Yliniemi, E., Luck, E., Shulze, A., Forster, K., 1994. Crustal structure of the transition zone between Precambrian and Variscan Europe from new seismic data along LT-7 profile (NW Poland and eastern Germany). C. R. Acad. Sci. Paris Geophys. 319, 1489 – 1494. Hurtig, E., Cermak, V., Haenel, R., Zui, V. (Eds.), 1992. Geothermal Atlas of Europe. Hermann Haack Verlagsgesellschaft, Geographisch Kartographische Anstalt Gotha, Potsdam. Kinsman, D.J.J., 1975. Rift basins and peculiarities of sediment accumulation in the conditions of sagged margins of continents. In: Fisher, G., Judson, S. (Eds.), Petroleum and Global Tectonics. Princeton Univ. Press, Princeton, NJ. Kopp, M.L., Shcherba, I.G., 1998. Caucasus basin in the Paleogene. Geotectonika 2, 29 – 50 (in Russian). Marquering, H., Snieder, R., 1996. Shear-wave velocity structure beneath Europe, the northeastern Atlantic and western Asia from waveform inversions including surface-wave mode coupling. Geophys. J. Int. 127, 283 – 304. McKenzie, D.P., 1978. Some remarks on the development of sedimentary basins. Earth Planet. Sci. Lett. 40, 25 – 32. Moskalenko, V.N., 1975. Earth’s crust of Mediterranean Sea. In: Belyaevsky, N.A. (Ed.), Earth’s Crust of Continent Margins and Inner Seas. Nauka, Moscow, pp. 122 – 141 (in Russian). Nikishin, A.M., Ziegler, P.A., Stephenson, R.A., Cloetingh, S., Furne, A.V., Fokin, P.A., Ershov, A.V., Bolotov, S.N., Korotaev, M.V., Alekseev, A.S., Gorbachev, V.I., Shipilov, E.V., Lankreijer, A., Bembinova, E.Yu., Shalimov, I.V., 1996. Late Pre-
T.P. Yegorova, V.I. Starostenko / Tectonophysics 346 (2002) 5–21 cambrian to Triassic history of the East European Craton: dynamics of sedimentary basin evolution. Tectonophysics 268, 19 – 46. Nikishin, A.M., Cloetingh, S., Brunet, M.-F., Bolotov, S.N., Ershov, A.V., Stephenson, R.A., 1999. Scythian Platform, Caucasus and Black Sea region: mesozoic – cenozoic tectonic history and dynamics. In: Crasquin-Soleau, S., Barrier, E. (Eds.), Peri-Tethys Memoir 3, Mum. Mus. Natn. His. Nat., Paris, vol. 177 (in press). Panza, G.F., Mueller, St., Calcagnile, G., 1980. The gross features of the lithosphere- and asthenosphere system in Europe from seismic surface waves and body waves. Pure Appl. Geophys. 118, 1209 – 1213. Pavlenkova, N.I., Romanyuk, T.V., 1991. Seismo-gravity modelling for the crust of Siberia. Geol. Geophys. 1, 34 – 42. Podoba, N.V., Ozerskaya, M.A. (Eds.), 1975. Physical Properties of Sedimentary Cover of East-European Platform. Nauka, Moscow (in Russian). Pristley, K., Baker, C., Jackson, J., 1994. Implications of earthquake focal mechanism data for the active tectonics of the south Caspian basin and surrounding regions. Geophys. J. Int. 118, 11 – 141. Rusakov, O.M., 1990. The thickness and density of sedimentary cover in the Indian Ocean. Geophys. J. 3 (8), 390 – 399. Sclater, J.G., Christie, P.A., 1980. Continental stretching: an explanation of the post Mid-Cretaceous subsidence of the Central North Sea basin. J. Geophys. Res. 85, 3711 – 3739. Spakman, W., van der Lee, S., van der Hilst, R., 1993. Travel-time tomography of the European-Mediterranean mantle down to 1400 km. Phys. Earth Planet. Inter. 79, 3 – 74. Starostenko, V.I., Manukyan, A.G., 1987. Gravimetric problems for spherical celestial bodies and their application to the study of the Nectaris Basin Mascon on the Moon. Boll. Geofis. Teor. Appl. 29 (115), 201 – 219. Stephenson, R.A., 1997. Intraplate tectonics and basin dynamics of the EEP: Prypyat – Dnieper – Donets Basin review—problems and prospects. Geophys. J. 17, 170 – 175.
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Stephenson, R.A., Wilson, M., de Boorder, H., Starostenko, V.I. (Eds.), 1996. EUROPROBE: Intraplate Tectonics and Basin Dynamics of the Eastern European Platform. Tectonophysics Spec., vol. 268. Wilson, M., Lyashkevich, Z.M., 1996. Magmatism and the geodynamics of rifting of the Pripyat – Dnieper – Donets rift, Eastern European Platform. Tectonophysics 268, 65 – 81. Wood, R., Barton, P., 1983. Crustal thinning and subsidence in the North Sea. Nature 302, 134 – 136. Yegorova, T.P., Starostenko, V.I., 1999. Large-scale 3-D gravity analysis of the lithosphere below the transition zone from Western Europe to East-European Platform. Tectonophysics 313, 41 – 58. Yegorova, T.P., Kozlenko, V.G., Pavlenkova, N.I., Starostenko, V.I., 1995. 3-D density model for the lithosphere of Europe: construction method and preliminary results. Geophys. J. Int. 121, 873 – 892. Yegorova, T.P., Kozlenko, V.G., Starostenko, V.I., 1997a. Density heterogeneities of the European upper mantle inferred from 3-D large-scale gravity modelling. In: Fuchs, K. (Ed.), Upper Mantle Heterogeneities from Active and Passive Seismology. Kluwer Academic, Amsterdam, pp. 249 – 256. Yegorova, T.P., Starostenko, V.I., Kozlenko, V.G., Pavlenkova, N.I., 1997b. Three-dimensional gravity modelling of the European Mediterranean lithosphere. Geophys. J. Int. 129, 355 – 367. Ziegler, P.A., 1983. Crustal thinning and subsidence in the North Sea. Nature 304 (5926), 561. Ziegler, P.A., 1990. Geological Atlas of Western and Central Europe, 2nd edn. Shell International Petroleum Maatschappij, Geol. Soc. London, Elsevier, Amsterdam. Zielhuis, A., Nolet, G., 1994. Shear-wave velocity variations in the upper mantle beneath central Europe. Geophys. J. Int. 117, 695 – 715. Zonenshain, L.P., Le Pichon, X., 1986. Deep basins of the Black Sea and Caspian Sea as remnants of Mesozoic back-arc basins. Tectonophysics 123, 181 – 211.
