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Tectonophysics 382 (2004) 51 – 84 www.elsevier.com/locate/tecto Slab behaviour and its surface expression: new insights from gravity modelling in the SE-Carpathians Blanka Sperner a,*, Dumitru Ioane b, Robert J. Lillie c a Geophysical Institute, Karlsruhe University, Hertzstr. 16, D-76187 Karlsruhe, Germany b Faculty of Geology and Geophysics, University of Bucharest, Bucharest, Romania c Department of Geosciences, Oregon State University, Corvallis, OR 97331-5506, USA Received 21 March 2003; accepted 16 December 2003 Abstract We use lithosphere-scale gravity models to calculate gravity anomalies resulting from oceanic subduction, continental collision, slab steepening, delamination, and break-off. Local isostasy was assumed for determining vertical movements caused by mass changes related to these tectonic processes. Our results show that subduction is accompanied by basin subsidence on the upper plate caused by the heavy lithospheric root of the subducting slab. The basin evolution goes parallel with the slab evolution and shows considerable modifications when the processes at depth change (slab steepening, delamination, break-off). Characteristic gravity anomaly curves were acquired for the different tectonic scenarios. These curves together with other data (e.g. basin evolution on the upper and the lower plate) were used for the reconstruction of the tectonic evolution of the SECarpathians which includes Tertiary subduction and collision followed by slab steepening and delamination. D 2004 Elsevier B.V. All rights reserved. Keywords: Gravity anomalies; Subduction; Isostasy; Vertical movements; Basins; Carpathians 1. Introduction Subduction of oceanic lithosphere induces largescale mass transfers as heavy lithospheric mantle replaces lighter asthenosphere. The density contrast between the two materials is high enough to produce a slab-pull force thought to be an important driving mechanism for plate tectonics (e.g. Conrad and Lithgow-Bertelloni, 2002). It can be expected that * Corresponding author. Fax: +49-721-71173. E-mail address: [email protected] (B. Sperner). 0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2003.12.008 this heavy slab also has an effect on the surface gravity field, not only at the plate boundary, but also above the upper-plate region underlain by the heavy lithospheric root of the subducted slab. Another consequence is vertical movement due to isostatic adjustments. We use gravity modelling to study these effects in general and then to study an example of young subduction in the SE-Carpathians. Our type of gravity modelling differs from standard gravity modelling in two ways: it includes the tectonic evolution of the study area and it considers isostatic adjustments. The inversion of gravity data yields non-unique solutions. This means that the integration of informa- 52 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 tion from other fields like geology and geophysics is mandatory. By considering the tectonic evolution of a region, gravity modelling can be used to test different evolution models. This has the additional advantage that the density distribution at depth is explained by the tectonic evolution. No ‘‘mysterious’’ bodies with density contrasts that best fit the gravity data are presumed (as is sometimes the case in standard gravity modelling). For isostatic balancing, local (Airy) isostasy was assumed for the different stages of tectonic evolution. The resulting vertical movements are assumed to be followed by erosion (of uplifted areas) and sedimentation (in subsided areas). Previous gravity modelling studies concentrated on the crustal evolution of continental collision zones (e.g. Lillie, 1991; Lillie et al., 1994). We consider greater depths and thus include the influence of the lithospheric root created by the subducted slab. We first study processes related to the evolution of the slab in general, namely subduction, steepening, delamination and break-off. Then, we apply this method to the SE-Carpathians, test different models for the tectonic evolution of the region, and compare the results with measured data. In the SE-Carpathians, Miocene oceanic subduction was followed by continental collision at ca. 10 Ma. Post-collisional shortening was small (soft collision), so that the structures and basins which developed during subduction are well preserved. Intermediatedepth seismicity (70 – 200 km) is concentrated in a nearly vertical column (Bonjer et al., 2000; Oncescu and Bonjer, 1997), suggesting that the subducted slab is still hanging beneath the SE-Carpathians (Sperner et al., 2001). Data from seismic experiments (refraction, tomography) provide detailed information about the lithospheric structure (Hauser et al., 2001; Martin et al., 2001; Wortel and Spakman, 2000). 2. Modelling approach Both free-air and Bouguer gravity anomaly data are used in this paper. 2.1. Free-air anomaly (FAA) The free-air correction accounts for gravitational effects related to the latitude and elevation of gravity stations with masses between the station and the reference level being preserved. FAA data can give some indication of the degree to which a region approaches isostatic compensation. 2.2. Bouguer anomaly (BA) The Bouguer correction accounts for the gravitational effects of masses between gravity stations and the reference level. For stations on land, this is often performed by subtracting the gravity effect of an infinite slab with a specific density (typically 2670 kg m 3). For stations at sea, the opposite is done by virtually filling the ocean with crustal material, i.e. by adding the effects of an infinite slab with a density equal to the difference between water and crustal material. Therefore, oceans are always characterized by positive BA values. On land, BA data often show a strong correlation with the depth to the Moho because the density contrast between crust and mantle is large and topographic effects have been accounted for by the Bouguer correction. Thus, undulations in the Moho have large gravity effects at the surface. 2.3. Density contrasts We use density contrasts because gravity anomalies, and especially the shape of the anomaly curves, are the result of lateral density changes which depend on density contrasts and not on absolute density values. Additionally, the use of density contrasts has the advantage that the depth-dependent increase of density is inherent (density increases uniformly in all units so that the contrast remains the same). 2.4. Density contrast values We used the crust as a reference level and assigned it a relative density of zero (Table 1). The density contrast between crust and lithospheric mantle was set to 300 kg m 3 (Bielik, 1999; Lillie et al., 1994). For the density contrast between lithospheric mantle and asthenosphere, values in the literature range from 0 to 200 kg m 3 (e.g. Buiter et al., 1998; Jones et al., 1996; Lachenbruch and Morgan, 1990; Lillie et al., 1994; Royden and Karner, 1984). Estimations from seismic tomography data, which use the velocity – density relationships of, e.g. Birch (1964) or Cermak B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Table 1 Density contrasts used in the gravity models Topography (young sediments) Topography (crystalline crust) Water Young sediments Continental crust Oceanic crust, at depth < 30 km Oceanic crust, subducted below 30 km Lithospheric mantle Asthenosphere [kg m 3] Symbol 2470 2670 1640 200 0 0 300 300 270 qTsed qTc qW qSed qCc qOc qEclo qLM qA Values are relative to crustal material (qCc, qOc). et al. (1990), lie in most cases between 10 and 50 kg m 3. Therefore, we assume an average value of 30 kg m 3 (i.e. 270 kg m 3 relative to the crust). The density distribution in the down-going slab is depthand time-dependent due to the eclogitisation of the oceanic crust at greater depths. Eclogitisation is, amongst other things, controlled by temperature and is thus dependent on the subduction velocity. For the sake of simplicity, in our models we neglect the influence of subduction velocity and we assume that the oceanic crust undergoes eclogitisation at a depth of 30 km. Thus, we set the density for the oceanic crust at depths larger than 30 km to the same value as the lithospheric mantle (i.e. a density increase of 300 kg m 3). Depth-dependent density increase in the slab (crust and lithospheric mantle) is already taken into account through the use of relative instead of absolute densities. This means that density in the slab increases in parallel with the density increase in the surrounding asthenospheric material. 