the southern urals. decoupled evolution of the thrust belt ... · a layer of dense maﬁc rocks,...

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

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

lithospheric weakening

Eugene V. Artyushkov a,*, Michael A. Baer a, Peter A. Chekhovich b,Nils-Axel Morner c

a Institute of Physics of the Earth, Russian Academy of Sciences, B. Gruzinskaya 10, 123810, Moscow, Russiab Institute of the Lithosphere of Marginal Seas, Russian Academy of Sciences, 109180, Moscow, Russia

c Stockholm University, Institute of Paleogeophysics and Geodynamics, Kraferiket 24, S-10691, Stockholm, Sweden

Abstract

An analysis is presented of the mechanisms of tectonic evolution of the southern part of the Urals between 48Nand 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 shelfon the margin of the East European platform in the early Late Carboniferous. It is commonly supposed that strongshortening 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 collisionalevents at the end of the Late Carboniferous, in the second half of the Early Permian, and at the start of the MiddleTriassic. A low magnitude of the crustal uplift in the regions of collision indicates a synchronous density increasefrom rapid metamorphism in mafic rocks in the lower crust. This required infiltration of volatiles from theasthenosphere 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 formedin the Late Permian and Early Triassic when there was no collision. Their moderate height precluded asthenosphericupwelling to the base of the crust, which at that time was ~65–70 km thick. The mountains could be formed due todelamination of the lower part of mantle root with blocks of dense eclogite and/or retrogression in a presence offluids 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 andsubsurface loads in thrust belts. Most of tectonic subsidence on the Uralian foreland occurred in a form of shortimpulses, 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 subsidencemechanism. To ensure the subsidence, a rapid density increase was necessary. It took place due to metamorphism inthe lower crust under infiltration of volatiles.

The absence of flexural reaction on the Uralian foreland on collision in thrust belt together with narrow-wavelengthbasement deformations under the nappe indicate a high degree of weakening of the lithosphere. Such deformationstook 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: S0040-1951 ( 00 ) 00044-5

272 E.V. Artyushkov et al. / Tectonophysics 320 (2000) 271–310

Nomenclature 1c volume per cent of gabbroic intrusionsd thickness of lithospherehc thickness of crusth0c initial thickness of crusth1c thickness of crust after collision or stretchingher thickness of eroded rockshde thickness of delaminated eclogiteshe thickness of eclogiteshgi total thickness of gabbroic intrusionshgg thickness of garnet granuliteshml thickness of mantle lithospherehn thickness of nappehor thickness of sediments overridden by nappehpg thickness of pyroxene granuliteshs depth of sedimentary basinhths depth of sedimentary basin formed by thermal relaxation of lithospherehw depth of waterL0 minimum width of deflection of elastic lithosphericP pressureT temperatureTa temperature of asthenosphere (1300°C )Te effective elastic thickness of lithosphereTM temperature at MohoVP P-wave velocityVgbP P-wave velocity in gabbroVprP P-wave velocity in peridotitesa thermal expansivity (3×10−5 K−1)b intensity of stretchingDhc increase in thickness of crust due to collisionDhlc thinning of lower crust by stretchingDhml decrease in thickness of mantle lithosphere due to asthenospheric upwellingDhs thickness of sediments pushed out by nappeDf uplift of shortened crustDfde uplift of crust due to delamination of eclogitesDfgi uplift of crust due to gabbroic intrusionsDferg uplift of crust due to retrogression of eclogites to garnet granulitesDfggrg uplift of crust due to retrogression of garnet granulites to pyroxene granulitesDfuw uplift of crust due to asthenospheric upwellinge intensity of compaction of sediments overridden by nappef altitude of crustal surfacef0n minimum altitude of nappe in a case of no density changes in lithospherej tectonic subsidence in shortened region due to density increase in lithospherera density of asthenosphere (3220 kg m−3)rc density of crust (2830 kg m−3)re density of eclogiterer density of eroded rocksrgb density of gabbro (2930 kg m−3)rgg density of garnet granulitesrlc density of lower crustrm density of mantle (3350 kg m−3)rn density of napperpg density of pyroxene granulitesrs density of sedimentsrw density of water (1030 kg m−3)1 In the nomenclature, symbols without assigned values have been assumed to be variable.