Lithosphere structure of European sedimentary basins from regional three-dimensional gravity modelling T.P. Yegorova*, V.I. Starostenko Institute of Geophysics, National Academy of Sciences of Ukraine, 32 Pr. Palladina, 252680 Kiev, Ukraine Received 2 October 2001
Abstract A three-dimensional (3D) density model, approximated by two regional layers—the sedimentary cover and the crystalline crust (offshore, a sea-water layer was added), has been constructed in 1 averaging for the whole European continent. The crustal model is based on simplified velocity model represented by structure maps for main seismic horizons—the ‘‘seismic’’ basement and the Moho boundary. Laterally varying average density is assumed inside the model layers. Residual gravity anomalies, obtained by subtraction of the crustal gravity effect from the observed field, characterize the density heterogeneities in the upper mantle. Mantle anomalies are shown to correlate with the upper mantle velocity inhomogeneities revealed from seismic tomography data and geothermal data. Considering the type of mantle anomaly, specific features of the evolution and type of isostatic compensation, the sedimentary basins in Europe may be related into some groups: deep sedimentary basins located in the East European Platform and its northern and eastern margins (Peri-Caspian, Dnieper – Donets, Barents Sea Basins, Fore – Ural Trough) with no significant mantle anomalies; basins located on the activated thin crust of Variscan Western Europe and Mediterranean area with negative mantle anomalies of 150 to 200 10 5 ms 2 amplitude and the basins associated with suture zones at the western and southern margins of the East European Platform (Polish Trough, South Caspian Basin) characterized by positive mantle anomalies of 50 – 150 10 5 ms 2 magnitude. An analysis of the main features of the lithosphere structure of the basins in Europe and type of the compensation has been carried out. D 2002 Elsevier Science B.V. All rights reserved. Keywords: Sedimentary basins; Europe; Earth’s crust; Upper mantle; Gravity modelling
1. Introduction Gravity modelling of a three-dimensional (3D) structure of the Earth’s crust and upper mantle is a powerful tool for studying the lithosphere in its whole volume. It is linked with the seismic method, whose data serve as a basis for constructing the density
*
Corresponding author. Fax: +380-44-450-2520. E-mail address: [email protected] (T.P. Yegorova).
models. So far, the most used method in gravity modelling is the construction of a two-dimensional (2D) gravity model on the DSS profiles. The velocity models using the correlation functions between velocity and density are converted into the density equivalents, from which the gravity effect is calculated. Using 2D modelling, velocity –density models have been obtained on most of the DSS profiles in Europe. Very often, it turns out that these models, constructed even for the same profiles, have strong discrepancies. This disagreement is caused by differences in the
0040-1951/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 0 1 ) 0 0 2 2 5 - 6
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methodologies used by different authors in interpretation of data of both seismic and gravity methods. Additional factor that should be considered, especially in gravity modelling, is regarding the 3D character of the model. As a rule, main features of the observed gravity field are usually explained by distribution of density heterogeneities in the crust, generally without taking into account heterogeneous upper mantle. The 3D gravity modelling with using the unique methodology, allows to minimize the shortages of 2D modelling and to receive comparable (on a uniform ground) models of different tectonic units of the crust. An important result that might be obtained in the course of 3D modelling, is the information about density heterogeneities of the uppermost mantle received from residual gravity anomalies (obtained by subtraction of calculated crustal effect from the observed field). It is known that over young tectonic regions, such mantle density heterogeneities create gravity
anomalies reaching some hundreds 10 5 ms 2 (Artemjev et al., 1994; Yegorova et al., 1995, 1997a). Usually, such mantle gravity anomalies are not manifested in the observed field, being compensated by crustal heterogeneities. Our regional 3D gravity model for the lithosphere of Europe and part of the North Atlantic (Fig. 1) is based on generalized velocity model (Giese and Pavlenkova, 1988) and consists of two regional layers of variable thickness with laterally varying average velocity/density—the sedimentary cover and the crystalline crust. Recent publications discuss the results of 3D large-scale gravity modelling for two most interesting portions of this continent-wide model: for the Alpine – Mediterranean region (Yegorova et al., 1997b) and for the transition zone from Variscan Western Europe to East European Platform (Yegorova and Starostenko, in press). For these models, a 3D gravity influence has been calculated. Residual grav-
Fig. 1. Tectonic setting of Europe. The picture shows the region of the present large-scale 3D gravity modelling. The areas with crosses indicate Precambrian Shields on the East European Platform. The areas of Caledonides, Variscides and Alpides in Europe are shown by gray dark, medium and light shadings. A solid line indicates the position of the section of the 3D gravity model (Fig. 8).
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ity anomalies, obtained by removal of the calculated effect from the observed field, are generally caused by density heterogeneities of the uppermost mantle (Yegorova et al., 1997a). Here we present the 3D large-scale gravity model (in 1 averaging) for the lithosphere of Europe and part of the North Atlantic with the emphasis on the sedimentary basins. Although the velocity structure and geometry of most of these basins are known, new additional information might be obtained by considering them all together on the base of 3D regional model. It allows to study the type and depth of isostatic compensation, inner crustal structure and to reveal heterogeneities in the upper mantle. It has a special value since most of sedimentary basins are not manifested in the averaged observed gravity field by intensive anomalies, being
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compensated by the heterogeneities of the crust and upper mantle.
2. Initial data 2.1. Gravity data The initial observed gravity field that we used, averaged on 1 1 grid, was based on the Bouguer anomaly (BA) and free-air anomaly (FAA) data obtained by ZNIIIGAiK (the Central Research Institute of Geodesy, Airsurvey and Cartography of the former Soviet Union). We have used a combined reduction joining the BA on land and FAA offshore (on the coastal line, values of BA and FAA became
Fig. 2. Generalized map of the observed gravity anomaly field for Europe and North Atlantic (Yegorova et al., 1995). The map uses Bouguer anomalies on land and free-air anomalies offshore, averaged over 1 1 grid. Data courtesy of the Central Research Institute of Geodesy, Airsurvey and Cartography of the Former Soviet Union (ZNIIGAiK). Contour interval 10 10 5 ms 2. A solid line indicates the location of the section of the 3D gravity model. AB: Alboran Basin; AL: Alps; BB: Bay of Biscay; BS: Black Sea; BSH: Baltic Shield; BSB: Barents Sea Basin; CA: Calabrian Arc; CC: Caucasus; DDB: Dnieper – Donets Basin; FUT: Fore – Ural Trough; PGB: Polish – German Basin; NSB: North Sea Basin; PB: Pannonian Basin; PT: Polish Trough; PY: Pyrenees; PCB: Peri-Caspian Basin; SCB: South Caspian Basin; TS: Tyrrhenian Sea, TTZ: Teisseyre – Tornquist Zone; UR: Urals, USH: Ukrainian Shield. A solid line indicates the position of the section of the 3D gravity model (Fig. 8).