2.5. Local isostasy For each step, the model was kept in local isostatic equilibrium (Airy isostasy). This means that mass changes are compensated directly beneath the mass anomaly. For the resulting vertical mass movements, the shear strength of the material is assumed to be zero (no rigidity). This simplification makes the calculations much easier and faster in comparison with flexural bending models and thus allows the testing of a broader suite of models. The Airy approximation is further justified because previous studies showed that flexural rigidities in the study area are low (low effective elastic thickness of the lithosphere: a few 53 kilometres according to Artyushkov et al. (1996) and 14 km according to Matenco et al. (1997)). Fig. 1 shows two examples for local isostasy, one with mass deficit (crustal thickening; Fig. 1b) and one with mass excess (lithospheric thickening; Fig. 1d) in the subsurface. In both cases, Bouguer and free-air anomalies are the same for the uncompensated stage (Fig. 1b,d). The broad free-air anomalies indicate that isostatic equilibrium is not realized and the Bouguer anomalies reveal the uncompensated mass deficit and mass excess for the two models. Note that the gravity anomalies of the two uncompensated models are different by an order of magnitude, the crustal thickening effect (Fig. 1b) being far greater than that of the lithospheric thickening (Fig. 1d). Reasons for the smaller anomalies in the lithospheric thickening model are the greater depth of the lithospheric root and the lower density contrast between lithospheric mantle and asthenosphere. Isostatic surface effects are uplift in the case of mass deficit (Fig. 1c) and subsidence in the case of mass excess (Fig. 1e). In the crustal thickening model, the Bouguer anomaly is slightly reduced (due to the rise of the Moho), while the freeair anomaly shows a typical ‘‘edge effect’’: the uplifted area (topography) leads to a sharp increase in the gravity anomaly, while the negative effect of the crustal root is, due to its larger depth, of smaller amplitude, but larger wavelength. Isostatic equilibrium is indicated when the areas under the FAA curve (positive above zero line, negative below zero line) sum to zero, which is the case in the example in Fig. 1c. In the lithospheric thickening model, isostatic compensation leads to subsidence and thus the freeair anomaly flips to negative values in the middle part of the model (Fig. 1e). The free-air anomaly is again characterized by negative– positive couples typical of the edge effect (see above). The Bouguer anomaly curve shows only little changes due to the only slightly deeper Moho. 2.6. Sedimentation During the subduction and collision process, the sediment volume on the two continental margins was kept constant. Newly formed basins, like those on the continental margin of the upper plate, were filled with sediments when there was a nearby supply area with topographic relief. Continental collision provides such 54 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 1. Effects of local isostasy due to crustal and lithospheric thickening respectively. (a) Starting model with no gravity anomalies. (b, c) Crustal thickening results in a mass deficit in the lithospheric mantle (due to the lower density of the crust) which is compensated by uplift (heavy asthenospheric material is added at the bottom). The Bouguer anomaly is negative, with a slight decrease after uplift. Without uplift, the free-air anomaly is the same as the Bouguer anomaly. The large negative free-air anomaly in (b) indicates that the model is not in isostatic equilibrium. After isostatic adjustment, smaller free-air anomalies remain due to the edge effect caused by the different depths of the mass anomalies (see text for details). (d,e) During lithospheric thickening, asthenosphere is replaced by higher-density lithospheric mantle. The resulting mass excess is compensated by subsidence. The Bouguer anomaly is positive in both models. The initial free-air anomaly is also positive, indicating that the model is not in isostatic equilibrium. Subsidence leads in parts to negative free-air anomalies, again showing an edge effect. Note the different scales of the gravity plots in (b,c) and (d,e). B. Sperner et al. / Tectonophysics 382 (2004) 51–84 uplifted areas, so that synorogenic sediment supply was assumed. Basins were filled to zero-level, but topographic heights were kept constant to provide information about the maximum topography possible for the respective tectonic evolution (i.e. no erosion and thus no volume constancy). 2.7. Software Modelling was done with the commercial software package GM-SYS (v. 4.04 and 4.6) of Northwest Geophysical Associates (NGA; http://www.nga.com/). 3. Models The starting configuration for our models consisted of two continents with 30-km-thick crust separated by an ocean with 8-km-thick crust (Fig. 2a). We assumed an old, thermally equilibrated lithosphere with a horizontal lithosphere –asthenosphere boundary at 120-km depth. This implies a thickness of the oceanic lithosphere of 116.6 km which correlates with an age of 80 Ma (e.g. Turcotte and Schubert, 2002). The width of the transition zone between continent and ocean was set to 100 km and the thickness of the water – sediment package was limited to 10 km, with the maximum sediment thickness at the continent – ocean boundary. The compensation level was set to a depth of 550 km, thus providing enough space for the subducting slab. Models are shown on a lithospheric scale (no vertical exaggeration) and on a crustal scale (vertical exaggeration of 4:1). For easier description, the orientation of the cross-sections is assumed to be west –east and the subduction direction is towards the west. All models have a 2D-geometry, i.e. they extend to infinity in the dimension perpendicular to the profiles shown. 3.1. Oceanic subduction 3.1.1. Local isostasy Oceanic subduction leads to the development of a lithospheric root beneath the upper plate (Fig. 2). The mass excess at depth results in basin subsidence on the upper plate. Due to the origin of this basin, a subsurface load, we refer to it as a load 55 basin. Continued subduction leads to an increase in basin width; the basin depth also increases as long as the lithospheric root continues to thicken. From a certain point on, when the maximum thickness of the lithospheric root is reached, the basin will only widen without deepening (Fig. 3). The maximum thickness of the lithospheric root (hLR) depends on the thickness of the oceanic lithosphere (hOL) and on the dip angle (a) of the slab (Fig. 3a): hLR ¼ hOL ðcosðaÞÞ1 ð1Þ The lithospheric root with density qLM replaces asthenospheric material with density qA, so that the mass excess DmLR can be calculated: DmLR ¼ hLR ðqLM qA Þ ð2Þ For the mass calculations, the two other dimensions of the mass volume are constant and thus are assumed to be 1 m each; they have no influence on the result, so they are not shown in the formulas. In the same way, the mass deficit of the water-filled basin on the upper plate can be calculated (again, asthenospheric material is replaced because it disappears out of the model at its lower boundary). Subsidence hw of a water-filled basin causes a mass deficit Dmw of Dmw ¼ hw ðqw qA Þ ð3Þ For isostatic equilibrium, the sum of the mass anomalies must equal zero: hLR ðqLM qA Þ hw ðqA qw Þ ¼ 0 and thus the basin depth can be calculated as hw ¼ hLR ðqLM qA ÞðqA qw Þ1 ð4Þ The maximum basin depth then is given by substituting Eq. (1) into Eq. (4) hw ¼ hOL ðcosðaÞÞ1 ðqLM qA ÞðqA qw Þ1 ð5Þ Inserting the values in Table 1 for the density contrasts, the formula can be simplified to hw ¼ hOL ðcosðaÞÞ1 30=1910 56 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 2. Gravity anomalies (BA, Bouguer anomaly, FAA, free-air anomaly) and surface effects of oceanic subduction. The heavy lithospheric root causes subsidence in the upper plate (load basin). Initial ocean width is 700 km, slab dip is 30j; density contrasts are given in Table 1 (V.E., vertical exaggeration). B. Sperner et al. / Tectonophysics 382 (2004) 51–84 57 Fig. 3. Basin evolution on the upper plate during subduction (same configuration as in Fig. 2). (a) With ongoing subduction the basin increases in width and depth until a certain stage is reached (stage 2). From then on the thickness of the lithospheric root does not increase any more (hLR2 = hLR3, maximum thickness of the lithospheric root), so that the basin only widens while maintaining constant depth. (b) Uppermost part of (a), with tenfold vertical exaggeration (V.E.). (c) Basin depth increases with increasing convergence until the maximum depth is reached. Sedimentary fill produces a basin about four times deeper than with water fill (caused by the lower density contrast between sediment and asthenosphere compared to the one between water and asthenosphere). 58 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 A similar relationship exists for the maximum depth of a sediment-filled basin (hSed) with density qSed: hSed ¼ hOL ðcosðaÞÞ1 ðqLM qA ÞðqA qSed Þ1 ¼ hOL ðcosðaÞÞ1 30=470 ¼ 4:06hw ð6Þ Thus, sedimentary infill instead of water increases the depth of the basin by a factor of about 4 due to the lower density contrast between young sediments and asthenosphere (470 kg m 3) in comparison to the contrast between water and asthenosphere (1910 kg m 3). Fig. 3b shows the increase of basin depth with ongoing subduction until it reaches its maximum depth, which in our model is 2.1 km for water fill and 8.6 km for sediment fill for a slab dip a of 30j and a thickness of the oceanic lithosphere hOL of 116.6 km. A steeper slab shows a larger maximum basin depth due to its thicker lithospheric root (Fig. 4a). The limit for the maximum thickness of the lithospheric root (and thus for the maximum basin depth) is controlled by the down-dip length of the subducted slab (otherwise, it would increase to infinity; Fig. 4b). With increasing dip angle, the lithospheric root not only increases in thickness, but it is also located closer to the plate boundary (Fig. 5). In the same way, the upper-plate basin shifts to a more eastern position so that it lies on the continent – ocean transition zone (Fig. 5b,c) or even on the downgoing plate (Fig. 5c,d). Due to the high-mass excess in the subsurface, this region is pulled down to large depths. In the most extreme case of a vertical slab, the basin is completely situated on the down-going plate and has a symmetric shape. 