Late Cenozoic burial history and dynamics of the Northern Caucasus molasse basin: implications for foreland basin modelling

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Tectonophysics 313 (1999) 219–241 www.elsevier.com/locate/tecto

Late Cenozoic burial history and dynamics of the Northern Caucasus molasse basin: implications for foreland basin modelling Andrei V. Ershov a,b,Ł , Marie-Franc¸oise Brunet b , Maxim V. Korotaev a , Anatoly M. Nikishin a , Sergei N. Bolotov a a

b

Geological Faculty, Moscow State University, Vorobievy Gory, 119899 Moscow, Russia ESA 7072 UPMC–CNRS, case 129, Universite´ Pierre et Marie Curie, 4 place Jussieu, 75252 Paris Cedex 05, France Received 18 November 1997; accepted 14 August 1998

Abstract The collisional history of the Caucasus segment of the Alpine–Himalayan fold belt started at the end of the Eocene. The associated development of the Northern Caucasus foreland basin occurred during two syn-collisional stages, each displaying different subsidence patterns. The 34–16 Ma (Maikopian or pre-foreland stage) displays a long-wavelength subsidence of a broad area whereas the 16–0 Ma (foreland stage) displays asymmetrical foreland subsidence. The first is correlated with the termination of subduction in the southern area and can be associated with mantle flow induced by the re-equilibration of the subducted slab. The along-strike configuration of the second, molasse basin, stage contradicts the hypothesis in which topographical loading is considered to be the main control on foreland subsidence. There is a clear anti-correlation between basin depths and orogen heights: the deepest parts of the basin are at the tips of the orogen, where mountain heights are negligible while the area adjacent to the highest mountains (Central Caucasus) is uplifted. The influence of other types of loading in foreland basin development has therefore been investigated and it has been possible to find a good fit with both gravity and basin architecture by including only two additional model parameters into the model. These are crustal and lithospheric thickening=thinning. The results demonstrate that crustal thickening and removal of lithospheric roots are responsible for supporting the high Central Caucasus Mountains and uplift of adjacent areas. The subsidence of the basins at the orogen tips is explained by loading of lithospheric roots. Both effects are important in the geographically intermediate areas. In general, it is concluded that topography should not be considered as the main control on foreland subsidence, but only as one of several counterbalancing mechanisms. The existing flexural model therefore needs improvements such that it can be related directly to ‘subsurface loading’ by real structures and processes during collision.  1999 Elsevier Science B.V. All rights reserved. Keywords: foreland basin; Caucasus; collision; flexural model; gravity; subsidence; lithospheric root

1. Introduction We use the term foreland basin in its general sense as the basin adjacent to a fold-and-thrust belt (DickŁ Corresponding

author. E-mail: [email protected]

inson, 1974), making no difference between ‘retro’ and ‘pro’ foreland or foreland and hinterland. Price (1973) introduced the concept of isostatic flexural response of the lithosphere to orogenic loading as being a first-order control on foreland basin formation. One of the main things supporting this concept

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 7 - 3

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is the gravity signature of mountain belts, which typically show the presence of high-amplitude negative anomalies in adjacent, undeformed areas. This is difficult to explain in the framework of local isostasy and it is considered as evidence for regional compensation of the orogenic load, especially in view of the successful application of a flexural model to the response of oceanic lithosphere to surface loading (e.g. Turcotte and Schubert, 1982). Jordan (1981) and Beaumont (1981) modelled foreland basin subsidence as the flexural response of lithosphere to the topographic load, in the framework of purely elastic and visco-elastic models, respectively. The concept was subsequently applied in numerous modelling studies. It was often recognised, however, that topographic loading itself was not sufficient to explain the observed subsidence of many foreland basins. Additional factors were often introduced to fit the observed data, including subsurface crustal loading (e.g. Karner and Watts, 1983), the effects of the bending moment of the associated subducted slab (e.g. Sheffels and McNutt, 1986), and loading of lithospheric roots (e.g. Brunet, 1986). In this paper we present the Late Cenozoic burial history and results of flexural modelling of the Northern Caucasus foreland basin (Fig. 1). This basin is well studied, as it is one of the oldest oil-productive regions in the world. The geological

history of the Caucasus was described in English by Gamkrelidze (1986), Zonenshain and Le Pichon (1986), Philip et al. (1989), and Adamia et al. (1992). A description of the main stages of the geological history and a reconstruction of the palaeogeography and palaeotectonics of the Scythian Platform (Northern Caucasus, Fore-Caucasus and Crimea) are presented briefly in Nikishin et al. (1998a,b). Further, Mikhailov (1993) published a 2D kinematical restoration and palaeotectonic analysis of the evolution of the Terek–Caspian trough (cf. Fig. 1). Although the tectonic structure of the region appears to be rather straightforward, a simple geodynamical model explaining the subsidence of the molasse basin does not exist. The Fore-Caucasus foreland basin appears similar to other foreland basins from the viewpoint of shape, lithofacies, and synchronism with orogenic processes and 2D flexural modelling of some sections through the basin shows good agreement with observations (Ruppel and McNutt, 1990; Ershov et al., 1998). However, a paradox appears when the basin architecture is considered along strike of the Caucasus orogenic belt: the area adjacent to the highest part of the orogen did not subside and, in turn, the deepest subsidence occurred near the lowest part. Indeed, difficulties with the standard flexure due to the orogen loading model in the North Caucasus led some researchers to declare

Fig. 1. Schematic map of the investigated area showing positions of backstripped pseudo-wells (circles), seismic sections (black lines), and crustal-scale synthetic sections (grey lines).

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it as unrealistic (Artyushkov, 1993) and to investigate alternative subsidence mechanisms (e.g. Artyushkov, 1993; Mikhailov et al., 1997). This paper begins with a synthesis of the available regional-scale geological and geophysical data in order to reconstruct the tectonic structure and geological history of the region. This forms the basis for considering various possible geodynamical models. To fit the observations we need at least a partial reconsideration of the conventionally adopted first-order processes that control foreland basin subsidence.

2. Structural setting The Caucasus region can be subdivided into the Great Caucasus Orogen, the Fore-Caucasus foreland basins, to the north, and the Transcaucasus molasse basins, to the south (Fig. 1). The Great Caucasus Orogen is a northwest-trending, nearly linear mountain belt, comprising an eastern and central-western segment. It lies along the deformed margin of Scythian Platform, which extends along the southern edge of the Russian Platform. The Fore-Caucasus basins of the Scythian Platform include a western and an eastern basin, separated by the Stavropol High. These are bounded on the north by the uplifted Karpinsky Swell, formerly a part of the Dnieper– Donets–Karpinsky rift system (e.g. Nikishin et al., 1996). The deepest part of the Eastern Fore-Caucasus basin is the Terek–Caspian depression. The deepest part of the Western Fore-Caucasus basin is the Indol–Kuban depression. The Transcaucasus basins include the western (Rioni) and the eastern (Kura) basins bounded by the arcuate Lesser Caucasus Mountains and opening respectively towards the Black Sea and the South Caspian Sea. Representative crustal sections of the Caucasus and its molasse basins are displayed in Fig. 2. Lines A and B are based on deep seismic sounding (DSS) sections (Krasnopevtseva, 1984) Volgograd– Nakhichevan and Stepnoe–Bakuriani, in the eastern and central parts of the Fore-Caucasus, respectively. Geological data from field observations in the orogenic area, and from seismic and well data in the basins were used to construct the upper parts of the sections. Section C was constructed on the ba-

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sis of a Moho map of the area (Volvovsky et al., 1989), the Black Sea geological section of Robinson et al. (1996), and a seismic section in the Western Fore-Caucasus basin (see below). The Great Caucasus Orogen is upthrust mainly to the south, towards the Transcaucasus basins, with local retro-thrusting to the north (Fig. 2). It was formed by Alpine deformation of the Scythian and Transcaucasus plates and an intermediate depression. The southern edge of the Scythian plate is expressed in the central-northern area by the main thrust (Fig. 2, line A). The northern boundary of the Transcaucasus plate is overridden by the Scythian plate and folded sediments of the southern slope of the Great Caucasus. These sediments were initially lying in an intermediate basin (the Great Caucasus trough). The northern slope of the Great Caucasus is characterised by a comparatively simpler tectonic style (Fig. 2): a monocline transition to the Scythian Platform (central part, line B), retro-thrusting into the molasse basin (eastern part, line A), and a steep monocline transition to the molasse basin (western part, line C). The Eastern and the Western Fore-Caucasus basins have an asymmetrical shape with deepening towards the orogen and shallowing towards the foreland (Fig. 2). The Transcaucasus basins are faulted by numerous south-verging thrusts. In general, the sediments of the former consist of three main units: Jurassic–Eocene platform cover, Oligocene–Early Miocene shale, and Middle Miocene–Quaternary molasse. The last does not occur in the Central Fore-Caucasus area (Fig. 2, line B), which has been recently uplifted.

