The Southern Urals. Decoupled evolution of the thrust belt and its foreland: a consequence of metamorphism and lithospheric weakening

Tectonophysics 320 (2000) 271–310 www.elsevier.com/locate/tecto

The Southern Urals. Decoupled evolution of the thrust belt and its foreland: a consequence of metamorphism and lithospheric weakening Eugene V. Artyushkov a, *, Michael A. Baer a, Peter A. Chekhovich b, Nils-Axel Mo¨rner c a Institute of Physics of the Earth, Russian Academy of Sciences, B. Gruzinskaya 10, 123810, Moscow, Russia b Institute of the Lithosphere of Marginal Seas, Russian Academy of Sciences, 109180, Moscow, Russia c Stockholm University, Institute of Paleogeophysics and Geodynamics, Kra¨feriket 24, S-10691, Stockholm, Sweden

Abstract An analysis is presented of the mechanisms of tectonic evolution of the southern part of the Urals between 48N and 60N in the Carboniferous–Triassic. A low tectonic activity was typical of the area in the Early Carboniferous — after closure of the Uralian ocean in the Late Devonian. A nappe, ≥10–15 km thick, overrode a shallow-water shelf on the margin of the East European platform in the early Late Carboniferous. It is commonly supposed that strong shortening and thickening of continental crust result in mountain building. However, no high mountains were formed, and the nappe surface reached the altitude of only ≤0.5 km. No high topography was formed after another collisional events at the end of the Late Carboniferous, in the second half of the Early Permian, and at the start of the Middle Triassic. A low magnitude of the crustal uplift in the regions of collision indicates a synchronous density increase from rapid metamorphism in mafic rocks in the lower crust. This required infiltration of volatiles from the asthenosphere as a catalyst. A layer of dense mafic rocks, ~20 km thick, still exists at the base of the Uralian crust. It maintains the crust, up to ~60 km thick, at a mean altitude ~0.5 km. The mountains, ~1.5 km high, were formed in the Late Permian and Early Triassic when there was no collision. Their moderate height precluded asthenospheric upwelling to the base of the crust, which at that time was ~65–70 km thick. The mountains could be formed due to delamination of the lower part of mantle root with blocks of dense eclogite and/or retrogression in a presence of fluids of eclogites in the lower crust into less dense facies. The formation of foreland basins is commonly attributed to deflection of the elastic lithosphere under surface and subsurface loads in thrust belts. Most of tectonic subsidence on the Uralian foreland occurred in a form of short impulses, a few million years long each. They took place at the beginning and at the end of the Late Carboniferous, and in the Late Permian. Rapid crustal subsidence occurred when there was no collision in the Urals. Furthermore, the basin deepened away from thrust belt. These features preclude deflection of the elastic lithosphere as a subsidence mechanism. To ensure the subsidence, a rapid density increase was necessary. It took place due to metamorphism in the lower crust under infiltration of volatiles. The absence of flexural reaction on the Uralian foreland on collision in thrust belt together with narrow-wavelength basement deformations under the nappe indicate a high degree of weakening of the lithosphere. Such deformations took also place on the Uralian foreland at the epochs of rapid subsidences when there was no collision in thrust belt. * Corresponding author. Fax: +7-095-255-60-40. E-mail address: [email protected] ( E.V. Artyushkov) 0040-1951/00/$ - see front matter © 2000 Elsevier Science B.V. All rights reserved. PII: S0 0 4 0- 1 9 51 ( 0 0 ) 0 00 4 4 -5

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Nomenclature 1 c volume per cent of gabbroic intrusions d thickness of lithosphere h thickness of crust c h0 initial thickness of crust c h1 thickness of crust after collision or stretching c h thickness of eroded rocks er h thickness of delaminated eclogites de h thickness of eclogites e h total thickness of gabbroic intrusions gi h thickness of garnet granulites gg h thickness of mantle lithosphere ml h thickness of nappe n h thickness of sediments overridden by nappe or h thickness of pyroxene granulites pg h depth of sedimentary basin s hth depth of sedimentary basin formed by thermal relaxation of lithosphere s h depth of water w L minimum width of deflection of elastic lithospheric 0 P pressure T temperature T temperature of asthenosphere (1300°C ) a T effective elastic thickness of lithosphere e T temperature at Moho M V P-wave velocity P Vgb P-wave velocity in gabbro P Vpr P-wave velocity in peridotites P a thermal expansivity (3×10−5 K−1) b intensity of stretching Dh increase in thickness of crust due to collision c Dh thinning of lower crust by stretching lc Dh decrease in thickness of mantle lithosphere due to asthenospheric upwelling ml Dh thickness of sediments pushed out by nappe s Df uplift of shortened crust Df uplift of crust due to delamination of eclogites de Df uplift of crust due to gabbroic intrusions gi Dfe uplift of crust due to retrogression of eclogites to garnet granulites rg Dfgg uplift of crust due to retrogression of garnet granulites to pyroxene granulites rg Df uplift of crust due to asthenospheric upwelling uw e intensity of compaction of sediments overridden by nappe f altitude of crustal surface f0 minimum altitude of nappe in a case of no density changes in lithosphere n j tectonic subsidence in shortened region due to density increase in lithosphere r density of asthenosphere (3220 kg m−3) a r density of crust (2830 kg m−3) c r density of eclogite e r density of eroded rocks er r density of gabbro (2930 kg m−3) gb r density of garnet granulites gg r density of lower crust lc r density of mantle (3350 kg m−3) m r density of nappe n r density of pyroxene granulites pg r density of sediments s r density of water (1030 kg m−3) w 1 In the nomenclature, symbols without assigned values have been assumed to be variable.

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Weakening of the lithosphere can be explained by infiltration of volatiles into this layer from the asthenosphere and rapid metamorphism in the mafic lower crust. Lithospheric weakening allowed the formation of the Uralian thrust belt under convergent motions of the plates which were separated by weak areas. © 2000 Elsevier Science B.V. All rights reserved. Keywords: collision; crustal subsidence; eclogitization; lithospheric weakening; mountain building; Urals

1. Introduction The formation of foreland basins is commonly considered as a result of plate collision and flexing of the elastic lithosphere towards convergent plate boundaries without strong density changes in the lithosphere (Quinlan and Beaumont, 1984; Malinverno and Ryan, 1986; Royden, 1993; Stewart and Watts, 1997). Mountain building in thrust belts is explained by a synchronous isostatic response to thickening of the crust from plate collision (Molnar and Tapponier, 1975; Miyashiro et al., 1982; Zonenshain et al., 1990) with its possible lowering by slab pull (Royden, 1993). In this scheme, vertical crustal movements in fold belts result from horizontal plate motions. It is also commonly believed that the formation of thrust belts is associated with no significant changes in the density of the lithosphere and the thickness of its elastic part T , which can be e strongly reduced only due to steep bending or strong heating of this layer in some places ( Kusznir et al., 1991; Burov and Diament, 1995). The crustal subsidence and uplift, however, widely occurred in plate interiors without any significant lithospheric stretching and far from convergent boundaries. They have taken place, for example, during the subsidence in the West Siberian, Peri-Caspian, Volga–Urals, Timan– Pechora and Vilyuy basins (Artyushkov and Baer, 1986a; Artyushkov, 1993). In the Neogene, many mountain ranges and plateaus were formed by a strong crustal uplift without significant compressive deformations in East Siberia, north-eastern Asia and Africa (Nikolaew and Neumark, 1977; Partridge and Maud, 1987; Makarov, 1990; Ollier, 1991; Summerfield, 1991). The crustal uplift and subsidence in plate interiors require density changes in the crust and/or mantle. This poses the

question as to whether or not vertical crustal movements of this type also occur in fold belts. It has been shown that the Neogene foredeep of the East Carpathians was formed without significant lithospheric stretching and at the epochs when no collision took place in the adjacent thrust belt (Artyushkov et al., 1996). At times of strong collision, the crustal surface in the thrust belt remained at a low altitude. The Carpathian mountains began to grow 8 m.y. after the end of the collision. These vertical crustal movements required density changes in the lithosphere. No flexural reaction occurred on the Carpathian foreland at the epochs of collision. This indicates a drastic weakening of the lithosphere that ensured a strong crustal subsidence under the nappe without a synchronous subsidence on the adjacent foreland. In this paper, we analyse the tectonic development of the southern part of the Urals after closure of the Uralian ocean — since the start of the Carboniferous. This area, 1300 km long, includes the Southern and Middle Urals (Fig. 1). For the sake of simplicity, it will be called ‘the Southern Urals’. In the preceding publications, attention has been focused on a description and timing of collisional events in the area (e.g. Ruzhentsev, 1976; Zonenshain et al., 1984; Ivanov et al., 1986; Brown et al., 1997; Puchkov, 1997, 1999). Here, we consider another problem: what vertical crustal movements took place near to collisional boundaries in the Southern Urals, and could these movements result from plate motions, or were they caused by deep seated processes? The main questions to be answered are: 1. What were the modes of crustal subsidence on the Uralian foreland and their role in the basin formation? 2. Were the major collisional events in thrust belt associated with strong synchronous crustal subsidence on its foreland?

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5. Did a density increase in the lithosphere play an important role in the subsidence? 6. Was collision associated with a synchronous formation of mountain ranges? 7. Did considerable density changes occur in the lithosphere at times of collision? 8. Was mountain building associated with a strong synchronous shortening of the crust? 9. What was the role of density changes in the lithosphere in mountain building? We first consider the geological data necessary to obtain answers to these questions and methods of analysis of the data. Then, we formulate the answers to questions 1–9. On this basis, we determine possible mechanisms of the crustal subsidence on the Uralian foreland and of mountain building in thrust belt, and consider lithospheric rheology in these regions.

2. Tectonic environment and crustal structure at the end of Early Carboniferous

Fig. 1. Structure of the Southern Urals (compiled using the data by Peive and Yanshin, 1979; Puchkov, 1993; Yazeva and Bochkarev, 1993). 1=the Uralian foredeep with its main depressions in the Southern Urals: A — Aktyubinsk, B — Bel’sk, YS — Yuryusan’–Solikamsk; 2=Zilair–Lemva zone: strongly shortened deposits of the European continental slope of the Uralian ocean and the adjacent eastern margin of the East European platform; 3=outcrops of the Precambrian crystalline basement; 4–6=ensimatic island arcs and ophiolitic blocks of the Uralian ( UO-4) and Kazakhstan ( KO-5) oceanic basins that are separated by the East Uralian microcontinent (EUM-6); 7=eastern and south-western boundaries of the area where the Uralian thrust belt is exposed to the surface; 8= boundaries of the main tectonic units exposed to the surface (a) and overlain by the Mesozoic–Cenozoic cover (b); 9=Main Ophiolitic Suture of the Urals.

3. Did the lithosphere in thrust belt and on its foreland preserve a high effective elastic thickness? 4. Was lithospheric stretching responsible for a considerable part of the subsidence?

At that time, the Southern Urals included the continental margin of the East European continent (Baltica), the East Uralian microcontinent ( EUM ), and the undeformed eastern part of the Magnitogorsk arc (MA) between them (Fig. 2a). The Kazakhstan oceanic basin probably still existed to the east of EUM (Didenko et al., 1994; Puchkov, 1999; Puchkov et al., 1999). Its closure was associated with subduction under EUM that resulted in island-arc volcanism in this region. Subduction could synchronously occur under the Kazakhstan continent. This continent collided with EUM most likely at the beginning of the Late Carboniferous — in the middle of the Serpukhovian age. In the Early Carboniferous, the Sakmara– Magnitogorsk nappe (SMN ), up to ~15 km thick, was lying on the European continental slope ( Fig. 2a). The nappe, the adjacent eastern margin of the East European platform and the Magnitogorsk arc were overlain by several-kilometre-thick deposits of the Late Devonian–Early Carboniferous Zilair basin. Since the middle of the Early Carboniferous and until the end of the Serpukhovian, during a period of time, ~30 m.y.

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Fig. 2. Evolution of the Southern Urals in the late Early Carboniferous and early Late Carboniferous (profiles of Figs. 2, 8 and 13 are compiled using the data by Khvorova, 1961; Geology of the USSR, v. XIII, 1964; Bochkarev, 1973; Kamaletdinov, 1974; Chuvashov, 1975; Ruzhentsev, 1976; Khain, 1977; Samygin, 1980; Chuvashov et al., 1984; Chuvashov and Misens, 1980; Melamud, 1981; Seliverstov and Denisov, 1982; Sigov and Romashova, 1984; Ivanov et al., 1986; Chuvashov and Puchkov, 1990; Kazantsev et al., 1992; Puchkov, 1993, 1996, 1997, 1999; Didenko et al., 1994; Puchkov and Svetlakova, 1993; Yazeva and Bochkarev, 1993; Misens, 1995a,b; Popov and Rapoport, 1996; Seravkin, 1997; Puchkov et al., 1999). (a) Structure of the Southern Urals in the late Early Carboniferous ( late Visean). (b) Closure of the Kazakhstan oceanic basin in the middle of the Serpukhovian, rapid formation of the Western and Eastern flysch basins at the beginning of the Late Carboniferous near to the transition from the Serpukhovian to the Bashkirian. (c) Superposition of the Sakmara–Magnitogorsk nappe onto the eastern margin of the East European continent in the late Bashkirian–early Moscovian–event of collision C 1. SMN — Sakmara–Magnitogorsk nappe; MA — Magnitogorsk arc. In Figs. 2, 8 and 13, the depth of water, sediment and nappe thicknesses, and the altitude of topography are strongly exaggerated to be visible in this scale.