3.1.2. Free-air anomaly The free-air anomaly is significantly affected by the basin subsidence in the upper plate. Subsidence results in a sharp and pronounced negative peak in the FAA curve. The slab itself causes a broad, but only small rise of the FAA data, leading to the typical ups and downs near the basin margin (edge effect; see above). Steeper slabs result in deeper, but narrower basins with a corresponding effect on the negative FAA peak (Fig. 5). Fig. 4. Dependence of basin subsidence on the slab dip angle (same configuration as in Fig. 2 except dip angle). (a) A larger slab dip angle (a2>a1) results in a larger maximum basin depth. The amount of convergence to reach this maximum depth is smaller than for flatter slabs (c2 < c1). (b) Correlation between the slab dip angle and the maximum thickness of the lithospheric root (left axes) and the maximum basin depth (right axes). The upper limit of these two parameters is given by the length of the subducted slab (400 km in the example shown). 3.1.3. Bouguer anomaly The influence of oceanic subduction on the Bouguer anomaly is only small (Figs. 2 and 5). The lithospheric root causes a moderate rise of the BA data, which is partly balanced by the down-warping of the continental crust. 3.2. Continental collision 3.2.1. Local isostasy Continental collision is preceded by the collision of the two sedimentary wedges lying on the margins of B. Sperner et al. / Tectonophysics 382 (2004) 51–84 59 Fig. 5. Different slab dip angles (a) and their influence on gravity anomalies and on load basin location relative to the plate boundary (same configuration as in Fig. 2, with 300 km convergence). (a, b) a = 30j, 45j: The basin is located on continental crust of the upper plate. The transition zone underlies parts of the basin, but the basin axis lies on the continent. (c) a = 60j: The basin axis lies on the transition zone. (d) a = 75j: The basin sits on the plate boundary. The high in the middle of the basin is caused by the down bended, not eclogised (i.e. ‘‘light’’) part of the oceanic crust. The location of the basin is also dependent on the thickness of the slab: a larger thickness shifts the basin axis toward the ocean. the two continents (Fig. 6a – c). Internal shortening of this sediment wedge is assumed to be limited to 150 km which is about the shortening inside the Carpathian accretionary wedge (see below, Section 4.3). When this limit is reached, the whole sediment package starts to thrust over the down-going plate (Fig. 6c – f). In our example, the maximum thickness of the sedimentary wedge is 13.3 km and its maximum height above sea level reaches 2.1 km. Topography of the continental crust caused by crustal 60 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 thickening during collision reaches a maximum level of 2.8 km. The lithospheric root initially causes load basin subsidence in the upper plate (Fig. 6a – c). However, as soon as the continents collide, the crust thickens and uplift begins, first in the accretionary wedge (Fig. 6d) and later in the eastern part of the basin (Fig. 6e). With ongoing underthrusting, the uplifting region propagates towards the west (Fig. 6e –f). At the same time, the lithospheric root also shifts westward, forcing the basin to move in the same direction. Thus, the westward-propagating basin is followed by an uplifting region, similar to a wave that first goes down and then up again. This means that a basin with constant width and geometry migrates westward, but is later uplifted as part of an expanding mountain range. Note that in our model, subsidence of this basin has nothing to do with (back)thrusting and loading by the mountain range; it is a pure drag effect of the lithospheric root. 3.2.2. Free-air anomaly The load basin on the upper plate is characterized by a negative free-air anomaly (Fig. 6c– f), while a positive anomaly develops east of the basin in the uplifted region (Fig. 6d –f). 3.2.3. Bouguer anomaly Closure of the ocean basin and uplift of the sediments on the continental margins removes the oceanrelated positive Bouguer anomaly (Fig. 6a– c). A new positive anomaly develops due to the lithospheric root beneath the upper plate. Crustal thickening during continental collision results in a negative Bouguer anomaly which broadens with ongoing underthrusting (Fig. 6d – f) (see also Lillie, 1991). In the following sections, we will discuss the evolution of the subducted slab after continental collision. We assume a soft continental collision with only 100 km of underthrusting and study the different slab processes which have been proposed to take place after collision (slab steepening, break-off and delamination). A more detailed discussion about the 61 gravity effects of hard continental collision (i.e. collision accompanied by intense shortening and underthrusting) is given by Lillie (1991) and Lillie et al. (1994). 3.3. Slab steepening Numerical modelling by Dvorkin et al. (1993) showed that the hydrodynamic suction between upper and lower plate, induced by corner flow during downwards movement, stabilizes the slab at a constant dip angle. However, several parameters and processes can lead to a steepening of the slab. One important factor is the along-strike length of the slab: for narrow slabs, asthenospheric material can flow around the corners of the slab, thus reducing the hydrodynamic suction near the plate contact and leading to slab steepening and rollback (Dvorkin et al., 1993). Another factor is the end of down-dip motion after continental collision. In this case, the hydrodynamic suction is decreasing because corner flow stops. Thus, the subducted lithosphere starts to steepen under its own weight (for the same reason slab break-off only takes place after continental collision; Yoshioka and Wortel, 1995). To study the surface effects of this important process, we modelled a slab that first subducted at a 20j dip angle and then (after soft continental collision) steepened to 75j. The down-dip length of the slab is 300 km. The other parameters are the same as in the previous models. 3.3.1. Local isostasy As in the previous models, subduction leads to the evolution of a basin on the upper plate. The maximum basin depth is near the western margin of the basin (Fig. 7a). We assumed that the basin was filled with sediments from the surrounding uplifted regions created during continental collision (Fig. 7b). Steepening of the slab (in our example from 20j to 75j) has two effects (Fig. 7c): (a) the lithospheric root is now narrower, but thicker, and its location is further to the east. Thus, the basin axis also shifts towards the east and the basin deepens; (b) the region west of the Fig. 6. Collision of sedimentary wedges and of continents (slab dip: 20j). The sedimentary wedges thrust over each other until a predefined maximum amount of internal shortening is reached (150 km in our models). Afterwards, the sedimentary wedge as a whole thrusts over the down-going plate. Continental collision and underthrusting are accompanied by crustal thickening leading to surface uplift on the upper plate. At the same time, the load basin on the upper plate moves westwards and its eastern parts are uplifted. 62 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 7. Effects of slab steepening after continental collision. (a) Subduction with an angle of 20j was followed by 100 km of continental underthrusting, resulting in basin subsidence on the upper plate and shortening and uplift of the sediments on the continental margins. (b) The basin was filled with sediments which were delivered from the uplifted surrounding regions (caused by continental collision). (c) Steepening of the slab to 75j results in a new basin with an axis located east of the old one, while the region west of it rises above sea level. (d) Same as (c), but without asthenospheric upwelling during slab steepening. Differences between (c) and (d) concerning isostatic movements and gravity anomalies are minor. new basin is now missing the formerly present lithospheric root so that it rises above sea level. As shown in Fig. 7c, we assume that the steepening slab leaves a gap in the lithospheric mantle which is filled by upwelling asthenospheric material. The opening of such a gap is dependent on the coupling between the subducted slab and the overlying lithospheric mantle of the upper plate. Gvirtzman and Nur (1999) assume a decoupling between slab and upper plate beneath the Calabrian arc and the Kuriles. We tested the gravitational effect by doing the same model with and without a gap. The results show only small differences in the calculated isostatic movements and gravity anomalies (Fig. 7d). Effects on other parameters, like heat flow and volcanism, might be more severe. B. Sperner et al. / Tectonophysics 382 (2004) 51–84 3.3.2. Free-air anomaly The FAA curves mainly reflect the surface movements. Negative anomalies exist where basins subside and positive anomalies where uplift above sea level occurs (Fig. 7a– d). 3.3.3. Bouguer anomaly Filling the load basin with sediments results in a negative Bouguer anomaly reflecting the downbended Moho (Fig. 7a,b). The crustal keel of the underthrusted plate is responsible for a second negative BA farther east. After slab steepening, the old load basin rises by 1.16 km and thus reduces the negative Bouguer anomaly, but it cannot balance the mass deficit caused by the loss of the lithospheric root, so that the negative Bouguer anomaly is stronger than before (Fig. 7c). The new load basin reduces the local high in the BA curve (Fig. 7c). 