3. Brief overview of the Cenozoic Caucasus collision The Cenozoic evolution of the region is defined by successive collisions of Eurasia with Tethyan terranes and Arabia (cf. Fig. 3). The polyphase Great Caucasus orogeny began at the end of the Eocene (Milanovsky, 1968), prior to which a deep-water basin (the Great Caucasus trough) was situated in the same place. Zonenshain et al. (1990) and Nikishin et al. (1998b) suggested that this basin was similar to the present Black Sea and the South Caspian

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Fig. 3. Present-day tectonic setting of the Caucasus region. The plate velocities (mm=year, in a Eurasia fixed framework) and shortening values are shown on the basis of geodetic observations (Reilinger et al., 1997).

basins. That is, like the modern basins, it would have been a back-arc basin, related to subduction south of the Pontides–Transcaucasus Jurassic–Eocene volcanic belt, with oceanic or very thin continental crust (Zonenshain et al., 1990; Dercourt et al., 1993). Cessation of subduction-related volcanism in the Pontides–Transcaucasus belt occurred at the end of

the Eocene when the north-dipping subduction zone evolved into a collisional system (Sengor and Kidd, 1979; Zonenshain et al., 1990). The central segment of the Great Caucasus Orogen developed since this time (Milanovsky, 1968; Nesmeyanov, 1992). The zone of Caucasus underthrusting therefore was probably also formed at this time. The main collisional

Fig. 2. Crustal-scale synthetic sections through the Eastern Caucasus (A), Central Caucasus (B) and Western Caucasus (C) areas. Positions of the synthetic sections are shown in Fig. 1 (grey lines). The first two sections are drawn on the basis of deep seismic sections Volgograd–Nakhichevan (A) and Stepnoe–Bakuriani (B) (modified after Krasnopevtseva, 1984). The upper parts are constructed on the basis of field geological data in the area of the orogen and from seismic and well data in basins. Section (C) is constructed from the section of Robinson et al. (1996) in the Black Sea, the Moho map of Volvovsky et al. (1989) and the seismic section in the Western Caucasus basin. Stars mark the epicentres of large earthquakes.

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phase began in the Middle Miocene (Milanovsky, 1968; Sengor and Yilmaz, 1981). Its effects included the lateral expulsion of the rigid Anatolian Plate (to the west) and Iran (to the east), distributed deformation in the broad area between the Arabian margins and the Caucasus orogen (Fig. 3), closure of the Great Caucasus trough followed by orogeny in the Caucasus–Crimea region and formation of the molasse basins, uplift of a peripheral lithospheric bulge, intracratonic folding=thrusting inside the East-European Platform up to 1200 km north of the Great Caucasus (Nikishin et al., 1997), and possibly to the rapid syn-compressional subsidence of the Black Sea, Precaspian (Nikishin et al., 1998a; Brunet et al., 1999, this issue) and South Caspian basins (Brunet et al., 1997). Geodetic measurements indicate that about half of the recent Arabia-relative-to-Eurasia approach velocity is accommodated by thrusting in the Caucasus area (Reilinger et al., 1997).

4. Late Cenozoic burial history of the Fore-Caucasus basins 4.1. Method of burial history restoration Three representative seismic sections (A, B and C; Fig. 1) were used to restore a 2D burial history of the basin. There is no good seismic section through the Central Fore-Caucasus and we have used the section on its eastern boundary (section B). Section B is not ‘archetypal’ for the Central Fore-Caucasus (as sections A and C are for the Eastern and Western parts), but, nevertheless, it reflects the main features of subsidence of the central area. The sections were subdivided into a set of pseudowells and a 1D reconstruction was made for each pseudo-well. We used the standard backstripping technique (e.g. Steckler and Watts, 1978) to restore the 1D burial history. The compaction correction was made on the basis of porosity–depth dependencies, obtained by a least-square approximation of porosity data for the Western Caucasus basin (Ershov et al., 1998). The two main difficulties encountered were the choice of appropriate absolute dates and palaeobathymetric corrections. The first difficulty arises from the fact that a common time scale does not exist for the northern

intracontinental Tethyan basins for the Middle–Late Cenozoic. In effect, there is no overall consensus on absolute ages of Late Eocene–Quaternary regional stratigraphic units in the area. Numerous erosional events led to the absence of some parts of the section, and the use of palaeomagnetic data is problematic. For absolute dating of Maikopian units (Oligocene– Early Miocene) we have used the biostratigraphical correlation based on the micropalaeontological zonation of Popov et al. (1993). For the Late Miocene– Quaternary (13.7–0.5 Ma) radiometric ages, based on the analysis of 238 U fission tracks in volcanic glass from tuffs, were adopted (Chumakov et al., 1992; Chumakov, 1993). Fig. 4 shows a chronostratigraphical chart of the regional units in comparison with the global time scales of Harland et al. (1989) and Odin (1994). The lithostratigraphy of different segments of the Northern Caucasus molasse basin are also shown. The presence of clinoformal complexes on the eastern and central sections allows the restoration of palaeobathymetry for the corresponding stages. A special algorithm, described elsewhere (Ershov et al., 1998), was used. The basic principles are as follows. For each prograding sedimentary body (clinoform), the shelf-to-slope-transition point was defined and the level of the shelf was positioned horizontally at eustatic sea level. This allows restoration of the palaeobathymetry of the previously deposited clinoform under the shelf of the next clinoform. The procedure is applied for each clinoform from top to bottom, therefore restoring the palaeobathymetry for each in turn. Corrections for compaction and regional isostatic compensation were made at each step. Eustatic sea level changes were adopted from Haq et al. (1987) and calibrated at a maximum of 250 m. Reconstructions of the Late Cenozoic burial history along three seismic sections crossing the Eastern, eastern-central and western segments of the Fore-Caucasus basin and nine pseudo-wells extracted from these sections are presented in the Figs. 5 and 6. 4.2. Overview of basin evolution The Fore-Caucasus basin originated in the Triassic as a rift basin (Nikishin et al., 1998a,b). Most

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Fig. 4. Chronostratigraphical charts and tectonic subsidence rates (m=m.y.) of some representative wells for the different parts of the Fore-Caucasus basin for Oligocene–Quaternary time. See text for explanation of absolute dates.

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Fig. 5. 2D burial history restoration along the regional seismic sections (black lines in Fig. 1) for four selected Oligocene–Quaternary time slices. The lithological composition of the sections and average tectonic subsidence rates are shown in Fig. 4. The 1D burial histories for some pseudo-wells are shown in Fig. 6.