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long, all the Southern Urals were situated near to sea level and characterized by a low tectonic activity. Deposition of shallow-water carbonates took place in this area with some sands and coals in the east (Chuvashov et al., 1984; Chuvashov and Puchkov, 1990). This indicates a high crustal thickness all over the Urals, which is typical of continental areas.

3. Method of analysis Two basic modes of crustal subsidence in sedimentary basins have been identified in our preceding studies (e.g. Artyushkov and Baer, 1986a; Artyushkov et al., 1991). This is a slow sediment loaded subsidence at a rate of 10–100 m m.y.−1 with a long duration of ≥200–300 m.y. and rapid subsidence at a rate of ~1 km m.y.−1, which, in many cases, resulted in the formation of deepwater basins during ~1 m.y. Using the stratigraphic records for the Uralian foreland basins, we will determine the epochs of occurrence of these two types of subsidence and estimate their input into the formation of the basins. This will give an answer to question 1 in Section 1. To answer question 2, we will compare the epochs of strong and rapid crustal subsidence in the foreland basins with the major events of collision in thrust belt. The effective elastic thickness of the lithosphere T is related to the characteristic width of lithe ospheric deflection L as ( Turcotte and Schubert, 0 1982; Artyushkov et al., 1996): (T ) ~0.05[(L ) ]4/3. (1) e km 0 km According to this equation, the crustal subsidence in a foreland basin, ≥100 km wide, at a time of collision indicates deflection of the lithosphere with a high T under surface and subsurface loads in e thrust belt. An additional requirement is basin deepening towards the thrust front. The absence of strong subsidence in the foreland region at a time of collision, will show a high degree of weakening of the lithosphere and decoupling of the thrust belt from its foreland. We will also pay attention to steep bending of the lithosphere in regions, several tens of kilometres wide, in the

foreland basins and under a superimposed nappe as another indication of lithospheric weakening. All these data will contribute to an answer to question 3. Extensive lithospheric stretching is associated with normal faulting in the basement and the sedimentary cover. A continuity of sedimentary beds in the present foredeep will indicate the absence of significant stretching. The method for revealing a presence or absence of stretching in a shortened sedimentary cover of past basins is described in Artyushkov and Baer (1983). In the case of rapid subsidence with a formation of deepwater basin on a shallow-water shelf, a conformable position (parallelism) of shallow water and deep-water strata indicates the absence of block rotation and lithospheric stretching. Using these approaches, we will find an answer to question 4. Those subsidences that occurred without lithospheric stretching and at times when there was no collision required a density increase in the lithosphere as a mechanism (question 5). An analysis of the character of correlative deposits in sedimentary basins is a classic method for estimating the altitude of the adjacent land (Penck, 1924; Meshcheryakov, 1965; King, 1967; Khain, 1977; Timofeev, 1979; Makarov, 1980; see also Artyushkov et al., 1996 for more detail ). Denudation of high mountains produces kilometres of coarse molasse with a large volume of conglomerates. Erosion of a low and smooth topography, ≤0.3–0.5 km high, results in deposition of sands, clays and turbidites in the adjacent basins. In this way, we will determine whether or not collisional events in the Southern Urals were associated with a synchronous formation of high mountains (question 6). After estimating the altitude of the shortened regions, we can consider possible density changes in the lithosphere at the epochs of collision (question 7) in the following way. Suppose that a nappe with the density, r , and thickness, h , was supern n imposed onto a region with the initial depth of water, h . Assume that, on its way, the nappe w pushed out a layer of sediments with the density, r , and thickness, Dh , and overrode sediments s s with thickness, h . The minimum mean altitude or of the nappe surface (f0) will be reached in a state n

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of local isostasy. If nappe superposition resulted in no density changes in the lithosphere below the sedimentary cover, f0 =[(r −r )/r ]h −[(r −r )/r ]Dh n m n m n m s m s −[(r −r )/r ]h −eh . (2) m w m w or Here, r =3350 kg m−3 and r =1030 kg m−3 are m w the densities of the mantle and water, respectively, and e is the intensity of compaction of the overridden deposits. If altitude (2) were considerably larger than that estimated from the character of deposition in the adjacent basin, a density increase would be necessary to lower the lithospheric surface. Continental collision can also result in shortening of the lithosphere, which is uniform over the depth. Designate by r the density of the crust and c by Dh the increase in its thickness after collision. c Suppose that before collision the crustal surface was at or above sea level. Then, in a state of isostasy, the surface of shortened crust would be raised by: Df=(r −r )/r Dh . (3) m c m c If the height of the shortened region were considerably smaller than Eq. (3), a density increase would be necessary to maintain the crustal surface at a low altitude. To answer question 8, we will compare the epochs of mountain building in the Urals with those of collision. Mountain building synchronous to collision can indicate thickening of the crust with the isostatic rebound as a mechanism. The formation of high mountains at the epochs of no strong collision requires a density decrease in the lithosphere (question 9).

4. Crustal subsidence and collision in the early Late Carboniferous 4.1. Rapid formation of flysch basins between the Serpukhovian and Bashkirian Two deep-water basins were formed in the Southern Urals during ~1 m.y. by rapid crustal subsidence between the Serpukhovian and

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Bashkirian ages of the Late Carboniferous ( Figs. 2b–4a). The western basin, ~30–50 km wide with a water depth of 0.5–1 km, was formed on the western margin of the Zilair basin and the adjacent eastern margin of the East European platform ( Khvorova, 1961; Chuvashov, 1975; Melamud, 1981; Misens, 1995b). In the deepest outer (western) part of this basin, Serpukhovian shallow-water limestones (C s legend 3 in column 2 IV in Fig. 5) were abruptly overlain by Bashkirian turbidites (C b legend 7). 2 At the same time another basin, ~400 km wide, was formed in the east, on the Magnitogorsk arc and on the EUM (middle part of Fig. 2b) (Chuvashov et al., 1984). Serpukhovian shallowwater limestones were abruptly overlain by Bashkirian highly organic pelagic limestones. During the following ~8–10 m.y. in the Bashkirian and Moscovian ages, the deepest, western part of the Eastern flysch basin was filled with ~2 km of pelagic limestones and turbidites. Taking into account the isostatic subsidence under their load, which increases the basin depth by about 2.5 times, the initial water depth was ~800 m. The deposits of the Western flysch basin are now incorporated into the frontal part of the Uralian thrust belt. The strata of Bashkirian deepwater deposits are conformable to Serpukhovian shallow-water carbonates at the transition, which is only 1–5 m thick (Fig. 6). This indicates deposition on an almost flat basin floor and precludes lithospheric stretching and block rotation at the epoch of rapid subsidence. 4.2. Superposition of the Sakmara–Magnitogorsk nappe onto the European shallow-water shelf in the Bashkirian–Early Moscovian At the time of rapid subsidence between the Serpukhovian and Bashkirian, the Eastern and Western flysch basins were separated by a shallowwater shelf, ~200 km wide (Fig. 2b). About 2 m.y. later, conglomerates and olistoliths appeared on the western margin of the Eastern flysch basin (Chuvashov et al., 1984). They were derived from the island chain, ~50 km wide, which began to grow in the rear part of the Sakmara–

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Fig. 3. Main events of tectonic development of the foreland basins and western part of thrust belt of the Southern Urals in the Carboniferous–Triassic. Triangles above the horizontal axis indicate events of collision with their numbers corresponding to collisional events in western Urals since the Late Carboniferous. The width of the triangles equals the duration of collision. Their height

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Fig. 4. Schematic profiles of western and central parts of the Southern Urals. (a) In the early Late Carboniferous after the formation of the Western and Eastern flysch basins and superposition of the Sakmara–Magnitogorsk nappe towards the west for 70–80 km. (b) In the late Early Permian after collision of the East European continent with the East Uralian microcontinent at the end of the Carboniferous; shortening of the Bashkir shelf, filling with deposits of the Western flysch basin and the inner part of the late Late Carboniferous Uralian foredeep.

Magnitogorsk nappe above the lower part of the European continental slope. The crustal uplift was associated with the onset of collision in the middle of the Bashkirian (Ruzhentsev, 1976; Chuvashov

et al., 1984). At that time, the Magnitogorsk arc began to move further westwards, overriding the eastern buried slope of the East European continent and pushing the Sakmara–Magnitogorsk

approximately equals an increase in the surface load DP in the shortened region (the method of calculating DP is described in Artyushkov et al., 1996). Changes in the altitude of the crustal surface in regions of collision are schematically shown by a grey line near the horizontal axis. The subsidence curves below the horizontal axis are labelled according to the place of occurrence. They are plotted using the corresponding stratigraphic columns of Fig. 5, with the curve numbers corresponding to the numbers of columns in Fig. 5. The method of compilation of the curves is described in Section 10.1. Two rapid subsidences most likely occurred in the western part of the Western flysch basin in the Late Carboniferous (curve IV ); however, only their cumulate magnitude is known. The second subsidence is tentatively shown in this curve by dashed line. Time scale after Odin (1994).

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Fig. 5. Stratigraphic columns of the past and present foreland basins of the Southern Urals (compiled using the data by Khvorova, 1961; Geology of the USSR, v. XIII, 1964; Chuvashov, 1975; Unified Stratigraphic and Correlational Schemes of the Urals, 1980; Melamud, 1981; Semenovich et al., 1982; Zhivkovich and Chekhovich, 1985; Skrypiy and Yunusov, 1989; Misens, 1995b). Events of rapid water loaded (uncompensated ) crustal subsidence and the onset of rapid sediment loaded (compensated) subsidence are shown by arrows pointing downwards ( legends 13 and 14, respectively). The onsets of collisional events in the adjacent western part of thrust belt (C 1, C 3, C 4 and C 6) are indicated by horizontal arrows ( legend 15). A — Archean, PR — Proterozoic, R — Riphean, V — Vendian, O — Ordovician, S — Silurian; D — Devonian, D fm — Famennian; C — Carboniferous, C t — Tournaisian, C v — 3 1 1 Visean, C s — Serpukhovian, C b -Bashkirian; C m — Moscovian, C k — Kasimovian, C g — Gzhelian; P — Permian, P a — 2 2 2 2 2 1 Asselian, P s — Sakmarian, P ar — Artinskian, P k — Kungurian, P u — Ufimian, P kz — Kazanian, P t — Tatarian; T — Triassic, 1 1 1 2 2 2 J — Jurassic. Note different scales for the Paleozoic and Mesozoic deposits ( left) and the Proterozoic deposits (right).

nappe in the same direction (collisional event C 1 in Figs. 3 and 5). At the same time, deposition of conglomerates derived from the frontal part of the nappe began on the eastern slope of the Western flysch basin ( legend 1 C b in column III in Fig. 5) 2 2 ( Khvorova, 1961). During the following period of ~7 m.y., until

the middle of the Moscovian age, the nappe overrode the upper part of the European continental slope, ~40 km wide, and the adjacent shelf in the Bel’skaya zone, 30–40 km wide, the total magnitude of superposition being ~70–80 km (Figs. 2c and 4a). After the Serpukhovian, no more deposits were formed in the above two regions. The nappe

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Fig. 6. Profile across the shortened deposits of the Western flysch basin that are superimposed onto the inner part of the Uralian foredeep (modified after Kamaletdinov, 1974). The transition from the Serpukhovian shallow-water limestones to Bashkirian deepwater turbidites is indicated by arrows. Parallelism of the strata at this transition indicates the absence of significant lithospheric stretching during the subsidence.

did not reach the eastern margin of the Western flysch basin (Fig. 2c). An undeformed band of the Bashkir shelf still remained between the basin and the nappe front and was ~20–30 km wide in the south (Fig. 4a), widening to ~100–150 km in the north. Shallow-water deposition proceeded with some interruptions in this region during the following ~50 m.y. — until the end of the Early Permian (carbonates 3, sands 2, conglomerates 1 and evaporites 10 of C b-P k in columns I, II of 2 1 Fig. 5) ( Khvorova, 1961; Chuvashov, 1975; Unified Stratigraphic and Correlational Schemes of the Urals, 1980).