3.4. Slab break-off The fate of subducted lithosphere after continental collision is still under discussion (e.g. Davies, 1980; Davies and von Blanckenburg, 1995; Schott and Schmeling, 1998). Most old subduction and collision zones show no evidence for the existence of subducted material (e.g. no seismicity) so that the slab somehow has to disappear. One possibility is slab break-off followed by gravitational sinking of the detached lithosphere into the deeper mantle (e.g. Levin et al., 2002). Such a process has been proposed for the Alps (von Blanckenburg and Davies, 1995); Wortel and Spakman (1992) assume that break-off beneath the Hellenic arc started at one lateral end of the slab and now propagates laterally along the arc. The slab gap is visible in seismic tomography and it shows that break-off started in that region of the Hellenic arc where continental collision already occurred. Numerical modelling by Yoshioka and Wortel (1995) point in the same direction: slab break-off only takes place when the down-dip movement stops, so that extensional forces inside the slab can increase to values large enough to initiate break-off. The region where slab break-off will preferentially take place is the transition zone between oceanic and continental crust. Bending forces will be concentrated in this region due to the buoyancy of the continental crust which resists the downward pull of the slab. Addi- 63 tionally, this region is already weakened by extensional fault zones which developed during the opening of the ocean. Thus, in our models we assume slab break-off to occur at the continent –ocean boundary after soft continental collision took place. The detached slab sinks into the deeper mantle so that no isostatic or gravitational effect remains. 3.4.1. Local isostasy Removal of the lithospheric root results in uplift of the overlying region. Fig. 8 shows an example where slab steepening took place before break-off, so that the uplift occurred in two steps, first in the western part of the load basin (due to slab steepening; Fig. 8a), later in the eastern part (due to slab break-off; Fig. 8b). 3.4.2. Free-air anomaly Slab break-off reduces the edge effects in the FAA data produced during slab steepening and thus brings the FAA curve closer to the zero line (Fig. 8a,b). Small anomalies remain due to the thickened crust (additional sediments) and the removal of parts of the lithospheric mantle during steepening and break-off. 3.4.3. Bouguer anomaly In comparison with the BA curve for the steep slab (Fig. 8a), slab break-off shifts the curve to slightly more negative values because the additional mass of the subducted lithosphere has been removed (Fig. 8b). Uplift cannot completely balance this effect because the crust had been thickened by sediments, so that a crustal root is left over. 3.5. Delamination In contrast to slab break-off, which occurs along a plane perpendicular to the surface of the lithosphere, delamination takes place along a plane parallel to it. One triggering mechanism for delamination of the lithospheric mantle from the crust is the weight of an overthickened lithospheric mantle caused by shortening in a continental collision zone (e.g. Tibet, Alboran) (Bird, 1978; Seber et al., 1996), another is the slab pull of subducted lithosphere (Girbacea and Frisch, 1998; Sacks and Secor, 1990). In our model, we assume 100-km delamination along the Moho (Fig. 8c). The starting model and thus the preceding evolution is the same as for the break-off model 64 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 (Fig. 8a,b), so that the effects of both processes can be compared. The final position of the slab is vertical. 3.5.1. Local isostasy Delamination removes the whole lithospheric mantle from the westernmost part of the lower plate, so that this region experiences uplift (Fig. 8c). Further eastward, the vertically hanging slab causes subsidence and basin formation. With ongoing delamination, the slab and thus the basin migrate towards the east, while the previously down-warped region west of the slab will rise. In the case that sediments had been deposited into the basin, the topography after delamination will be higher than before (due to the larger crustal thickness). 3.5.2. Free-air anomaly The FAA curve is similar to the steep –slab curve with a shift of the minimum towards the east according to the eastward shift of the slab during delamination (Fig. 8a,c). Additionally, the steeper dip and the greater vertical extent of the slab cause a more pronounced minimum. 3.5.3. Bouguer anomaly In the delaminated region, BA data are shifted to more negative values due to the removal of the heavy lithospheric slab (similar to slab break-off; Fig. 8b). The region above the slab experiences subsidence accompanied by down-bending of the Moho, which also results in a shift to more negative values. The BA curves for break-off and delamination are similar, but the influence of the lithospheric slab is still visible in the easternmost part where the BA data for delamination remain at a higher level (positive values) as compared with the break-off data (Fig. 8e). If sediments were deposited into the basin above the slab, the thickness of the crust increases resulting in an additional down-bending of the Moho and thus a more 65 pronounced negative Bouguer anomaly; the FAA curve changes only slightly (Fig. 8e). 4. Gravity modelling in the SE-Carpathians (Vrancea region) To date, gravity modelling of the Carpathians has mainly covered the northern parts (i.e. Western Carpathians) (Bielik, 1995; Bielik, 1999; Lillie and Bielik, 1992; Lillie et al., 1994; Szafian et al., 1999) or concentrated on large-scale structures (Szafian et al., 1997). None of this work considered the tectonic evolution of a subducting slab (including slab steepening, delamination and break-off) and its influence on vertical movements at the surface as well as on the gravity data. We have chosen the SE-Carpathians as a study area because intermediate-depth seismicity and seismic tomography indicate that cold subducted material is still present in the upper mantle (Martin et al., 2001; Oncescu, 1984; Sperner et al., 2001). Furthermore, abundant data from different fields (e.g. structural geology, sedimentology, paleomagnetics, seismics) are available, allowing the reconstruction of the Neogene tectonic evolution. The Neogene to recent tectonic evolution of the SE-Carpathians has for a long time been debated and discussions are still going on, especially concerning the existence and the type of Miocene subduction (e.g. Girbacea and Frisch, 1998; Girbacea and Frisch, 1999; Linzer, 1996a; Linzer, 1996b; Pana and Erdmer, 1996; Pana and Morris, 1999). Here we present our preferred model for the tectonic evolution, test it using gravity modelling, and compare it with other models for the recent platetectonic configuration. The time scale used in this paper is based on the Central Paratethys stages as compiled by Rögl (1996) (Table 2). Fig. 8. Effects of slab break-off and delamination. (a, b) The subducted and steepened slab (a) breaks off at the ocean – continent boundary and sinks into the deeper mantle (b) so that no gravitational effect remains. As a result, the region above the removed slab raises. The BA curve is shifted to slightly more negative values because lithospheric mantle is replaced by less-dense asthenosphere. The FAA curve approaches the zero line because the high-density slab is removed. (c) Delamination along the Moho causes uplift above the region of the removed slab, while subsidence occurs above the new position of the slab. The BA curve is similar to that for slab break-off, while the FAA curve shows a pronounced minimum at the slab location (strong edge effect). (d) Infill of sediments into the basin above the slab changes the FAA curve only slightly, but results in a stronger negative Bouguer anomaly due to the deepening of the Moho. (e) For better comparison, the BA and FAA curves for the different tectonic situations are plotted in one diagram. 66 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Table 2 Correlation between Mediterranean and Paratethys stages after Rögl (1996) 4.1. Neogene tectonic evolution of the Carpathian region Subduction beneath the Carpathians had begun by late Cretaceous when the intra-Carpathian blocks (North Pannonian block and Tisia – Dacia block in Fig. 9) and the Adriatic microplate moved north(west)ward. In the Alps, continental collision had already started at the end of the Cretaceous (Stampfli et al., 2002), while subduction continued beneath the Carpa- B. Sperner et al. / Tectonophysics 382 (2004) 51–84 67 Fig. 9. Tectonic map of the Carpathian – Pannonian region showing Tertiary – Quaternary structures and the location of intermediate-depth earthquakes in the SE-Carpathians (Vrancea region; geology based on Horvath, 1993). The blow-up (lower map) shows the relationship between topography and main basins in the SE-Carpathians. 68 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 thians. An oceanic embayment in the European foreland, which was still open when collision took place in the Alps, provided additional space for the northeast- and eastwards movement of the intra-Carpathian blocks (Fig. 10). The Miocene movements of these blocks can be reconstructed by taking into consider- Fig. 10. Geodynamic evolution of the Carpathian – Pannonian region during the Miocene as derived from kinematic, sedimentary, and paleomagnetic data (Sperner et al., 1999). Opposite rotation of two crustal blocks (a) resulted in an oblique collision with the European foreland that started in the north and successively propagated towards the southeast and south (b). TDB, Tisia – Dacia block. B. Sperner et al. / Tectonophysics 382 (2004) 51–84 ation data from structural geology, paleomagnetism, and sedimentology (e.g. Balla, 1985; Csontos et al., 1992; Fodor et al., 1999; Linzer et al., 1998; Royden et al., 1983). Continental collision in the Carpathians first took place in the western part (18 Ma), then shifted to the north (13 Ma) and finally reached the SE-Carpathians (9 Ma; Fig. 10). Intermediate-depth seismicity is only recorded from the southeastern corner of the Carpathians (Vrancea region; Fig. 9), so that we assume that the subducted slab is only preserved in this region and is still hanging beneath the collisional orogen. In all other regions of the Carpathians, the slab has already broken off and sunk into the deeper mantle (e.g. Sperner, 1996; Sperner et al., 2001). The fact that the subducted slab, as defined by the intermediate-depth seismicity in the Vrancea region, is not hanging beneath the upper plate or the suture zone, but is located beneath the accretionary wedge and the foredeep (Fig. 11), has led to a number of different interpretations (e.g. Csontos, 1995), which we have tested by means of gravity modelling. 