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Fig. 6. Late Cenozoic burial histories (basement subsidence, sea level and palaeobathymetry at the left and the rate of ‘air-loaded’ tectonic subsidence at the right) for nine pseudo-wells (PW) extracted from the modelled seismic sections (Fig. 5) in the Eastern (EFC), Central (CFC) and Western (WFC) Fore-Caucasus. Locations of wells are shown in Figs. 1 and 5. Time scale is after Odin (1994) and Chumakov et al. (1992) (Mkp D Maikopian; Srm D Sarmatian; Me D Meotian; Pn D Pontian; Km D Kimmerian; Ak D Akchagylian).

of the Triassic sedimentary succession was subsequently eroded during the Cimmerian Orogeny and only folded remnants of Triassic sediments are preserved. Deposition of the platform cover of the

basin began in the Middle Jurassic. Thereafter, until the Eocene, the basin formed the shelf margin of the deep-water back-arc mentioned above. During the Paleocene–Eocene (65–34 Ma) the Fore-Cau-

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casus basin was filled by shallow-water carbonates and marls. The subsidence regime was quiet and characterised by a decreasing subsidence rate, probably caused by thermal cooling of the lithosphere after Jurassic and Early Cretaceous phases of lithospheric thinning (Nikishin et al., 1998a,b). The Eocene=Oligocene boundary is marked by the subsidence of the whole basin (Figs. 5 and 6) and usually it is considered as the beginning of the orogenic stage (Koronovsky et al., 1987; Scherba, 1993), which has continued until the present. Sediment lithologies (Fig. 4) and subsidence patterns (Fig. 5) indicate that there have been two main sub-stages of basin evolution during this time: an Oligocene–Early Miocene (34–16 Ma) (Maikopian) stage and a Late Miocene– Quaternary (16–0 Ma) stage. 4.3. Oligocene–Early Miocene (pre-foreland) stage Note on Fig. 5 that sections A and B are situated entirely in the shelf area whereas the southern part of section C continues into the northernmost part of the Great Caucasus trough (probably a deep-water terrace). This fact explains some of the observed differences in the sedimentation patterns of the western (C) and eastern-central (A, B) sections and they will be discussed separately. The beginning of the Oligocene is indicated by an increase of the subsidence rate in the eastern and the central parts of the basin (Fig. 6). It is fixed by the change of Late Eocene shallow-water limestones (e.g. Scherba, 1993; Lozar and Polino, 1997) to Oligocene deep-water clays. Later, the basin was filled by clinoforms (Fig. 5), prograding from the northeast (initially) and northwest (Kunin et al., 1990; Kosova, 1994). Palaeowater depths given by our palaeobathymetric estimations are around 500– 800 m in the central part of the basin and as much as 1200 m in its southern part. In general, subsidence was homogeneous along the basin with a slight tilting to the southeast. Once the basin was filled by the clinoforms, sedimentation continued in shallow-water conditions (Fig. 5). The overall thickness of Maikopian sediments was near 1.7 km in the northern parts, increasing to the southeast. The Oligocene–Early Miocene (Maikopian) sediments in the northern parts of the sections were

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later eroded (Fig. 5, Section 4.4) and the present Maikopian thickness in the northern part is less than its initial value. The reduction of present-day thickness to the north can lead to the erroneous conclusion that the northern part of the basin was uplifted during this time (e.g. Maslyaev, 1990; Alekseev, 1990; Popov and Stolyarov, 1996). However, the erosional nature of this thinning is clearly seen on the seismic sections. The thickness of the Maikopian layer (as well as thicknesses of the shelf parts of every clinoform it comprises) is about constant in the central-northern parts of the sections (Fig. 5) and, by our suggestion, was about the same in the subsequently eroded parts. The western section (Fig. 5) demonstrates a different subsidence pattern in Oligocene–Early Miocene time. Sediments filled a deep-water depression in the southern part of the section. The seismic record shows parallel layering with onlap on the slope surface. The same pattern is characteristic for the Maikopian sediments of the Shatsky Ridge, an underwater uplift further to the south in the Black Sea (e.g. Robinson et al., 1996). It is impossible to discriminate between subsidence before and during the Maikopian stage in this part of the section. Tectonic subsidence during the Maikopian could have been either positive or zero (or even negative). To be definitive, we have accepted some value, but it is quite arbitrary. Subsidence on the shelf was not large; overall thickness of Maikopian sediments in the northern part of the section is near 200 m. Its increase to the south may be associated (at least in part) with the flexural response of lithosphere to the loading of sediments, deposited in the deep-water southern part. 4.4. Middle Miocene–Quaternary (foreland) stage This stage began with a short Chokrakian– Karaganian (15.8–14.7 Ma) event of rapid subsidence, which was most pronounced in the southern areas, adjacent to the northern slope of Great Caucasus. It is also marked by lithofacies changes (Fig. 4). The main changes in lithological composition (Fig. 4) and in basin configuration (Fig. 5) occurred in Middle–Late Sarmatian time (12.2–9.3 Ma). Thereafter, the western and eastern segments of the Great Caucasus were exposed as continental

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areas (Koronovsky et al., 1987; Nesmeyanov, 1992; Scherba, 1993). The Central Caucasus had been uplifted since the Eocene (Nesmeyanov, 1992; Lozar and Polino, 1997). Major changes in the along-strike basin configuration also occurred at this time. At the end of the Sarmatian, the Central Fore-Caucasus began to be uplifted (Fig. 5; Scherba, 1993) while the eastern and western areas continued to subside (Figs. 4–6). The basin became asymmetrical, deepening towards the orogen. It was filled by carbonate sediments and clastic material derived from the Great Caucasus. The Sarmatian succession comprises clinoforms, prograding from the south to the north (Fig. 5). In contrast to the previous stage, the Middle Miocene–Quaternary subsidence patterns of the western and eastern areas are generally similar (Fig. 5), although some differences exist. Sedimentation in the western basin was continuous during the Late Miocene–Early Pliocene (9.3–3.4 Ma), while in the eastern basin periods of sedimentation are punctuated by hiatuses and erosional episodes (Fig. 4). During most of this time the eastern basin was in continental conditions. These events may be associated with periods of opening and closing of the palaeo-Caspian and palaeo-Black seas (Nevesskaya et al., 1984), but the differences between the western and eastern basins also show some regional uplift of the eastern basin. The most significant sea-level drop occurred at the end of the Miocene (6–5.2 Ma). It is related to the Messinian event recorded in the Mediterranean Sea (Chumakov, 1993). This led to the final separation of the palaeo-Caspian Sea from the palaeo-Black Sea (Nevesskaya et al., 1984) and a significant decrease of the palaeo-Caspian sea level. The level of the palaeo-Black sea also fell but not so significantly. As a result, Early Pliocene (5.2–3.4 Ma) sediments are absent in the eastern basin (Fig. 4) while in the northern part of the western basin, deep river incisions were formed (Fig. 5). The rapid rise of sea level in the Late Pliocene (3.4–1.8 Ma) led to clinoformal sedimentation in the eastern basin, when clinoforms filled already existing accommodation space. Sediments were derived from the Great Caucasus Orogen (primarily) and the peripheral bulge (secondarily) to the north of the basin (Fig. 5). The time of the formation of this bulge is difficult to establish precisely but it probably began

to form during Sarmatian–Pontian times (12.2–5.2 Ma; Ershov et al., 1998). In any case, it existed by the beginning of the Akchagylian (3.4 Ma). The amount of erosion in the northernmost part of the eastern section by the Akchagylian was more than 1 km. The distance between the axes of the Great Caucasus and the peripheral bulge is nearly 450 km. Its position roughly corresponds to the position of Karpinsky Swell. In the Quaternary, the whole eastern basin, including the peripheral bulge, subsided and has been covered by sediments. The subsidence=sedimentation patterns in the western basin do not display as much heterogeneity as in the eastern one. During Middle Miocene– Quaternary times, the western basin underwent asymmetrical subsidence. The peripheral bulge occurs only as an area with a decreasing thickness of sediments. The central Fore-Caucasus area was uplifted and its southern part has been partly eroded.