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Fig. 3. The 3D density model for the crust of Europe and North Atlantic. Data averaging is over 1 1 grid. The model consists of a sea-water layer with q = 1.03 103 kg m 3 and two regional layers with laterally varying average density: a sedimentary cover and a crystalline crust. (a) Sea-bottom topography, km; (b) depth to the bottom of the sedimentary layer (Giese and Pavlenkova, 1988), km; (c) depth to the bottom of the crystalline crust (the Moho), km, according to Giese and Pavlenkova (1988). A solid line indicates the position of the section of the 3D gravity model in Fig. 8. For the abbreviation of tectonic units, see Fig. 2.
equal). Accordingly, offshore, a layer of sea water was included in the model. A map of the observed gravity field using such a combined reduction for the whole European continent and the North Atlantic was published by Yegorova et al. (1995). This map is demonstrated also in Fig. 2. Quiet and non-anomalous gravity field with BA values fluctuating around zero ( ± 0 20 10 5 ms 2), observed over Western European and East-European Platforms, means that in general, the gravity field here is compensated. A specific feature of the observed field gob of Europe is the gravity low over Western Scandinavia where the lowest BA values in the north of the continent, reaching 80 10 5 ms 2, are found. The orogens of the Alps, Caucasus and Carpathians are also marked by great distinctive gravity lows. Generally anomalous gravity field pattern is seen over the Alpine – Mediterranean region and North Atlantic. As a rule, deep sedimentary basins are not manifested in the observed field by strong gravity anomalies. This
means that in the whole, the sedimentary basins are supposed to be compensated units, and the characteristic type of compensation depends on the structure and evolution of the lithosphere. 2.2. Structure of the crustal model according to seismic data Our gravity model is based on the information about the crustal structure of Europe (thickness of the sedimentary cover, crustal thickness, average crustal seismic velocity and characteristic types of the crust) generalized by Giese and Pavlenkova (1988). These data have also been used in the compilation of the Geothermal Atlas of Europe (Hurtig et al., 1992). In the Alpine – Mediterranean region, we have taken into account recent seismic data on Moho topography obtained by investigations of the EGT profile (Blundell et al., 1992). Thus, the crustal model adopted for Europe and the northern part of Northern Atlantic
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comprises two regional layers of variable thickness— the sedimentary cover and the crystalline crust. It assumes lateral variation of average density inside the layers. The model layers are bounded by reliable seismic boundaries of first order. They are (1) the depth to the ‘‘seismic basement’’ (Hsb), that is, the surface of the crustal layer specified by the compressional wave velocity Vp = 5.8 –6.2 km s 1 and (2) the total crustal thickness or depth to the Moho discontinuity. Offshore, a sea-water layer was added in the model. Averaged on 1 1 grid in a plane, these structure maps are shown in Fig. 3a – c. On the map of depth to the basement (Fig. 3b), one can clearly see sedimentary basins of different depths. Very deep basins are located in the southern and southeastern part of the East European Platform (EEP)—Dnieper – Donets Basin and Peri-Caspian Basin of more than 18 km depth—and in the vicinity of its northern (Barents Sea Basin) and southern margins (Black Sea and South Caspian Basins of more than 10 and 16 km depth, respectively). The shapes of these basins are regular with the predominance of round and oval basins (Peri-Caspian, South Caspian, Barents Sea Basins) as well as linear rift basins (Dnieper – Donets Basin). From the east, the EEP is confined by meridional Fore – Ural Trough with the thickness of sediments of more than 4 – 6 km. From NW and W, the Baltic Shield and the EEP are bounded by a series of sedimentary troughs associated with the NW trending Permo-Mezosoic European Basin (Ziegler, 1990). This series of basins begins with the North Sea (thickness > 8 km), followed by the Polish – German Basin and Polish Trough of more than 4 –6 km thickness. Nearly the same depth (4 –6 km) is typical for the Western Mediterranean (Alboran and Tyrrhenian Seas) and Pannonian Basins. The map of Moho depth in Europe (Fig. 3c) reveals the division of the continental lithosphere into two major blocks: thick (40 –44 km) consolidated crust of the East European Platform and thin activated Phanerozoic crust of West Europe. From the analysis of both maps for the basement (Fig. 3b) and for the Moho (Fig. 3c), one can see that very deep sedimentary basins in Europe are located over thick crustal block of the East European Platform. Below these basins, the thickness of the crystalline crust becomes thinner because of the basement subsidence and the Moho shallowing to a depth of 36– 38 km. However,
the Polish Trough, where the Moho boundary deepens down to a depth of 50 km, is assumed to be an exception from this regularity. Beneath the basins in Variscan Western Europe and Alpine –Mediterranean belt, the Moho shallowing is estimated in the range of 16 –28 km.
3. Density parameterization of the model The two regional layers (the sedimentary layer and the crystalline crust) in our continent-wide 3D crustal model were assigned laterally varying average density values. Offshore, the model was supplemented by a sea-water layer with density of q = 1.04 103 kg m 3. Since the initial data on the velocity structure of the model (Giese and Pavlenkova, 1988; Hurtig et al., 1992) have no information on seismic velocity distribution in the sedimentary cover, we paid particular attention to data acquisition on the density (q) and velocity (Vp) distributions in this layer. The lateral variation of average density q of the sedimentary cover was estimated by a number of approaches. The main one deals with estimations of mean interval density according to laboratory data measurements. For that purpose, densities of dry sediments of the East European Platform (Podoba and Ozerskaya, 1975) were recalculated into the average densities, taking into account 100% water saturation. We adopted the increase of average density with depth q(h) obtained for the sediments on the EEP as a base function for evaluating the densities of sediments of Western Europe. According to Fig. 4, the average density of sediments q on the EEP (slanting shading) gradually increases with depth according to exponential function. It varies from 2.15 –2.25 103 kg m 3 on the depositional surface to 2.6 103 kg m 3 at the depth level 4– 5 km. At greater depths, the densities approach values typical for the crystalline rocks. For density parameterization of the sediments in the regions of Urals, Caspian Sea and South Caspian Basin, we have used laboratory data measurements from Dortman (1976). The second approach consists of P-wave velocity conversion into its density equivalent by using appropriate correlation density/velocity (q/V) functions. For the Black Sea Basin, density values for the sediments have been derived from P-wave velocity data using the
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Fig. 4. Dependence of sediment density with depth. (1) Distribution of average density with depth for sediments of the East European Platform (Podoba and Ozerskaya, 1975) adopted as a basic distribution for the gravity modelling; (2) the function used by Granser (1987) for the Pannonian Basin; (3 – 5) areas of density variations obtained by Kinsman (1975) for limestones (3), clays (4) and sandstones (5); (6) density distribution for oceanic sediments according to Rusakov (1990).
correlation function q =1.8822 0.0871V + 0.1104V 2 0.01132V 3 (Balavadze et al., 1975). For the Western Mediterranean, where sediments are distinguished as ‘‘uncompacted’’(V = 2.0 km s 1), ‘‘normal’’ or ‘‘semicompacted’’(V = 3.0 km s 1 ) and ‘‘compacted’’(V = 4.5 km s 1) (Moskalenko, 1975), we have estimated the average P-wave velocity for the whole sedimentary-cover thickness. These data were recalculated into density values using the mentioned above q/ V function. Assuming that sediment compaction with depth is similar in the Atlantic and Pacific Oceans (Rusakov, 1990), a generalized density/depth function for young oceanic sediments (Rusakov, 1990), shown in Fig. 4 by a solid line, was used to estimate q of sediments in the North Atlantic and North Sea. By this approach, only approximate estimates of average density of the sedimentary layer can be made. Nevertheless, its expediency is substantiated by the evaluation of the sensitivity of the method to changes in density. For example, a density decrease of 0.1 103 kg m 3 in
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the density of North Sea sediments of 6 km thickness will create a gravity effect of 25 10 5 ms 2 less than in the model ( 250 10 5 ms 2 instead of 225 10 5 ms 2) that should be compared with the residual gravity anomaly of 200 10 5 ms 2 magnitude. The resulting distribution of average density of the sedimentary cover for the study region is shown in Fig. 5a. The estimates of average density in the crystalline crust (Fig. 5b) were derived by using the conversion function of the type q = 2.7 + 0.27 (Vp 6.0) (Pavlenkova and Romanyuk, 1991) for q units of 103 kg m 3 when Vp is in km s 1. The average crustal Pwave velocities were taken from Giese and Pavlenkova (1988). The crystalline crust of very high density reaching 2.90 103 kg m 3 underlies extremely deep sedimentary basins over the EEP—the Dnieper – Donets and Peri-Caspian Basins (Fig. 5b). The same high densities are also typical for the whole crystalline crust of the Urals. The crystalline crust of the Variscan West Europe is characterized by lower average densities than that of EEP; however, under the Western Mediterranean and Pannonian Basins, a high-density crystalline crust (up to 2.83 103 kg m 3) is evident. The same high-density domain has been revealed in the crystalline crust beneath the Black Sea.