4.2. Available data The Romanian gravity maps were compiled at the Institute of Geology and Geophysics (Bucharest) and 69 are based on 78,000 point observations (mean coverage 0.33 data points per km2) (Ioane and Radu, 1995). The normal gravity was calculated with the International Gravity Formula (Cassinis, 1930) using a reduction density of 2670 kg m 3. The Bouguer correction was done with an infinite slab; for the terrain correction the Schleusener method was used (Schleusener, 1940). Both the Bouguer and the free-air gravity anomaly map are shown in Fig. 12. For topography, the GTOPO30 data were used which are available online at http://edcwww.cr.usgs.gov/ landdaac/gtopo30/gtopo30.html. Sediment thickness and other geological information come from maps and profiles published by the Geological Institute in Bucharest (e.g. Dumitrescu and Sandulescu, 1970; Sandulescu et al., 1978; Stefanescu, 1985). Vertical movement rates have been studied for several decades by using repeated levelling data (Popescu and Lazarescu, 1988; Zugravescu et al., 1998). They show that the SE-Carpathians are uplifting, while the foreland subsides. Fission-track data can be used to estimate the beginning of uplift, but they have to be handled with care because they identify the time of exhumation (erosion) and not of uplift, i.e. the acquired ages are in many cases younger than the beginning of uplift. For the northern and central part of the Eastern Fig. 11. Profile through the Vrancea region in the SE-Carpathians (modified after Radulescu et al., 1976; Stefanescu, 1985; for location, see Fig. 9). Note that the earthquake epicentres are located beneath the accretionary wedge and the foredeep, and not beneath the Miocene suture. Shown on the left side is the estimation of the slab dip angle (a) from the distance between the plate boundary and subduction-related volcanism (240 km) and the depth of the magma generating window (90 km): a = arctan (90/240) = 21j. 70 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 12. Free-air (a) and Bouguer (b) anomaly maps of Romania (Ioane and Atanasiu, 1998). From 78,000 point observations, mean values were calculated for blocks of 5V 7.5Vsize. For details about the data processing, see Ioane and Atanasiu (1998). Profiles run through the seismically active Vrancea region (white star) where the existence of a subducted slab is postulated. For map location, see Fig. 9. B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Carpathians, Sanders et al. (1999) conclude from fission-track data that erosion rates increased in Late Badenian –Sarmatian time (11– 14 Ma) and that this increase is related to the thrusting of the accretionary wedge onto the European foreland. In the southern part, a major erosion phase took place during the Pliocene (1.8 – 5.3 Ma), while at the same time rapid subsidence and sedimentation occurred in the foreland (Sanders et al., 1999). Moho maps have been published in different versions, some based on gravity data with consideration of isostasy and geology (Socolescu et al., 1964), others with additional information from refraction seismic data (Radulescu and Diaconescu, 1998). Using these maps for our gravity modelling would result in circular arguments (gravity data were used to construct a Moho map which was then used to do gravity modelling). We thus rely on the new refraction seismic profiles of Hauser et al. (2001, 2002), which give Moho information that is independent of gravity data, but have the disadvantage that they do not give map-scale information for all of Romania. The depth of the lithosphere – asthenosphere (L/A) boundary has been estimated from heat flow (Radulescu and Diaconescu, 1998), refraction seismic (Enescu, 1992) or seismological and magnetotelluric data (Praus et al., 1990). The results show large variations, e.g. ranging between 100 and 200 km for the East European platform. In our models, we assumed an initial depth of the L/A boundary of 90 km for the whole model. This depth is reported for the Transylvanian basin (western part of our modelling profile) (Radulescu and Diaconescu, 1998) and we assume that the lithosphere of the East European platform had about the same thickness near the continental margin; thickening towards the old and cold, 200-km-thick lithosphere in the inner parts of the East European platform is not taken into account. Creation of oceanic lithosphere started at ca. 160 Ma when the opening of the Alpine Tethys was initiated (Stampfli and Borel, 2002). Parts of this ocean were already subducted during Cretaceous time, but the Carpathian part survived until Oligocene time (ca. 30 Ma). We can thus assume an age of the oceanic lithosphere of 100 Ma or more and an associated thickness of ca. 90 km according to the age –thickness relationship of, for example, Forsyth (1977). 71 Seismic tomography results are available at different scales ranging from regional (Mediterranean region; Wortel and Spakman, 2000) to local (Vrancea region; Martin et al., 2001). They give information about the dimension and orientation of the subducted lithosphere beneath Vrancea. Together with data from structural geology and other fields, this information can be used to reconstruct the (plate-)tectonic evolution of the Carpathian – Pannonian region (Csontos et al., 1992; Fodor et al., 1999; Linzer et al., 1998; Sperner and the CRC 461 team, in press). This evolution serves as the basis for our gravity modelling. 4.3. Parameters and boundary conditions for the gravity models For the Carpathian gravity models, we used the same density contrasts as in the previously discussed models (Table 1). For the density contrast between lithospheric mantle and asthenosphere, velocity data from seismic tomography can be used for an independent estimation. Results of seismic tomography show a high-velocity body that encloses most of the intermediate-depth earthquakes and can be interpreted as the subducted slab (Fan et al., 1998; Martin et al., 2001; Wenzel et al., 1998a). Therefore, the velocity difference between high-velocity body and background model (representing the surrounding asthenosphere) can be used to estimate the density contrast between lithospheric mantle and asthenosphere. The mean velocity contrast for the high-velocity body is about 2% (Martin et al., 2001) and the velocity of the background model ranges between 7.9 and 8.3 m s 1 for depths of 70 to 300 km. Thus, the velocity difference Dvp ranges between 0.158 and 0.166 m s 1. Using the velocity –density relationship of Cermak et al. (1990) for ultra-basic rocks (i.e., for lower crust and upper mantle; q = 170vp + 1850; Dq = 170Dvp) the resulting density contrast is ca. 27 kg m 3 which is about the same density contrast that was used in the models described previously (30 kg m 3; Table 1). For the starting model, we assumed a distance of 200 km between the European continent and the Tisia– Dacia block. This distance is based on three observations. First, minimum amounts of Tertiary shortening, as recorded by cross-section balancing of 72 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 the Eastern Carpathian accretionary wedge, range from 100 km (Burchfiel and Bleahu, 1976) through 130 km (Roure et al., 1993) to 150 –160 km (Morley, 1996). Roure et al. (1993) and Morley (1996) restored only parts of the accretionary wedge, so that the total amount of shortening must be larger. Secondly, paleomagnetic data indicate an early to middle Miocene clockwise rotation of the Tisia – Dacia block by about 60j (Panaiotu, 1998; Patrascu et al., 1994). For a reconstruction of the early Miocene position, the Tisia – Dacia block has to be rotated backwards around the northwestern edge of Moesia (Fig. 10; note that the early Miocene (late Egerian, 20 – 22 Ma) position of the Tisia – Dacia block was even more southeastwards than the Ottnangian to Karpatian position shown; Sperner et al., 1999). Thus, the distance between the Tisia – Dacia block and the European foreland (i.e. the width of the ocean) was about the length of the northern margin of Moesia, i.e. about 200 km. Thirdly, seismic tomography shows that a high-velocity body exists beneath the Vrancea region that reaches depths of about 350 km (Wortel and Spakman, 2000). This high-velocity body can be interpreted as the subducted slab indicating that the assumed 200 km oceanic subduction is a minimum value. Post-collisional processes like lengthening due to slab pull or mantle delamination along the Moho could have increased the originally shorter slab to its present-day length of 350 km. The thickness of continental crust and lithospheric mantle were defined to be 30 and 60 km, respectively. The thickness of the oceanic crust was assumed to be 8 km and the depth to the L/A boundary was fixed to 90 km (see above). The maximum depth of water plus sediments on the continental margin was limited to 6 km in the starting model. The dip angle of the down-going plate can be estimated from the distance between the plate boundary and subduction-related calc-alkaline volcanism. This distance is about 240 km (Fig. 11). Assuming the depth of the magma-generating window to be about 90 km, the mid-Miocene dip angle of the subducting slab must have been about 20j. All other parameters, like the width of the transition zone between oceanic and continental crust (100 km) or the maximum shortening in the sediments on the continental margin (150 km), are the same as in the models discussed in Section 3. The assumed maxi- mum shortening is in agreement with the minimum amount of Tertiary shortening estimated from balanced cross sections (see above). 