5. Discussion of geodynamical models The subsidence patterns of the two stages of development of the Fore-Caucasus basin are different from the viewpoints of both size and shape. The Oligocene–Early Miocene (Maikopian) subsidence pattern in the Eastern and Central Fore-Caucasus is uniform, slightly inclined to the south. Its characteristic wavelength is much larger than the size of the backstripped sections. In contrast, the characteristic size of the Late Miocene–Quaternary subsidence is about the length of the sections. At this time, the basin has a characteristic flexural shape, deepening towards the orogen and with a forebulge to the north. The difference between subsidence patterns was presumably caused by differences in the processes controlling basin development. The characteristic wavelength of Maikopian subsidence is too great to be explained by the flexural response of lithosphere to vertical loading at its edge. (This is evident in comparison with the flexural subsidence of the next stage.) The cause of this broad subsidence must lie in larger-scale processes, such as subduction-induced mantle flow. Similar subsidence patterns are seen in the Late Cretaceous basin of western North America (Cross and Pilger, 1978; Mitrovica et al., 1989) and the Late Tertiary Taranaki

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basin, New Zealand (Holt and Stern, 1994). In these basins this phase of subsidence has been linked to active subduction processes. Long-wavelength tilting of the continent occurs with the beginning of subduction; cessation of subduction leads to the complementary uplift of the area. In the Caucasus we have a different situation: the Maikopian subsidence is contemporaneous with termination of subduction, and the new zone of underthrusting began to form when a deep-water broad basin already existed. But the cited studies indicate in a general way that mantle flows associated with subduction can produce significant long-wavelength deflection of the surface. The dip of the subducted slab depends on the subduction rate (e.g. Hsui et al., 1990). Change in subduction rate leads to a change in the dip of the slab. This can lead to the dragging “of the mantle wedge above the slab downward, resulting in a depression at the surface”, as it was stated by Hsui et al. (1990). Termination of subduction is followed by a rapid change of subduction rate and, therefore, some ‘rotation’ of the slab to a new equilibrium dip and associated mantle flows above it. We suggest that the Maikopian subsidence was induced by such a mantle flow after the termination of ‘southern’ subduction. This change in slab dip can be responsible for 1 km of excess subsidence in some back-arc areas (Hsui et al., 1990) and we have a similar amplitude of tectonic subsidence during the Maikopian in the Fore-Caucasus basin. This mechanism produces ‘dynamically supported’ subsidence, which should revert to uplift after the termination of slab down-rotation. Such a mechanism could be responsible for the postMaikopian uplift of the Central Fore-Caucasus area. To the east, uplift was masked by the foreland-type subsidence of the next stage. The central area did not undergo foreland subsidence and we will discuss this in the next section. Probably some trace of uplift can be seen in the erosion of peripheral bulge (Fig. 5), where foreland subsidence did not fully mask post-Maikopian uplift. The amplitude of this erosion (around 1 km) is too great to be explained only by elastic flexure (Ershov et al., 1998; as compared with the weaker amplitude of uplift of the peripheral bulge of the western basin, Fig. 5). We did not model the Maikopian subsidence, as the size of the investigated area is smaller than the characteristic

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length of the subsidence. To constrain properly such a model it would be necessary to consider a larger region. The beginning of the foreland stage is marked by the formation of an asymmetry in the basin subsidence pattern. The main subsidence events are simultaneous with stages of orogenic folding (Scherba, 1993) and a correlation exists between the formation of the orogenic belt and the foreland basins. The subsidence pattern is typical for foreland basins formed as a result of elastic flexure of the lithosphere in response to orogen loading. Such basins have been recognised broadly across the Alpine– Himalayan orogenic belt (e.g. Karner and Watts, 1983; Lyon-Caen and Molnar, 1983, 1989; Brunet, 1986). Ruppel and McNutt (1990) analysed gravity anomalies along three selected sections in the Eastern, Central and Western Caucasus and obtained good agreement with observations using a flexure model. Ershov et al. (1998) showed that the shape of the Eastern Caucasus basin can be satisfactorily matched by the shape of an elastic flexural response of the lithosphere to orogenic loading for any time slice during the foreland stage (12.5–0 Ma). But a more general consideration of basin development reveals that the situation is not so simple. Consider the along-strike structure of the molasse basin (Fig. 7). The map shows the amplitude of vertical movements during the Late Sarmatian– Quaternary (i.e. depths to the bottom of the molasse formation in the basins combined with an estimate of orogen uplift). The first evident, but surprising, observation is that the deepest basins are on the tips of the orogen, whereas the area adjacent to the highest mountains (Central Caucasus) is uplifted. This fact immediately rejects topographical loading as the main control on foreland subsidence and necessitates a consideration of other possible loading mechanisms. The sense of loading must be positive (to produce uplift of the central area) as well as negative. Possibilities include a change of crustal=lithospheric thickness due to shortening and stacking of material or the removal of lithospheric roots. Thickening of the crust results in additional buoyancy forces whereas thickening of the sub-crustal lithosphere that is colder and heavier than the surrounding asthenosphere (Brunet, 1986; Royden, 1993), causes

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Fig. 7. Map of neotectonic movements (after Milanovsky, 1968) during Late Miocene–Quaternary times, i.e. base of the upper Sarmatian layer in the basin combined with estimates of uplift during this time in the orogenic areas.

downward loading. The latter may be especially important during the initial stages of collision when the orogen is not high; subsequent removal of the lithospheric root could remove this loading resulting in uplift. The importance of these two factors during orogenesis has been recognised for many orogens (e.g. Burg and Ford, 1997 and references therein). In the next section we introduce the simple model of elastic response of the lithosphere to the different types of loading, including the effects of lithospheric and crustal thickening=thinning during the collision. The results of 2D flexural and gravity modelling along several sections within the study area are then presented.

6. Flexural modelling of the foreland stage 6.1. Modelling method and data The insufficiency of loading from topography only in many foreland basins was recognised since the earliest modelling studies. Subsurface loading was introduced to improve the fit of model predictions with the observations. It was usually included in the model as a body of anomalous den-

sity with a simple configuration, but interpreted in terms of structures such as ‘obducted’ crustal blocks (Karner and Watts, 1983), subducted slab (Royden and Karner, 1984), lithospheric roots (Brunet, 1986), and so on. Here, we have used a somewhat different approach (Appendix A). All structural (or geometrical) types of subsurface loading are included directly in the model. In such a way we do not consider a generalised subsurface load but instead model the ‘best fitting’ geometrical structure of the lithosphere. However, the possibility to use subsurface loading in a common way to simulate loading due to the phase transitions or intrusion of heavy material (referred to as ‘intracrustal loading’) has been retained, for the case in which such loads appear to be unavoidable. Four kinds of data can be used to constrain the model. These are: (1) topography data, (2) Bouguer gravity anomalies, (3) basin configuration, or structure on the base of the molasse formation, and (4) Moho depth. The last is unreliable in the study area; several published maps show large difference (more than 5 km) and it has not been used in the present study. Basin subsidence data were extracted from the neotectonic map of Milanovsky (1968) shown in

A.V. Ershov et al. / Tectonophysics 313 (1999) 219–241

233

Fig. 8. Bouguer gravity anomalies for the investigated region (data of the Geological Survey of Russia) with the position of modelled sections.

Fig. 7. Bouguer gravity data (Fig. 8) were taken from the database of the Geological Survey of Russia. Note that these data differ from those used by Ruppel and McNutt (1990), which were obtained by the inverse transformation of an old map of isostatic anomalies (with unknown values of isolines) made by Artemiev and Balavadze (1973). Topography along the modelled sections was extracted from the 50 spaced ETOPO-5 data set of U.S. NGDC.