273E.V. Artyushkov et al. / Tectonophysics 320 (2000) 271–310

Weakening of the lithosphere can be explained by infiltration of volatiles into this layer from the asthenosphere andrapid metamorphism in the mafic lower crust. Lithospheric weakening allowed the formation of the Uralian thrustbelt under convergent motions of the plates which were separated by weak areas. © 2000 Elsevier Science B.V. Allrights reserved.

Keywords: collision; crustal subsidence; eclogitization; lithospheric weakening; mountain building; Urals

1. Introduction question as to whether or not vertical crustalmovements of this type also occur in fold belts.

It has been shown that the Neogene foredeep ofThe formation of foreland basins is commonlythe East Carpathians was formed without signifi-considered as a result of plate collision and flexingcant lithospheric stretching and at the epochs whenof the elastic lithosphere towards convergent plateno collision took place in the adjacent thrust beltboundaries without strong density changes in the(Artyushkov et al., 1996). At times of strong colli-lithosphere (Quinlan and Beaumont, 1984;sion, the crustal surface in the thrust belt remainedMalinverno and Ryan, 1986; Royden, 1993;at a low altitude. The Carpathian mountains beganStewart and Watts, 1997). Mountain building into grow 8 m.y. after the end of the collision. Thesethrust belts is explained by a synchronous isostaticvertical crustal movements required density changesresponse to thickening of the crust from platein the lithosphere. No flexural reaction occurred oncollision (Molnar and Tapponier, 1975; Miyashirothe Carpathian foreland at the epochs of collision.et al., 1982; Zonenshain et al., 1990) with itsThis indicates a drastic weakening of the lithospherepossible lowering by slab pull (Royden, 1993). Inthat ensured a strong crustal subsidence under thethis scheme, vertical crustal movements in foldnappe without a synchronous subsidence on thebelts result from horizontal plate motions. It isadjacent foreland.also commonly believed that the formation of

In this paper, we analyse the tectonic develop-thrust belts is associated with no significantment of the southern part of the Urals after closurechanges in the density of the lithosphere and theof the Uralian ocean — since the start of thethickness of its elastic part Te, which can beCarboniferous. This area, 1300 km long, includesstrongly reduced only due to steep bending orthe Southern and Middle Urals (Fig. 1). For the

strong heating of this layer in some places (Kusznirsake of simplicity, it will be called ‘the Southern

et al., 1991; Burov and Diament, 1995). Urals’. In the preceding publications, attention hasThe crustal subsidence and uplift, however, been focused on a description and timing of colli-

widely occurred in plate interiors without any sional events in the area (e.g. Ruzhentsev, 1976;significant lithospheric stretching and far from Zonenshain et al., 1984; Ivanov et al., 1986; Brownconvergent boundaries. They have taken place, for et al., 1997; Puchkov, 1997, 1999). Here, we con-example, during the subsidence in the West sider another problem: what vertical crustal move-Siberian, Peri-Caspian, Volga–Urals, Timan– ments took place near to collisional boundaries inPechora and Vilyuy basins (Artyushkov and Baer, the Southern Urals, and could these movements1986a; Artyushkov, 1993). In the Neogene, many result from plate motions, or were they caused bymountain ranges and plateaus were formed by a deep seated processes? The main questions to bestrong crustal uplift without significant compres- answered are:sive deformations in East Siberia, north-eastern 1. What were the modes of crustal subsidence onAsia and Africa (Nikolaew and Neumark, 1977; the Uralian foreland and their role in the basinPartridge and Maud, 1987; Makarov, 1990; Ollier, formation?1991; Summerfield, 1991). The crustal uplift and 2. Were the major collisional events in thrust beltsubsidence in plate interiors require density associated with strong synchronous crustal sub-

sidence on its foreland?changes in the crust and/or mantle. This poses the


5. Did a density increase in the lithosphere playan important role in the subsidence?

6. Was collision associated with a synchronousformation of mountain ranges?

7. Did considerable density changes occur in thelithosphere at times of collision?

8. Was mountain building associated with a strongsynchronous shortening of the crust?

9. What was the role of density changes in thelithosphere in mountain building?

We first consider the geological data necessary toobtain answers to these questions and methods ofanalysis of the data. Then, we formulate theanswers to questions 1–9. On this basis, we deter-mine possible mechanisms of the crustal subsidenceon the Uralian foreland and of mountain buildingin thrust belt, and consider lithospheric rheologyin these regions.