4.4. Subduction Miocene oceanic subduction beneath the Tisia – Dacia block lead to the development of a basin on the upper plate with a maximum water depth of 1.28 km. This depth corresponds to a sediment thickness of 5.20 km at the final stage of subduction (Fig. 13a). The equivalent to this basin in reality is the Transylvanian basin (for location, see Fig. 9), which is known to be a cold basin with normal crustal thickness (ca. 30 km) and without large-scale extensional structures despite its up to 5-km-thick Paleogene to Upper Miocene sediment fill (Fig. 14). These features led to the development of various, sometimes complicated models for the evolution of the Transylvanian basin (e.g. Ciulavu and Dinu, 1998; Nielsen et al., 1999). Royden et al. (1982) offered a simple solution by assuming that the steepening of the slab resulted in a suction force on the upper plate. Our solution is even simpler: the pure existence of the slab is a reason enough to cause a downward pull of the upper plate (Fig. 13a). Slab steepening occurs later in our model and has a different effect (see below). Sediments for the basin fill were sourced in the early Miocene from the north and northwest (De Broucker et al., 1998). Later, sediment supply also came from the Southern Carpathians in the south and from the Apuseni mountains in the west. Both of these mountain ranges show uplift and erosion from late Badenian time (15 Ma; Sanders et al., 1999). 4.5. Continental collision The youngest significant thrusting movements in the SE-Carpathians, marking the end of the convergence between Tisia – Dacia block and European foreland, took place during the late Sarmatian (Roure et al., 1993). Younger shortening is related to local events (Girbacea and Frisch, 1998). Continental collision of the Tisia – Dacia block with the European foreland started earlier (Badenian) when the transition zone entered the subduction zone (Zweigel et al., 1998). Crustal thickening in the SE-Carpathians only reaches values of about 40 km (Hauser et al., 2001) B. Sperner et al. / Tectonophysics 382 (2004) 51–84 73 Fig. 13. Gravity model for the SE-Carpathians. We started the model with 200 km distance between the two continents. Thicknesses are 8 km for oceanic crust, 30 km for continental crust, and 60 km for continental lithospheric mantle. (a) Oceanic subduction resulted in the evolution of a basin on the upper plate, the Transylvanian basin. The basin was filled during the Paleogene and lower Miocene with sediments from the Apuseni mountains in the west and the Southern Carpathians in the south. (b) Slab steepening led to a southeastwards migration of the depocentre in the Transylvanian basin, so that middle and upper Miocene sediments have their maximum thickness in the southeastern part of the basin (see also Fig. 14). (c) Delamination caused a shift of subsidence further southeastward into the foredeep and uplift of the Eastern Carpathian flysch belt. Note that sedimentary filling of the subsiding area results in a 15-km-deep basin and a Moho at 45-km depth accompanied by gravity anomalies much larger than the ones measured in the SE-Carpathians. (d) Adaptation of Moho (40 km) and basin depth (ca. 10 km Neogene and Quaternary sediments) in combination with a more realistic (smoothed) topography in the foreland results in gravity anomalies which are close to the measured values. FAA, free-air anomaly (calculated: dashed line; measured: filled circles); BA, Bouguer anomaly (calculated: solid line; measured: open circles). 74 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 14. Profile trough the Transylvanian basin (after Ciupagea et al., 1970) (for location, see Fig. 9). Between the Lower and Middle Miocene, the axis of maximum basin depth migrated southeastward, consistent with the pattern that results from steepening of the subducted slab (Fig. 13a,b). indicating a soft collision where only the two transition zones thrusted over each other, but where no continental subduction took place. Thus, in our model convergence stops after 100 km of underthrusting (Fig. 13a). Fig. 13b shows the steepening of the subducted slab after continental collision stopped. The steepening process might already have been initiated during an earlier stage when subduction velocity was slowed due to the entrance of transition-zone crust into the subduction zone. In our model, slab steepening results in a shift of the basin axis towards the southeast (Fig. 13b). The same evolution can be seen in the sediments of the Transylvanian basin: the depocentre, as indicated by the thickest sediment pile, migrates towards the southeast (Fig. 14). 4.6. Post-collisional evolution Strong intermediate-depth seismicity in the SECarpathians indicates that the subducted slab is still hanging beneath the Vrancea region (Fig. 11). Earthquake epicentres are located in a narrow, nearly vertical column extending from 70 to 200 km depth (Bonjer et al., 2000; Oncescu and Bonjer, 1997). Focal mechanisms show, in most cases, a downward extension, while the compression axes are subhorizontal with no preferred orientation (Oncescu, 1987). These features, as well as the high deformation rate of the seismogenic volume (6.3 10 15 s 1; Wenzel et al., 1998b), demonstrate that the slab is still mechanically coupled with the overlying lithosphere (Sperner et al., 2001). Slab break-off, as sometimes proposed for the SE-Carpathians (e.g. Fuchs et al., 1979), would result in a sinking of the slab into the deeper mantle without producing strong earthquakes. One of the most interesting points is the location of the earthquakes relative to the surface structures: they are located beneath the accretionary wedge and the foredeep (Fig. 11). For a Wadati – Benioff zone with the earthquakes near the upper surface of the slab and a slab being located at the suture zone, one would expect the earthquakes beneath the upper plate or the suture zone. Thus, different plate-tectonic configurations, called ‘‘model 1’’ to ‘‘model 4’’, can be discussed to explain this situation (Fig. 15). Model 1 assumes a subduction zone beneath the suture zone between Tisia – Dacia block and European platform (Fig. 15a). As already mentioned, this model can not explain the location of the earthquakes because they are situated ca. 100 km further east than expected. A subduction zone beneath the foredeep (model 2; Fig. B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 15. Possible plate-tectonic configurations for the recent situation in the SE-Carpathians in which earthquake hypocentres (black dots) are located beneath the accretionary wedge, not beneath the Miocene suture. 75 76 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 Fig. 16. Gravity models for three of the plate-tectonic configurations of Fig. 15. (a, b) Subduction beneath the Miocene suture (Fig. 15a) as well as subduction beneath the foredeep (Fig. 15b) results in a lateral shift of the gravity anomaly minimum between measured and modelled data for both Bouguer and free-air anomaly. (c) Subduction beneath the Miocene suture followed by delamination (Fig. 15d) gives a better fit between measured and modelled gravity anomaly data concerning the position of the minimum. Nevertheless, the modelled anomalies are larger than those measured, indicating that the basin and the Moho are not as deep as assumed in the model. However, adjustments have been made in order to improve the fit (Fig. 13d). B. Sperner et al. / Tectonophysics 382 (2004) 51–84 15b) would fit better for the earthquake locations, but now the accretionary wedge lies completely on the upper plate. This would be reasonable if large backthrusts (towards the west) could be found, especially at the western end of the accretionary wedge. However, the majority of the nappes have a movement direction towards the east and no indications are found that the whole accretionary wedge was (back)thrusted onto the Tisia – Dacia block. Thus, the accretionary wedge needs to be located on the lower plate. Model 3 combines model 1 and 2 by assuming two subduction zones (Fig. 15c), but this means that twice the amount of convergence is needed. Furthermore, two different accretionary wedges as well as two different foredeeps should develop. This is not the case. In model 4, subduction took place at the same locality as in model 1 (Fig. 15d). However, continental collision was followed by delamination along a horizontal tear either at Moho depth or somewhere deeper in the lithospheric mantle (Girbacea and Frisch, 1998). Thus, the subducted slab rolled back towards the east until it reached its present location about 100 km from the Miocene suture zone. The different models can be tested with gravity modelling. We excluded model 3 from this test due to its unrealistic, double amount of convergence. For