4.3. Narrow-wavelength deformations of the lithosphere It follows from the thickness of the deposits that the Eastern flysch basin had a steep western slope and a smooth eastern slope (Figs. 2c and 4a) (Chuvashov et al., 1984). The basin axis with up to 1.5–2 km of Bashkirian–Moscovian conglomerates was located ~30 km to the east of the western basin margin. Large olistoliths slid down the basin slope into this region from the Sakmara– Magnitogorsk nappe. The Western flysch basin was only 30–50 km wide (Fig. 4a). Its western

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slope was very steep ( Khvorova, 1961) and could include a boundary scarp near to the platform margin. In the deepest part of the basin, finegrained and deep-water turbidites include layers of large olistoliths, up to ~100 km thick, that were derived from the Bashkir platform (C b and 2 C m in column IV ) (Misens, 1995a,b). Deposition 2 of shallow-water sands and conglomerates took place on the eastern basin slope (C b-m in column 2 III ). Thus, the eastern and western flysch basins included steep slopes, which were only L ~20–30 km wide. According to Eq. (1), this 0 shows that during the formation of the basins, T e decreased to ~3–5 km, i.e. the elastic layer in the lithosphere practically disappeared. Lithospheric weakening also follows from the absence of crustal subsidence at the front of the nappe where a shallow-water shelf came into existence (Figs. 2c and 4a), and from a steep basement bending under the nappe that reached ~10–15 km at a distance L ~20–30 km from its front (Fig. 4a). 0

shallow-water shelf with a depth of water h ~0.1 km. On its way, the nappe pushed out w Dh ~1–1.5 km of sediments with the density s r #2550 kg m−3 and piled them up at its front. s The nappe overrode h ~2–3 km of sediments. or After erosion of 2–5 km of the deposits from the nappe, it still preserved a large thickness, h ~15–18 km (Fig. 4a). Substituting these values n into Eq. (2) with the nappe density r =2700 kg m−3 and the intensity of compaction n of the overridden deposits e~0.2, we obtain: f0~1.8–2.5 km. Since the altitude of the nappe at n the end of collision C 1 was no more than ~0.5 km, a tectonic (unloaded ) crustal subsidence by j~1.3–2 km should have occurred due to the appearance of subsurface loads at the time of nappe superposition.

4.4. Low altitude of the nappe surface

5. Superposition of the eastern part of the Magnitogorsk arc onto the western slope of the East Uralian microcontinent in the late Late Carboniferous

In the process of superposition, the Sakmara– Magnitogorsk nappe, ~100–150 km wide, emerged above sea level ( Figs. 2c and 4a). The volume of the middle Bashkirian to middle Moscovian coarse clastic deposits is ~5 times larger in the Eastern flysch basin than in the Western flysch basin. This shows that a higher topography existed in the rear eastern part of the nappe (Fig. 4a). About 3–5 km of rocks have been eroded from this region, 50–70 km wide, the level of erosion reaching volcano-sedimentary rocks and ophiolites of the Sakmara–Magnitogorsk nappe. The eroded surface did not reach these rocks in the frontal part of the nappe, which was at a lower altitude. About 2 km of rocks have been eroded from this region. Deposition of conglomerates and turbidites in the Eastern flysch basin had ceased by the late Moscovian when shallow-water carbonates and evaporites began to form. This indicates planation of the topography at the end of collisional event C 1 and a low altitude of the nappe surface above sea level (most probably, ≤0.5 km). The Sakmara–Magnitogorsk nappe overrode a

At the end of convergence C 1, in the middle of the Moscovian age, the eastern (buried ) edge of the East European continent and the western edge of the East Uralian microcontinent ( EUM ) were still ~100 km apart with the undeformed part of the Magnitogorsk arc between them ( Fig. 2c). In the present structure, the eastern part of the arc overlies the western slope of EUM ( Fig. 7). Shallow-water deposition took place on this part of the arc until the end of the Moscovian (Chuvashov et al., 1984); however, no Kasimovian deposits have been found in the region. This most likely indicates that superposition of the arc onto the microcontinent slope occurred around the transition from the Moscovian to the Kasimovian event of collision C 2 ( Fig. 8a) (Seliverstov and Denisov, 1982). During ~4 m.y., the arc overrode the western slope of the microcontinent for ~50– 70 km in the south and ~100 km in the north ( Yazeva and Bochkarev, 1993). The arc was broken into tectonic slices and shortened by about a factor of two. On its way, this nappe ( East Magnitogorsk nappe) pushed out Dh =2–3 km of s the deposits and incorporated them into its frontal

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Fig. 7. Tentative scheme of the crustal structure along the URSEIS’95 profile; see Fig. 1 for location (modified after Skrypyi and Yunusov, 1989; Kazantsev et al., 1992; Berzin et al., 1996; Echtler et al., 1996).

Fig. 8. Evolution of the Southern Urals in the late Late Carboniferous–Early Permian. (a) Shortening of the eastern part of the Magnitogorsk arc and its superposition as the East Magnitogorsk nappe onto the western margin of the East Uralian microcontinent (EUM ) around the boundary between the Moscovian and Kasimovian-collisional event C 2; rapid formation of a deep-water basin in the present Uralian foredeep in the late Gzhelian. (b) Shortening of the Sakmara–Magnitogorsk nappe and the Bashkir shelf in the late Early Permian collision C 4; collision in the central and eastern parts of EUM (event C 5). (c) Planation of the topography in the Uralian thrust belt by the end of the Early Permian.

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part. Dh =5–7 km of rocks have been eroded er from the nappe surface later — in the Permian and Triassic (Sigov and Romashova, 1984). The present thickness of the nappe h1 reaches the n maximum value above the western edge of EUM where h1 =hmax~27 km (Fig. 7). After collision n n C 2, an increase in the crustal thickness due to convergence was reached in this region Dh =h1+Dh −Dh ~30 km. c n er s The East Magnitogorsk nappe was superimposed onto a shallow-water shelf. Without density changes in the lithosphere, the surface of the nappe, up to ~30 km thick, would reach a very high altitude. No coarse clastic material has been brought from the East Magnitogorsk nappe into the Western flysch basin in the Kasimovian and early Gzhelian (C k+g in column III in Fig. 5). 2 1 Fine-grained turbidites with minor carbonates were predominantly deposited on its eastern slope. Only fine-grained sediments and carbonates were formed in shallow marine basins that bounded the Urals in the south (Chuvashov et al., 1984). This shows that after collision C 2, the East Magnitogorsk nappe was hardly higher than ~0.5 km. Hence, a density increase occurred in the lithosphere below the nappe synchronously to superposition. In the early Gzhelian, a layer of large blocks, olistoliths, 100–150 m thick, was formed on the eastern slope of the Western flysch basin. This layer was composed of Visean and Moscovian limestones derived from the Bashkir shelf ( legend 6, C g in columns III, IV, Fig. 5) (Misens, 1995b). 2 1 At the same time conglomerates, up to ~100 m thick, were formed on the eastern basin slope in the northern part of the Southern Urals (Chuvashov, 1975). These data indicate a short impulse of convergence C 3 on the Bashkir shelf with erosion of several hundred metres of carbonates (see the hiatus between C k and C g in 2 2 2 column I and between C m and C g in column 2 2 2 II ). According to the small amount of erosion on the Bashkir shelf and small volume of synchronous coarse deposits in the adjacent flysch basin, collision was weak, and the altitude of eroded topography was low. The total shortening of the region during this period of time could not exceed ~10 km.

6. Evolution of the Uralian foredeep in the Late Precambrian–Carboniferous 6.1. Basin structure and deposition in the Late Proterozoic The thrust belt of the Southern Urals is bounded by the Uralian foredeep, 40–130 km wide (Fig. 1). The depth to the crystalline basement in this basin, on the eastern margin of the East European platform and in the Peri-Caspian basin is shown in Fig. 9. Over most of the Southern Urals, the foredeep is up to 12–15 km deep. The age of the crystalline basement varies over the basin. In most cases, this boundary coincides with the base of the Riphean (Fig. 10). In some regions, the basement, shown in Fig. 9, refers to the top of metamorphosed deposits of the Lower Riphean ( Fig. 11). The thickness of the Phanerozoic deposits varies along the foredeep, and the deposits are considerably thicker in the south. Thus, in the profile of Fig. 10, the thickness of the Paleozoic and Triassic reaches 10 km, while in the profile of Fig. 11, the Phanerozoic deposits are only 2–3 km thick. We will describe the evolution of the foredeep on the example of the southern part of Bel’sk depression, using the profile of Fig. 10 and the corresponding stratigraphic columns V–VII of Fig. 5. Their position is shown above the profile. Upper Permian and Triassic deposits in the foredeep are strongly deformed by salt-dome tectonics. Their average thicknesses in columns V–VII are given according to Unified Stratigraphic and Correlational Schemes of the Urals (1980). In the west, the foredeep is bounded by the East European platform. Column VIII refers to the eastern platform margin. In the east, the foredeep is bounded by the nappe, which includes shortened deposits of the Western flysch basin (right-hand side of Fig. 10). Slow deposition took place on the eastern margin of the East European continent in the Riphean (R) and Vendian ( V ) in the Late Proterozoic. In the Uralian thrust belt, deposits of this age are exposed to the surface in many places ( Figs. 1 and 11). The thickness of R+V varies along the foredeep and increases towards the thrust belt (Figs. 10 and 11). In the profile of Fig. 11, at

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of superposition of the Phanerozoic Uralian nappe over the Proterozoic sedimentary basin, whose depth increased eastwards. 6.2. Slow subsidence in the Devonian– Carboniferous and the crustal uplift in the middle of the Gzhelian age in the latest Carboniferous

Fig. 9. Depth to the crystalline basement in the foredeep of the Southern Urals and on the adjacent East European platform and Peri-Caspian basin (km) (modified after Semenovich et al., 1982).

the nappe front, these deposits are 12–13 km thick, and maybe even 15–17 km thick, if layer R (?) is 1 included. Similar thicknesses of R+V can be seen in the thrust belt in column II in Fig. 5. Hence variations in the thickness of R+V deposits at the nappe front can mostly reflect a variable magnitude

Almost no deposition took place in the Uralian foredeep and on the eastern platform margin in the Cambrian-Silurian (see the hiatus between V and S, D in columns V–VIII ). A slow crustal subsidence resulting in deposition of 1–2 km of shallow-water limestones with minor sands occurred under cratonic conditions in the Devonian–Carboniferous. Near to the end of the Late Carboniferous, in the middle of the Gzhelian, a short regression occurred for ~1 m.y. on a shallow-water shelf in the Uralian foredeep to the west of the Western flysch basin (Melamud, 1981). In a band, ~50–100 km wide, this resulted in erosion and hiatus increasing eastwards towards the Urals. Up to 500 m of limestones had been eroded in the eastern part of the foredeep (column V in Fig. 5). The lower Gzhelian, Kasimovian and upper part of the Moscovian are missing. Fifty kilometres further to the west, in columns VI and VII, only 50–100 m of the deposits had been washed out. No regression or erosion took place on the present platform margin (column VIII ). Strong lateral non-uniformity of erosion precludes a eustatic origin of regression and indicates that it resulted from a tectonic crustal uplift. This uplift cannot be recognized in the western part of the flysch basin where deep-water deposition proceeded until the very end of the Carboniferous (C g −C g in column IV ). 2 1 2 2 6.3. Rapid crustal subsidence in the late Gzhelian The crustal uplift in the middle of the Gzhelian was followed by rapid subsidence in the late Gzhelian. A deep-water basin, 40–100 km wide and up to ~1 km deep, was formed on the eastern platform margin to the west of the flysch basin during ~1 m.y. (Figs. 4b and 8a) ( Khvorova, 1961; Chuvashov, 1975; Melamud, 1981; Chuvashov et al., 1990). This was the initial fore-

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Fig. 10. Section of the southern part of the Bel’sk depression in the foredeep of the Urals (modified after Melamud, 1981; Semenovich et al., 1982).

deep of the Southern Urals. In its deepest eastern and middle parts, eroded shallow-water limestones were abruptly overlain by deep-water marls, black limestones and argillites ( legend 8 in columns V and VI in Fig. 5; see also Figs. 10 and 12). On the outer western foredeep margin, the subsidence was gradual. This basin part was separated by a normal fault from the adjacent East European platform. During the period of time, ~30 m.y. long, until the Artinskian age of the Early Permian, 0.7– 1.5 km high reefs were formed in this region (C g−P ar in column VII ) (Geology of the USSR, 2 1 v. XIII, 1964; Melamud, 1981). A marginal reef can also be seen in the left-hand side of profile of Fig. 12. Rapid subsidence probably also occurred in the deepest western part of the flysch basin. It cannot be recognized in this region, however, since a deepwater basin had already existed there before the Gzhelian subsidence. No intensive subsidence occurred further to the east — on the eastern slope of the Western flysch basin and on the Bashkir shelf. After a short hiatus, shallow-water deposition resumed on the Bashkir shelf in the late

Gzhelian (conglomerates 1, sands 2 and limestones 3 of C g in columns I, II ) and on the eastern 2 2 slope of the flysch basin at the beginning of the Early Permian (P a — Asselian in column III ). 1 After the rapid subsidence in the Gzhelian, a continuity has been preserved of the Upper Carboniferous shallow-water carbonates in the foredeep (9 in Fig. 12), which underlie the upper Gzhelian deep-water marls and argillites (8). Hence, the subsidence occurred without significant lithospheric stretching. In Fig. 11, except at some minor thrust faults, Riphean and Vendian strata are continuous in the Uralian foredeep, 15–20 km deep, which also indicates no stretching. In the profile of Fig. 10, a continuity of the strata in the foredeep can be reliably traced at the base of the Devonian. Several minor thrust faults exist in its inner part. The only normal fault occurs between the foredeep and platform margin in the west. This ensures a negligible mean lithospheric extension in the foredeep. At the end of the Carboniferous, collision C 2 took place only in the eastern Urals, far from the foredeep (Fig. 8a). No strong subsidence occurred

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Fig. 11. Section across the Uralian foredeep and frontal part of thrust belt (modified after Skrypiy and Yunusov, 1989). R — Riphean, V — Vendian, D — Devonian, C — Carboniferous, P — Permian, P s — Sakmarian, P k — Kungurian, P u — Ufimain, P k — 1 1 2 2 Kazanian, P t — Tatarian. 2

Fig. 12. Upper part of the sedimentary cover on the outer margin of the Uralian foredeep; see map in Fig. 6 for location (modified after Kamaletdinov, 1974). The transition from Upper Carboniferous shallow-water limestones (9) to upper Gzhelian deep-water marls, black limestones and argillites (8) is shown by arrows. A continuity of the Upper Carboniferous shallow-water limestones (9) precludes significant lithospheric stretching since the end of the Carboniferous.