2. Tectonic environment and crustal structure at theend of Early Carboniferous

At that time, the Southern Urals included thecontinental margin of the East European continent(Baltica), the East Uralian microcontinent(EUM ), and the undeformed eastern part of the

Fig. 1. Structure of the Southern Urals (compiled using the data Magnitogorsk arc (MA) between them (Fig. 2a).by Peive and Yanshin, 1979; Puchkov, 1993; Yazeva and The Kazakhstan oceanic basin probably stillBochkarev, 1993). 1=the Uralian foredeep with its main existed to the east of EUM (Didenko et al., 1994;depressions in the Southern Urals: A — Aktyubinsk, B —

Puchkov, 1999; Puchkov et al., 1999). Its closureBel’sk, YS — Yuryusan’–Solikamsk; 2=Zilair–Lemva zone:was associated with subduction under EUM thatstrongly shortened deposits of the European continental slope

of the Uralian ocean and the adjacent eastern margin of the resulted in island-arc volcanism in this region.East European platform; 3=outcrops of the Precambrian crys- Subduction could synchronously occur under thetalline basement; 4–6=ensimatic island arcs and ophiolitic Kazakhstan continent. This continent collided withblocks of the Uralian (UO-4) and Kazakhstan ( KO-5) oceanic

EUM most likely at the beginning of the Latebasins that are separated by the East Uralian microcontinentCarboniferous — in the middle of the(EUM-6); 7=eastern and south-western boundaries of the area

where the Uralian thrust belt is exposed to the surface; 8= Serpukhovian age.boundaries of the main tectonic units exposed to the surface (a) In the Early Carboniferous, the Sakmara–and overlain by the Mesozoic–Cenozoic cover (b); 9=Main Magnitogorsk nappe (SMN), up to ~15 km thick,Ophiolitic Suture of the Urals.

was lying on the European continental slope(Fig. 2a). The nappe, the adjacent eastern marginof the East European platform and theMagnitogorsk arc were overlain by several-kilome-3. Did the lithosphere in thrust belt and on its

foreland preserve a high effective elastic tre-thick deposits of the Late Devonian–EarlyCarboniferous Zilair basin. Since the middle of thethickness?

4. Was lithospheric stretching responsible for a Early Carboniferous and until the end of theSerpukhovian, during a period of time, ~30 m.y.considerable part of the subsidence?


Fig. 2. Evolution of the Southern Urals in the late Early Carboniferous and early Late Carboniferous (profiles of Figs. 2, 8 and 13are 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; Kazantsevet 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 lateEarly Carboniferous ( late Visean). (b) Closure of the Kazakhstan oceanic basin in the middle of the Serpukhovian, rapid formationof the Western and Eastern flysch basins at the beginning of the Late Carboniferous near to the transition from the Serpukhovianto the Bashkirian. (c) Superposition of the Sakmara–Magnitogorsk nappe onto the eastern margin of the East European continentin 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 tobe visible in this scale.


long, all the Southern Urals were situated near to foreland basins and under a superimposed nappeas another indication of lithospheric weakening.sea level and characterized by a low tectonic

activity. Deposition of shallow-water carbonates All these data will contribute to an answer toquestion 3.took place in this area with some sands and coals

in the east (Chuvashov et al., 1984; Chuvashov Extensive lithospheric stretching is associatedwith normal faulting in the basement and theand Puchkov, 1990). This indicates a high crustal

thickness all over the Urals, which is typical of sedimentary cover. A continuity of sedimentarybeds in the present foredeep will indicate thecontinental areas.absence of significant stretching. The method forrevealing a presence or absence of stretching in ashortened sedimentary cover of past basins is3. Method of analysisdescribed in Artyushkov and Baer (1983). In thecase of rapid subsidence with a formation of deep-Two basic modes of crustal subsidence in sedi-

mentary basins have been identified in our preced- water basin on a shallow-water shelf, a conform-able position (parallelism) of shallow water anding studies (e.g. Artyushkov and Baer, 1986a;