4. Modelling methodology and main results For the purposes of our regional gravity analysis, a two-part calculation scheme was developed. It consists of two main elements: a database and a computer program. A database contains all information about the model structure and parameterization (in 1 averaging), results of the computations for every stage of the modelling as well as the resulting maps. The kernel of the computer program block is solution of the direct gravity problem from a 3D mass distribution taking into the account the spherical configuration of the Earth (Starostenko and Manukyan, 1987). The gravity effect of the model, approximated by a set of 1 1 parallelepipeds, has been calculated by this program. For gravity effect calculations, we have used anomalous densities of the layers obtained by subtracting the upper mantle density (taken to be 3.3 g cm 3 ) from the average density of the layer: Dq = q 3.3 103 kg m 3. As a result of the mod-
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Fig. 6. Gravity inf luence of the 3D density model, obtained by summing up the effects of all three layers: the sea-water layer, the sedimentary cover and the crystalline crust. Isoline interval 50 10 5 ms 2. A solid line indicates the position of the section of the 3D gravity model in Fig. 8. For the abbreviation of tectonic units, see Fig. 2.
elling, the maps of gravity influence have been obtained for each the model layer (sea water, sediments, consolidated crust). Total gravity effect gcr, obtained by summing up the effects of the layers, is shown in Fig. 6. Since the range of anomalous density of the model layers varies from 2.27 103 kg m 3 (sea-water layer) up to 0.47 103 kg m 3 (crystalline crust), the values of the calculated gravity effect reach large negative values (up to 900 10 5 ms 2 over the Alps; see Fig. 6). Therefore, it is necessary to choose a certain background level (a ‘‘norm’’ gnor) to normalize calculated gravity effect gcr. Proceeding from our experience of 3D gravity modelling (Yegorova et al., 1995; 1997a,b), we have used the average value of gravity influence gcr for the
crust of the EEP as that value gnor. The EEP is the stable Precambrian core of the European continent, which is characterized by an observed gravity field with Bouguer anomalies up to ± 20 10 5 ms 2. For the 3D crustal model considered and the q/V function, the average value of gcr for the EEP ( gnor) is equal to 790 10 5 ms 2. The distribution of residual anomalies Dgr, obtained by subtracting the calculated effect of the model gcr (Fig. 6), normalized by constant gnor, from the observed gravity field (Fig. 2): Dgr = gob ( gcr gnor), is demonstrated in Fig. 7. After the normalizing procedure, the residual anomalies, shown in Fig. 7, are in the range + 200 to 200 10 5 ms 2. Since our model has been constructed on the base of a simplified velocity model for the crust of the
Fig. 5. Distribution of the average density in the model layers, 103 kg m 3. (a) Density in the sedimentary cover obtained with the dependence in Fig. 4; (b) in the crystalline crust, average densities were calculated from average velocities by using the conversion function q = 2.7 + 0.27 (Vp 6.0). A solid line shows the location of a section of the 3D density model (Fig. 8). For the abbreviation of tectonic units, see Fig. 2.
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Fig. 7. Residual gravity anomalies Dgr obtained by subtracting the modelled gravity effect gcr (Fig. 6), normalized by constant gnor = 790 10 5 ms 2, from the observed gravity field gob in Fig. 2: Dgr = gob ( gcr gnor) (isoline interval 50 10 5 ms 2). A solid line indicates the position of the section of the 3D gravity model in Fig. 8. For the abbreviation of tectonic units, see Fig. 2.
European continent, we are obliged to make accuracy evaluations of our modelling by choosing an optimal isoline interval for maps of gravity effects from the model layers. The sedimentary cover, with an average thickness of about 2 km (contour interval is 1 km) and with 0.05 103 kg m 3 isoline interval of average density, should be presented by a gravity effect map with a contour interval not less than 50 10 5 ms 2 . The crystalline crust of the study area, described by an average thickness of 35 –40 km (with 4 km isoline interval) and by 0.1 km s 1 isoline interval of average velocity and correspondingly by that of 0.02 103 kg m 3 of average density, is featured by a gravity effect map no less than 35 10 5 ms 2. Therefore, we have defined the main isoline interval for our gravity maps as 50 10 5 ms 2 which is in correspondence with a rms of about 20 – 25 10 5 ms 2. If the model presented reflects the real structure of the Earth’s crust, the residual anomalies should be caused by density heterogeneities in the uppermost
mantle. However, since structure of the crust is investigated incompletely because of the rare and irregular set of the DSS profiles in Europe, it is quite probable the existence of density heterogeneities in the crust (especially in the lower crust poorly investigated by seismic methods), not taken into account in our initial model. The gravity influence of the ‘‘unknown’’ heterogeneities of the upper and middle crust, not exceeding 50 10 5 ms 2, has been filtered using a 50 10 5 ms 2 contour interval in the gravity maps. We therefore presume that large residual anomalies are caused mainly by heterogeneities in the uppermost mantle, though heterogeneities of the lower crust can also contribute to the residual gravity anomalies. It is necessary to consider the origin of the residual anomalies for each specific geologic unit on the basis of comparison of configuration and magnitude of the anomalies with contours of geological units, by invoking the data from other geophysical methods, first of all, the results of seismology and geothermal data.