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in a wide region between the foredeep and the area of collision. The Bashkir shelf to the east of the foredeep remained near sea level or slightly above it (C g and P a in columns I, II of Fig. 5). 2 2 1 These features show that the rapid crustal subsidence in the foredeep in the late Gzhelian did not result from collision. In the absence of stretching, this could be caused only by a density increase in the lithosphere.

7. Evolution of the Uralian foredeep and thrust belt in the Early Permian 7.1. Collision in western Urals in the late Early Permian and deposition in the foreland basins Three collisional phases took place in the second half of the Early Permian. They occurred in the middle of the Sakmarian (C 4a), Artinskian (C 4b) and Kungurian (C 4c) ages (Chuvashov, 1975; Chuvashov and Misens, 1980; Misens, 1995a,b). These phases were each 2–3 m.y. long and separated by longer time intervals ~5–6 m.y. Under the pressure of the advancing Sakmara– Magnitogorsk nappe, the Bashkir shelf was shortened by 10–20 km ( Figs. 4b and 8b). This resulted in strong folding and slicing with a complete overriding of the shelf by the nappe in some places. Most likely, the Sakmara–Magnitogorsk nappe was also shortened at that time with thrusting in the crystalline basement ( Fig. 7). Each phase of collision was accompanied by deposition of conglomerates on the western margin of the Bashkir shelf and on the eastern slope of the Western flysch basin (columns I–III in Fig. 5). Olistostromes were formed further to the west — in the deep-water western part of the flysch basin (column IV ) and in the adjacent eastern part of the foredeep (column V ). Coarse material is mostly represented by cherts, mafic volcanites and ultramafic rocks that were derived from the Sakmara– Magnitogorsk nappe (Geology of the USSR, v. XIII, 1964; Chuvashov, 1975; Melamud, 1981). Conglomerates deposited at each phase of collision C 4 were 200–500 m thick. In addition, at that time, 1.5–3 km of turbidites had been formed in the outer western part of the flysch basin and in the eastern part of the foredeep. These are

indications of intensive erosion all over western Urals. Folds and cordilleras, ≥1 km high, were probably formed in this region during the collisional phases. Piles of conglomerates are, however, separated by layers of sands, carbonates and evaporites formed during the epochs of stability 5– 6 m.y. long between the collisional phases. They can be seen, for example, in the interval P s−P k 1 1 in columns I–III. This shows that a rugged topography formed at the phases of collision became rapidly eroded, and between them, the nappe was ≤0.5 km above sea level. Near to the end of the Artinskian, the innereastern part of the Uralian foredeep became filled with ~3.5 km of turbidites with minor pelagic marls and limestones (column V ). The top of the Artinskian (P ar) in this column is represented by 1 shallow-water limestones. The central part of the foredeep (column VI ) remained at deep water until the Kungurian age of the Early Permian. Only 150–200 m of black pelagic limestones and marls were formed in this region during ~30 m.y. after the rapid subsidence in the Gzhelian. The height of the reefs on the western basin margin (0.7– 1.5 km) is a good constraint for a minimum water depth in the deep-water central part of the foredeep. In the Kungurian, the Uralian foredeep and the Peri-Caspian basin became semi-isolated from the ocean, and during the following ~7 m.y. the central and outer parts of the foredeep have been filled with salt (P kg in columns VI, VII ) 1 ( Khain, 1977). No significant additional subsidence occurred in the foreland region at the collisional phases C 4a–C 4c. A slow shallow-water deposition proceeded on the eastern slope of the flysch basin in the Sakmarian–Artinskian (P s–P ar in column 1 1 III ), and only 1.5 km of the deposits were formed during this period of time, 20 m.y. long. The adjacent western deep-water part of the flysch basin had been filled with turbidites by the late Artinskian and became shallow water after phase C 4b (conglomerates at the top of P ar in column 1 IV ). Only about 300 m of shallow-water salts and sands were formed in this region in the Kungurian age (P k), which includes collisional phase C 4c. 1 Thus, the collision in the late Early Permian did not result in the formation of high mountains and strong crustal subsidence in the Uralian foredeep.

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7.2. Continental collision on the East Uralian microcontinent in the late Early Permian Shallow-water deposition proceeded in the central and eastern parts of EUM until the end of the Kasimovian age of the Late Carboniferous (Chuvashov et al., 1984). After a hiatus, ~40 m.y. long, marine transgression occurred in the eastern part of the microcontinent in the Kazanian age of the Late Permian. Between these two epochs, continental collision took place in the eastern and central parts of EUM (event C 5) ( Kamaletdinov and Kazantseva, 1983; Puchkov and Svetlakova, 1993). The shortened region includes numerous granitic plutons ( Keilman, 1974; Perfiljev, 1979) that were emplaced into the upper part of the crystalline crust. Many of them were formed along thrust faults, most probably, synchronously with the deformations (Popov and Rapoport, 1996). The plutons are dated as the second half of the Early Permian (Smirnov, 1997). Hence, shortening of EUM could probably occur at the time of collision C 4 in western Urals ( Fig. 8b). Before collision C 5, the central part of EUM was approximately at sea level, and it can be supposed that the crustal thickness at that time was h0~40 km. The mean present crustal thickness c in this region is h1 ~50 km. No further collision c took place on EUM after collisional event C 5, and ~5 km of rocks have been eroded from this region in the Late Permian and Triassic (Sigov and Romashova, 1984). Hence, it is probable that collision C 5 resulted in an increase in the crustal thickness in the central part of EUM by Dh ~15 km. Taking this value in Eq. (3) with the c mean density of the crust r =2830 km m−3, we c obtain an average altitude of the mountains that would be formed in the absence of a density increase in the lithosphere as Df~2.3 km. This is comparable with the average altitude of the Alps. No deposition took place near to EUM on the western margin of the Kazakhstan continent in the Early Permian. Clastic material that had been brought into this region from the Urals became eroded and removed by rivers into the northern part of the West Siberian basin (Bush et al., 1995). Hence, the data on the Kazakhstan cannot be used for estimates of the altitude of eastern Urals after

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collision C 5. However, at the end of the Early Permian, salts with a small amount of fine-grained terrigenous material were forming on the Uralian foreland in the west (P k in columns I–VII, Fig. 5). 1 At the start of the Late Permian, thin layers of carbonates were formed between the Urals and Kazakhstan. This shows that the topography produced by collision C 4, C 5 in western and eastern Urals was rapidly eroded, and the thrust belt became smooth, remaining at a moderate altitude ≤0.5 km (Chuvashov et al., 1984).

8. Deposition in the Uralian foredeep, mountain building and collision in the Late Permian–Triassic 8.1. Rapid crustal subsidence in the Uralian foredeep and mountain building in thrust belt in the Late Permian Since the start of Late Permian and until the Middle Triassic, terrestrial deposition prevailed in the Uralian foredeep. Sands, clays and coastal marine carbonates were forming during the Ufimian and Kazanian (P u and P kz in columns 2 2 V–VII ). The sediment loaded subsidence occurred at a rate ≤300 m m.y.−1, and minor conglomerates appeared in the deposits. This indicates the onset of crustal uplift in the Uralian thrust belt in the Ufimian (Geology of the USSR, v. XIII, 1964; Khain, 1977). At that time, the Western flysch basin and the shortened Bashkir shelf slightly emerged above sea level, and no deposits younger than the Kungurian deposits occur in these regions (P k in columns I–IV ). 1 A strong acceleration of the subsidence took place in the inner-eastern and central parts of the foredeep in the Tatarian age (3 km of P t in 2 column V and 2 km in column VI ). Rapid deposition of coarse molasse took place in these regions in the first half of the Tatarian. The cumulate thickness of conglomerates is 0.5–1 km. Deposition of coarse molasse was caused by erosion of the rapidly rising Uralian thrust belt (Fig. 13a). The material had been transported to the foredeep by rivers that were flowing across the slightly elevated region of the former Bashkir shelf and Western flysch basin. Conglomerates were deposited in piedmont fluvial fans in a band, several

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Fig. 13. Evolution of the Southern Urals in the Late Permian to the Middle Triassic. (a) Mountain building in thrust belt and rapid sediment loaded subsidence in the Uralian foredeep in the Late Permian and Early Triassic. (b) Last phase of convergence (C 6) with an advance of the Sakmara–Magnitogorsk nappe and shortening of the Western flysch basin in the early Middle Triassic.

tens of kilometres wide (Geology of the USSR, v. XIII, 1964). The thickness of the Lower Tatarian conglomerates rapidly decreases westwards. Sands, clays and lacustrian carbonates are predominant in the inner part of the foredeep in the upper Tatarian (the upper part of P t in column 2 V ). No significant changes in the climatic conditions took place in the Southern Urals during the Late Permian and Early Triassic. Hence, it is probable that at the Late Permian time, the highest altitude of the Uralian mountains was reached in the early Tatarian. At the end of the Late Permian, weathered mantles were forming in eastern Urals (Bochkarev, 1973). This indicates the end of the first epoch of mountain building and planation of the topography. 8.2. Second impulse of mountain building in the Early Triassic A new impulse of intensive uplift occurred in the Uralian thrust belt in the Early Triassic. This gave rise to deposition of up to 1 km of conglomerates in the central part of the foredeep (molasse T in column VI ). According to the large size of 1

boulders (up to ~50 cm), at that time, the Uralian mountains probably reached the highest altitude. Coarse molasse includes the material derived from the area, 200–300 km wide, namely carbonates from the shortened Bashkir shelf, cherts, diabases and ultramafic rocks from the Sakmara– Magnitogorsk nappe, metamorphic rocks from the Uraltau, juspers from all the Magnitogorsk zone, and granitoids from its eastern part. In the Early Triassic, a slight crustal uplift occurred in the inner part of the Uralian foredeep, and the deposition ceased in this region (column V ). The central and outer parts of the foredeep proceeded to subside until the Late Triassic ( T 1 and T in columns VI, VII ). The rate of subsi2–3 dence in the central basin part decreased severalfold as compared to the Late Permian. Since the late Late Carboniferous and until the Triassic, 1 km of sediments had been deposited on the eastern margin of the East European platform. During the same epoch, up to 7 km of deposits had been formed in the Uralian foredeep (Fig. 5). This led to the formation of a deep foreland basin. No extensive compression occurred in the Uralian foredeep at the epoch of mountain

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building in the Late Permian–Early Triassic. Preservation of a thin layer of Kungurian deposits on the former Bashkir shelf and in the Western flysch basin (P k in columns I–IV, Fig. 5) shows 1 that no shortening occurred in these regions, which were at a low altitude above sea level. The front of the Uralian nappe remained stable (Fig. 13a) ( Kamaletdinov, 1974; Melamud, 1981). By the Late Permian, the nappes in eastern Urals had already been cut by granitic intrusions (Popov and Rapoport, 1996; Smirnov, 1997). Furthermore, during mountain building, since the start of the Late Permian slight extension with normal faulting took place in the Uralian thrust belt (Bochkarev, 1973; Chuvashov et al., 1984; Popov and Rapoport, 1996; Ivanov, 1997). The intensity of extension increased in the Early Triassic, which resulted in graben formation and basaltic magmatism (Bochkarev and Nesterov, 1987; Surkov et al., 1987). Thus, mountain building in the Urals could not be produced by shortening of the crust in the thrust belt or underthrusting under it of the lithosphere from the foreland basin. This required a density decrease in the lithosphere. Conglomerates, up to several hundred metres thick, occur in the lower part of the sedimentary pile in Early Triassic to Jurassic grabens in eastern Urals (Bochkarev, 1973; Ivanov, 1997; Puchkov, 1997). Since the end of the Early Triassic, only sands, clays and coals were forming in the grabens (Bochkarev, 1973; Ivanov, 1997; Puchkov, 1997), which shows that the Uralian mountains became eroded.