6.2. Discussion of the results Modelling has been carried out on a number of sections orthogonally traversing the Caucasus Orogen. The results of only five of them are presented here (see Figs. 7 and 8 for location). Initial data and modelling results are presented in Fig. 9 and in Table 1. The effect of topographical loading only (i.e., without subsurface loads) is shown by dotted lines in

Table 1 Best-fit parameters for the five modelled sections (Figs. 7 and 8)

Broken end position Orogen core: outer part Orogen core: inner part EET Crustal thickening: x0 Crustal thickening: ∆x Crustal thickening: ∆z Lith. Thickening: x0 Lith. Thickening: ∆x Lith. Thickening: ∆z

Section 1

Section 2

210 200–285 210–270 25=60

150 130–250 140–240 60 180 150 4 180 150 6

0 220 300 32

Section 3 70

60 130 250 5 130 350 60

Section 4

Section 5

70 70–150

100 110–160

50 110 100 5 110 200 28

40 140 100 2.5 140 100 45

Results are also presented in Fig. 9. All values are in km. The two EET values for section 1 correspond to the two parts of the plate: 210  x  380 (EET D 25 km) and x > 380 (EET D 60 km).

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Fig. 9. Results of the flexural and gravity modelling for the sections shown in Figs. 7 and 8, including the best-fit model (upper panel), observed and modelled topography and basin configuration (middle panel), and observed and modelled Bouguer gravity anomalies (lower panel). Solid lines are observations, dotted lines indicate the effect of topographical loading only, dashed lines are from the best-fitting model. See Table 1 for best-fitting model parameters and the text for further discussion.

Fig. 9, in order to be able to compare its influence and relative importance to the total load including the subsurface loads. Effective elastic thickness (EET) values were chosen independently for the best-fitting topographical loading-only models. In general, we may conclude that there is no good model approx-

imation of observed gravity and basin subsidence, when only topographical loading is included, for any of the modelled sections. Other types of loading=unloading are necessary to achieve a satisfactory fit. In achieving one, we have tried to use as few additional parameters as

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235

Fig. 10. Sketch showing the main subsidence-driving mechanisms (for the foreland stage of basin evolution) in different areas of the Caucasus.

possible. It was possible to obtain a reasonable result by considering only three parameters for the broken plate case: EET (constant along x, with the exception of section 1 for which two segments with different EET were necessary), thickening of the crust, and thickening of the lithosphere. Acceptable models with zero values for intraplate force and bending moment were found. Varying these two parameters could improve the fit, but only in a second-order way. We have preferred simplicity (and clarity) instead of slightly improved accuracy. The position of the end of the broken plate as well as the boundaries of the basins and orogen were chosen on the basis of geological data for each section. These were not subsequently varied during the modelling process. Taking into account the influence of the ‘core’ of the orogen was important for achieving agreement between observed and modelled gravity anomalies near the orogen on sections 1 and 2. The configurations of the best-fit model for each section are shown in the upper parts and the model results with dashed lines in the lower parts of Fig. 9. The results indicate: (1) loading by a lithospheric root is the principal reason for subsidence in the western and easternmost areas (Fig. 9a,d,e); (2) removal of the lithospheric root and crustal thickening is responsible for the uplift of the Central Caucasus and adjacent basins (Fig. 9c); and (3) that the geographically intermediate area of the Eastern Cau-

casus (Fig. 9b) occupies an intermediate position in the modelling between these two extremes. The two extreme modelling cases (cf. Fig. 10) correspond to the two main stages of orogen formation through time: the initial stage (easternmost and westernmost areas, Fig. 9 a,e) and the final stage (central area, Fig. 9c). The Central Caucasus area began to form first, probably in the Eocene (Nesmeyanov, 1992; Lozar and Polino, 1997); with time it spread laterally and the recent orogenic tips underwent their first folding=thrusting event only in pre-Akchagylian time (5.2–3.5 Ma) (Nesmeyanov, 1992; Scherba, 1993). This fact gives us a guideline to interpret the results for section 2 (Fig. 9b). If this section were considered separately from the others, it would be possible to say that the main load producing the foreland basin is the topography. But considered in the context of sections 1 and 3 (easternmost and central sections), another possible scenario emerges: subsidence was induced by the loading of lithospheric root at the initial stages (as now for sections 1, 4, and 5) but, later, continued crustal thickening and the removal of the lithospheric root induced the elevation of the topography. This hypothesis is supported by some geological data. The rate of the uplift of the Eastern Caucasus, as derived from the elevation of river terraces, significantly increased (by an order of magnitude) in the Late Quaternary (Nikitin, 1987). This corresponds

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to the general uplift of the eastern basin area, adjacent to the orogen (Maslyaev, 1990), and subsidence of the eastern part of peripheral bulge (Fig. 5), i.e. the possible ‘unflexing’ of the flexural basin. This concept is in agreement with the observation that folding=thrusting events are contemporaneous with episodes of the basin subsidence, but that the uplift of the orogen appears to be later (e.g. Scherba, 1993). We conclude in general, therefore, that a significant factor in controlling basin subsidence and orogen uplift in the Caucasus is the interplay between loading=buoyancy due to changes in the geometrical structure of lithosphere (i.e., the appearance and removal of lithospheric roots, crustal thickening, etc.). Topography in itself appears to be a secondary effect of this interaction as well as in driving basin subsidence. The EET values inferred for the eastern area (near 60 km) are in general agreement with the gravity modelling results of Ruppel and McNutt (1990; 40– 50 km). We found it impossible to discriminate EET in the central area. The ‘best-fit’ presented in Fig. 9c was obtained with EET D 60 km. On the other hand, it is also easily possible to obtain a good fit in the framework of local isostasy, taking into account post-Maikopian platform uplift (inherited from the dynamically supported pre-foreland subsidence). It follows that it is possible to obtain a good fit for any intermediate value of EET. Our inferred EET in the western area (about 40 km) is different from ‘no flexural rigidity’ reported by Ruppel and McNutt (1990) for the ‘western part of the range’. Note, however, that the section used by Ruppel and McNutt to model the ‘western area’ lies in our central area (which corresponds to the accepted subdivision of the Caucasus and Fore-Caucasus). In this case, for the reasons stated above, there is no contradiction between the two studies. Finally, we point out that the inferred best-fit EET values may be not very precise. This follows from the fact that the position of the peripheral bulge was found to be more or less coincident with the boundary of the Scythian Platform, i.e. within the area of change of mechanical properties. In this case, the effect of localisation of the flexural bulge on the mechanical heterogeneity discussed by Waschbuch and Royden (1992) could take place.

6.3. Implications for foreland basin modelling In the Caucasus case (as in most of the other orogenic areas) topography itself does not explain fully the observed subsidence of the foreland basin. Moreover, in the Caucasus area we observe a clear anticorrelation between topography and basin subsidence. In general, we conclude that topography should not be considered as the main control of foreland subsidence. Topographical loading may certainly be significant, but it is not the primary factor. It is itself a result of thickening of the crust and lithosphere and the redistribution of densities within the crust and mantle during and after collision (e.g. Burg and Ford, 1997), and the consequent buoyancy forces resulting from the interplay between these primary factors. Growth of topography continues until it equilibrates the forces induced by the structural changes within the lithosphere. Topography is only one of the compensation mechanisms and, in this respect, it is similar to the foreland basin subsidence itself. Uplift of the orogen and subsidence of the foreland basin are only two different counterbalances to the same loading processes. In contrast, the conventional model of foreland basin subsidence is based on the assumption that topographical loading is the primary driving mechanism. Modelling results are usually expressed with an estimation of the amount of additional subsurface loading necessary to construct an acceptable model. It means, in our opinion, that the actual primary controls of orogenic evolution are considered only as ‘additional’ factors. We would argue that it is necessary to produce a more sophisticated model that includes the primary factors explicitly. The primary or ‘input’ factors are the crustal structure of orogen as inferred from deep seismic data and the inferred evolution of the lithospheric root. The secondary observations, for comparison with model ‘output’ include basin subsidence, the topography, and gravity anomalies. In other words, it is necessary to clarify the meaning of ‘subsurface loading’ and relate it to real structures and processes during collision. These arguments are not intended as a refutation of the existing model. It allows a first-order fit of model to data. Moreover, it easily incorporates crustal thickening and lithospheric roots (as we have demonstrated in this paper, for instance). But, in our

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opinion, for further advances in the modelling of foreland basins, it is necessary to develop a more realistic model, taking into account the structure of the orogen as it is inferred from deep seismic data. It seems that recent progress in the fields of deep seismic exploration and experimental rock mechanics, as well as an increase in computational power, make it possible.

