The geodynamic evolution of the Precaspian Basin (Kazakhstan) along a north–south section

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Tectonophysics 313 (1999) 85–106 www.elsevier.com/locate/tecto

The geodynamic evolution of the Precaspian Basin (Kazakhstan) along a north–south section Marie-Franc¸oise Brunet a,Ł , Yuri A. Volozh b , Mikhail P. Antipov b , Leopold I. Lobkovsky c a

ESA 7072 CNRS–Universite´ Pierre et Marie Curie, Ge´otectonique case 129, 4 Place Jussieu, 75252 Paris Cedex 05, France b Geological Institute, Russian Academy of Sciences, Pyzhevski per. 7, 109017 Moscow, Russia c Centre for Sedimentary Basin Research, Krasikova st. 23, 117851 Moscow, Russia Received 13 November 1997; accepted 14 August 1998

Abstract Several hypotheses exist for the origin and evolution of the Precaspian Basin. There are more than 20 km of sediments deposited, yet there is little consensus on the causes of the subsidence. Except for the presence of a thick salt layer (Lower Permian), the main problem is the chronostratigraphic interpretation of the sediments in the centre of the basin, where the calibration of seismic data with well data from the basin margins is problematic at deep levels. The age of the deepest sediments could be either Riphean or Devonian. A generalised cross-section of the Precaspian Basin, roughly perpendicular to the central elongated E–W depocentre, is used to represent the basin’s evolution. Several evolutionary steps from the end of the Riphean until the Present are demonstrated. Tectonic subsidence analysis indicates that there are six main phases of evolution: (1) subsidence during an active rifting phase in Riphean times; (2) rifting during the Vendian–Ordovician (poorly dated); (3) significant subsidence during the Late Devonian in an extensional context possibly due to back-arc rifting; (4) acceleration of subsidence during the Late Carboniferous–Permian, synchronous with or just following the closure of the Uralian Ocean, a major subsidence phase in the Dniepr–Donets Basin and possible subduction to the south of the basin; (5) renewed rifting during the Triassic coincident with a general phase of extension in Eurasia and the opening of the Neo-Tethys; and (6) neotectonic subsidence resulting from crustal down-bending in a generally compressional setting. The nature of the crust underlying the basin is not well known. It could be Riphean–early Palaeozoic or Devonian oceanic crust or continental crust attenuated during the several episodes of supposed rifting between the Riphean and the Triassic. Estimations of crustal thickness in the basin vary and they depend on the interpretation of a high-velocity layer situated at the base of the crust. The presence and general distribution of this layer is confirmed by gravity data. It may be considered as the uppermost mantle, as oceanic crust metamorphosed into eclogite at depth during collision and then exhumed and emplaced at the base of the crust, or as lower crust transformed in situ (into eclogites?). The crustal thinning factor leading to the observed present crustal thickness — assuming an initial thickness of 40 km — is 3.3 in the two first cases and 2 in the latter, if the crust is considered to be continental. The geophysical and subsidence data are discussed in terms of basin-forming mechanisms such as: (1) intracontinental rifting of early Palaeozoic, Devonian or Permo–Triassic ages; (2) oceanisation of continental crust during Riphean or Devonian time; (3) eclogititisation at the base of the crust in the upper mantle or by metamorphism of subducted oceanic crust; and (4) down-bending due to compressional forces or mantle flow induced by subduction–collision processes in Carboniferous and Recent times.  1999 Elsevier Science B.V. All rights reserved. Keywords: Precaspian basin; Kazakhstan; subsidence; eclogites; crust; salt Ł Corresponding

author. Tel.: C33 1 4427 5168; Fax: C33 1 4427 5085; E-mail: [email protected]

0040-1951/99/$ – see front matter  1999 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 9 9 ) 0 0 1 9 1 - 2

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1. Introduction The Precaspian Basin (PCB) is a large (600 ð 900 km), complex and composite basin in the southeast corner of the Russian Platform (Fig. 1), part of the stable East European Craton (EEC). Located in western Kazakhstan, it contains more than 20 km of sedimentary fill. The size and petroleum potential of the PCB are such that it is considered to be one of the most important sedimentary basins in the world. It acquired its present shape mostly during the Permo–Carboniferous in relation with orogenies along its eastern (Urals) and southern (Emba–Karpinsky) margins. During its long Riphean to Present history, it recorded the evolution of these systems together with intracontinental deformations affecting the craton. The ultimate aim of this study, funded by the International Peri-Tethys Programme, is to present an improved model for the origin and evolution of the PCB using geodynamical reinterpretation and new subsidence analysis of existing data. In this paper, we report on the results obtained so far on the basis of a north–south generalised cross-section and subsidence curve for the basin centre, and review the

different hypotheses on the mechanisms which may have caused the origin and the evolution of this large basin.

2. Regional setting 2.1. Sedimentary fill The PCB clearly differs from adjacent areas by its greater depth (>20 km) to crystalline basement (Fig. 2). It has an east–west-elongated shape, including the main Central Precaspian Depression (CPD) and a NE–SW-trending extension towards the southwest known as the Sarpinsky Basin. The PCB is bounded to the north by the shallower parts of the Russian Platform with basement depths of 3–8 km. The basin is bordered to the east by the Ural Mountains (basement at 0–5 km), to the southeast by the Northern Ustyurt Block (basement at 5–11 km), and to the southwest by the Karpinsky inverted rift (basement at 8–20 km). The sedimentary succession of the PCB is divided into a pre-salt complex and post-salt sediments by a conspicuous, several kilometres thick saliferous

Fig. 1. Location of the PCB and the main regional tectonic units.

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Fig. 2. Basement depth map (in km) of the PCB (modified from Volozh, 1991), with position of the section and central point C (cross) used in the subsidence study.