Artyushkov et al., 1991). This is a slow sediment deep-water strata indicates the absence of blockrotation and lithospheric stretching. Using theseloaded subsidence at a rate of 10–100 m m.y.−1

with a long duration of ≥200–300 m.y. and rapid approaches, we will find an answer to question 4.Those subsidences that occurred without lithosphe-subsidence at a rate of ~1 km m.y.−1, which, in

many cases, resulted in the formation of deep- ric stretching and at times when there was nocollision required a density increase in the litho-water basins during ~1 m.y. Using the strati-

graphic records for the Uralian foreland basins, sphere as a mechanism (question 5).An analysis of the character of correlativewe will determine the epochs of occurrence of

these two types of subsidence and estimate their deposits in sedimentary basins is a classic methodfor estimating the altitude of the adjacent landinput into the formation of the basins. This will

give an answer to question 1 in Section 1. To (Penck, 1924; Meshcheryakov, 1965; King, 1967;Khain, 1977; Timofeev, 1979; Makarov, 1980; seeanswer question 2, we will compare the epochs of

strong and rapid crustal subsidence in the foreland also Artyushkov et al., 1996 for more detail ).Denudation of high mountains produces kilome-basins with the major events of collision in thrust

belt. tres of coarse molasse with a large volume ofconglomerates. Erosion of a low and smoothThe effective elastic thickness of the lithosphere

Te is related to the characteristic width of lith- topography, ≤0.3–0.5 km high, results in depos-ition of sands, clays and turbidites in the adjacentospheric deflection L0 as (Turcotte and Schubert,

1982; Artyushkov et al., 1996): basins. In this way, we will determine whether ornot collisional events in the Southern Urals were

(Te)km~0.05[(L0)km ]4/3. (1)

associated with a synchronous formation of highmountains (question 6).According to this equation, the crustal subsidence

in a foreland basin, ≥100 km wide, at a time of After estimating the altitude of the shortenedregions, we can consider possible density changescollision indicates deflection of the lithosphere with

a high Te under surface and subsurface loads in in the lithosphere at the epochs of collision (ques-tion 7) in the following way. Suppose that a nappethrust belt. An additional requirement is basin

deepening towards the thrust front. The absence with the density, rn, and thickness, hn, was super-imposed onto a region with the initial depth ofof strong subsidence in the foreland region at a

time of collision, will show a high degree of water, hw. Assume that, on its way, the nappepushed out a layer of sediments with the density,weakening of the lithosphere and decoupling of

the thrust belt from its foreland. We will also pay rs, and thickness, Dhs, and overrode sedimentswith thickness, hor. The minimum mean altitudeattention to steep bending of the lithosphere in

regions, several tens of kilometres wide, in the of the nappe surface (f0n) will be reached in a state


of local isostasy. If nappe superposition resulted Bashkirian ages of the Late Carboniferous(Figs. 2b–4a). The western basin, ~30–50 kmin no density changes in the lithosphere below the

sedimentary cover, wide with a water depth of 0.5–1 km, was formedon the western margin of the Zilair basin and the

f0n=[(rm−rn)/rm ]hn−[(rm−rs)/rm ]Dhs adjacent eastern margin of the East Europeanplatform ( Khvorova, 1961; Chuvashov, 1975;−[(rm−rw)/rm ]hw−ehor . (2)Melamud, 1981; Misens, 1995b). In the deepest

Here, rm=3350 kg m−3 and rw=1030 kg m−3 are outer (western) part of this basin, Serpukhovianthe densities of the mantle and water, respectively, shallow-water limestones (C2s legend 3 in columnand e is the intensity of compaction of the overrid- IV in Fig. 5) were abruptly overlain by Bashkirianden deposits. If altitude (2) were considerably turbidites (C2b legend 7).larger than that estimated from the character of At the same time another basin, ~400 km wide,deposition in the adjacent basin, a density increase was formed in the east, on the Magnitogorsk arcwould be necessary to lower the lithospheric and on the EUM (middle part of Fig. 2b)surface. (Chuvashov et al., 1984). Serpukhovian shallow-

Continental collision can also result in shorten- water limestones were abruptly overlain bying of the lithosphere, which is uniform over the Bashkirian highly organic pelagic limestones.depth. Designate by rc the density of the crust and During the following ~8–10 m.y. in theby Dhc the increase in its thickness after collision. Bashkirian and Moscovian ages, the deepest, west-Suppose that before collision the crustal surface ern part of the Eastern flysch basin was filled withwas at or above sea level. Then, in a state of ~2 km of pelagic limestones and turbidites.isostasy, the surface of shortened crust would be Taking into account the isostatic subsidence underraised by: their load, which increases the basin depth by

about 2.5 times, the initial water depth wasDf=(rm−rc)/rmDhc . (3)~800 m.