7. Summary The Late Cenozoic collisional history of the Northern Caucasus basin started with broad subsidence of its central and eastern areas. The deepwater Maikopian basin had water depths of 500–800 m in its central part and up to 1200 m in the south. The Western Fore-Caucasus basin did not undergo significant subsidence at this time. The characteristic wavelength of the subsidence is greater than the lengths of the studied sections (about 200 km) and also greater than the characteristic length of elastic flexure of the basin lithosphere. The cause of such subsidence should lie in large-scale processes, such as mantle convection. Analysis of Late Eocene–Early Oligocene collisional settings allows this subsidence to be associated with the termination of subduction in the southern areas. If so, its cause may be the mantle flow induced by the reequilibration of the subducted slab after the cessation of subduction. This subsidence is therefore ‘dynamically supported’ and it should lead to uplift upon its cessation. However, in the Northern Caucasus basins the implied uplift stage appears to have been masked by the foreland subsidence of the next stage of basin development. Evidence for it can be seen in the more complex subsidence and sedimentation patterns of the eastern basin in comparison with the western one. (In particular, the amplitude of total uplift of the eastern peripheral bulge was more than 1 km, whereas the western bulge was not even elevated above sea level.) The beginning of the next stage of the basin evolution is fixed by lithological changes and a subsidence event in the southern areas in the Middle Miocene (about 16 Ma). This event marks the beginning of the main collisional phase of the Great Caucasus. This led to the uplift of the western and the eastern

237

segments of the Great Caucasus Orogen above sea level in the Late Miocene (11 Ma) and also to significant changes of along-strike basin configuration at this time. The Central Fore-Caucasus area began to uplift and separated the western and the eastern basins. Subsidence in both these basins was asymmetrical with deepening towards the orogen, typical for foreland basins. But the analysis of the alongstrike basin configuration shows that topographical loading cannot be considered as the main control on foreland subsidence. We have carried out flexural and gravity modelling of Middle Miocene–Quaternary subsidence. The results show that a combination of crustal and lithospheric thickening and removal of a lithospheric root can explain the observed 3D subsidence pattern (Fig. 10). The removal of the lithospheric root, in combination with crustal thickening, can further explain uplift of the Central Fore-Caucasus. Loading due to the presence of lithospheric root had induced subsidence at the tips of the orogen and, more generally, subsidence of the basin. Highly elevated topography appeared only at the final stages of evolution, following the removal of the lithospheric root. The responsible mechanism (e.g. conductive heating, convective removal or delamination) demands further investigation. The flexural modelling implies an effective elastic thickness (EET) for the lithosphere of nearly 60 km for the Eastern Fore-Caucasus and 40 km for the Western Fore-Caucasus. It was impossible to constrain EET in the Central Fore-Caucasus; any value in the range 0–60 km gives an equally good fit.

Acknowledgements This work was funded by the Peri-Tethys Programme (grants 95–96=66) and became possible with a post-doctoral position for the first author granted by the French Ministry of Foreign Affairs. EUROPROBE (European Science Foundation) and the International Lithosphere Programme supported communications and discussions. We are grateful to A.F. Morozov and the Russian Geological Survey for granting access to the data and support of our work. The manuscript was improved after the constructive remarks of H. Lyon-Caen, K. Sobornov

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and A. Lankreijer; thanks to them. We thank also J.-P. Burg and R.A. Stephenson who helped us to reshape the final version, and A.S. Alekseev, E.V. Artyushkov, M.A. Goncharov, J. Dercourt, S.S. Kosova, E.E. Milanovsky, B.P. Nazarevich, M.Yu. Nikitin, N.V. Koronovsky, D.I. Panov, and P.A. Ziegler for fruitful discussions.

right and left sides. For the broken plate the calculations were performed for two plates. For the edges adjacent to the failure (at x f ), the assignment of momentum M gives the condition on the second derivative (in a net ‘broken plate’ case it is zero): D Ð w00 .xf / D M and the isostatic constraint Z 1 q Ð dx D 0 0

gives the following condition at the broken edge (Sheffels and McNutt, 1986):

Appendix A. Flexural and gravity model A scheme showing the configuration of the gravity and elastic models and the applied forces is presented in Fig. 11. The elastic model is based on the finite-difference numerical solution of equation: .Dw00 /00 C .Pw0 /0 D q where D.x/ is flexural rigidity, P is intraplate stress, w.x/ is vertical deflection of lithosphere, and q is total loading of different types. The effective elastic thickness (EET; Te ) is expressed through the flexural rigidity as: DD

[D.xf / Ð w00 .xf /]0 D P Ð w0 .xf / The numerical scheme was tested on known analytical solutions. Total loading is defined by the difference between the weight of the material column from the surface to a conditional compensation level in the asthenosphere in the unperturbed area and the weight of such a column in the investigated area. The former is equal to (cf. Fig. 11 for the definition of symbols): g [²ast .z c

∆z cr

∆z lit / C ²lit ∆z lit C ²cr ∆z cr ]

the latter is equal to (excluding intracrustal loading qic ):

E Te3 12.1 ¹ 2 /

(1)

where E is Young modulus and ¹ is the Poisson ratio. The definition of loading is described as follows. The region modelled was extended by several wavelengths to the right and to the left (up to 1500 km) to avoid the influence of boundary conditions; the elastic flexure of the lithosphere was calculated over such a broad region, but the comparison with observations was performed within the initial area only. The zero deflection w.0/ D w.1/ D 0 and zero flexure w00 .0/ D w00 .1/ D 0 were used as boundary conditions for the continuous plate on the

g [²ast .z c

∆z cr

∆z lit

∆z cr th

∆z lit th

w/

C²lit .∆z lit C ∆z lit th / C ²cr .∆z cr C ∆z cr th / C ²sed w

Here, g is the acceleration of gravity, ²ast , ²lit , ²cr , ²sed , ²or are densities of asthenosphere, subcrustal lithosphere, crust, sediments and orogen (material above zero), respectively; w is the flexural deflection of the plate; h is topography; ∆z cr , ∆z lit are the thicknesses of undeformed crust and lithosphere; ∆z cr th , ∆zlit th are the changes of crust and lithosphere thickness during collision; and z c is the reference isostatic compensation level

3 3

3

3

E1 3

3

c

²or h]

0

Fig. 11. Schematic illustration of model configuration and symbolism.