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series of Lower Permian age (Fig. 3). While the shallow Cenozoic and Mesozoic sections have been drilled throughout the basin, the upper Palaeozoic has only been extensively drilled in the marginal zones where it is 2.5–6 km deep. Lower Palaeozoic and Riphean strata have been reached by only a few wells on the northwestern margin of the basin and in the Pachelma rift (northwest of PCB; Fig. 1). Consequently, the development of a geological model for the basin requires the accurate calibration of seismic markers with well data in the marginal zones and the tracing of these markers into the deeper, undrilled regions. However, the tracing of the deep-seated seismic markers from the margins to the centre of the basin is uncertain, especially at very steep slopes where seismic reflections converge and seismic facies change. 2.2. Main seismic horizons Three important seismic horizons with basin-wide significance for the pre-salt section are known as P1, P2 and P3. Fig. 4 shows one example of seismic data from the Aralsor–Khobda zone, situated in the central part of the north–south section shown in Fig. 5 (near point C). At each side, a column presents the interpretation of the seismic reflectors in the pre-salt section (i.e. below reflector P1). The P1 horizon has been penetrated in many wells and corresponds to the base of the Kungurian evaporites. The P2 regional seismic marker is well calibrated. It has been drilled in more than 30 fields and recognised as a regional mid-Carboniferous unconformity. Additional markers may be picked in the Devono– Carboniferous sequences (Figs. 3 and 4). Of the three principal seismic horizons, P3 is the most difficult to interpret as it is located at a depth greater than 6 km and has been reached in very few wells. The precise age of the P3 horizon is a matter of dispute and could represent anything from the top of basement to the top of Devonian according to different authors. Drilling at Zhanazhol (eastern margin) and Mynsualmas (southeastern margin) has shown that the upper Frasnian (the deepest section drilled in these wells) is some 4 km above P3, suggesting that the horizon could represent a pre-Middle Devonian unconformity. Nevertheless, the drilling

data obtained from some areas in the east of the PCB strongly support a definable stratigraphic position for P3 (Akhmetshina et al., 1993). This reflector was penetrated by Hole G-5 (East Akzhar) (located in Fig. 2) in the zone of the Aktyubinsk–Astrakhan uplifts. Within this area two holes, G-1 (Baktygaryn) and G-4 (Kumsai), penetrated thick Devonian carbonate deposits but drilling was stopped at 200– 300 m above the P3 horizon. The presence of conodonts species Ozarkodina remscheidenis remscheidenis Ziegler in carbonates from Hole G-5 (intervals 5738–5745 m and 5745–5751 m) is indicative of the lowermost portion of the Lower Devonian Lochkovian stage. The presence of foraminifers Tubeporina tenue Sabirov in Hole G-1 (interval 6204–6212 m) determines the stratigraphic interval as the Lower Devonian Pragian stage. On the basis of these data the main reflector P3 is restricted to the base of the Lower Devonian (lower Lochkovian) carbonate deposits. Northward into the centre of the CPD or southward into the Tugarakchan trough, P3 and a refracting horizon on top of the basement (Pf) gradually diverge, so that 6–7 km of section separates P3 from the presumed top of basement in the central Precaspian Depression, while Pf is 3–4 km below P3 in the Tugarakchan trough. In the Aralsor–Khobda zone, the deeper horizons P4 and P5 can be traced. The age of the section between horizons P4 and P5 is believed to be Vendian– Mid-Ordovician and that of the section below P5 is supposed to be Riphean (Volozh, 1991). The pre-salt section, in the CPD, is up to 10– 12 km thick, mainly due to the contribution of its Riphean–early Palaeozoic component (Fig. 5). The seismic character of the Devonian–Lower Permian interval is different in the central part of the basin (assumed to be terrigenous) and in the marginal areas (carbonates). The mid-Carboniferous and Early Permian (Artinskian) sections are up to 2 km thick in the basin centre, but they are absent in the marginal zones. 2.3. Depth structural maps at the main horizons Depth structure maps of the lowermost Devonian seismic reflector P3, the pre-Moscovian reflector P2, and the pre-Kungurian reflector P1 (Figs.

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Fig. 3. Synthetic column in the centre of the PCB, showing thickness of sediments, palaeowater depths, ages (time scale of Odin, 1994) and main seismic horizons. The seismic velocities and densities (after Nevolin and Kunin, 1977, modified) are used respectively for the time–depth conversion of the seismic lines, and in the calibration of the densities and laws of evolution of the sediments porosity as a function of depth in the subsidence study.

90 M.-F. Brunet et al. / Tectonophysics 313 (1999) 85–106 Fig. 4. Example of seismic data in the Aralsor–Khobda zone, central part of the north–south section. P1 a: Artinskian; P1 C3 : Gzhelian–Sakmarian; C2 : Moscovian–Kazimovian; C1 V–C2 : Upper Visean–Bashkirian; D3 –C1 : Frasnian–Tournaisian; D1–2 : Lower–Middle Devonian; O3 –S: Upper Ordovician–Silurian.

Fig. 5. Simplified sketch of the regional north–south cross-section (seismic line Zhambay–Uralsk with seismic trace numbers shown along the top), located in Fig. 2, after the compilation by Yu.A. Volozh, V.I. Kozlov and Yu.G. Yurov (unpublished). C is the point of calculation of the subsidence curve presented in Fig. 12. Same abbreviations as Fig. 4; R–V: Riphean–Vendian; indication of seismic refraction velocities in the crust and high-velocity layer (HVL with v D 8:1 km=s).

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6–8) illustrate the morphological variations in the four principal tectonic zones of the PCB, i.e. the Northwest Monocline (NM), the CPD, the North Caspian–Aktyubinsk Uplift Zone (NCAZ) and the South-East Marginal Depression (SEMD). The NM is a less than 100 km wide structural zone that stretches over more than 1300 km from Volgograd in the southwest to Orenburg in the northeast. The monocline is a gently south-dipping slope at the P3 level (Fig. 6) and merges with the relatively undeformed rocks of the Russian Platform in the north. The structural map of the P2 horizon shows a sharp, east–west-trending break or step. The structural picture of the P1 horizon (Fig. 8) shows a similar step, shifted slightly to the north in comparison to the P2 step. This margin is generally considered as a passive margin whereas the western (Aralsor) and eastern (Kzyljar) margins are considered by some authors as transform margins, being the boundaries of a large ‘pull-apart basin’ (Kleshev and Shein, 1994). The CPD is by far the deepest part of the PCB. In its central area, the P3 horizon lies at a depth of more than 13 km (Fig. 6). There may be as much as an additional 8 km of Riphean to Silurian sedimentary section below P3. On its southern flank, the CPD is separated from the Mezhdurechensk Monocline by the Elton–Inder fault (Fig. 5), which exhibits 0.5– 1.5 km vertical displacement controlling the distribution of the Riphean–lower Palaeozoic sediments but influencing younger sediments. This monocline may be considered as the transition zone between the PCB and the NCAZ, an early Palaeozoic structural high. Devonian rocks in the NCAZ unconformably overlie Riphean basement, while lower Palaeozoic deposits onlap the uplift to the north and south. The southern boundary of the NCAZ consists of a series of faults that separate the basement high from the SEMD. At the higher structural levels of P2 and P1 (Figs. 7 and 8), the anticlinal geometry of the NCAZ is less conspicuous. The southern and eastern slopes of the basement structure appear to be buried beneath Devonian to mid-Carboniferous sediments that merge with those of the SEMD. The NCAZ thus takes at P2 and P1 the shape of a north- to northwest-dipping monocline, interrupted by several large

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Fig. 6. Depth of the horizon P3 (bottom of Devonian unconformity) (after Volozh, 1991). (a) Contour of P3 in km below sea level. (b) Faults.