If the height of the shortened region were consider- The deposits of the Western flysch basin areably smaller than Eq. (3), a density increase would now incorporated into the frontal part of thebe necessary to maintain the crustal surface at a Uralian thrust belt. The strata of Bashkirian deep-low altitude. water deposits are conformable to Serpukhovian

To answer question 8, we will compare the shallow-water carbonates at the transition, whichepochs of mountain building in the Urals with is only 1–5 m thick (Fig. 6). This indicates depos-those of collision. Mountain building synchronous ition on an almost flat basin floor and precludesto collision can indicate thickening of the crust lithospheric stretching and block rotation at thewith the isostatic rebound as a mechanism. The epoch of rapid subsidence.formation of high mountains at the epochs of nostrong collision requires a density decrease in the 4.2. Superposition of the Sakmara–Magnitogorsklithosphere (question 9). nappe onto the European shallow-water shelf in the

Bashkirian–Early Moscovian

At the time of rapid subsidence between the4. Crustal subsidence and collision in the early LateCarboniferous Serpukhovian and Bashkirian, the Eastern and

Western flysch basins were separated by a shallow-water shelf, ~200 km wide (Fig. 2b). About 2 m.y.4.1. Rapid formation of flysch basins between the

Serpukhovian and Bashkirian later, conglomerates and olistoliths appeared onthe western margin of the Eastern flysch basin(Chuvashov et al., 1984). They were derived fromTwo deep-water basins were formed in the

Southern Urals during ~1 m.y. by rapid crustal the island chain, ~50 km wide, which began togrow in the rear part of the Sakmara–subsidence between the Serpukhovian and


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


Fig. 4. Schematic profiles of western and central parts of the Southern Urals. (a) In the early Late Carboniferous after the formationof 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 theCarboniferous; shortening of the Bashkir shelf, filling with deposits of the Western flysch basin and the inner part of the late LateCarboniferous Uralian foredeep.

Magnitogorsk nappe above the lower part of the et al., 1984). At that time, the Magnitogorsk arcbegan to move further westwards, overriding theEuropean continental slope. The crustal uplift was

associated with the onset of collision in the middle eastern buried slope of the East European conti-nent and pushing the Sakmara–Magnitogorskof the Bashkirian (Ruzhentsev, 1976; Chuvashov

approximately equals an increase in the surface load DP in the shortened region (the method of calculating DP is described inArtyushkov et al., 1996). Changes in the altitude of the crustal surface in regions of collision are schematically shown by a grey linenear the horizontal axis. The subsidence curves below the horizontal axis are labelled according to the place of occurrence. They areplotted using the corresponding stratigraphic columns of Fig. 5, with the curve numbers corresponding to the numbers of columnsin Fig. 5. The method of compilation of the curves is described in Section 10.1. Two rapid subsidences most likely occurred in thewestern 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).


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 ofrapid water loaded (uncompensated) crustal subsidence and the onset of rapid sediment loaded (compensated) subsidence are shownby arrows pointing downwards ( legends 13 and 14, respectively). The onsets of collisional events in the adjacent western part ofthrust 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, D3fm — Famennian; C — Carboniferous, C1t — Tournaisian, C1v —Visean, C2s — Serpukhovian, C2b -Bashkirian; C2m — Moscovian, C2k — Kasimovian, C2g — Gzhelian; P — Permian, P1a —Asselian, P1s — Sakmarian, P1ar — Artinskian, P1k — Kungurian, P2u — Ufimian, P2kz — Kazanian, P2t — Tatarian; T — Triassic,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 the middle of the Moscovian age, the nappe over-rode the upper part of the European continentalin Figs. 3 and 5). At the same time, deposition of

conglomerates derived from the frontal part of the slope, ~40 km wide, and the adjacent shelf in theBel’skaya zone, 30–40 km wide, the total magni-nappe began on the eastern slope of the Western

flysch basin ( legend 1 C2b2 in column III in Fig. 5) tude of superposition being ~70–80 km (Figs. 2cand 4a). After the Serpukhovian, no more deposits(Khvorova, 1961).