E2

A.V. Ershov et al. / Tectonophysics 313 (1999) 219–241 inside the asthenosphere (with positive down). The difference between these two expressions defines the total loading force: q

qic D g [ ²ast .∆z cr th C ∆z lit th C w/ C ²lit ∆z lit th C²cr ∆z cr th C ²sed w

q

²or h]

This can be rewritten in the form: qic D .²ast ²sed /gw C .²lit ²ast /g∆z lit th ð g∆z cr th

.²ast

²cr /

²or gh

The first and last terms are those defining the well-known buoyancy forces due to flexure and topographical loading. The middle two terms represent loading=buoyancy due to crustal and lithospheric thickening. With time the lithosphere is heated and melted, the lithosphere–asthenosphere boundary flattens and the ∆z lit th value decreases and may even become negative due to simple heating or convective removal of lithospheric root or its delamination. As a result, the balance is violated and this induces uplift of the orogen (Brunet, 1986; Burg and Ford, 1997). We suppose that the distribution of ∆z cr th and ∆z lit th along the x-axis has a bell-like shape defined by: ∆z.x/ D ∆z max cos2 .³.x x0

x0 /=∆x/ ;

∆x=2 < x < x0 C ∆x=2

∆z D 0; x < x0

∆x=2 or x > x0 C ∆x=2

Therefore, only three parameters are necessary to determine thickening: the maximum value (∆z max ), the central point (x0 ) and the range (∆x). The area under the orogen is often composed of rocks denser than sediments; we name this area the ‘orogen core’. It induces the appearance of a local maximum in the Bouguer anomalies over the orogen. We have distinguished an outer and inner area of the orogen core (Fig. 11). The effect of the orogen core was taken into account by replacing the sediment density parameter in the first term of the above equation .²ast ²sed /gw by the density of the orogen (²or ) in the outer area and by the density of the crust (²cr ) in the inner area (Fig. 11). The placement of the orogen core was chosen mainly on the basis of geological data but the position of its inner part was fitted according to the Bouguer gravity data. Its influence on basin subsidence is of second order, but it is the only possibility to obtain a local Bouguer maximum over the orogen if intracrustal loading is not invoked. Thus, several parameters can be varied: the effective elastic thickness (EET), Te .x/, representing the elastic properties of lithosphere; intraplate force P; momentum M; intracrustal subsurface loading qic .x/; and thickening of crust (∆z cr th ) and lithosphere (∆z lit th ). The modelling method as well allows a choice between a continuous and a broken plate. One problem is that the peripheral bulge in the eastern part of the basin is buried at present. The implied subsidence occurred only recently, during the Quaternary. The cause may be long-wavelength subsidence of the whole region, interactions with the Precaspian basin (viz. Quinlan and Beaumont, 1984), or the onset of unflexing of the basin. In any case, it can be taken into account by allowing a constant (along x) shift of the deflection.

239

The gravity anomalies depend on the geometry and density distribution of the lithosphere before and after the thickening=flexure. Bouguer gravity anomalies were calculated by the direct integration of the gravity effect of all anomalous masses. The configuration of the anomalous masses was defined by the subtraction of a standard lithosphere (crust, including pre-foreland sediments 45 km, mantle part 100 km) from the flexed=thickened one (Fig. 11). The densities that were used are as follows: sediments 2.5 g=cm3 ; bulk orogenic rock 2.67 g=cm3 ; crust 2.8 g=cm3 ; subcrustal lithosphere 3.3 g=cm3 ; asthenosphere 3.2 g=cm3 .

References Adamia, S., Akhvlediani, K.T., Kilasonia, V.M., Nairn, E.M., Papava, D., Patton, D., 1992. Geology of the republic of Georgia: a review. Int. Geol. Rev. 34, 447–476. Alekseev, M.N. (Ed.), 1990. Paleogeographic Atlas of the Shelf Region of Eurasia for Mesozoic and Cenozoic. The Robertson Group, Geological Institute, Academy of Sciences, USSR. Artemiev, M.E., Balavadze, B.K., 1973. Isostasy of the Caucasus (in Russian). Geotectonics 6, 20–32. Artyushkov, E.V., 1993. Physical Tectonics (in Russian). Nauka, Moscow, 456 pp. Beaumont, C., 1981. Foreland basins. Geophys. J. R. Astron. Soc. 65, 291–329. Brunet, M.F., 1986. The influence of the Pyrenees evolution on adjacent basins. Tectonophysics 12, 343–354. Brunet, M.F., Rebetsky, Y.L., Nikitina, Y.S., 1997. First steps of modelling in the South Caspian Sea. Annual Meeting of the Peri-Tethys Programme, 10–12 june 1997, Rabat, pp. 25–26. Brunet, M.F., Volozh, Y.A., Antipov, M.P., Lobkovsky, L.I., 1999. The geodynamic evolution of the Precaspian basin (Kazakhstan) along a north–south section. Tectonophysics 313, 59–83, this issue. Burg, J.-P., Ford, M., 1997. Orogeny through time: an overview. In: Burg, J.-P., Ford, M. (Eds.), Orogeny through Time. Geol. Soc. Spec. Publ. 121, 1–17. Chumakov, I.S., 1993. Radiometric scale for Late Cenozoic of Paratethys (in Russian). Priroda 12, 68–75. Chumakov, I.S., Byzova, S.L., Ganzei, S.S., 1992. Radiometric Scale for Late Cenozoic of Eastern Paratethys (in Russian). Nedra, Moscow, 124 pp. Cross, T.A., Pilger, R.H., 1978. Tectonic controls of Late Cretaceous sedimentation, western interior, USA. Nature 274, 653– 657. Dercourt, J., Ricou, L.E., Vrielynck, B. (Eds.), 1993. Atlas Tethys, Paleoenvironmental Maps. Gauthier-Villars, Paris, 307 pp., 14 maps, 1 pl. Dickinson, W.R., 1974. Subduction and oil migration. Geology 2, 421–424. Ershov, A.V., Brunet, M.-F., Nikishin, A.M., Bolotov, S.N., Korotaev, M.V., Kosova, S.S., 1998. Evolution of the eastern Fore-Caucasus basin during the Cenozoic collision: burial history and flexural modelling. In: Crasquin-Soleau, S., Barrier,

240

A.V. Ershov et al. / Tectonophysics 313 (1999) 219–241

E. (Eds.), Peri-Tethys Memoir 4: Epicratonic Basins of PeriTethyan Platforms. Me´m. Mus. Natl. Hist. Nat. Paris 179, 111–130. Gamkrelidze, I.P., 1986. Geodynamic evolution of the Caucasus and adjacent areas in Alpine time. Tectonophysics 127, 261– 267. 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. Holt, W.E., Stern, T.A., 1994. Subduction, platform subsidence, and foreland thrust loading: the Late Tertiary development of Taranaki Basin, New Zealand. Tectonics 13, 1068–1092. Hsui, A.T., Tang, X.-M., Tokso¨z, M.N., 1990. On the dip angle of subducting plates. Tectonophysics 179, 163–175. Jordan, T.E., 1981. Thrust loads and foreland basin evolution, western United States. Am. Assoc. Pet. Geol. Bull. 65, 2506– 2520. Karner, G.D., Watts, A.B., 1983. Gravity anomalies and flexure of the lithosphere at mountain ranges. J. Geophys. Res. 88, 449–477. Koronovsky, N.V., Kozhevnikov, A.V., Panov D.I., Vishnevsky, L.E., Guschin, A.I., Nikitin, M.Yu., Voronina, A.A., Stafeev, A.N., Panina, L.V., Kopaevich, L.F., Stor, M.A., Belyakov, G.A., 1987. History of geological development and formation of structure of central part of Terek–Caspian foreland (in Russian). In: Milanovsky, E.E., Koronovsky, N.V. (Eds.), Geology and Mineral Resources of Great Caucasus. Nauka, Moscow, 147–174. Kosova, S.S., 1994. The forming of the Cenozoic sequences within East Precaucasus (in Russian). In: Leonov, Yu.G., Antipov, M.P., Morozov, A.F., Solodilov, L.N. (Eds.), Tectonics and Magmatism of East-European Platform, Proceedings of International EUROPROBE Workshop, Opalikha, May 12–15, 1993, Moscow, pp. 94–104. Krasnopevtseva, G.V., 1984. Deep Structure of Caucasus Seismoactive Region (in Russian). Nauka, Moscow, 107 pp. Kunin, N.Ya., Kosova, S.S., Blokhina, G.Yu., 1990. The subdivision of Cenozoic deposits of Eastern Fore-Caucasus (in Russian). Izv. Akad. Nauk SSSR, Ser. Geol. 71–81. Lozar, F., Polino, R., 1997. Early Cenozoic uprising of the Great Caucasus revealed by the reworked calcareous nannofossils. Terra Nova 9 (1), 141. Lyon-Caen, H., Molnar, P., 1983. Constraints on the structure of the Himalaya from an analysis of gravity anomalies and a flexural model of the lithosphere. J. Geophys. Res. 88, 8171– 8192. Lyon-Caen, H., Molnar, P., 1989. Constraints on the deep structure and dynamic processes beneath the Alps and adjacent regions from an analysis of gravity anomalies. Geophys J. Int. 99, 19–32. Maslyaev, G.A., 1990. The specifics of the Fore-Caucasus structure formation at the Cenozoic (in Russian). Geotectonica 4, 52–60. Mikhailov, V.O., 1993. Crustal control on the Terek–Caspian