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Fig. 7. Depth of the horizon P2 (middle Carboniferous unconformity) Pre-Moscovian (after Volozh, 1991). Same legend as Fig. 6.

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uplifts and carbonate build-ups. At the level of the P1 horizon, the inner slope of the NCAZ is dissected by several steep-sided, deep (to 1 km) canyons; erosion may reach the Lower Carboniferous to Early Devonian substrate. At the end of the Carboniferous to beginning of Permian, the basin was deep (i.e., in the interpretation of Volozh, 1991). These features may be incised by sediments bypass and=or erosion during a sea-level low stand prior to the Kungurian evaporite deposition (Volozh, 1991) (Fig. 9). The canyons are well documented in the south, especially in the area situated between the Volga and Ural rivers, and the limited amount of seismic data in the eastern PCB seems to confirm the existence of analogous features in that part of the basin as well (Volozh et al., 1991). The SEMD separates the PCB from the adjacent Ural orogen, the Turan plate to the east and the Scythian plate to the west. From southwest to northeast, the SEMD can be divided into three principal troughs (which were foredeeps during a part of their history): the North Caspian (north of the Caspian Sea), the South Emba and the PreMugodzhar troughs, of which the latter forms the transition to the Ural Foredeep. The outermost zones of these troughs have undergone moderate to severe deformation in late Palaeozoic and younger times. This deformation is expressed as north- and east-vergent thrusts and folds, associated with regional plate convergence.

3. Data used and uncertainties A generalised north–south cross-section has been chosen to illustrate the evolution of the basin. It is oriented roughly perpendicular to the central elongated depocentre where the basin fill is up to 20 km. From north to south it crosses three of the four main structural domains described above: the NM, the CPD, and the NCAZ. The SEMD is situated to the southeast and not crossed by this section. Data uncertainty and particular features such as the presence of salt makes the subsidence analysis difficult.

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3.1. Problems linked to the sediments 3.1.1. Chronostratigraphy The main characteristic of the pre-salt complex is its subdivision into the three units separated by seismic horizons P2 and P3. Unfortunately the central part of the PCB is only known from seismic data for which correlation with well data from the basin margins is problematic. Consequently, the chronostratigraphic interpretation in the centre of the PCB is equivocal. From the same datasets, different hypotheses can be proposed for the age of sediments at these deep levels. There are two main interpretations based on an age preference of either Riphean (e.g. Volozh, 1991) or Devonian (Shein et al., 1989; Zonenshain et al., 1990) for the deepest sediments. In this paper we have chosen the hypothesis of Riphean age for the oldest sediments deposited in the PCB following the interpretation of Volozh (1991), reinforced by our choice of the P3 horizon as near the base of the Devonian (Akhmetshina et al., 1993) and considering that a thick (about 8 km) unit of older carbonate and terrigenous rocks is present between the top of the basement and the P3 horizon in the CPD. Meanwhile, we will discuss below the geodynamical implications of the other possible interpretations of the chronostratigraphy. 3.1.2. Lithology and palaeobathymetry In the deep levels of the basin centre known only by seismic data, lithological interpretations have been generally made on the basis of seismic velocities (Fig. 3). According to the velocity models, the lower pre-salt complex consists of four alternating high- and low-velocity layers. The basal, low-velocity, unit is interpreted as Riphean clastics unconformably overlying crystalline basement. These are overlain by a high-velocity (carbonate?) unit possibly of Riphean age, bounded at its top by seismic marker P5. The unit lying between P5 and P4 consists of Vendian and lower Palaeozoic (to mid-Ordovician) clastics. Overlying P4 is another high-velocity, presumably carbonate, unit interpreted as having a Late Ordovician to Silurian age. The Lower–Middle Devonian to Lower

Fig. 8. Depth of the horizon P1, Pre-Kungurian, basis of the salt (after Volozh, 1991). (a) Contour of P1 in km below sea level. (b) Erosional canyons. (c) Faults. (d) Turbidites.

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Carboniferous is believed to be basinal shales and marls. Immediately below the salt, the upper part of the section contains mid-Carboniferous to Lower Permian strata consisting of pro-delta shales, turbidites and carbonates and thick Artinskian clastics with deep erosional channels. Above the salt layer, there are Upper Permian– Triassic red beds overlain by a carbonate–terrigenous series for the Jurassic to Present. We have used porosity–depth relationships (Brunet, 1989, 1991) to fit density observations at each level (Fig. 3), with particular attention to the apparently relatively uncompacted deep levels of Devonian to Early Permian age (having low velocities of 3–4 km=s, and correspondingly high porosities greater than 15%). It is thought that the presence of the salt layer has prevented normal burial compaction processes occurring in these units. We have also inferred palaeogeographical environments for the deep parts of the central depression, using seismic facies and velocity analysis, and the profile of the section (which shows a succession of reefs on the upper margin and a scarp and slope towards the basin). Shallow water conditions prevailed in the basin from the beginning of its evolution until the Middle Devonian. From the end Devonian– beginning Carboniferous, the basin deepened reaching deep-marine basinal conditions which prevailed until the beginning of the Permian (Volozh, 1991). Other interpretations exist in which the basin is shallow during this time with a part of the Upper Carboniferous and even mid-Carboniferous being absent (mainly) on the southern and eastern margins, having been eroded at the end Carboniferous– beginning Permian (Krylov et al., 1994; Nevolin and Fedorov, 1995; Fries and Liger, 1995). In this case, the marginal highs are interpreted as erosional features and not as atolls. Accordingly, the western part of Temir and west Zhanazhol are interpreted as having a continental environment during the Carboniferous while platform carbonates were deposited in Zhanazhol and the eastern part of Temir (J.L. Liger, pers. commun., 1998). However, recent sedimentological studies on the eastern margin of the