the other three models, we assumed that the tectonic evolution was the same until continental collision (as described in Sections 4.4 and 4.5), but the postcollisional evolution and the location of the slab are different. In model 1, with subduction beneath the Miocene suture, misfit between modelled and measured Bouguer anomaly data occurs in the foredeep and in the Transylvanian basin (Fig. 16a). Model 2 shows a better fit for the load basin (which is in this case not equivalent to the Transylvanian basin), but the foredeep and the accretionary wedge of the model are now located too far east, so that they produce a misfit in this region (Fig. 16b). Model 4 gives the best-fitting results with respect to the Bouguer anomaly data (Fig. 16c) and it also best explains the vertical movements in the SE-Carpathians in the last few million years. Fission track data indicate that the southern part of the Eastern Carpathian accretionary wedge started to uplift at about 5 Ma (Sanders et al., 1999). At the same time, the southern part of the foredeep basin (Focsani) subsided at its maximum rate (Artyushkov et al., 1996; Matenco et al., 2003). These 77 opposite vertical movements in close proximity are best explained by delamination of the lithospheric root beneath the accretionary wedge (causing uplift) and by a heavy lithospheric root still hanging beneath the foredeep (causing subsidence). The fact that delamination only occurred a few million years ago is also responsible for the still low heat flow above the delaminated region. According to Bird (1979), heat from asthenospheric material at lower-crustal levels takes several million years to reach the surface (maximum surface heat flow after ca. 9 Ma for a 30-kmthick crust) and produces a heat flow anomaly of only 15– 25 mW m 2. Our idealized delamination model results in modelled gravity anomaly curves which fit the measured curves concerning the position of minima and maxima, but they fail to explain the magnitude of gravity anomalies (the measured anomalies are smaller than the modelled ones; Fig. 13c). For a better fit, we adapted the model geometry to reality by reducing the depth of the basin and of the Moho (40 km as seen from the refraction seismic data of Hauser et al., 2001) and by smoothing the topography (instead of the vertical ‘‘wall’’ in Fig. 13c). As these processes (erosion, basin subsidence) are still ongoing, isostasy might not necessarily be achieved so that we made the adaption without mass balancing. The resulting model fits the measured data much better, at least at a larger scale (Fig. 13d). Since our intention was to study large-scale lithospheric processes, we are satisfied with this result. Smaller misfits are caused by smaller-scale structures that are not the topic of this study. 5. Discussion 5.1. Modelling assumptions 5.1.1. Time In our models, time is not included because timedependent deformation processes are not considered. Isostatic adjustments are assumed to happen instantaneously. Thus, our results have no time scale. Nevertheless, sequences for sedimentation and erosion can be determined. However, it should be kept in mind that the sedimentation or erosion processes might be interrupted by the next event before they reached their 78 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 final state. In the case of a sediment basin, which should, for example, reach a maximum depth of 4 km due to slab pull, subsidence might be interrupted after 2 km because slab break-off eliminated the driving process for the subsidence. This example makes it clear that our results give maximum values for the case that enough time was available to complete each distinct process. 5.1.2. Coupling between slab and overlying lithosphere In our models, we assume a strong coupling between subducted slab and overlying lithosphere so that the gravitational effect of the mass excess in the lithospheric root is transferred to the upper plate and thus causes subsidence. We make this assumption of strong coupling even for slabs which descend deeper than the lithospheric mantle (e.g. Fig. 2c). In this case, slab and overlying lithosphere are separated in the deeper regions by asthenospheric material (mantle wedge) which behaves ductilely at geological deformation rates (ca. 10 15 s 1). This means that the gravitational effect of the slab might be (partly) transformed into internal deformation of the mantle wedge instead of causing downward movements of the overlying crust. Thus, our calculations only give maximum values for possible vertical movements at the surface. 5.1.3. Model geometry We used 2D-models with infinite extent in the direction perpendicular to the profiles. This approach is useful to get general results for any subduction or collision zone. For our example from the SE-Carpathians, this approach gives only a first approximation because seismic tomography shows that the presentday slab has only a small lateral extent (Martin et al., 2001). This small lateral extent has no influence on the vertical movements along the profile because isostatic disequilibrium was compensated locally. However, the calculated gravity field would be different, not only because of the reduced mass excess at depth (smaller slab), but also because the basins evolving above the slab would be laterally limited. This would result in a reduced mass deficit. Thus, for modelling the present-day gravity field in the SECarpathians, a 2.5D model with limited lateral extent of the slab and the basins would be more appropriate. On the other hand, it has to be taken into account that the Carpathian slab was not always as confined as it is today. During subduction, it had a larger lateral extent which was successively reduced during collision (slab break-off) (Sperner et al., 2001). This could be an explanation for the fact that the Transylvanian basin, which evolved as a load basin on the upper plate during subduction, is broader than the Focsani basin (foredeep basin on the lower plate; Fig. 9) that mainly developed after continental collision when the slab was already reduced in its lateral extent. 5.2. General results Our modelling results give maximum values for vertical surface movements related to oceanic subduction, continental collision, slab steepening, breakoff, and delamination. Due to the fact that we used purely kinematic models, no time-dependence is included and consequently we cannot predict any rates for subsidence or uplift. Nevertheless, our models reveal characteristic sedimentary, structural and erosional patterns that depend on the tectonic evolution of the study area. These patterns can be used to identify subcrustal processes like oceanic subduction, slab steepening or break-off. We used this method to study the Tertiary tectonic evolution of the SE-Carpathians and found results that are in agreement with other data (e.g. uplift data from fission-track analysis). For a more sophisticated test of the general results, the method should be applied to other subduction or collision zones (e.g. Aegean region, Dinarides, Apennines). A characteristic of most of the studied processes is wave-like vertical movements at the surface. In some cases, like oceanic subduction, this wave simply produces an expanding basin (only subsidence), while other cases, like slab steepening, result in more complicated movement patterns (subsidence followed by uplift). In addition to different vertical movement directions (up or down), the wavelength and amplitude of the wave might also vary with time. Slab steepening, for example, results in a lithospheric root limited to a smaller area, but with a larger thickness. The result is a smaller, but deeper basin at the surface. In contrast, in the case of delamination, the area remains constant, but the slab length increases so that the basin size remains constant while the depth B. Sperner et al. / Tectonophysics 382 (2004) 51–84 increases with ongoing delamination. In both cases (slab steepening and delamination), the basin axis migrates parallel to the horizontal component of slab movement (rollback) so that depocentre migration might be an indication of processes acting in the mantle (if other causes for this migration can be excluded). In contrast to the sedimentary record, which might reflect the (long-term) tectonic evolution of an area, gravity data give only information about the presentday situation. Nevertheless, they are useful for testing different tectonic evolution scenarios by comparing the resulting (final-stage) gravity data with measured values. In general, the gravity data (especially the free-air anomaly) reflect vertical surface movements because basins and uplifted regions have a more pronounced influence on gravity than deep-seated structures like a subducted slab. The Vrancea slab, as revealed by seismic tomography, has a modelled positive gravity effect of about 20 mGal (Hackney et al., 2002), while the measured gravity anomaly (Bouguer and free-air) reaches negative values of about 120 mGal. These negative anomalies originate mainly from the Focsani basin. 5.3. SE-Carpathians 5.3.1. Type of subduction For a long time, it has been debated whether subduction beneath the Carpathians was oceanic or continental (e.g. Girbacea and Frisch, 1998; Girbacea and Frisch, 1999; Linzer, 1996a,b; Pana and Erdmer, 1996; Pana and Morris, 1999). Doubts about oceanic subduction are mainly based on the fact that no ophiolithes are preserved from the Miocene Carpathian subduction. However, slab retreat produced a low-stress subduction zone in the Carpathians so that the entire oceanic lithosphere disappeared into the mantle; only the sedimentary cover was (partly) scraped off and thrust onto the European foreland. Indications for slab retreat come from the occurrence of back-arc extension in the Pannonian basin, while at the same time compression took place along the Carpathian arc (thrusting in the accretionary wedge). This simultaneous occurrence of compression and extension is typical for slab retreat and low-stress subduction zones (e.g. Doglioni, 1992; Royden, 1993). 