9. Collision in the Middle Triassic and postcollisional vertical crustal movements At the start of the Middle Triassic, over a few million years, shortened deposits of the Bashkir shelf were pushed westwards for 15–30 km and overrode the inner-eastern part of the Western flysch basin, collision C 6 (Fig. 13b) ( Kamaletdinov, 1974; Melamud, 1981; Zhivkovich and Chekhovich, 1985; Brown et al., 1997). The sedimentary cover of the flysch basin was shortened by 20–50% and superimposed as a nappe onto the innermost part of the Uralian foredeep, 15–20 km wide (see Fig. 6, and right-hand side of

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Fig. 10). The nappe pushed out Permian deposits from the latter region, and most of them had been eroded. A minor part of these deposits became incorporated into the frontal part of the nappe. Nappe superposition was associated with extensive erosion on its surface. Sands and clays with minor conglomerates were formed in the Middle Triassic in the adjacent Uralian foredeep ( T in columns VI, VII, Fig. 5). A small volume 2 of coarse clastics indicates that rapid erosion of the nappe surface limited its average altitude at a level of ~0.5 km. No more than ~0.5 km of sediments was deposited in the central part of the foredeep in the Middle and Late Triassic ( T in 2–3 column VI ). Thus, the last collisional event in the Urals resulted in no mountain building, and no significant subsidence in the foreland region. By the end of the Triassic, a planation surface was formed over most of the Southern Urals at an altitude of several hundred metres (Sigov and Romashova, 1984). At that time, the thrust belt was surrounded by land areas. Since the Jurassic and until the Eocene, the Southern Urals remained at an altitude 200–300 m above sea level. Hard Precambrian blocks of the Uraltau and Bashkir Anticlinorium were staying 0.5 km above the planation surface. Since the Oligocene, the crustal uplift has taken place in the east of the East European platform and in West Siberia. Its magnitude is ~200 m in the former region and 100–150 in the latter region. The Southern Urals were also involved in a domelike uplift with magnitudes ~400–500 m ( Khain, 1977). All over these regions, the crustal uplift occurred without any significant compressive deformations. Using the above description of data, we can formulate the answers to questions 1–9 in Section 1.

10. Basic regularities of vertical crustal movements: no direct correlation with collision 10.1. Slow and rapid subsidences: their role in the formation of foreland basins To discuss the main regularities of the crustal subsidence on the Uralian foreland, we use the

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subsidence curves of Fig. 3. These are compiled based on the stratigraphic columns of Fig. 5. The numbers of the curves correspond to those of the columns. For the periods of shallow-water deposition, the amount of subsidence is taken as the sediment thickness. The depth of water loaded basins is usually rather uncertain. For the rapid water-loaded subsidences, we conventionally take their magnitude as 80% of the sediment thickness that filled the deep-water basins. The remainder of 20% is attributed to a slow subsidence during the following period of time until the recovery of shallow-water deposition. We presume that the duration of the rapid subsidences was 1 m.y. This approximation is very schematic. However, it allows us to distinguish between the epochs of slow and rapid subsidence and to compare the epochs of rapid subsidence with the main collisional events. As can be seen from Figs. 3 and 5, a slow sediment loaded subsidence at a rate ~10–30 m m.y.−1 took place on the Uralian foreland over most of time. In the Silurian–Carboniferous, during 100–140 m.y., it resulted in the formation of 0.5–2.5 km of predominantly shallowwater deposits. At the beginning of the Late Carboniferous, the deep-water Western flysch basin was formed by rapid crustal subsidence (curve IV ). Another deep-water basin was rapidly formed in the Uralian foredeep at the end of the Late Carboniferous (curves V, VI ). The depth of water in these two basins was ~1 km. In curves IV–VI, the magnitude of the rapid subsidence that formed the basins is shown as ~2–3 km. This corresponds to 80% of the sediment thickness that filled the basins. It is very likely that at the end of the Late Carboniferous, rapid subsidence also occurred in the flysch basin (curve IV ). The third rapid subsidence took place in the Uralian foredeep in the Late Permian (curves V, VI ). The subsidence by 2.5–3 km was sediment-loaded and reached a rate of ~0.6 km m.y.−1. The magnitudes of the total subsidence since the start of the Carboniferous are rather similar in the west (in the Uralian foredeep and in the western part of the Western flysch basin) and east (in the eastern part of the flysch basin and on the Bashkir shelf ). In the west (curves IV–VI ), short

impulses of rapid tectonic subsidence ensured the major input into the formation of the foreland basins. In the east, a slow sediment loaded subsidence was predominant (curves I–III ). However, the subsidence rate strongly varied in time and sometimes reached 100–200 m m.y.−1 in the Late Carboniferous and Permian. 10.2. Absence of strong crustal subsidence on the Uralian foreland at the epochs of collision in thrust belt As can be seen from Figs. 3 and 5, the rapid crustal subsidences on the Uralian foreland occurred when there was no strong collision in the adjacent western part of thrust belt. Rapid subsidence in the Western flysch basin in the early Late Carboniferous preceded superposition of the Sakmara–Magnitogorsk nappe (curve IV at the epoch of collision C 1). The nappe began to move 2 m.y. later, and its front was initially ~150 km from the basin (see Sections 4.1 and 4.2). The first rapid subsidence in the Uralian foredeep at the end of the Carboniferous took place at the end of collisional event C 3 (see curves V and VI in the Gzhelian). This event was very weak (shortening by ≤10 km) and could not result in the formation of the deep-water basin, ≥50 km wide. The second rapid subsidence in the foredeep (same curves) took place in the Late Permian. It was synchronous to mountain building in the thrust belt; however, no collisional events occurred at that time. However, collision in the western Urals did not significantly influence the subsidence on the Uralian foreland. Strong collisional events C 1, C 4 and C 6 were accompanied by only minor vertical displacements in curves I–VI. It can be suggested that the collisional events are determined incorrectly, and collision actually took place at the epochs of rapid subsidence on the foreland, which are timed precisely from the stratigraphic records. The magnitude of the subsidence due to deflection of the elastic lithosphere should increase towards the thrust belt. However, the areas of rapid subsidence on the foreland were separated by stable blocks from the nappe. At the time of rapid subsidence in the Western flysch basin in the early Bashkirian (curve IV in Fig. 3),

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the basin was separated from the nappe by the Bashkir shelf ( Fig. 2b), where no strong subsidence took place (curves I, II ). No strong subsidence occurred on the eastern slope of the flysch basin and on the western margin of the deformed Bashkir shelf at the end of the Carboniferous when rapid subsidence took place in the foredeep and in the western part of the Western flysch basin (Fig. 8a) (compare curves I–III with curves IV–VI in the Gzhelian). At the epoch of rapid subsidence in the foredeep in the Late Permian (Fig. 13a), the flysch basin between the foredeep and thrust belt was stable and remained above sea level (compare curves V, VI with curves III, IV in the Late Permian). Deepening of the Paleozoic–Triassic foreland basins away from thrust belt can also be clearly seen in Fig. 5. Thus, irrespective of timing of continental collision in western Urals, it could not be responsible for the formation of foreland basins. 10.3. High degree of weakening of the lithosphere indicated by its narrow-wavelength deformations The absence of flexural reaction of the lithosphere on the Uralian foreland on collision in thrust belt indicates either a strong decrease in the effective elastic thickness of this layer, T , or its e disruption by large normal faults. The magnitude of crustal subsidence under the Sakmara– Magnitogorsk nappe increased from zero at its front to 10–15 km at a distance, L ~20–30 km 0 further to the east (Fig. 4a and b). According to Eq. (1), these values of L correspond to very low 0 T values of ~3–5 km. In fact, at a dip angle of e the basement ~25–30°, the deformations were beyond the limit of elasticity. A system of normal faults could arise in this region. Large narrow-wavelength variations of the nappe thickness have been preserved until present. Thus, in the Kimpersay massif, the nappe thickness increases from h ~3–5 km to h ~10–15 km at an n n interval, L ~15 km wide (Saveljev and Saveljeva, 0 1991). The thickness of the East Magnitogorsk nappe on the western slope of the East Uralian microcontinent reaches 25–28 km at a distance, L ~75 km from the nappe front ( Fig. 7). 0 Strongly differentiated vertical crustal move-

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ments also took place during the crustal subsidence on the Uralian foreland, resulting in large lateral variations in thicknesses of synchronous deposits. Thus, in columns II and IV of Fig. 5, shallowwater Silurian beds were initially horizontal. At the end of the Early Permian, the base Silurian became 3–4 km deeper in column IV than in column II. The distance between these two regions was L ~30–40 km. Column V in Fig. 5 includes 0 3 km of Tatarian deposits (P t), while no subsi2 dence occurred at L ~20–30 km further to the 0 east in column IV. In the western part of the Uralian foredeep (Fig. 10), the initially flat layer of Late Carboniferous shallow-water limestones is bent down by 2.5 km in a band, L ~25 km wide. 0 The lithosphere of the Uralian foreland is disrupted by numerous faults (Fig. 9). Steep bending and faulting of the lithospheric layer indicate its high degree of weakening under the Uralian nappe and on the foreland. 10.4. Absence of significant lithospheric stretching on the foreland The depth of a sedimentary basin formed by uniform lithospheric stretching by b times in a state of isostasy is: h =[(r −r )/(r −r )]h0 (1−1/b), s m c m s c

(4)

where h0 is the thickness of the prestretched crust. c The thickness of Phanerozoic sediments in the Uralian foredeep reaches ~10 km ( Fig. 10). At the mean value of r =2500 kg m−3 and h0= s c 40 km, the value of b necessary to form this basin by stretching would reach ~1.7. Lithospheric stretching of such an intensity should result in extensive normal faulting with a formation of strongly tilted fault blocks. Except at the steep boundary fault and minor thrust faults, the Riphean strata are continuous in the foredeep ( Figs. 10 and 11). This means that the basin was formed without any significant stretching in the upper crust at least. Stretching of only the lower crust can be presumed, however, to be a cause of the subsidence. Designate the density of the lower crust by r , and its thinning from stretching by lc Dh . In a state of local isostasy and after a lc

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Fig. 14. Crust and mantle structure along the Krasnouralsk (A) and Troitsk (B) profiles (modified after Beloussov et al., 1991; Druzhinin et al., 1997).

complete thermal relaxation, Dh =[(r −r )/(r −r )]h . (5) lc m s m lc s h #7 km of the deposits were formed in the fores deep after the rapid subsidence at the end of the Carboniferous. Taking r #2500 kg m−3and s r =2930 kg m−3, which is typical of gabbro in lc the lower crust, we have: Dh #14 km. The occurlc rence of garnet granulites in the lower crust increases its mean density, r . Then, according to lc Eq. (5), a higher degree of thinning of the crust will be necessary to ensure a subsidence of the same magnitude. Thinning of the crust by 14 km with an addition of 7 km of deposits will raise the base of the crust by 7 km. Downwarping of the Moho with respect

to the East European platform actually occurs beneath the Uralian foredeep ( Figs. 7 and 14). This means that the subsidence in the foredeep took place without any significant stretching of the lower crust. As shown in Section 4.1, no significant stretching took place in the Western flysch basin at the time of initial rapid subsidence between the Serpukhovian and Bashkirian. There are no data that could reliably resolve the presence or absence of stretching on the Bashkir shelf where only a slow subsidence took place. However, it is extremely unlikely that a strong extension could have occurred in this region near to the front of the nappe and in the absence of stretching on the Uralian foreland further to the west.