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PCB support the presence of a basinal facies during the Carboniferous on the eastern slope of the basin (Ensepbaev et al., 1998). These authors document argilites in the basin and turbidites on the slope. At latest Carboniferous–earliest Permian times, different collisional events consolidated the east, southeast and southwest margins of the PCB leading to its isolation. The basin seems to have been almost closed to marine influence at this time, with sea level possibly dropping inside the basin by more than 1000 m. Such a situation may be compared to the Mediterranean Messinian event. From reconstructions based on the measured heights of buried sedimentary–erosional escarpments, Volozh (1991) proposed a maximum depth of 2500 m below sea level for the Kungurian depression and 800–1000 m below sea level for the Kazanian. These evaluations can be reduced (to a maximum of 1800 m of water depth) considering isostatic readjustment due to the weight of posterior sediments filling the depression. Thus, during Kungurian time, salt was deposited in the PCB in restricted marine environments. The presence of mainly halite in the middle of the depression testifies that it was due to sea water evaporation (Volchegursky et al., 1995). To allow for the precipitation of several kilometres of salt, the basin must have been regularly supplied by marine water. The regular intercalation of shale within the salt layers and of sulphates, chlorures, and salt in the carbonated deposits of the basin margins seems to corroborate the occurrence of periodical marine invasions. The climate was arid at the end of the Permian and conditions almost continental, with deposits of red terrigenous series interlayered with carbonates (especially in the Tatarian). The occasional links with the open sea probably were to the north, to the central part of the Russian Platform (Volchegursky et al., 1995). Then, for the remainder of its evolution, the PCB returned to shallow water or continental environments. 3.1.3. Salt The 4.5 km of Permian evaporites were deposited during a short period of less than 10 Myr. The

Fig. 9. Evolution along the north–south cross-section (Fig. 5) at various times; the present stage is virtual, reconstructed without salt movements (after Volozh et al., 1991). For stratigraphic abbreviations, see Figs. 3–5.

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Fig. 10. Depth of the reconstructed Permian salt surface in km, without salt movements (after Volozh, 1991). The ornamented line indicates the boundary of the PCB and the grey area the Salt Basin extension.

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boundaries of the salt distribution are diachronous, beginning at the earliest Artinskian in the southwest (Sarpinsky Basin) and at the early Kungurian over the rest of the basin and ending from the late Kungurian in the eastern margin to the late Kazanian over the rest of the area (Volozh et al., 1997). The lower unit (Artinskian–Ufimian) reaches 2.5 km thickness and the upper one (Kazanian) 2 km. The presence of this several kilometres thick layer of salt introduces a severe problem in the analysis of basin subsidence. Over most of the PCB, strong Permian to Recent salt tectonics has led to the formation of many salt domes and salt-related sub-basins. The depressions between the domes are filled by postsalt sediments. Their weight in turn has probably amplified the salt movements. Thus, inside the salt area, the subsidence values derived from the saliferous and post-salt sediments deposited in interdome depressions reflect, at least for a considerable part, rather the effect of halokinesis than the results of tectonic mechanisms. The spatial distribution of salt displacements has been reconstituted by the Russian co-authors in great detail, salt dome by salt dome, throughout the basin on the basis of the available seismic data. It led to the reconstitution (using sediment volumes) of the N–S section presented in this paper (Fig. 9a) and of a depth map of the top of the salt without salt movements (Fig. 10). These results are used for the current subsidence modelling. We did not take into account the detailed subsidence history of the post-salt layer; we only considered the thicknesses of the main reconstituted post-salt layers of sediments (Figs. 3 and 9a). 3.2. Nature and thickness of the crust Below the central part of the PCB, the crust is composed of a 12–16 km thick layer (Fig. 11a) with refraction velocity of 6.8–7 km=s. Lower velocities, more typical of the upper continental crust, have not been recognised and the nature of the crustal layer is not well known. Some authors have suggested that it is oceanic crust; for others it is continental crust that has been attenuated during several rifting phases. We will review the different hypotheses about the nature and age of the crustal layer below, in the discussion of geodynamic mechanisms.

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In some parts of the basin, the intermediate velocity crustal layer is underlain by a high-velocity layer (HVL), as shown in Fig. 11b. Volozh et al. (1975) identified two reference seismic boundaries in the lower part of the crust: a refracting boundary with velocity 8.0–8.1 km=s, corresponding to the Moho surface, and a reflecting (lower) boundary. Beneath most of the basin, both horizons coincide in depth and correspond to one single discontinuity. In some areas, the refracting horizon shallows. However, the principal reference lower reflecting horizon remains practically at the same depth level and, when crust and HVL are considered together, the Moho is roughly flat (cf. Fig. 5). The layer between the refracting and lower reflecting boundaries is up to almost 10 km thick. Areas where the HVL is present correspond to gravity anomaly maxima in the Aralsor and Khobda zones (up to 40 mGal). The HVL does not form a continuous layer but makes up individual massifs or lenses in the lower part of the crust (Fig. 11b), which correspond approximately to the loci of minimum crustal thickness of the CPD. Volozh et al. (1975) proposed an eclogitic nature for this HVL. They based their hypothesis on the presence of the gravity maxima and an analysis of seismic waves that indicated density variations with depth in what was otherwise a zone of almost uniform velocity (8.0–8.2 km=s). The inferred density jumps have values of C0.3–0.5 g=cm3 between the intermediate (or ‘true’) crust and the HVL, and 0.2–0.3 g=cm3 between the HVL and upper mantle. The lower reflector identified by Volozh et al. (1975) may correspond to a boundary in the upper mantle identified in recent studies by Kostyuchenko et al. (1996). These authors report a seismic discontinuity in the upper mantle, although its depth is not precisely fixed, between layers with velocities 8.4– 8.7 and 8.0–8.1 km=s. If the HVL is composed of eclogites, the problem is to know how they can exist at the observed depth and when they were formed. Artyushkov (1993) proposed that the sole mechanism responsible for rapid subsidence of the PCB was the eclogitisation of basaltic rocks in the lower crust. However, he did not take into account the observed thickness of the present crust, which also contributes to the subsidence of the basin, and the likelihood that eclogites cannot be formed at this depth where pressures are

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Fig. 11. (a) Map of crustal thickness without the high-velocity layer. (b) Thickness of the high-velocity layer.