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The general decrease of average residual anoma-lies from 0 50 10 5 ms 2 on the east to 100 150 10 5 ms 2 on the west can be explained by a decrease in the upper mantle density below the continent. The division of the continent in two main blocks of different structure and activity of the crust and upper mantle—activated Phanerozoic Western Europe and stable Precambrian East European Platform—has been revealed from numerous geophysical investigations: seismic data (Giese et al., 1976; Fuchs, 1979; Guterch et al., 1984; Blundell, 1990), seismotomographic investigations with using P- and S-waves (Panza et al., 1980; Babuska et al., 1984; Spakman et al., 1993; Zielhuis and Nolet, 1994; Marquering and Snieder, 1996) and geothermal data (Cermak, 1982, 1993). In the area of the Teisseyre –Tornquist Zone (TTZ), a positive anomaly of 50 100 10 5 ms 2 has been revealed against the background of the gradient zone of residual anomalies. The results on this part of regional 3D gravity model in the transition zone from Western Europe to East European Platform are presented by Yegorova and Starostenko (in press). Against the regional background of general decrease of the residual gravity field from east to the west, positive and negative anomalies have been distinguished over specific tectonic units. The negative Dgr anomalies with maximal amplitude (more than 250 10 5 ms 2) are found in the North Atlantic. Negative mantle anomalies of 150 200 10 5 ms 2 are typical for the Western Mediterranean Basin, Pannonian Basin, Polish –German Basin, North Sea and the Iberian Peninsula. Local mantle gravity highs are found mainly over the Alpine orogeny. Alps and Adriatic plate/Dinarides are marked by relative Dgr highs with more than + 50 10 5 ms 2 amplitude. Local highs are also observed above the Calabrian Arc and the Pyrenees. The strongest mantle anomaly with more than 150 10 5 ms 2 amplitude is found in the Caucasus area. The regional character of this anomaly is confirmed by its extension over the Kura and South Caspian Basins (Yegorova et al., 1995). The orogens are considered as structures formed as the result of the collision of the continental plates and microplates, which led to the conservation of a rigid, cool, highvelocity/density lithospheric block in the upper mantle. Interpretation of the 3D gravity modelling results for the Alpine –Mediterranean area has been done earlier by Yegorova et al. (1997b).
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5. Imaging the sedimentary basins by 3d largescale gravity analysis The main results of the 3D gravity analysis showing the contribution of each of the layer into the model are demonstrated in the section (Fig. 8) crossing the 3D model along the Tyrrhenian Sea-Pannonian BasinTTZ-Dnieper – Donets Basin-Peri-Caspian Basin. Sedimentary basins in Europe regarding their manifestation in the calculated gravity fields, type of isostatic equilibrium state and geotectonic evolution might be related to three major groups. (1) Deep sedimentary basins located in the East European Platform and in its northern and eastern margins—the Peri-Caspian, Dnieper – Donets, Barents Sea Basins and Fore –Urals Trough. These basins are characterized by transitional type (from continental to oceanic) of the crust (Belousov and Pavlenkova, 1984, 1989). According to our gravity modelling, mantle gravity anomalies over these basins are absent or slightly positive (reaching 50 10 5 ms 2 amplitude). These deep basins have originated on the thick continental crust of the EEP. Rapid subsidence of the crust and sediment accumulation occurred in the basins during Late Devonian – Carboniferous. The common characteristics of these basins are the great thickness of the sedimentary cover, the uplift of the Moho boundary (up to the 36-km depth) below the most subsided part of the basins, strong increase of average velocity/ density in the consolidated crust and normal values of the terrestrial heat flow. The increase of average velocity in the crystalline crust is the result of the reduction of thickness of ‘‘granitic’’ layer (with Vp velocity up to 6.5 km s 1) or its total elimination. It is much more probable that the destruction of the continental crust is affected by the action of hot melted upper mantle material on the crust. As a driving process resulting in a rapid increase of velocity (density) in the crystalline crust below the basins, an ‘‘oceanisation’’ or ‘‘basification’’ of the crust—intrusion of mantle magmas of basic and ultrabasic composition into the crust is suggested by Belousov (1990). As a result of density increase of the crust due to ‘‘basification’’ or metamorphism, the crustal blocks sink into the upper mantle; they may experience eclogitization (Belousov, 1990; Artyushkov, 1992, 1993). These authors believe that the destruction of continental crust occurs in situ, without extension
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Fig. 8. A density section of the 3D density model of the European crust. A section crosses the continent through Tyrrhenian Sea, Pannonian Basin, Dnieper – Donets Basin and Peri-Caspian Basin (for location, see Figs. 2 and 3, 5 – 7). The numbers in the crustal section are density values (in 103 kg m 3) derived from the dependence in Fig. 4 (for sediments) and from the conversion function q = 2.7 + 0.27 (Vp 6.0) (for the crystalline crust), where Vp was taken from Giese and Pavlenkova (1988). The gravity effects (middle part of the figure) were derived for each the model layer (sea water, sediments, crystalline crust). Upper part of the figure characterizes the thermal regime of the region: terrestrial heat f low density (simplified version of Hurtig et al., 1992) and calculated temperatures at 40 km depth (Cermak, 1993). The crustal section was complemented by the thickness of the lithosphere taken from Panza et al. (1980) and Babuska and Plomerova (1992), numbers in the brackets are the shear-wave velocity (in km s 1) at 80 km depth according to seismotomography study of Marquering and Snieder (1996).
being essential. As it is shown by Artyushkov and Beer (1983), to make the crust subside for several kilometers, it should be extended to three or more times its size. Belousov (1990) considers that there is just no
room for that on the continents, where basins are surrounded by stable structures. Now, there are new data available on a particular role of extension in formation of deep basins on the
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EEP. These data have been obtained in the framework of EUROPROBE project ‘‘GeoRift’’ focused on the well-documented Dnieper– Donets Basin (DDB), the largest and deepest Devonian – Carboniferous rift in Europe (Stephenson et al., 1996). It was shown by Nikishin et al. (1996) that origin and evolution of the sedimentary basins on the EEP was governed by repeatedly changing regional stress field. Intraplate magmatism was controlled by changes in stress field and by mantle hot-spot activity. Based on geochemical study of magmatic rocks of the DDB, combined with estimations of the amount of extension (b = 1.1– 1.3), Wilson and Lyashkevich (1996) suggest that magmatism was trigged by the upwelling of a thermally and geochemically anomalous mantle plume from the deep mantle. They concluded that magmatism, rifting and domal basement uplift were contemporaneous at several localities within the EEP, suggesting that the thermal and geodynamic evolution of the platform could have been influenced by a cluster of mantle plumes during Late Devonian. The mass compensation of these very deep basins (up to 20 km) is reached on the depth level of he bottom of the crust (the Moho), generally due to rapid increase of average density (up to 2.9 103 kg m 3) in the crystalline crust (see Fig. 5b) and a slight Moho uplift (up to 36 km). Fig. 8 shows that the gravity effect of the sediments of Peri-Caspian Basin is practically compensated by the effect of the highdensity thin crystalline crust. Slight residual anomaly (up to 50 10 5 ms 2) may be indicative of the higher densities of the sedimentary cover (especially in its lower part) than those taken in the model (2.52 103 kg m 3), though an additional consolidation of the subcrustal lithosphere below the basin also can take place as a result of phase transition in the lower crust –upper mantle. (2) Basins of active rift and taphrogenic regime mainly in Western Europe, manifested by negative mantle gravity anomalies of 150 200 10 5 ms 2 amplitude. As a rule, these basins (Pannonian Basin, Tyrrhenian and Alboran Basins, North Sea, Rhine Graben and Polish – German Basin) are characterized by crust of transitional and oceanic types with thinner (15 – 28 km) crust and lower average velocities/densities in the consolidated crust in comparison with the EEP. The crust below the Western Mediterranean Basin is of oceanic type due to its considerable
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thinning. The lithosphere beneath these basins, as it is seen from Fig. 8, is characterized by low-velocity heterogeneities in the upper mantle (Spakman et al., 1993; Marquering and Snieder, 1996) and high-temperature regime caused by the upwelling of the heated substance of the asthenosphere up to the depth from 30 (Western Mediterranean Basin) to 50 – 60 km (North Sea, Rhine Graben, Pannonian Basin) (Panza et al, 1980; Babuska and Plomerova, 1992). It is generally accepted that the subsidence in these young basins has happened in the result of stretching the lithosphere, during which the continental crust becomes thinner and finally ruptures [shown by Sclater and Christie (1980), Wood and Barton (1983) and Ziegler (1983) on the example of the North Sea Basin]. It is considered also that a fast extension, which has taken place during some epochs, originated an initial pulse for basin subsidence (McKenzie, 1978). Subsequent slow cooling of the crust and upper mantle caused an additional subsidence. As factors leading to the subsidence of the crust, there are also considered the density increase of the lower crust and its fracture in the result of: phase transitions of gabbro – eclogite, processes of ‘‘basification’’, subcrustal erosion, thermoelastic compression, etc. Negative mantle anomalies, revealed over these basins, are caused by thermal density decrease of the asthenosphere. If the density decrease due to mantle heating is assumed to occur below the Moho to a depth of 200 km, the anomalies with 150 to 200 10 5 ms 2 amplitude [as it has been shown on the example of Western Mediterranean and Pannonian Basins by Yegorova et al. (1997b)] are caused by a density decrease of 0.045 103 kg m 3. This corresponds with the calculated thermal density decrease values of 0.01 103 kg m 3 for each 100 C. The maximal depth of this layer is evaluated as 200 km. The choice of the lower depth level is based on the results of seismotomographic study (Spakman et al., 1993), showing that the strongest anomalies are located above this level. The compensation of these basins is realized as follows. Mass deficit of sediments is overcompensated by strong Moho uplift (Fig. 8), the full mass compensation is reached on the depth level in the upper mantle, more probably related to the bottom of the asthenosphere, due to thermal expansion of the asthenosphere substance.