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10.5. Density increase in the lithosphere as a cause of crustal subsidence on the Uralian foreland The crustal subsidence in the Uralian foreland basins occurred without any significant lithospheric stretching and was independent of collision in the thrust belt (see Sections 10.2 and 10.4). Under such circumstances, a density increase in the underlying lithosphere was necessary to ensure subsidence. 10.6. Continental collision without a synchronous formation of high mountains All the events of continental collision in the Southern Urals (C 1–C 6) were accompanied by deposition of moderate volumes of coarse clastic material in the adjacent basins. Deposition of only sands, clays, carbonates and evaporites took place in these basins after the end of each collisional event. This shows that after rapid erosion of local folds and cordilleras formed at the epochs of collision, the areas remained at a low altitude ≤0.5 km above sea level. The occurrence of several epochs of strong thrusting in the Southern Urals long before the onset of mountain building has been described by many authors ( Khain, 1977; Perfiljev, 1979; Chuvashov et al., 1984; Zonenshain et al., 1990; Artyushkov, 1993; Khain and Lomize, 1995; Puchkov, 1997). Continental collision orogeny is, however, commonly considered as a synonym of mountain building. We considered the above stratigraphic data in detail to confirm that collision in the Urals did not result in the formation of high mountains. 10.7. Density increase in the lithosphere at the epochs of collision Events of collision in the Southern Urals began when the crust was near, or ≤0.5 km above, sea level. Collision produced a large increase in the crustal thickness, which, at a constant lithosphere density, must result in the formation of high mountains. Then, during the following ~10– 20 m.y. a thick layer of rocks would be eroded from the mountains with a recovery of approximately the initial crustal thickness. In particular,

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thick nappes superimposed onto shallow-water regions would be eroded almost completely. In fact, after the end of collisional events, the crustal surface remained at an altitude ≤0.5 km, with only a partial erosion of the nappes. In some regions, thick nappes have been preserved until the present. Collisional event C 1 (Fig. 2c) occurred soon after closure of the Kazakhstan oceanic basin to the east of EUM in the middle of the Serpukhovian ( Fig. 2b), which could be associated with the appearance of slab pull. At that time, however, no strong subsidence occurred on this microcontinent near the collisional front. This indicates delamination of the subducted slab synchronously with plate collision. A density increase in the underlying lithosphere was necessary to maintain the Sakmara–Magnitogorsk nappe at a low altitude. Similarly, a density increase should have occurred in the lithosphere during the following epochs of strong collision that took place in the absence of slab pull. Shortening of the crust in the Southern Urals had been predominantly completed before the onset of mountain building in the Late Permian. Only ~15–30 km of shortening took place in the Middle Triassic, after the end of mountain building. Hence, the Uralian crust had been almost completely formed by the Late Permian. The present crustal thickness in the Urals reaches 50– 60 km ( Figs. 7 and 14), and up to ~5–10 km of rocks have been eroded since the Late Permian (Sigov and Romashova, 1984). At the end of the Early Permian, the crustal thickness reached h ≥60–65 km. These values of h are typical of c c high mountains. However, deposition of salts with minor sands and shales on the Uralian foreland at the end of the Early Permian indicates that the Uralian thrust belt was no higher than ~0.5 km (Section 8.1). Therefore, at that time, the lithosphere in the Urals was considerably denser than in low cratonic areas where the mean crustal thickness was h ~40 km. c 10.8. Mountain building without synchronous shortening of the crust According to the appearance of large volumes of coarse molasse in the Uralian foredeep, moun-

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11. Mechanisms of the crustal subsidence and lithospheric weakening

us take, for example, d −d ~50 km. Then, at a 2 1 mean sediment density, r =2650 kg m−3, it fols lows from Eq. (6) that hth~4.7 km. The sediment s thickness on the Uralian foreland reaches ≥15 km. Hence, thermal relaxation could be responsible for only about a third of the total subsidence. Furthermore, the characteristic cooling time of the lithosphere, 150–200 km thick, is ~90–170 m.y. This phenomenon could not influence the subsidence of the crust, ~1000 m.y. old, in the Paleozoic and Early Mesozoic when the Uralian thrust belt was evolving. Erosion of the lower crust by convective flows in the mantle can occur when the asthenosphere wells up to the base of the crust. In cratonic areas with a thick mantle lithosphere, such an upwelling must be associated with a crustal uplift by ≥1 km. Only slight uplifts took place on the Uralian foreland in the Phanerozoic. Furthermore, as in the case of stretching of the lower crust, subcrustal erosion should result in a large uplift of the Moho boundary. Under the Uralian foredeep, the Moho is warped down with respect to the adjacent platform margin ( Figs. 7 and 14). These features preclude the subsidence from subcrustal erosion.

11.1. Possible role of thermal relaxation and subcrustal erosion

11.2. Crustal subsidence from metamorphism in the lower crust

A density increase in the lithosphere can result from thermal relaxation (Sleep, 1971; McKenzie, 1978), subcrustal erosion ( Keen, 1985), and metamorphism in mafic rocks in the lower crust (Haxby et al., 1976; Artyushkov and Baer, 1983; Artyushkov et al., 1991; Baird et al., 1995). Cooling of the lithosphere, which was initially at sea level, produces a sedimentary basin with a depth:

In the absence of strong thermal relaxation and subcrustal erosion, a large density increase beneath the Uralian foreland could be produced only by metamorphism in mafic rocks in the lower crust. In the case of local isostasy, the formation of a layer of garnet granulites with density, r , and gg thickness, h , from a gabbro with density, r , gg gb results in the subsidence, h , expressed by: s h =(r /r )[(r −r )/(r −r )]h . (7) s m gb gg gb m s gg In the profiles of Fig. 14, a layer, ~8–13 km thick, with P-wave velocities, V ~7.4–7.5 km s−1, is P located in the lowermost crust beneath the foredeep. At such a velocity, garnet granulites can have a density, r ~3250–3300 kg m−3 (Sobolev gg and Babeiko, 1994). Similar P-wave velocities are typical of the lowermost crust below the Uralian thrust belt, where the density has been estimated as ~3240 kg m−3, according to the gravity data

tain building in thrust belt took place in the Late Permian–Early Triassic ( Fig. 13a) ( Khain, 1977). At that time, the nappe front remained stable. Slight extension with the formation of grabens and basaltic volcanism occurred in the rising thrust belt. Thus, mountain building in the Urals was associated with no synchronous shortening of the crust. 10.9. Mountain building as a result of a density decrease in the lithosphere The Southern Urals are wide, and their crust was (and still is) close to the isostatic equilibrium (Artemjev et al., 1994; Do¨ring et al., 1997). Mountain building in the Late Permian–Early Triassic took place without any additional thickening of the crust. Under such circumstances, a considerable density decrease in the lithosphere was necessary to ensure the formation of the mountain range.

hth =a(T /2)[r /(r −r )](d −d ). (6) s a m m s 2 1 Here, a=3×10−5 K−1 is the thermal expansivity, T =1300°C is the temperature of the asthenospha ere, and d and d are the initial and final lithosphe1 2 ric thickness, respectively. The subsidence on the Uralian foreland began in the Riphean on the lithosphere that was already rather cool and thick. Hence, the lithospheric thickness could not be considerably increased during the subsidence. Let

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(Do¨ring et al., 1997). Taking r =2400 kg m−3, s r =2930 kg m−3, r ~3250–3300 kg m−3, h = gb gg gg 8–13 km in Eq. (7), we obtain: h ~4–5 km. The s sediment thickness that has been formed in the foredeep after the end of thermal relaxation (~10 km) is ~5 km larger, and an additional density increase was necessary to ensure this subsidence. P-wave velocities under the Moho beneath the foredeep are V ~8.0–8.1 km s−1. Such velociP ties can be typical of both the mantle and denser eclogites with r ~3500 kg m−3 (ibid., Manghnani e et al., 1974; Christensen and Mooney, 1995). Hence, it cannot be precluded that the Moho beneath the foredeep in this profile is underlain by eclogites. Let us take the density of sediments and metasediments in the lower part of the sedimentary cover as r =2700 kg m−3. Then, according to Eq. s (7), to ensure an additional subsidence of ~5 km at r =r ~3500 kg m−3, the thickness of eclogitic gg e layer below the Moho should also be h =h ~5 km. e gg In the profile of Fig. 14a, the Moho boundary beneath the foredeep is at a depth of ~45–48 km. According to their composition, mafic eclogites pertain to the crust. If we add ~5 km of eclogites to the crust, its thickness will be 50–53 km. Then, the thickness of the crystalline crust that is overlain by ~15 km of sediments will be ~35–38 km, as on the adjacent platform margin. In this case, the formation of the Uralian foredeep can be explained by thermal relaxation in the lithosphere and metamorphism in the lower crust with a preservation of the material in the crystalline crust. The high value of V ~8.5 km s−1 occurs under P the Moho in the Uralian foredeep in profile of Fig. 14b. The composition of rocks with such velocities is a special problem. These might be eclogites with a low mean atomic number. 11.3. Rapid metamorphism and crustal subsidence due to infiltration of volatiles from the asthenosphere At the epochs of slow crustal subsidence on the Uralian foreland, the subsidence evolved at a rate ~10–100 m m.y.−1. The rapid subsidences produced water-loaded basins ~1 km deep during ~1 m.y. This corresponds to an isostatic subsi-

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dence under a sediment load by ~3 km at a rate of ~3 km m.y.−1, which is 1.5–2.5 orders of magnitude higher than that of slow sediment loaded subsidence. The increase in the subsidence rate should be associated with a drastic increase in the rate of gabbro–garnet granulites–eclogite transformation. Consider possible causes of this phenomenon. Typical diagrams of the gabbro–garnet granulite–eclogite transformation are shown in Fig. 15. Fig. 15b is more representative of the continental crust. Due to a very low reaction rate, the experiments were carried out at high temperature, and only extrapolations can be used for a medium temperature typical of the lower crust in cratonic areas. If these extrapolations are correct, at pressure, p≥0.5–1 GPa and temperature, T ≤400– 600°C, the gabbro in the lower crust is metastable and should transform into dense eclogites. In cool and dry mafic rocks, the transformation is very slow and could be responsible for only a slow sediment loaded subsidence on the Uralian foreland. An increase in the rate of the transformation by 1.5–2.5 orders of magnitude requires heating of the rocks by several hundred degrees or the presence of a small amount of water-containing fluid (CO +H O) to catalyse the reaction (Ahrens 2 2 and Schubert, 1975). Several tens of millions of years are necessary to heat the lower crust in a thick cratonic lithosphere after an increase in the heat flow from the asthenosphere. This is incompatible with a short duration, ~1 m.y., of rapid crustal subsidences. Numerous examples are known for rapid metamorphism in mafic rocks in the presence of volatiles (Austrheim, 1987, 1998; Wayte et al., 1989; Rubie, 1990; Walther, 1994; Austrheim et al., 1997). The lower crust includes hydrous minerals such as biotite, muscovite, epidot and amphibol (Spear, 1993). Segregation of water occurs at T~700–950°C, and it is commonly supposed that volatiles appear in the lower crust due to the heating of rocks. Several tens of million years are also necessary to cause dehydration reactions in a cratonic lower crust by heating from the asthenosphere. This mechanism cannot produce strong crustal subsidences with a duration of ~1 m.y. To explain this, infiltration of volatiles into the

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Fig. 15. Experimental p–T diagrams of the gabbro–eclogite transformation. (a) Oceanic tholeites (modified after Ito and Kennedy, 1971), (b) Quartz tholeite (modified after Ringwood and Green, 1966). 1 — isolines of density (g cm−3); 2 — extrapolation of the curves.

lower crust from the asthenosphere has been suggested (Artyushkov et al., 1991; Artyushkov, 1993). Rapid subsidence with a formation of deepwater basins occurred in many areas, e.g. in West Siberia in the Late Jurassic and in the PeriCaspian, Volga–Urals and Timan–Pechora basins in the Late Devonian. Volatiles can be brought to the base of the lithosphere by small fluid-containing plumes (Artyushkov and Hofmann, 1997, 1998). Their emergence to the lithospheric base is manifested by slight crustal uplifts, ~100 m, of a short duration, ~1 m.y., that preceded rapid crustal subsidences in the above regions, as well as in many others (Artyushkov and Baer, 1986a; Artyushkov, 1993). A slight uplift also occurred before the rapid subsidence on the Uralian foreland at the end of the Late Carboniferous (Section 6.2). In many cases, slight uplifts that preceded rapid crustal subsidences without lithospheric stretching and collision nearby were associated with slight alkaline basaltic volcanism (e.g. the Volga–Urals, Timan–Pechora and Peri-Caspian basins) and plutonism (e.g. West Siberia) (ibid.). This can be another indication of infiltration of volatiles into the lithosphere. Infiltration of volatiles from the asthenosphere also follows from the geochemical data on the regions of rapid crustal subsidence. Sediments in the Volga–Urals and Timan–Pechora have very high contents of Se, As, Mo, Hg, U and Re (Pushkarev et al., 1994; Pisotsky, 1999). High

concentrations of these elements occur in layers with both high and low contents of organic matter. This precludes a significant influence of organic substances on the concentration of the microelements. Infiltration of volatiles from the mantle has been suggested as a mechanism of enrichment. Oils and bitumens in these regions include Nd and Sr, with the isotopic characteristics typical of lamproites in Australia and Spain derived from the mantle: e =−(9−12), and 87Sr/86Sr=0.708– Nd 0.719. In the Peri-Caspian basin, carbonates with a low content of organic matter, which were formed at the time of rapid subsidence in the Visean age, have a content of uranium (Pisotsky, 1999) that exceeds that typical of marine carbonates by one order of magnitude ( Taylor and McLennon, 1985). This has been interpreted as a result of transportation of uranium with fluids from the mantle. We suggest that the rapid crustal subsidences on the Uralian foreland, as well as rapid additional subsidences in the areas of collision in thrust belt, which maintained it at a low altitude, resulted from a gabbro–garnet granulites–eclogite transformation catalysed by infiltration of volatiles from the asthenosphere into the lower crust. 11.4. Weakening of the lithosphere as a result of infiltration of volatiles and rapid metamorphism Several mechanisms have been proposed to explain the high degree of weakening of the litho-