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far too low (e.g. Ringwood, 1982). One way around the latter objection is to suppose that the eclogites were initially formed at greater depths=higher pressures in a different tectonic setting, such as a plate convergent one and, subsequently, were transported to their present depth. This kind of mechanism could occur as a result of rapid uplift and denudation related to post-orogenic extensional collapse events (e.g. Le Pichon et al., 1997) or by transport of the high-pressure metamorphic rocks to shallower levels by buoyancy forces combined with retrothrusting (e.g. Behrmann and Ratschbacher, 1989; Platt, 1993; Chemenda et al., 1995). If the eclogites inferred to lie beneath the PCB were formed as a result of collisional tectonics followed by exhumation, then there are two possible times at which the eclogitisation may have occurred. There are two main collisional phases during the history of the PCB. The first is at the end of Riphean–Vendian collision with exhumation taking place during the subsequent rifting phase. Although the timing of this phase of rifting is not well known because of the scarcity of data, most of the consequent subsidence would have occurred in Cambrian–Ordovician times and not later as observed. Nevertheless, the geometrical arguments used to support this possibility are as follows. The inferred eclogitic lenses are roughly limited to the southeast by a deeply penetrating fault bordering the CPD (Elton–Inder fault; Figs. 2–5). This structure dips to the southeast and is sealed by lower Palaeozoic sediments and is thought to be the base of the NCAZ overthrust onto the CPD, possibly representing an end of Riphean–early Vendian ‘suture’ marking the closure of the Pre-Ural Ocean (Volozh, 1991). The area of the present PCB would have been situated at this time on the continental margin of the Pre-Ural Ocean, adjacent to the oceanic crust subducted during the collision. The model supposes that subducted oceanic crust transformed into eclogites at depths of more than 60 km but was subsequently transported to a position beneath the present CPD and, in so doing, inducing an important component of its subsidence (Garagash et al., 1997). The mechanism transporting the eclogitic lens, involving a displacement of at least 300 km from the inferred suture, is not clear. The second possibility is that eclogites formed

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during the Uralian collision; indeed, it could be suggested that they are related to Middle Devonian eclogites observed in the Urals (380 Ma; Matte et al., 1993). In this case, an excessively thick Devonian crust should have been subsequently affected by an important crustal thinning. But the presence of thick older sediments, not extended, seems to contradict this hypothesis, unless the formation of eclogites can be explained by other mechanisms. Other explanations for eclogitisation, not involving collisional tectonics, have also been suggested. Lobkovsky et al. (1996) and Ismail-Zadeh (1998), for example, proposed that eclogites can form in the uppermost mantle as a result of mantle flow and explained the middle Frasnian–Famennian and postDevonian subsidence phase of the PCB in such a manner. There may also be models whereby the HVL is not composed of eclogites; for example, it could comprise ultramafic mantle rocks emplaced near or at the bottom of the crust. However, the same question as discussed in the previous paragraphs about eclogitisation arises about the timing of the emplacement of the HVL. The lack of a confident identification of the true base of the crust and its relationship with the HVL is, in our view, a significant problem hampering a better understanding of basin forming mechanisms. Estimations of crustal thickness in the basin vary; they depend greatly on the chosen interpretations of the HVL. If the HVL is considered to be part of the upper mantle or if eclogitic lenses have been brought to the base of the crust (but are not part of it), then the Moho depth, in the central PCB, is a depth of 32– 36 km. With a reference crustal thickness of 40 km (and taking into account the sedimentary layer), the crustal thinning factor in such a case is 3.3. However, if the HVL is interpreted to be an eclogitic part of the lower crust transformed in situ by a process to be defined, the depth of the Moho below the central basin would be 40–44 km and the thinning factor would not exceed 2. These values of the thinning factor will be used in the subsequent modelling analysis. Another possibility that is considered is that the PCB crust is either oceanic beneath the whole CPD or that it comprises two separate oceanic sub-basins in the areas of the Aralsor and Khobda anomalies.

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4. Discussion of subsidence evolution The evolution of tectonic basin subsidence along the north–south section is reconstructed for ten steps from the end of the Riphean until the Present, focusing on the pre-salt and salt history (Fig. 9). In the schematic present stage, the Lower Permian salt layer is restored as if no salt movements had taken place. It can be compared to the actual present-day section complete with salt domes in Fig. 5. In this work we have mainly adopted the hypothesis that sedimentation commenced during the Riphean and the interpretation of seismic data giving a comparable stratigraphy. A synthetic tectonic subsidence curve in the central part of the basin (point C, Fig. 12) has been drawn according to this hypothesis

Fig. 12. Tectonic subsidence curve of the synthetic stratigraphic column in the centre of the Precaspian Basin (Fig. 3), positioned in Figs. 2 and 5: 1 D curves of tectonic air-loaded subsidence, thin line without corrections, thick line with corrections of palaeowater depth and sea-level changes (Haq et al., 1987 calibrated at a maximum of 250 m, and Ross and Ross in Harland et al., 1989 for the Palaeozoic); 2 D curves of ‘total subsidence’ with sediments, cross without decompaction; thin line with decompaction; 3 D hypothesis of palaeowater depth variations; 4 D diagram of tectonic air-loaded subsidence rates (corrected for water depths and sea level) showing the main tectonic phases of the evolution of the basin.

(column detailed in Fig. 3) but not showing the older ages for reason of scale. This allows us to discuss the main geodynamic phases of the basin evolution. We have compared the subsidence curves in the centre of the basin with four different interpretations of the ages of the older sediments. Yanshin et al. (1980) and Volozh (1991) consider the older sediments as Riphean but with two variants. Nevolin and Fedorov (1995) followed the same approach but with Devonian sediments directly overlying Riphean rocks. In another hypothesis (Shein et al., 1989; Zonenshain et al., 1990; Bankovsky et al., 1991; Kleshev and Shein, 1994), Devonian sediments make up the lowermost part of the sedimentary column. This obviously leads to various geodynamic evolution scenarios of the basin in respect to the process of continental rifting and=or oceanisation in Riphean or Devonian times, a problem that requires further investigation. Emplacement of oceanic crust below the PCB has been proposed by many authors. This could have occurred during Riphean (Nevolin, 1978) or Riphean–early Palaeozoic (Brazhnikov, 1993; Rikhter, 1997), or even Devonian times (Shein et al., 1989; Zonenshain et al., 1990; Volchegursky et al., 1995). The earliest possible depositional stage in the basin evolution occurred during the Riphean when the eastern margin of the European Platform underwent pervasive rifting. The Pachelma rift (to the northwest of the PCB; Fig. 1) was formed at this time, linking the area of the PCB to grabens in the interior of the platform. If the older sediments are considered as Riphean, the subsidence history of the basin would begin with crustal thinning at this time. The basin is underlain during this rifting phase by continental crust on the passive continental margin of the Pre-Uralian Ocean extending towards the southeast. Alternatively, this crust could be almost oceanic and of Riphean age (Nevolin, 1978), in which case, if the oceanic crust had followed conventionally understood laws of oceanic lithosphere cooling, most of the subsidence of the basin should have occurred much earlier (in the Vendian–Cambrian) than observed. This step is followed at the end of the Riphean by closure of the Pre-Uralian Ocean, possibly accompanied by the formation of a foreland flexural basin and related accumulation of sediments. During Vendian–Ordovician times (Fig. 9j), a second rifting event occurred. This event is not clearly