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(3) Basins associated with the suture zones and located at the western and southern margins of the East European Platform. In the transition zone from Western Europe to the EEP, marked by the Polish Trough on the surface, a positive mantle anomaly of 50 10 5 ms 2 amplitude has been revealed. This anomaly coincides with the high-velocity inhomogeneity in the upper mantle recognised from seismic data (Marquering and Snieder, 1996). It can be explained by a cold lithosphere block sunk and frozen in the upper mantle due to the complex evolution of tectonic regimes in the suture zone between Variscan West Europe and Precambrian East European Platform (Yegorova and Starostenko, in press). However, recent interpretation of seismic data on the LT-7 profile has the Moho (M1) below the TTZ at 35 km, whereas a second reflector at the depth 50 –55 km may represent a ‘‘second Moho’’ (M2) (Guterch et al., 1994). These authors suggest the existence of a crust –mantle layer in the depth interval between M1 and M2. Basin modelling of sediment accumulation due to rifting in the Polish Trough (Dadles et al., 1995) indicates that in the central part of the TTZ, thickness of the crust has to be about 40 km. These authors suggest that Permo-Mesozoic basin development may be related partly to the intrusion of mantle material into and densification of the lower crust (Dadles et al., 1995). That might be another explanation of positive residual gravity anomaly revealed below the TTZ due to our regional gravity analysis. Strong positive mantle anomaly of 150 10 5 ms 2 amplitude has been also evident over the South Caspian Basin. The regional character of this anomaly is confirmed by its extension over the Kura Depression and Caucasus orogen. In contrast to the PeriCaspian Depression, the evolution of the South Caspian Depression is well under way. It is considered (Zonenshain and Le Pichon, 1986; Nikishin et al., in press) that the formation of the South Caspian Basin has happened in the Upper Jurassic – Lower Cretaceous as result of back-arc extension. Rapid basement subsidence has happened in the Middle Pliocene – Pleistocene. At present, the basin belongs to the region of compressional regime (Pristley et al., 1994; Kopp and Shcherba, 1998). The isostatic equilibrium in the South Caspian Basin is disturbed, since the sharp subsidence of the sedimentary basin bottom is accompanied only by a
slight uplift of the Moho. Revealed strong positive mantle anomaly assumes a subcrustal layer with increased density which underlies both the South Caspian Depression and the Caucasus mountain ranges. Therefore, we suggest that the upper-mantle structures beneath the depression and adjacent orogenic structures of the Caucasus have been inherited from a common pre-alpine source. These relatively heavy portions in the upper mantle under the South Caspian Depression and Caucasus may be interpreted as a result of collision between the Eurasian and Arabian plates. Thus, from the west and south, the EEP is framed by series of high-density domains, located in the uppermost mantle and lower crust, which have been formed probably as a result of collision and accretion on the junction zone of Precambrian EEP with young plates. The location, structure and role of the southern margin of the EEP in Late Paleozoic rifting and in governing the tectonic style and distribution of subsequent (primarily compressional) reactivation are the targets of ongoing national and international programmes of geophysical data acquisition on the EEP (Stephenson, 1997). Vast positive residual anomaly has been also revealed over the Bay of Biscay and adjoining part of the North Atlantic. The value of this anomaly is 50 10 5 ms 2 amplitude (Fig. 7) against the background of mantle anomalies of 100 150 10 5 ms 2. This residual anomaly may be caused by nonbalanced state of the crust of oceanic type. Namely, strong mass deficiency in the Bay of Biscay due to large thickness of both sea-water layer (>4 km) and sedimentary cover (up to 6 km) is not balanced by the Moho uplift up to the depth of 16 km. For reaching the total compensation of the crust, it is necessary to introduce additional high-density domains not only in the crystalline crust, but also in the upper mantle. Mantle origin of this anomaly is confirmed by the highvelocity domain found out in the upper mantle to the west from Iberian peninsula and Bay of Biscay (Marquering and Snieder, 1996). The presence of heavy block in the upper mantle below the Bay of Biscay corresponds with nonactivated regime of the upper mantle of this part of Northern Atlantic. According to Cermak (1982, 1993), the calculated temperatures in the upper mantle at 40 km depth are extremely low ( < 500 C), comparable with that below the Adriatic
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plate and the EEP. To a certain extent, it might be related with the most ancient lithospheric block that occurred in this part of Northern Atlantic (Marquering and Snieder, 1996). Thus, the positive residual anomaly, revealed in the Northern Atlantic to the west from Iberian peninsula and Bay of Biskay, confirms the presence of a cold, high-density/velocity lithospheric block of ancient consolidation traced in the upper mantle down to the depth of approximately 140 km.