trough evolution: constraints based on a new paleotectonic analysis method. Tectonophysics 238, 21–32. Mikhailov, V.O., Polino, R., Timoshkina, E.P., Panina, L.V., 1997. North Caucasus foredeep: subsidence history and model of formation. Terra Nova Abstracts, European Union of Geosciences 9, Strasbourg, p. 138. Milanovsky, E.E., 1968. Neotectonics of the Caucasus (in Russian). Nedra, Moscow, 482 pp. Mitrovica, J.X., Beaumont, C., Jarvis, G.T., 1989. Tilting of continental interiors by the dynamical effects of subduction. Tectonics 8, 1079–1094. Nesmeyanov, S.A., 1992. Neostructural Subdivision of the North-West Caucasus (in Russian). Nedra, Moscow, 255 pp. Nevesskaya, L.A., Voronina, A.A., Goncharova, I.A., Ilina, L.B., Paramonova, N.P., Popov, S.V., Chepalyga, A.L., Babak, E.B., 1984. The history of Paratethys (in Russian). Proc. 28th Int. Geol. Congr., Paleooceanology 3, Nauka, Moscow, pp. 91–101. Nikishin, A.M., Ziegler, P.A., Stephenson, R.A., Cloetingh, S.A.P.L., 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 324, (IIa), 49–57. Nikishin, A.M., Cloetingh, S., Bolotov, S.N., Baraboshkin, E.Yu., Kopaevich, L.F., Nazarevich, B.P., Panov, D.I., Brunet, M.-F., Il’ina, V.V., Kosova, S.S., Stephenson, R.A., 1998a. Scythian Platform: chronostratigraphy and polyphase stages of tectonic history. In: Crasquin-Soleau, S., Barrier, E. (Eds.), Peri-Tethys Memoir 3. Stratigraphy and Evolution of Peri-Tethyan Platforms. Me´m. Mus. Natl. Hist. Nat. Paris 177, 151–162. Nikishin, A.M., Cloetingh, S., Brunet, M.-F., Bolotov, S.N., Ershov, A.V., Stephenson, R.A., 1998b. Scythian Platform, Caucasus and Black Sea region: Mesozoic–Cenozoic tectonic history and dynamics. In: Crasquin-Soleau, S., Barrier, E. (Eds.), Peri-Tethys Memoir 3. Stratigraphy and Evolution of Peri-Tethyan Platforms. Me´m. Mus. Natl. Hist. Nat. Paris 177, 163–176. Nikitin, M.Y., 1987. Neotectonics of the Eastern Caucasus (in Russian). Bull. MOIP 62 (3), 21–36. Odin, G.S., 1994. Geological Time Scale (1994). C. R. Acad. Sci. Paris 318 (II), 59–71. Philip, H., Cisternas, A., Gvishiani, A., Gorshkov, A., 1989. The Caucasus: an actual example of the initial stages of continental collision. Tectonophysics 161, 1–21. Popov, S.V., Stolyarov, A.S., 1996. Paleogeography and anoxic environments of the Oligocene–Early Miocene Eastern Paratethys. Isr. J. Earth Sci. 45, 161–167. Popov, S.V., Akhmetiev, M.A., Zaporozhets, N.I., Voronina, A.A., Stolyarov, A.S., 1993. History of Eastern Paratethys

A.V. Ershov et al. / Tectonophysics 313 (1999) 219–241 in Late Eocene–Early Miocene (in Russian). Stratigr. Geol. Korrelyatsiya 1, 10–39. Price, R.A., 1973. Large-scale gravitational flow of supracrustal rocks, southern Canadian Rockies. In: De Jong, K.A., Sholten R. (Eds.), Gravity and Tectonics. Interscience, New York, pp. 491–502. Quinlan, G.M., Beaumont, C., 1984. Appalachian thrusting, lithospheric flexure, and the paleozoic stratigraphy of the eastern interior of North America. Can. J. Earth Sci. 21, 973–996. Reilinger, R.E., McClusky, S.C., Souter, B.J., Hamburger, M.W., Prilepin, M.T., Mishin, A., Guseva, T., Balassanian, S., 1997. Preliminary estimates of plate convergence in the Caucasus collision zone from global positioning system measurements. Geophys. Res. Lett. 24 (14), 1815–1818. Robinson, A.G., Rudat, J.H., Banks, C.J., Wiles, R.L.F., 1996. Petroleum geology of the Black Sea. Mar. Pet. Geol. 13, 195–223. Royden, L.H., 1993. The tectonic expression slab pull at continental convergent boundaries. Tectonophysics 12, 303–325. Royden, L.H., Karner, G.D., 1984. Flexure of the lithosphere beneath Apennine and Carpathian foredeep basins: evidence for an insufficient topographic load. Am. Assoc. Pet. Geol. Bull. 68, 704–712. Ruppel, C., McNutt, M., 1990. Regional compensation of the Great Caucasus mountains based on an analysis of Bouguer gravity data. Earth Planet. Sci. Lett. 98, 360–378. Scherba, L.G., 1993. Stages and Phases of Cenozoic Development of Alpine Fold Belt (in Russian). Nauka, Moscow, 231 pp.

241

Sengor, A.M.C., Kidd, W.S.F., 1979. Post collisional tectonics of the Turkish–Iranian plateau and a comparison with Tibet. Tectonophysics 55, 361–376. Sengor, A.M.C., Yilmaz, Y., 1981. Tethyan evolution of Turkey: a plate tectonic approach. Tectonophysics 75, 181–241. Sheffels, B., McNutt, M., 1986. Role of subsurface loads and regional compensation in the isostatic balance of the Ranges, California: evidence for intracontinental subduction. J. Geophys. Res. 91, 6419–6431. Steckler, M.S., Watts, A.B., 1978. Subsidence of the atlantic-type continental margin off New York. Earth Planet. Sci. Lett. 41, 1–13. Turcotte, D., Schubert, G., 1982. Geodynamics: Applications of Continuum Physics to Geological Problems. Wiley, New York, 124 pp. Volvovsky, I.S., Moskalenko, V.N., Neprochnov, Y.P., Shlezinger, A.E., 1989. Thickness and structure of the consolidated crust (in Russian). In: Belousov, V.V., Volvovsky, B.S. (Eds.), Structure and Evolution of the Earth’s Crust and Upper Mantle of the Black Sea. Nauka, Moscow, pp. 136–138. Waschbuch, P.J., Royden, L.H., 1992. Episodicity on foredeep basins. Geology 20, 915–918. Zonenshain, L.P., Le Pichon, X., 1986. Deep basins of the Black Sea and Caspian Sea as remnants of Mesozoic back-arc basins. Tectonophysics 123, 181–211. Zonenshain, L.P., Kuzmin, M.I., Natapov, L.M., 1990. Geology of the USSR, a Plate Tectonic Synthesis. Am. Geophys. Union, Geodyn. Ser. 21, 242 pp.