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sphere. Its extreme bending near thrust belts can result in yielding and a large decrease in T e ( Washbusch and Royden, 1992; Burov and Diament, 1995). Lithospheric weakening, however, occurred not only under the Uralian nappe, but also on the foreland at times when there was no collision in the thrust belt. Lithospheric stretching can also thin the elastic lithosphere (Desegaulx et al., 1991; Stewart and Watts, 1997). No rifting occurred in the Uralian foredeep for ~1000 m.y. before weakening of the lithosphere in the Carboniferous. A strong decrease in T takes place e during a high heat flow (Burov and Diament, 1995). The old cratonic lithosphere of the Uralian foreland was cool. Weakening of the lithosphere on the Uralian foreland occurred at the epochs of rapid crustal subsidence. Steep slopes of the crystalline basement, which are a few tens of kilometres wide, were formed during the rapid subsidences in many other areas (Artyushkov et al., 1996, submitted for publication; Artyushkov and Mo¨rner, 1997, 1998). This happened, for example, in the Carpathian and Caucasian foredeeps at times when there was no collision in the adjacent thrust belts, and the Peri-Caspian and Trans-Caspian intraplate areas. Thus, strong lithospheric weakening at times of rapid crustal subsidence is a widely occurring phenomenon. Rapid subsidences in sedimentary basins resulted from metamorphism in the lower crust under infiltration of volatiles. The same phenomena ensured an additional rapid subsidence in the regions of continental collision in the Southern Urals, which limited their altitude by ≤0.5 km. In all cases, lithospheric weakening occurred at times of infiltration of volatiles. Hence, the latter phenomenon can be supposed to be a cause of the former. As evidenced by large viscous deformations in the lower crust, rapid metamorphism in mafic rocks is associated with a strong viscosity decrease (Austrheim, 1991, 1998). At the epochs of rapid contraction of the lower crust from metamorphism, large deviatory stresses comparable with the lithostatic pressure can arise in the upper crust, which can result in a disappearance of the elastic core in this layer. In wet rocks, the viscosity is

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much lower than in dry rocks (Poirier, 1985; Strehlau and Meissner, 1987; Kohlstedt et al., 1995). Hence, the viscosity drop in the upper crust can also be caused by infiltration of volatiles. A strong viscosity decrease in the mantle lithosphere can be due to infiltration of volatiles through this layer from the asthenosphere into the crust. Thus, infiltration of volatiles can result in a high degree of weakening in the whole lithospheric layer (Artyushkov and Mo¨rner, 1997, 1998; Artyushkov et al., submitted for publication).

12. Crustal structure in thrust belt and mechanisms of mountain building in the Late Permian and Early Triassic 12.1. Composition of the transitional layer in the lower crust The crust in the Urals is now up to h ~60 km c thick (Figs. 7 and 14), as in the high mountains of the Tien Shan and Eastern Alps. Since the mean altitude of the Southern Urals is only 600–800 m, their crust should have a high mean density. In the profiles of Fig. 14, the transitional layer, up to ~20 km thick, with P-wave velocities of 7.4–7.5 km s−1 exists in the lowermost crust. These velocities are intermediate between those typical of gabbro in the basaltic layer (V =6.8– P 7.0 km s−1) and mantle peridotites under the Moho boundary (V =8.0–8.4 km s−1). In the P east, V in the transitional layer increases to P 7.8–8.0 km s−1. The lower crust of the Urals is characterized by a moderate temperature, T~600°C ( Khachay and Druzhinin, 1993; Kukkonen et al., 1997). At V =7.4–7.5 km s−1, it P can be composed of (1) a crust and mantle mixture, and (2) garnet granulites, or their mixture with eclogites. The formation of the crust and mantle mixture due to intrusions into mantle peridotites of layers of gabbro with a cumulate thickness, h , results gi in the crustal uplift by: Df =[(r −r )/r ]h , (8) gi m gb m gi In a volume that includes large blocks of peridotite and gabbro with P-wave velocities Vpr and Vgb , P P

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respectively, the average P-wave velocity will be: V #Vpr Vgb /[c( Vpr −Vgb )+Vgb ], (9) P P P P P P where c is the volume per cent of gabbro in the crust and mantle mixture. Taking Vgb= P 6.8 km s−1, Vpr=8.2 km s−1, we find that P V =7.4–7.5 km s−1 corresponds to a case when P gabbro constitutes c#0.45–0.53 of the total volume of rock. In the 20 km thick transitional layer, the total thickness of gabbroic intrusions will be h #9–10.6 km. Substituting these values gi into Eq. (8), we find: Df #1.1–1.3 km. The total gi thickness of the middle and upper crust above the transitional layer in the Urals reaches 40 km (Fig. 14). This is similar to the crustal thickness on the adjacent East European platform. According to Eq. (8), with ~9–10.6 km of gabbro in the transitional layer, the Urals would stand Df ~1.1–1.3 km above the platform. In fact, over gi most of the Urals, their elevation with respect to this area is only 0.3–0.5 km. The most likely time for the formation of the crust and mantle mixture with the associated crustal uplift would be the epoch of mountain building in the Late Permian–Early Triassic. Intrusion of a large volume of basaltic magmas into the mantle under the Moho boundary would most likely result in extensive basaltic volcanism. Intrusions of minor portions of basaltic magmas with T~1200°C into the overlying crust should result in granitic magmatism due to melting of metamorphosed sediments incorporated into the middle and lower crust at the preceding stages of collision. Most of the granitic intrusions had been formed in the Early Carboniferous and Early Permian — before the onset of mountain building. Slight basaltic volcanism began in the Southern Urals in the Middle Triassic — after the end of mountain building. At that time, eastern Urals represented only the western margin of a wide volcanic province that spread for ~1000 km further to the east into West Siberia. Thus, the occurrence of a large volume of gabbroic intrusions in the transitional layer is very unlikely. This layer would be most likely composed of dense garnet granulites formed from gabbro in the lower crust at the epochs of crustal subsidence and the following collision.

12.2. Possible crustal uplift in the future The lithospheric thickness in the Urals is estimated to be ~200 km (Steer et al., 1998). At a crustal thickness of ~60 km, the thickness of mantle lithosphere should be h ~140 km. ml Suppose that asthenospheric upwelling with a magnitude Dh ≤h will occur beneath the Urals in ml ml the future. Assume that the temperature in the mantle lithosphere increases linearly with depth from the Moho temperature, T , at the base of M the crust to the asthenospheric temperature, T , at a the base of the lithosphere. Then, a rapid replacement by the asthenosphere of the lower part of mantle lithosphere, Dh thick, results in the isoml static uplift by: Df =a(T −T )Dh2 /2h . (10) uw a M ml ml Suppose that, in the future, the asthenospheric upwelling will reach the base of the crust: Dh =h =140 km. Then, at T =1300°C and ml ml a T =600°C, we have: Df =1.5 km. M uw Garnet granulites are stable in the lower crust only at a moderate temperature. Upwelling of a hot asthenosphere to the base of the crust will result in their heating and retrogression to pyroxene granulites that are considerably less dense. This should produce an uplift at the surface (Artyushkov and Baer, 1986b; Artyushkov, 1993; Dewey et al., 1993; Artyushkov et al., 1996; Le Pichon et al., 1997). A transformation of garnet granulites in a layer, h thick, into pyroxene gg granulites with density, r , will raise the crustal pg surface by: Dfgg =[(r −r )/r ]h . (11) rg gg pg pg gg Taking h =20 km, r =3200–3250 kg m−3, r = gg gg pg 2960 kg m−3 (Bousquet et al., 1997), we have: Dfgg=2.0–2.3 km. The present mean altitude of rg the Southern Urals is f =0.6 km. The mean height p of the mountains formed due to asthenospheric upwelling to the crust will be f =f +Df +Dfgg =4.1–4.4 km. (12) f p uw rg Thus, in the future, the present Urals can become as high as eastern Tien Shan, where the crustal thickness is also ~60 km, but where asthenosph-

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eric upwelling reaches the base of the crust ( Yudakhin, 1983; Roecker et al., 1993). 12.3. Possible mechanisms of mountain building The amount of erosion in the Uralian thrust belt since the beginning of the Late Permian and up to the present reaches h =7–9 km on a regional er scale (Sigov and Romashova, 1984). Let us designate, by f , the average initial height of the moun0 tains and, by f , the average altitude of the area 1 after their erosion. Let r be the density of eroded er rocks. Then: f =[(r −r )/r ]h +f . (13) 0 m er m er 1 Taking h =7–9 km, r =2700 kg m−3 and f = er er 1 0.5 km, we find: f =1.9–2.2 km. The crustal uplift 0 that formed the Late Hercynian mountains in the Southern Urals occurred in two impulses: in the Tatarian age and in the Early Triassic. A considerable volume of rocks had been eroded during the first phase before the mountains reached their maximum height in the Early Triassic. Hence, it is most likely that the mean height of the mountains did not exceed 1.5 km. This is smaller than the mean height of the Alps (~2 km). Asthenospheric upwelling to the base of the Uralian crust, which was 65–70 km thick in the Late Paleozoic, would result in a formation of mountains, ≥4 km high. The Uralian mountains, however, had only a moderate height. This indicates that in the Late Permian and Early Triassic, the asthenosphere did not reach the base of the crust. Two mechanisms of mountain building of moderate intensity can be supposed. The first is convective replacement of the lower part of mantle lithosphere by a less dense asthenosphere (Artyushkov, 1983). This could result from weakening of the lower lithosphere under infiltration of volatiles from the asthenosphere (Artyushkov and Hofmann, 1997, 1998). According to Eq. (10), the magnitude of the uplift is proportional to a square of thickness of the delaminated layer: Df ~Dh2 . Hence, at uw ml Dh ≤0.3h , the uplift is small. However, it could ml ml be strongly increased if the delaminated volume included large blocks of dense eclogites that had sunk into it at the preceding epochs of collision

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(Artyushkov, 1993). This situation is supposed for the present Alps (Laubscher, 1990; Artyushkov, 1993; Austrheim et al., 1997; Marchant and Stampfli, 1997) and Himalayas (Sapin and Hirn, 1997). Delamination of eclogites of thickness h de and their replacement by the asthenosphere with density r produces the uplift: a Df =[(r −r )/r ]h . (14) de e a a de Taking r =3220 kg m−3, we find that the crustal a uplift by ~1 km in the Southern Urals could be ensured by delamination of h ~10 km of de eclogites. Second, as suggested by Richardson and England (1979), at times of collision, eclogites form from garnet granulites due to a pressure increase in the lower part of a thickened crust. Then, in a course of the following thermal reequilibration and heating of the crust, eclogites retrogress to less dense granulites, which produces the crustal uplift (ibid.; Artyushkov and Baer, 1986b; Dewey et al., 1993; Le Pichon et al., 1997). Thermal relaxation evolves gradually during ~40– 80 m.y. In the Urals, shortening of the crust began in the Devonian, and the shortened regions remained low for up to 120 m.y. Then, a rapid uplift began synchronously in all the regions, which were shortened at different times. Two peaks of mountain building, each several million years long, occurred in the Late Permian and Early Triassic. This sequence of crustal movements could not result from retrogression of eclogites due to a slow thermal relaxation that began at different times in different regions. Retrogression of eclogites is perhaps possible only in the presence of fluid (Heinrich, 1982; Austrheim, 1998). In a dry lower crust, they can be metastable for a very long time. Infiltration of volatiles will catalyse the reaction. The formation of garnet granulites from eclogite in a layer of the initial thickness, h , produces the crustal uplift by: e Dfe =[(r −r )/r ]h . (15) rg e gg gg e Taking r =3550 kg m−3 and r =3250–3300 kg e gg m−3, we find that for an uplift by Dfe ~1 km, rg expansion of eclogites is necessary in a layer of initial thickness h ~11–13 km. Retrogression of e eclogites under infiltration of volatiles from the

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asthenosphere can explain why a rapid uplift took place synchronously all over the Southern Urals, 1300 km long and several hundreds of kilometres wide, and occurred concomitantly in regions that were shortened at different epochs and remained low for long periods of time afterwards. Delamination of the lower part of mantle lithosphere and retrogression of eclogites in the lowermost crust both require infiltration of volatiles. Hence, these mechanisms probably operated synchronously.