6. Conclusions A large-scale 3D density model for the whole European continent has been constructed. The model, approximated by two regional layers—sedimentary cover and the consolidated crust (an offshore seawater layer was added), is based on a simplified velocity model (Giese and Pavlenkova, 1988; Hurtig et al., 1992). The gravity effects of this model were calculated with the program developed by Starostenko and Manukyan (1987), designed to solve the 3D direct problem, taking into account the spherical form of the Earth. Residual gravity anomalies, obtained by subtracting the gravity influence of the crustal model from the observed field, reach amplitudes of a few hundreds 10 5 ms 2. Assumed mantle origin of residual anomalies is supported by their clear correlation with both velocity heterogeneities of the subcrustal upper mantle obtained by seismic tomography and geothermal data. The sedimentary basins in Europe are manifested differently on the map of mantle gravity anomalies (Fig. 7). Taking into account its character, specific features of the evolution of the basin and type of the isostatic compensation, the sedimentary basins in Europe might be classified into three groups. (1) Deep sedimentary basins located in the Precambrian East European Platform and its northern and eastern margins—the Peri-Caspian, Dnieper– Donets, Barents Sea Basins and Fore – Ural Trough. These basins are characterized by great thickness of sediments, slight uplift of the Moho and by an increase of average velocity/density in the consolidated crust (as a result of the reduction of the ‘‘granitic’’ layer). No strong mantle gravity anomalies have been revealed here. The mass compensation of this very deep basins (up to 20 km) is reached on the Moho boundary due to
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strong ‘‘basification’’ of the crust, resulting in a rapid increase of velocity and density up to 6.8 km s 1 and 2.9 103 kg m 3, respectively, and a slight uplift of the Moho up to the depth of 36 km. (2) Basins of active rift and taphrogenic regime located mainly in Western Europe and Western Mediterranean. Negative mantle anomalies of 150 200 10 5 ms 2 have been obtained in the Pannonian Basin, the Tyrrhenian Sea, the Balearic Basin, the Polish – German Basin, the North Sea and Rhine Graben, which are characterized by thin (15 –28 km) crust of transitional and oceanic type. The lithosphere beneath these basins is characterized by lowvelocity domains and high-temperature regime caused by the rise of the asthenosphere up to the subcrustal depths. Negative mantle anomalies obtained over these basins are caused by thermal expansion of the material in the shallow asthenosphere. The compensation of these basins is reached in the upper mantle (probably on the bottom of the asthenospheric layer) due to thermal expansion of the asthenosphere substance. (3) Sedimentary basins located over suture zones at the western and southern margins of the EEP and marked by positive mantle anomalies of 50 – 150 10 5 ms 2 amplitude. Such anomaly has been revealed in the transition zone from Phanerozoic Western Europe to Precambrian East European Platform over the Polish Trough and correlates well with highvelocity domain in the upper mantle. Over the South Caspian Basin, the strongest positive anomaly (up to 200 10 5 ms 2), with its extension to the Caucasus, has been also obtained. Thus, the lithosphere of the East European Platform is surrounded from west to south by a series of upper mantle domains of high density formed as a result of collision and accretion of different lithospheric plates. Thus, the continent-wide 3D gravity modelling, based on the results of generalization of seismic data, may serve as unique structure basis for getting additional important information (revealing the main features of similarity and distinction) on the structure and evolution of main tectonic elements of the study region.
Acknowledgements This paper is a result of collaboration within the framework of the EUROPROBE project GEORIFT.
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The scientific investigations leading to this paper were undertaken as part of INTAS Grants 93-3346 and 970743, and the financial support of INTAS is gratefully acknowledged. We extend our appreciation to Dr. R.A. Stephenson (Free University, Amsterdam, The Netherlands), Dr. P. Szafian (Geomega, Budapest, Hungary) and an anonymous reviewer for their constructive comments and useful suggestions on improving the English of the manuscript. The authors thank Dr. V.G. Kozlenko (Institute of Geophysics, Kiev, Ukraine) and Prof. N.I. Pavlenkova (Institute of Physics of the Earth, Moscow, Russia) for precious help and fruitful discussions. We thank also the participants of the 9th Workshop of the ILP Task Force ‘‘Origin of Sedimentary Basins’’ and EUROCONFERENCE-EUROPROBE held in Oliana (Spain), September 1998, for the interest showed to the topic and stimulating discussions. References Artemjev, M.E., Kaban, M.K., Kucherinenko, V.A., Demyanov, G.V., Taranov, V.A., 1994. Subcrustal density inhomogeneities of Northern Eurasia as derived from the gravity data and isostatic models of the lithosphere. Tectonophysics 240, 249 – 280. Artyushkov, E.V., 1992. Role of crustal stretching in subsidence of the continental crust. Tectonophysics 215, 187 – 215. Artyushkov, E.V., 1993. Physical Tectonics. Nauka, Moscow (in Russian). Artyushkov, E.V., Beer, M.A., 1983. About role of vertical movements in formation the troughs on the continental crust in fold belts. Izv. Akad. Nauk SSSR, Ser. Geol. 9, 25 – 52. Babuska, V., Plomerova, J., 1992. The lithosphere in Central Europe—seismological and petrological aspects. Tectonophysics 207, 141 – 163. Babuska, V., Plomerova, J., Sileny, J., 1984. Large-scale oriented structures in the subcrustal lithosphere of Central Europe. Ann. Geophys. 2 (6), 649 – 662. Balavadze, B.K., Belokurov, V.S., Garkalenko, I.A., Malovitski, Ya.P., Mindeli, P.Sh., Moskalenko, V.N., Neprochnov, Yu.P., Neprochnova, E.A., Starshinova, E.A., Chekunov, A.V., 1975. Physical properties of the rocks. Black Sea Basin. In: Bulanze, Yu.D., Muratov, M.V., Subbotin, S.I., Balavadze, B.K. (Eds.), Earth’s Crust and Evolution of Black Sea Basin. Nauka, Moscow, pp. 254 – 269 (in Russian). Belousov, V.V., 1990. Tectonosphere of the Earth: upper mantle and crust interaction. Tectonophysics 180, 139 – 183. Belousov, V.V., Pavlenkova, N.I., 1984. The types of the Earth’s crust. J. Geodyn. 1, 167 – 183. Belousov, V.V., Pavlenkova, N.I., 1989. Types of the earth’s crust in Europe and the North Atlantic. Geotectonika 3, 3 – 34 (in Russian).