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defined on the tectonic subsidence curve because the data are too scarce. Hence, the tectonic subsidence rate is only averaged for this period. In the Late Cambrian, the Uralian region underwent intense rifting. This created a new (or renewed) passive margin at the eastern border of the EEC and oceanic crust began to develop to the east of the basin before the Middle Ordovician. After a major depositional hiatus at the level of the P3 seismic horizon (end of Silurian–earliest Devonian; Fig. 9i), subsequent rifting during the (Middle–Late) Devonian affected most of the eastern margin of the EEC (Fig. 9h). The intensity of rifting at this time must be much greater for the hypothesis that the oldest sediments in the basin are of Devonian age (Shein et al., 1989; Zonenshain et al., 1990; Bankovsky et al., 1991; Kleshev and Shein, 1994). In the Devonian oceanisation hypothesis, the oldest sediments should be Devonian, linked to the problem of chronostratigraphic interpretation of the seismic deep levels. During this period, the deposition of sediments could only keep up with subsidence near the margins of the PCB, while water depths elsewhere progressively increased. Basinal marls, carbonates and shales were deposited in most of the basin. To the south, these grade into deeper water siliceous facies and, to the west and north, they change into thicker shallow-marine carbonates and clastics. In most of the basin, the depositional patterns established at the end of the Devonian persisted through the Tournaisian (Fig. 9g). Beginning in the middle Frasnian, a thick terrigenous clastic series known as the Zilair Formation was deposited to the southeast and east of the outer shelf margin in a submarine trench or foredeep basin. Being blocked by the presence of a series of atolls (Volozh, 1991 and others) that had developed on the NCAZ, these sediments never spilled northwestward into the central PCB itself. This area of atolls is interpreted as erosional features by Fries and Liger (1995). On the subsidence curve constructed for the centre of the CPD (Fig. 12), it is not possible to date accurately the rifting phase; the Late Devonian is grouped with the Early Carboniferous (Tournaisian). So the noticeable acceleration of subsidence is dated as Late Devonian–Early Carboniferous. In the Dniepr– Donets Basin (the nearest basin where an extensive subsidence analysis has been performed; cf.

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van Wees et al., 1996), the main rifting phase is Frasnian–Famennian and leads to tectonic subsidence of up to 1500 m. The passive margin of the EEC existed until at least the Early Carboniferous. Meanwhile, subduction and continental accretion continued along the active margin of the Kazakhstan–Kirghizian Block, which gradually drifted westward. During the Late Palaeozoic, the EEC and the Kazakhstan–Kirghizian Block collided, creating the Urals. A similar situation developed to the south of the PCB where an assemblage of microplates began to encroach upon the EEC during the Carboniferous (Zonenshain et al., 1990) implying, therefore, the presence of a subduction zone. Krylov et al. (1994) assumed the existence of an island arc to the southeast of the basin until the Early Carboniferous. During the Artinskian (Early Permian) orogenic processes apparently took place along the southern and eastern margins of the EEC (collision of central Kazhakstan with the EEC and rapprochement with the Turan and Scythian blocks). The PCB was isolated during the Kungurian (until the Triassic) in the shape known today. Collision processes welded together these continental masses and resulted in mountain building at their margins, providing a source of abundant terrigenous material to be deposited in the adjoining depressions. We can see on the subsidence curves (Fig. 12) that the interpretation of the tectonic evolution of the PCB during Carboniferous–Permian times is strongly influenced by the palaeobathymetric interpretations. In our hypothesis of a deep depositional environment persisting from the end of the Devonian until the Permian, we have to assume significant tectonic subsidence during this period. In the interpretation of shallow to continental environment during the Carboniferous (Fries and Liger, 1995) most of the subsidence would have occurred during the Permian. In this case, Fries and Liger (1995) proposed that the PCB was initiated only during the Early Permian as a post-orogenic collapse successor basin. The subsidence curve (Fig. 12) displays an important event at the end of the Carboniferous–Permian and Triassic with some decrease in the tectonic subsidence rate subsequently. Tectonic subsidence phases are averaged because of the problem of chronostratigraphic resolution. The Late Carbonif-

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erous–Permian–Triassic phase of rapid subsidence has been interpreted as due to the loading of the Urals on the east of the basin and of the Karpinsky swell orogen on the southeast (Nikishin et al., 1996 and others). The loading of the Urals created a flexural foreland basin on its western side but which is separated from the east of the PCB by a zone of structural highs that is tilted towards the east by the progression of nappes to the west. But considering the usual wavelength of the flexural basin deformation, it seems unlikely that deformation over 900 km away induces subsidence near the western side of the PCB. Nevertheless this is a time of collision and two possibilities exist. The first is a phase of extension, as assumed in the Dniepr–Donets Basin where a late-early Visean rifting phase (with extreme subcrustal lithospheric thinning) induced 500 m of tectonic subsidence (van Wees et al., 1996). The extension may have resulted from back-arc opening behind a north-directed subduction zone situated to the south, prior to the accretion of the Ustyurt–Karpinsky block to the EEC (Zonenshain et al., 1990). The second possibility is subsidence induced by mantle flow, associated with Ural subduction, driven below the overriding plate by the descending slab. This mechanism has been proposed for long-wavelength tilting of several basins. For example, Mitrovica et al. (1996) used this model on the Russian Platform to model two phases of tilting from the Devonian to the Permian. It can explain the longwavelength (more than 1500 km) tilting of platforms combined with flexure closer to overthrust loads that produce narrower, superimposed, foreland basins. This type of subsidence, however, is dynamically supported; the cessation of subduction leads normally to an uplift of the area. A more accurate analysis of the complete subsidence and uplift behaviour is needed along an east–west section of the PCB to test this hypothesis. The post-salt evolution of the PCB involves a thick Permo–Triassic series. There are some indications, from the southern part of the PCB, that an extensional stress field prevailed during this time (Lamber and Stepakova, 1990). Permo–Triassic rifting occurred in the Mangyshlak (to the southeast; Fig. 1), in the Dniepr–Donets trough and in other basins from west of the Black Sea to the Scythian

Plate (Nikishin et al., 1998). This major continental rift system may correspond to a general collapse of the orogenic areas formed during the Variscan– Hercynian collisional phase. Half-grabens formed at this time were subsequently inverted during the Late Triassic and Jurassic, at the time of closure of Palaeo-Tethys. In the PCB itself, the accumulation of a thick sedimentary section continued practically without interruption during the late Palaeozoic and Mesozoic. The geological development of the PCB during the Cenozoic was comparatively quiescent in terms of basement tectonics. But throughout the Mesozoic and Cenozoic, halokinetic processes were very active, and this continuously changed the sedimentation and depocentre patterns. Yet basin subsidence accelerated during the Neogene, especially in the west-southwest of the area and during the Pliocene. This subsidence was possibly due to lithospheric synclinal downwarping in the contemporary compressional stress field (Nikishin et al., 1997).