79 Our modelling results show that continental subduction would result in an uplift of 2.8 km in the region where the crust was doubled due to underthrusting of the upper plate by the subducting plate (Fig. 6f; Section 3.2). However, topography of the SECarpathians is mainly in the range of 0.5 –1.5 km above sea level. Highest values occur in the Outer Carpathians, i.e. in the accretionary wedge, and not in the basement of the upper plate as should be the case if continental subduction had taken place. Furthermore, we would expect a large BA anomaly with long wavelength and high amplitude (Fig. 6f). Only the latter is valid for the SE-Carpathians with its strong but narrow BA anomaly (Fig. 12b). In addition, new refraction seismic data show that the Moho in the SECarpathians reaches a maximum depth of ca. 40 km (Hauser et al., 2001, 2002), thus indicating that only minor crustal thickening took place. Similar results were achieved by Lillie et al. (1994) when they compared the situation in the Eastern Alps and the Western Carpathians. They concluded from their modelling, in comparison with measured gravity and topography data, that the continental collision in the Carpathians was a ‘‘soft’’ one with only 50 km of underthrusting. In summary, we conclude that the evolution of the SE-Carpathians can be best explained by oceanic subduction followed by ‘‘soft’’ continental collision with only minor crustal thickening. 5.3.2. Correlation of sedimentation and tectonics We interpret the sedimentation in the Transylvanian basin (at least the Tertiary part of the basin) to be related to the subduction process beneath the Carpathian arc. Sedimentation in the Transylvanian basin shows a hiatus from the end of the Cretaceous to the beginning of the Eocene. This Palaeocene hiatus is probably related to a change in the large-scale tectonic situation. Sediments older than the hiatus correlate with the late Cretaceous subduction which took place beneath the Alpine – Carpathian arc. Then, at the end of the Cretaceous, the buoyant (thinned) continental margin of the European platform started to enter the subduction zone in the Eastern Alpine area (Stampfli et al., 2002). In the Carpathian region, oceanic lithosphere was still available in an embayment in the southern margin of the European continent (Sperner et al., 2001). However, subduction of this oceanic lithosphere was blocked by the continental collision in 80 B. Sperner et al. / Tectonophysics 382 (2004) 51–84 the Alps which is responsible for a slower northward movement of the Adriatic microplate and a nearly fixed position of the slab. In the late Eocene, slab break-off occurred in the Alps (Wortel and Spakman, 1992) and most probably also removed the Carpathian slab. Thus, the Carpathian slab was released and a new phase of subduction started. We interpret the sediments younger than the Palaeocene hiatus (i.e. Eocene and younger) to be correlated with this younger phase of subduction. Tertiary subduction started with a flat dip (ca. 21j) which can be calculated from the position of the subduction-related volcanism relative to the plate boundary (Fig. 11). At this time, thick sediments were mainly deposited in the western part of the Transylvanian basin (above the thickest part of the lithospheric root; Fig. 13a). During the Middle Miocene, the depocentre migrated SE-ward, which we interpret as a consequence of slab steepening with SE-ward migration of the lithospheric root (Fig. 13b). Slab steepening is proven by its present-day vertical situation which is visible from intermediatedepth seismicity and seismic tomography (Martin et al., 2001; Sperner et al., 2001). Removal of the lithospheric root from the western parts of the Transylvanian basin led to uplift and erosion (since 12 Ma) of the Apuseni mountains and neighbouring parts of the Transylvanian basin (Sanders et al., 1999). The NE –SW orientation of the present-day slab (Martin et al., 2001) and the distribution of Sarmatian sediments (13.0 – 11.5 Ma) in the Transylvanian basin (Sandulescu et al., 1978) are parallel to each other (Fig. 17). This indicates that slab orientation was already NE – SW during Sarmatian subduction retreat, as proposed in the geodynamic model of Sperner and the CRC 461 team (in press). Uplift and erosion of the southern part of the Eastern Carpathians started at about 5 Ma (Sanders Fig. 17. Sketch map of the SE-Carpathians showing the distribution of Miocene sediments in the Transylvanian basin (Sandulescu et al., 1978) and (preliminary) results from seismic tomography at 110 – 150 km depth (Martin et al., 2001). Note the parallelism (NE – SW) between the high-velocity body and the axis of Miocene sediments in the basin. B. Sperner et al. / Tectonophysics 382 (2004) 51–84 et al., 1999), while at the same time subsidence reached maximum rates in the southern part of the Focsani basin (Artyushkov et al., 1996; Matenco et al., 2003). We correlate these movements with a further SE-ward movement and steepening of the subducted slab towards its present-day vertical position (Fig. 13b). Its location beneath the foreland indicates delamination along the Moho or at a deeper level (Girbacea and Frisch, 1998) (Fig. 13c). The vertically hanging slab produced a deep, symmetrical basin in the foreland (Fig. 13c). This geometry is in accordance with the shape of the Focsani basin, which is characterized by symmetrical basin flanks and still active normal faults along the basin margins (Matenco et al., 2002). Uplift along its western flank during the Quaternary (Matenco et al., 2002) indicates that SE-ward rollback (delamination) is still active, thus moving the slab together with the ‘‘subsidence-uplift wave’’ further SE-ward. The symmetrical shape of the younger parts of the Focsani basin is in contrast to ‘‘normal’’ foredeep basins which show an asymmetric deepening of the basin towards the accretionary wedge (i.e. towards the plate boundary) due to flexural bending of the foreland. The latter is caused by loading of the end of the lower plate (slab pull) resulting in this asymmetrical shape. Our interpretation is that the state of (asymmetrical) flexural bending is, in the Carpathians, already over because the formerly separated lithospheric blocks were welded during continental collision. They now react as one plate which was loaded in the middle by the subducted lithosphere so that a symmetrical basin developed. As discussed previously (end of Section 4.6.), we adapted the Moho and the basin depth above the slab to better fit the gravity anomaly data. Several reasons are possible for the difference in Moho and basin depths between the idealized and reality-adapted model (Fig. 13c,d): (a) the slab might be shorter, (b) the density contrast between slab and asthenosphere might be smaller, (c) flexural rigidity might prevent a larger subsidence, (d) the limited lateral extent of structures like the Focsani basin and the subducted slab, and (e) the slab has already started to decouple from the overlying crust so that isostatic rebound led to uplift of the region above the slab (Sperner and the CRC 461 team, in press). The last possibility is a realistic scenario that is supported by levelling data 81 that show uplift for this region (Joó, 1992; Popescu and Lazarescu, 1988; Zugravescu et al., 1998). 6. Conclusions Isostasy is one of the fundamental processes shaping the surface of the Earth. Mass deficit or mass excess at the surface (e.g. ice sheets) or in the subsurface (e.g. crustal roots of mountain belts) are often responsible for vertical movements. Our modelling of oceanic subduction shows that the mantle part of the lithosphere can trigger significant surface movements on the upper (overriding) plate. After the end of subduction and depending on the post-collisional evolution of the slab (steepening, delamination, break-off), the lower plate might also be affected by mantle-induced surface movements. The often wavelike vertical surface movements result in uplifted areas and sedimentary basins with characteristic internal structure (e.g. depocentre migration) which can be used to reconstruct the tectonic evolution, especially with respect to lithospheric processes active during and after subduction. For our example from the SE-Carpathians, we propose a model that is consistent with the evolution of various basins (Transylvanian basin—load basin, Focsani basin—foredeep basin), with the present-day position and dip angle of the slab as seen in seismicity and seismic tomography, and with the gravity data. The model assumes Tertiary subduction which started with a shallow dipping slab (evolution of the Transylvanian basin on the overriding plate) followed by slab steepening and (later) delamination, so that the present-day position is subvertical and beneath the foredeep (evolution of the Focsani basin on the foreland). Our model gives a plausible and simple explanation for the evolution of the Transylvanian basin, a basin whose evolution was, until now, rather enigmatic. Acknowledgements This work was done under the auspices of the Collaborative Research Centre ‘‘Strong Earthquakes’’ (CRC 461) which is financed by the German Science Foundation (DFG). 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