Blundell, D.G., 1990. Seismic images of continental lithosphere. J. Geol. Soc. (London) 147, 895 – 913. Blundell, D., Freeman, R., Mueller, St. (Eds.), 1992. A Continent Revealed. Cambridge Univ. Press, Cambridge. Cermak, V., 1982. Crustal temperature and mantle heat flow in Europe. Tectonophysics 83, 123 – 142. Cermak, V., 1993. Lithospheric thermal regime in Europe. Phys. Earth Planet. Inter. 79, 179 – 193. Dadles, R., Narkiewicz, M., Stephenson, R.A., Visser, M.T.M., van Wees, J.-D., 1995. Tectonic evolution of the Mid-Polish Trough: modelling implications and significance for central European geology. Tectonophysics 252, 179 – 195. Dortman, N.B. (Ed.), 1976. Physical Properties of the Rocks and Mineral Resources (Petrophysics). Nedra, Moscow (in Russian). Fuchs, K., 1979. Structure, physical properties and lateral heterogeneities of the subcrustal lithosphere from long-range deep seismic sounding observations on continents. Tectonophysics 56, 1 – 15. Giese, P., Pavlenkova, N.I., 1988. Structural maps of the earth’s crust in Europe. Izv. Akad. Nauk SSSR, Fiz. Zemli 10, 3 – 14 (in Russian). Giese, P., Prodehle, G., Stein, A. (Eds.), 1976. Explosion Seismology in Central Europe, Methods and Results, Berlin. Granser, H., 1987. Three-dimensional interpretation of gravity data from sedimentary basins using an exponential density – depth function. Geophys. Prospect. 35, 1030 – 1041. Guterch, A., Grad, M., Materzok, R., Rajduel, J., Perchuc, E., Toporkiewicz, S., 1984. Structure of the Earth’s crust in the contact zone of the Paleozoic and Precambrian Platforms and the Carpathian Mts in Poland. Acta Geophys. Pol. XXXII, 25 – 41. Guterch, A., Grad, M., Janik, T., Materzok, R., Luosto, U., Yliniemi, E., Luck, E., Shulze, A., Forster, K., 1994. Crustal structure of the transition zone between Precambrian and Variscan Europe from new seismic data along LT-7 profile (NW Poland and eastern Germany). C. R. Acad. Sci. Paris Geophys. 319, 1489 – 1494. Hurtig, E., Cermak, V., Haenel, R., Zui, V. (Eds.), 1992. Geothermal Atlas of Europe. Hermann Haack Verlagsgesellschaft, Geographisch Kartographische Anstalt Gotha, Potsdam. Kinsman, D.J.J., 1975. Rift basins and peculiarities of sediment accumulation in the conditions of sagged margins of continents. In: Fisher, G., Judson, S. (Eds.), Petroleum and Global Tectonics. Princeton Univ. Press, Princeton, NJ. Kopp, M.L., Shcherba, I.G., 1998. Caucasus basin in the Paleogene. Geotectonika 2, 29 – 50 (in Russian). Marquering, H., Snieder, R., 1996. Shear-wave velocity structure beneath Europe, the northeastern Atlantic and western Asia from waveform inversions including surface-wave mode coupling. Geophys. J. Int. 127, 283 – 304. McKenzie, D.P., 1978. Some remarks on the development of sedimentary basins. Earth Planet. Sci. Lett. 40, 25 – 32. Moskalenko, V.N., 1975. Earth’s crust of Mediterranean Sea. In: Belyaevsky, N.A. (Ed.), Earth’s Crust of Continent Margins and Inner Seas. Nauka, Moscow, pp. 122 – 141 (in Russian). Nikishin, A.M., Ziegler, P.A., Stephenson, R.A., Cloetingh, S., Furne, A.V., Fokin, P.A., Ershov, A.V., Bolotov, S.N., Korotaev, M.V., Alekseev, A.S., Gorbachev, V.I., Shipilov, E.V., Lankreijer, A., Bembinova, E.Yu., Shalimov, I.V., 1996. Late Pre-
T.P. Yegorova, V.I. Starostenko / Tectonophysics 346 (2002) 5–21 cambrian to Triassic history of the East European Craton: dynamics of sedimentary basin evolution. Tectonophysics 268, 19 – 46. Nikishin, A.M., Cloetingh, S., Brunet, M.-F., Bolotov, S.N., Ershov, A.V., Stephenson, R.A., 1999. Scythian Platform, Caucasus and Black Sea region: mesozoic – cenozoic tectonic history and dynamics. In: Crasquin-Soleau, S., Barrier, E. (Eds.), Peri-Tethys Memoir 3, Mum. Mus. Natn. His. Nat., Paris, vol. 177 (in press). Panza, G.F., Mueller, St., Calcagnile, G., 1980. The gross features of the lithosphere- and asthenosphere system in Europe from seismic surface waves and body waves. Pure Appl. Geophys. 118, 1209 – 1213. Pavlenkova, N.I., Romanyuk, T.V., 1991. Seismo-gravity modelling for the crust of Siberia. Geol. Geophys. 1, 34 – 42. Podoba, N.V., Ozerskaya, M.A. (Eds.), 1975. Physical Properties of Sedimentary Cover of East-European Platform. Nauka, Moscow (in Russian). Pristley, K., Baker, C., Jackson, J., 1994. Implications of earthquake focal mechanism data for the active tectonics of the south Caspian basin and surrounding regions. Geophys. J. Int. 118, 11 – 141. Rusakov, O.M., 1990. The thickness and density of sedimentary cover in the Indian Ocean. Geophys. J. 3 (8), 390 – 399. Sclater, J.G., Christie, P.A., 1980. Continental stretching: an explanation of the post Mid-Cretaceous subsidence of the Central North Sea basin. J. Geophys. Res. 85, 3711 – 3739. Spakman, W., van der Lee, S., van der Hilst, R., 1993. Travel-time tomography of the European-Mediterranean mantle down to 1400 km. Phys. Earth Planet. Inter. 79, 3 – 74. Starostenko, V.I., Manukyan, A.G., 1987. Gravimetric problems for spherical celestial bodies and their application to the study of the Nectaris Basin Mascon on the Moon. Boll. Geofis. Teor. Appl. 29 (115), 201 – 219. Stephenson, R.A., 1997. Intraplate tectonics and basin dynamics of the EEP: Prypyat – Dnieper – Donets Basin review—problems and prospects. Geophys. J. 17, 170 – 175.
21
Stephenson, R.A., Wilson, M., de Boorder, H., Starostenko, V.I. (Eds.), 1996. EUROPROBE: Intraplate Tectonics and Basin Dynamics of the Eastern European Platform. Tectonophysics Spec., vol. 268. Wilson, M., Lyashkevich, Z.M., 1996. Magmatism and the geodynamics of rifting of the Pripyat – Dnieper – Donets rift, Eastern European Platform. Tectonophysics 268, 65 – 81. Wood, R., Barton, P., 1983. Crustal thinning and subsidence in the North Sea. Nature 302, 134 – 136. Yegorova, T.P., Starostenko, V.I., 1999. Large-scale 3-D gravity analysis of the lithosphere below the transition zone from Western Europe to East-European Platform. Tectonophysics 313, 41 – 58. Yegorova, T.P., Kozlenko, V.G., Pavlenkova, N.I., Starostenko, V.I., 1995. 3-D density model for the lithosphere of Europe: construction method and preliminary results. Geophys. J. Int. 121, 873 – 892. Yegorova, T.P., Kozlenko, V.G., Starostenko, V.I., 1997a. Density heterogeneities of the European upper mantle inferred from 3-D large-scale gravity modelling. In: Fuchs, K. (Ed.), Upper Mantle Heterogeneities from Active and Passive Seismology. Kluwer Academic, Amsterdam, pp. 249 – 256. Yegorova, T.P., Starostenko, V.I., Kozlenko, V.G., Pavlenkova, N.I., 1997b. Three-dimensional gravity modelling of the European Mediterranean lithosphere. Geophys. J. Int. 129, 355 – 367. Ziegler, P.A., 1983. Crustal thinning and subsidence in the North Sea. Nature 304 (5926), 561. Ziegler, P.A., 1990. Geological Atlas of Western and Central Europe, 2nd edn. Shell International Petroleum Maatschappij, Geol. Soc. London, Elsevier, Amsterdam. Zielhuis, A., Nolet, G., 1994. Shear-wave velocity variations in the upper mantle beneath central Europe. Geophys. J. Int. 117, 695 – 715. Zonenshain, L.P., Le Pichon, X., 1986. Deep basins of the Black Sea and Caspian Sea as remnants of Mesozoic back-arc basins. Tectonophysics 123, 181 – 211.