13. Discussions 13.1. Formation of foreland basins that was independent of collision in thrust belt According to classic schemes, foreland basins, including the Uralian basin, were formed due to deflection of the elastic lithosphere in response to collision in thrust belts (Quinlan and Beaumont, 1984; McNutt et al., 1988; Zonenshain et al., 1990; Royden, 1993; Stewart and Watts, 1997). The Uralian foreland basin was already up to ~10 km deep when collision began in the Urals ~410 m.y. ago. Since that time, the subsidence by up to 7– 8 km took place in the foredeep; however, it predominantly occurred at times when there was no collision in the thrust belt. Thus, the foredeep is at the right place, but the subsidence took place at the wrong times. Furthermore, at the epochs of major subsidences, the foreland basins deepened away from the thrust belt, which precludes their formation from collision. No significant lithospheric stretching took place on the Uralian foreland, and thermal relaxation could not considerably influence the subsidence of the lithosphere, which was ~1000 m.y. old. Under such circumstances, only a density increase due to metamorphism in mafic rocks in the lower crust could produce the subsidence up to 7–8 km. Most of the tectonic subsidence in the Uralian foredeep and Western flysch basin took place in the form of strong impulses, ~1 m.y. long. This required infiltration of volatiles from the asthenosphere that strongly increased the rate of metamorphism (Artyushkov et al., 1991). The subsidences of this

type also occurred in many intraplate basins, e.g. the West Siberian, Timan–Pechora and PeriCaspian basins (Artyushkov and Baer, 1986a; Artyushkov, 1993).

13.2. Collision without mountain building Strong collision on the continental crust is commonly supposed to result in a synchronous mountain building due to the isostatic response to thickening of the crust (Miyashiro et al., 1982; Zonenshain et al., 1990). In this approach, orogeny is equivalent to mountain building. Several phases of strong collision took place on the continental crust in the Southern Urals in the Late Carboniferous and Early Permian. However, the mean altitude of the shortened regions did not exceed ~0.5 km. The Urals are not unique. The occurrence of collision and mountain building at different stages of evolution of fold belts was first suggested by Stille (1920). He pointed out that collision orogeny results in the formation of only a short-lived relief of a moderate height, while high mountains form later without any strong collision and exist for a long time. A time lapse of 10–100 m.y. between collision and mountain building occurred in the Tien Shan (Schulz, 1948; Sokolov, 1949), Verkhoyansk range (Pushcharovsky, 1959), Caucasus (Milanovsky, 1968; Lukina, 1990), Alps (Artyushkov, 1993), and Carpathians (Artyushkov et al., 1996). This problem is considered in detail in Artyushkov (1993). The formation of only a low topography with a mean height ≤0.5 km at each epoch of collision in the Urals indicates a synchronous density increase in the lithosphere. The epochs of collision were short: from a few millions of years to ~10 m.y. A rapid density increase most likely resulted from metamorphism in the lower crust under infiltration of volatiles from the asthenosphere. This also occurred in the Alps, Carpathians and Tien Shan (Artyushkov, 1993; Artyushkov et al., 1996). Rock contraction is supported by a low present altitude of a thick Uralian crust (~60 km), and the presence of dense rocks with high P-wave velocities in its lower part.

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13.3. High degree of weakening of the lithosphere due to infiltration of volatiles from the asthenosphere and rapid metamorphism in the lower crust In the early Late Carboniferous, a thick nappe was superimposed onto the eastern margin of the East European platform, where the crust was ≥1000 m.y. old. The effective elastic thickness of an old lithosphere is T ≥50 km (Bechtel et al., e 1990; Burov and Diament, 1995; Stewart and Watts, 1997). A similar estimate was obtained for the present Urals ( Kruse and McNutt, 1988; McNutt et al., 1988). However, according to the absence of flexural reaction on the foreland and narrow-wavelength deformations under the nappe, at the time of nappe superposition, T in these e regions was no more than ~5 km. A strong decrease in T is commonly attributed to steep e bending of the lithosphere near to convergent boundaries ( Washbusch and Royden, 1992; Burov and Diament, 1995). However, narrow-wavelength vertical displacements also occurred on the Uralian foreland at times of rapid subsidence when there was no collision in the thrust belt. Similar lithospheric deformations took place at times of rapid subsidence in the intraplate Peri-Caspian basin and in the Transcaspian area far from active plate boundaries (Artyushkov and Mo¨rner, 1997, 1998; Artyushkov et al., submitted for publication). They occurred in the foredeeps of the Caucasus and Carpathians when there was no collision in these thrust belts (ibid.; Artyushkov et al., 1996). Thus, lithospheric weakening can occur independently of plate collision. Its occurrence at times of rapid subsidences probably indicates that the lithosphere was weakened by infiltration of volatiles from the asthenosphere and rapid metamorphism. 13.4. Mountain building without collision Most mechanisms of mountain building imply shortening and thickening of the lithosphere without any significant changes in its density (Molnar and Tapponier, 1975; England and McKenzie, 1983; Fleitout and Froidevaux, 1983; Royden, 1993, 1996; Brown and Beaumont, 1995; Ellis et al., 1995). In the absence of compression in the

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thrust belt, high mountains can be formed by injection of the crustal material into the lower crust from the foreland region ( Zhao and Morgan, 1985; Isacks, 1988). These mechanisms are inapplicable to the Urals where no thrusting occurred during mountain building in the Late Permian and Early Triassic. High mountains can be rapidly formed due to delamination of the mantle lithosphere with asthenospheric upwelling to the crust (Bird, 1979). Together with retrogression and expansion of metamorphosed mafic rocks in the heated lower crust, this would raise the Uralian crust, ≥60 km thick, to ≥4 km. This altitude is typical of the eastern Tien Shan, for example, where the crust, 60 km thick, is underlain by the asthenosphere (Roecker et al., 1993). The Late Uralian mountains were only ~1.5 km high. Hence, the asthenosphere remained far below the Moho. Mountain building can take place in a course of thermal re-equilibration after collision due to retrogression and expansion of eclogites in the lower part of a thick crust underlain by the mantle lithosphere (Richardson and England, 1979). Thermal relaxation evolves during ~40–80 m.y. (Le Pichon et al., 1997). Hence, this mechanism cannot explain a rapid mountain building in the Southern Urals that occurred synchronously in regions shortened at different epochs long before the uplift. The Uralian mountains are more likely to have formed due to infiltration of volatiles into the lithosphere, which triggered two processes. First, the rate of metamorphism in mafic rocks increased by several orders of magnitude all over the thrust belt, thus allowing rapid retrogression of eclogites in the lowermost crust to less dense garnet granulites. Second, a high degree of weakening of the lower part of the mantle root with dense block of eclogites ensured its rapid delamination and replacement by a less dense asthenosphere. 13.5. Lithospheric weakening as a trigger for the formation of thrust belts in the continental crust The cratonic lithosphere on the eastern margin of the East European platform was shortened and incorporated into the Uralian thrust belt. Strong shortening of cratonic lithosphere occurred in sev-

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eral other thrust belts, e.g. the Himalayas and Verkhoyansk range. The integral strength of an old lithosphere is ≥1014 N m (e.g. Dunbar and Sawyer, 1989). The forces acting in continental lithosphere and the forces driving plate motions are much lower: ≤1013 N m (Artyushkov, 1973; Harper, 1986). As a result, under normal conditions, continental cratons are stable, constituting rigid parts of drifting lithospheric plates. To allow extensive shortening of an old lithosphere under moderate forces acting in this layer, it should be weakened considerably. Narrow-wavelength deformations occurred during the rapid crustal subsidence in the Western and Eastern flysch basins of the Southern Urals in the early Late Carboniferous. They indicate lithospheric weakening that preceded the onset of continental collision by ~2 m.y. A considerable shortening of the continental crust in the Alpine and Verkhoyansk belts also occurred only in those basins where rapid subsidence had taken place (Artyushkov and Baer, 1984, 1986c). Basins where only a slow subsidence occurred remained as slightly deformed inner massifs within the shortened areas. In these thrust belts, as in the Urals, the rapid formation of foreland basins with a high degree of weakening of the lithosphere was not a consequence (an egg), but rather a cause (a hen) of formation of thrust belts. This explains why many thrust belts are adjacent to deep foreland basins where rapid crustal subsidence of a large magnitude took place. 13.6. Tectonic evolution of the Southern Urals as a result of infiltration of volatiles from the asthenosphere Rapid crustal subsidence on the Uralian foreland resulted from metamorphism in the lower crust catalysed by infiltration of fluids. Considerable shortening of the continental lithosphere in the thrust belt under convergent plate motions became possible due to weakening of this layer under infiltration of volatiles. Mountain building occurred due to rapid retrogression of dense mafic rocks in the lower crust and/or by delamination of weakened lower part of the mantle root. Both mechanisms required the presence of

fluids. Thus, the major tectonic events in the Southern Urals were caused by the appearance of volatiles in the lithosphere. In contrast to many other thrust belts, no postorogenic collapse occurred in the Urals, thus ensuring the preservation of a thick crust. This shows that after mountain building, the lithosphere became dry and has remained rigid until present. The presence of volatiles in a deep crust is commonly attributed to the dehydration of rocks (Heinrich, 1982; Austrheim, 1998). This is a phenomenon of a local extent that cannot occur synchronously in wide areas. Rapid crustal subsidences, mountain building, and a strong collision in the Southern Urals took place synchronously in regions ≥1000 km long and ≥100 km wide. Furthermore, a high degree of weakening occurred not only in the crust, but also in the mantle lithosphere. These phenomena can be better explained by infiltration of volatiles from small plumes emerging to the lithosphere and rapidly spreading along its base (Artyushkov and Hofmann, 1997, 1998). A synchronicity of rapid crustal subsidence and uplift, as well as of collision, was typical of many other areas. For example, in the Late Jurassic, within the limits of accuracy of paleontological data ~1 m.y., a deep-water basin, ~1000 km wide and ~2000 km long, was concomitantly formed all over West Siberia (Artyushkov and Baer, 1986a). The crustal uplift by ≥1 km took place during the last ~5 m.y. in many cratonic areas, ~1000 km in size, in Africa and north-eastern Asia (Artyushkov and Hofmann, 1997, 1998). Several epochs of collision, 1–10 m.y. long, took place synchronously in the Alpine belt in the Mesozoic and Cenozoic (Schwan, 1980). Thus, infiltration of volatiles from the asthenosphere can be a common cause of mobilization of continental lithosphere, including both strong vertical crustal movements and the formation of thrust belts under convergent motions of lithospheric plates. This is why, in the Southern Urals, collision C 1 began only 2 m.y. after the rapid crustal subsidence in the flysch basins in the early Late Carboniferous, and rapid subsidence in the foredeep in the Late Permian was synchronous to mountain building in thrust belt. It is quite probable that lithospheric

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stretching under divergent plate motions also occurs in regions weakened by infiltration of volatiles from the asthenoshere.

14. Conclusions Vertical crustal movements in thrust belts are commonly explained by horizontal plate motions. According to these ideas, foreland basins form due to deflection of the elastic lithosphere caused by plate collision. The occurrence of collision in thrust belts is predominantly controlled by the distribution of plate motions. High mountains arise as an isostatic response to thickening of the crust from collision. As follows from our analysis, this simple scheme is inapplicable to the Southern Urals. The Uralian foreland basins, past and present, were formed independently of collision in the thrust belt. In these basins, a predominant part of the crustal subsidence occurred from the gabbro–garnet granulites–eclogite transformation in the lower crust. Intensive convergence took place only at times of a high degree of weakening of the lithosphere due to infiltration of volatiles from the asthenosphere and rapid metamorphism in the lower crust. At all the epochs of collision on the continental crust, the crustal surface reached only a low altitude ≤0.5 km above sea level. The magnitude of the crustal uplift was strongly reduced by a synchronous density increase from phase transformations in the shortened regions. The Late Hercynian Uralian mountains of a moderate height were formed long after the preceding phases of collision. Mountain building could occur from infiltration of volatiles from the asthenosphere, which triggered two processes. The first process is rapid retrogression and expansion of dense eclogites in the lowermost crust that formed at the preceding stages of convergence. The second process is delamination of the lower part of mantle lithosphere with dense blocks of eclogite. Thus, as in intraplate areas, the formation of deep basins and mountain building in the Southern Urals is mostly due to deep-seated processes. These are metamorphism in the lower crust and weakening of the lithosphere by infiltration of volatiles from the asthenosphere. However, the Uralian

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mountains were formed only in those regions where the crust had been shortened and thickened at the preceding stages of convergence. Hence, mountain building occurred as a result of interference of deep-seated processes and plate motions. Similar regularities have been established for the East Carpathians (Artyushkov et al., 1996). An analysis of the development of other thrust belts and their foreland basins is necessary to understand whether or not these regularities are of a general character.

Acknowledgements We thank I.A. Basov, B.I. Chuvashov, V.S. Druzhinin, S.N. Kashubin, M. Lindstro¨m, A.A. Mossakovsky, A.S. Perfiljev, V.N. Puchkov, M.S. Rapoport, A.V. Rybalka, S.G. Samygin, A.A. Saveljev and Yu.K. Shchukin for valuable discussions. We are also very thankful to the referees H. Austrheim and J. Touret for their comments which helped to improve the presentation. A large part of this paper was prepared during the stay of E.V. Artyushkov at the Stockholm University in 1994– 1996. The support from the Russian Foundation for Fundamental Research and Peri-Tethys Programme is also acknowledged.

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