5. Summary and conclusions The history of the PCB has been long and complicated and is far from being well understood. This is largely due to the uncertainty of the interpretation of the stratigraphy of the basin fill and the nature of the underlying crust. The age of the deepest sediments, known only from seismic data and correlations from the margins, is either Riphean or Devonian. The velocity of the crust underlying the basin is high and it is interpreted either as thinned continental crust or as oceanic crust. The basin has been subjected to a succession of periods of crustal thinning and of compression related to subduction and collisional tectonics in the neighbouring Urals, Caucasus, Mangyshklak and Karpinsky areas. It has led to different phases of accelerated tectonic subsidence during Riphean, Vendian–Ordovician, LateDevonian, end of Carboniferous–Permian, Triassic and Pliocene times. During some phases of its evolution the PCB was situated on the continental margin of a neighbouring ocean and sometimes possibly in a back-arc opening setting. The PCB was never directly a foredeep basin of the collision belts of the Urals and the Caucasus, as the distance from these belts is larger than the

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wavelength of the flexural deformation due to thrust loading. However, it was influenced by compression in the broad foreland setting of these mobile belts. In this setting, mechanisms such as propagation of compressive forces and mantle processes could lead to accelerations of basin subsidence coinciding with the main tectonic events of regional importance. Different mechanisms can be envisaged to explain the various subsidence phases observed in the PCB. Crustal thinning, oceanisation, eclogitisation and the influence of compressive events must all be considered as possible factors in modelling the subsidence. Preliminary modelling results have shown that the total tectonic subsidence observed is in equilibrium with the underlying crustal column (with velocities of 6.8–7 km=s) and the HVL (at either the base of the crust or in upper mantle), taking into account the thickness and density of this layer (Brunet et al., 1995). The key modelling parameters are mainly concerned with the variables defining the nature of the crust underlying the basin. Thinned continental crust or oceanic crust (of either Riphean or Devonian age, metamorphosed or not) all must be considered when reconstructing the evolution of the PCB (e.g. Brunet et al., 1998).

Acknowledgements The data were collected and interpreted by the Russian team during several years and the synthesis work has been financially supported by the PeriTethys Programme (grants 94–96=61). We thank J.A. Mulock Houwer, J.-C. Rigenbach and R.A. Stephenson for their constructive review of the first version of the paper and for their help to improve the English, and F. Lethiers for checking the dating of fauna, and A. Chemenda for discussions about eclogites.

References Akhmetshina, L.Z., Bulekbaev, Z.E., Gibshman, N.V., 1993. Devonian of the eastern flank of the Precaspian syneclise (in Russian). Otechestvennaya Geol. 2, 42–48. Artyushkov, E.V., 1993. Physical Tectonics. Nauka, Moscow. Bankovsky, S.Yu., Ivanova, T.D., Shein, V.S., 1991. The facial– palaeogeographical mapping of the oil and gas-bearing regions

105

based on Late Palaeozoic plate reconstructions (on the example of the southern-east part of the Eastern Europe Continent) (in Russian). In: Geodynamics and Hydrocarbon Potential of the Sedimentary Basins of the USSR. VNIGNI, Moscow, pp. 95–105. Behrmann, J.H., Ratschbacher, L., 1989. Archimedes revisited: a structural test of eclogite emplacement models in the Austrian Alps. Terra Nova 1, 242–252. Brazhnikov, O.G., 1993. Oil-Bearing Potential of the Precaspian Depression from the Plate Tectonics Positions (in Russian). Doctorate Thesis, GANG, Moscow. Brunet, M.-F., 1989. Me´thode d’e´tude quantitative de la subsidence. In: Dynamique et me´thodes d’e´tude des bassins se´dimentaires. Technip, Rueil, pp. 87–98. Brunet, M.-F., 1991. Subsidence et ge´odynamique du bassin d’Aquitaine. Relations avec l’ouverture de l’Atlantique. The`se de doctorat d’Etat, Me´m. Sci. Terre Univ. P. et M. Curie, nº 91-8, Paris. Brunet, M.-F., Volozh, Yu.A., Mauffret, A., Lobkovsky, L.I., Purev, U., Cadet, J.-P., 1995. First elements of modelling the subsidence evolution in the Peri-Caspian basin. Proc. AAPG, International Conference and Exhibition Nice, pp. 10–11A. Brunet, M.-F. and the North-East Peri-Tethys Margin Modelling Team, 1998. Dynamics of the Peri-Tethyan basins in the Caucasus–Caspian area. European Association of Geoscientists and Engineers Meeting, Leipzig. Abstr. B03-46. Chemenda, A.I., Mattauer, M., Malavieille, J., Bokun, A.N., 1995. A mechanism for syn-collisional rock exhumation and associated normal faulting: results from physical modelling. Earth Planet. Sci. Lett. 132, 225–232. Ensepbaev, T., Izart, A., Cathelineau, M., Joltaev, G., Vachard, D., Boulekbaev, Z., 1998. La bordure Est du bassin Pre´caspien et le bassin d’avant-pays de l’Oural (Kazakhstan) au Carbonife`re et Permien infe´rieur. In: Crasquin-Soleau, S., Izart, A., Vaslet, D. and De Wever, P. (Eds.), Peri-Tethys: Stratigraphic Correlations 2, Geodiversitas. Editions Scientifiques du Muse´um, Paris 20(4), 687–700. Fries, G., Liger, J.L., 1995. Geodynamic evolution of the Eastern margin of the Precaspian basin from Devonian to Permian. Proc. AAPG, International Conference and Exhibition Nice, p. 23A. Garagash, I.A., Lobkovsky, L.L., Volozh, Yu.A., Brunet, M.-F., 1997. Numerical modelling of Precaspian Basin subsidence as a result of eclogite penetration beneath earth crust during Vendian collision. Proc. 2nd Peri-Tethys Workshop, Moscow State University, Moscow, pp. 12–13. Haq, B.U., Hardenbol, J., Vail, P.R., 1987. Chronology of fluctuating sea levels since the Triassic. Science 235, 1156–1167. Harland, W.B., Armstrong, R.L., Cox, A.V., Craig, L.E., Smith, A.G., Smith, D.G., 1989. A Geologic Time Scale. Cambridge University Press, Cambridge, 263 pp. Ismail-Zadeh, A.T., 1998. The Devonian to Permian subsidence mechanisms in basins of the East-European Platform. J. Geodyn. 26, 69–83. Kleshev, K.A., Shein, V.S., 1994. Geodynamic analysis of oil and gas basins in Russia and adjacent countries. In: Roure, F., Ellouz, N., Shein, V.S., Skvortsov, I. (Eds.), Geodynamic

106

M.-F. Brunet et al. / Tectonophysics 313 (1999) 85–106

Evolution of Sedimentary Basins, International Symposium. Technip, Rueil, pp. 1–18. Kostyuchenko, S.L., Egorkin, A.V., Solodilov, L.N., Zolotov, E.E., 1996. Genetic types of Precambrian rifts of the East-European platform’s Nizhne–Volga divergent belt based on deep studies (in Russian). Razvedka Okhrana Nedr 4–5, 46–54. Krylov, N.A., Avrov, V.P., Golubeva, Z.V., 1994. Geological model of the antesaliferous complex of the Precaspian basin and petroleum perspectives (in Russian). Geol. Nefti Gaza 6, 35–38. Lamber, G.I., Stepakova, G.I., 1990. Influence of tectonics on the sedimentation process in the Precaspian Depression during the Triassic (in Russian). Geol. Nefti Gaza 4, 12–14. Le Pichon, X., Henry, P., Goffe´, B., 1997. Uplift of Tibet: from eclogites to granulites — implications for the Andean Plateau and the Variscan belt. Tectonophysics 273, 57–76. Lobkovsky, L.I., Cloetingh, S., Nikishin, A.M., Volozh, Yu.A., Lankreijer, A.C., Belyakov, S.L., Groshev, V.G., Fokin, P.A., Milanovsky, E.E., Pevzner, L.A., Gorbachev, V.I., Korneev, M.A., 1996. Extensional basins of the former Soviet Union — structure, basin formation mechanisms and subsidence history. Tectonophysics 266, 251–285. Matte, P., Maluski, H., Caby, R., Nicolas, A., Kepehinskas, P., Sobolev, S., 1993. Geodynamic model and 39 Ar=40 Ar dating for the generation and emplacement of the high pressure (HP) metamorphic rocks in SW Urals. C. R. Acad. Sci. Paris 317 (II), 1667–1674. Mitrovica, J.X., Pysklywec, R.N., Beaumont, C., Rutty, A., 1996. The Devonian to Permian sedimentation of the Russian Platform: an example of subduction-controlled long-wavelength tilting of continents. J. Geodyn. 22, 79–96. Nevolin, N.V., 1978. Deep structure of the Pre-Caspian basin. Geotectonics 12, 192–199. Nevolin, N.V., Fedorov, D.L., 1995. Palaeozoic pre-salt sediments in the Precaspian petroliferous province. J. Pet. Geol. 18, 453–470. Nevolin, N.V., Kunin, N.Ya., 1977. Geological structure and oil and gas perspectives of salt dome basins of continents, based on the geophysical data (in Russian). Nedra, Moscow. Nikishin, A.M., Ziegler, P.A., Stephenson, R., 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.Y., Shalimov, I.V., 1996. Late Precambrian to Triassic history of the East-European Craton: dynamics of sedimentary basin evolution. Tectonophysics 268, 23–63. Nikishin, A.M., Brunet, M.-F., Cloetingh, S., Ershov, A.V., 1997. Northern Peri-Tethyan Cenozoic intraplate deformations: influence of the Tethyan collision belt on the Eurasian continent from Paris to Tian-Shan. C. R. Acad. Sci. Paris 234 (IIa),

49–57. Nikishin, A.M., Cloetingh, S., Brunet, M.-F., Stephenson, R.A., Bolotov, S.N., Ershov, A.V., 1998. Scythian Platform, Caucasus and Black Sea region: Mesozoic–Cenozoic tectonic history and dynamics. In: Crasquin-Soleau, S., Barrier, E. (Eds.), PeriTethys Memoir 3: Stratigraphy and Evolution of Peri-Tethyan Platforms. Me´m. Mus. Natl. Hist. Nat. Paris 177, 163–176. Odin, G.S., 1994. Geological Time Scale 1994. C. R. Acad. Sci. Paris 318 (II), 59–71. Platt, J.P., 1993. Exhumation of high-pressure rocks: a review of concepts and processes. Terra Nova 5, 119–133. Rikhter, 1997. Precaspian Depression — a relict of the Palaeozoic ocean (in Russian)? Nedra Povolzh’e Caspii 12, 3–19. Ringwood, A.E., 1982. Phase transformations and differentiation in subducted lithosphere: implications for mantle dynamics, basalt petrogenesis, and crustal evolution. J. Geol. 90, 611– 643. Shein, V.S., Payrazyan, V.A., Razmyshlyaev, A.A., Soloviev, B.A., 1989. Plate tectonic scheme and hydrocarbon potential of the Precaspian Depression. In: Prospecting of Oil and Gas Fields. VNIGRI, Moscow, pp. 28–41 van Wees, J.D., Stephenson, R.A., Stovba, S.M., Shymanovskyi, V.A., 1996. Tectonic variation in the Dniepr–Donets Basin from automated modelling of backstripped subsidence curves. Tectonophysics 268, 260–283. Volchegursky, L.F., Vladimirova, T.V., Kapustin, I.N., Natapov, L.M., 1995. Evolution of the Precaspian basin during Middle and Late Palaeozoic (in Russian). Otechestvennaya Geol. 2, 44–48. Volozh, Yu.A., 1991. Sedimentary basins of western Kazakhstan (applications of seismostratigraphic analysis) (in Russian). Doctorate Thesis, Geological Institute, Russian Academy of Sciences, Moscow. Volozh, Y.A., Sapozhnikov, R.B., Tsimmer, V.A., 1975. Structure of the crust in the Caspian Depression. Int. Geol. Rev. 19, 25–33. Volozh, Y.A., Pilifossov, V.M., Aznabaev, E.K., Puchkov, V., 1991. Geology and Hydrocarbon Resources of the Precaspian Basin. Technical Report, Earth Sciences and Resources Institute, University of South California, 2 Vols. Volozh, Y.A., Volchegurskii, L.F., Groshev, V.G., Shishkina, T.Yu., 1997. Types of salt structures in the Peri-Caspian Depression. Geotectonics 31, 204–217. Yanshin, A.L., Avrov, V.P., Volozh, Yu.A., V., Sapozhnikov, R.B., Shlezinger, A.Y., Yurov, Yu.G., 1980. Structure and paleotectonics of Palaeozoic subsalt formations in the southern part of the Precaspian trough. Geotectonics 14, 194–202. Zonenshain, L.P., Kuzmin, M.I., Natapov, L.M., 1990. Geology of the USSR: a plate-tectonic synthesis. Am. Geophys. Union, Geodyn. Ser. 21.