During the following period of ~7 m.y., until were formed in the above two regions. The nappe


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

did not reach the eastern margin of the Western 4.3. Narrow-wavelength deformations of thelithosphereflysch basin (Fig. 2c). An undeformed band of the

Bashkir shelf still remained between the basin andIt follows from the thickness of the depositsthe nappe front and was ~20–30 km wide in the

that the Eastern flysch basin had a steep westernsouth (Fig. 4a), widening to ~100–150 km in theslope and a smooth eastern slope (Figs. 2c and 4a)north. Shallow-water deposition proceeded with(Chuvashov et al., 1984). The basin axis with upsome interruptions in this region during theto 1.5–2 km of Bashkirian–Moscovian conglomer-following ~50 m.y. — until the end of the Earlyates was located ~30 km to the east of the westernPermian (carbonates 3, sands 2, conglomerates 1

and evaporites 10 of C2b-P1k in columns I, II of basin margin. Large olistoliths slid down the basinslope into this region from the Sakmara–Fig. 5) (Khvorova, 1961; Chuvashov, 1975;

Unified Stratigraphic and Correlational Schemes Magnitogorsk nappe. The Western flysch basinwas only 30–50 km wide (Fig. 4a). Its westernof the Urals, 1980).


slope was very steep ( Khvorova, 1961) and could shallow-water shelf with a depth of waterhw~0.1 km. On its way, the nappe pushed outinclude a boundary scarp near to the platform

margin. In the deepest part of the basin, fine- Dhs~1–1.5 km of sediments with the densityrs#2550 kg m−3 and piled them up at its front.grained and deep-water turbidites include layers

of large olistoliths, up to ~100 km thick, that The nappe overrode hor~2–3 km of sediments.After erosion of 2–5 km of the deposits from thewere derived from the Bashkir platform (C2b and

C2m in column IV ) (Misens, 1995a,b). Deposition nappe, it still preserved a large thickness,hn~15–18 km (Fig. 4a). Substituting these valuesof shallow-water sands and conglomerates took

place on the eastern basin slope (C2b-m in column into Eq. (2) with the nappe densityrn=2700 kg m−3 and the intensity of compactionIII ). Thus, the eastern and western flysch basins

included steep slopes, which were only of the overridden deposits e~0.2, we obtain:f0n~1.8–2.5 km. Since the altitude of the nappe atL0~20–30 km wide. According to Eq. (1), this

shows that during the formation of the basins, Te the end of collision C 1 was no more than~0.5 km, a tectonic (unloaded) crustal subsidencedecreased to ~3–5 km, i.e. the elastic layer in the

lithosphere practically disappeared. Lithospheric by j~1.3–2 km should have occurred due to theappearance of subsurface loads at the time ofweakening also follows from the absence of crustal

subsidence at the front of the nappe where a nappe superposition.shallow-water shelf came into existence (Figs. 2cand 4a), and from a steep basement bending underthe nappe that reached ~10–15 km at a distance 5. Superposition of the eastern part of the

Magnitogorsk arc onto the western slope of theL0~20–30 km from its front (Fig. 4a).East Uralian microcontinent in the lateLate Carboniferous4.4. Low altitude of the nappe surface

In the process of superposition, the Sakmara– At the end of convergence C 1, in the middleof the Moscovian age, the eastern (buried) edgeMagnitogorsk nappe, ~100–150 km wide,

emerged above sea level (Figs. 2c and 4a). The of the East European continent and the westernedge of the East Uralian microcontinent (EUM )volume of the middle Bashkirian to middle

Moscovian coarse clastic deposits is ~5 times were still ~100 km apart with the undeformedpart of the Magnitogorsk arc between themlarger in the Eastern flysch basin than in the

Western flysch basin. This shows that a higher (Fig. 2c). In the present structure, the eastern partof the arc overlies the western slope of EUMtopography existed in the rear eastern part of the

nappe (Fig. 4a). About 3–5 km of rocks have been (Fig. 7). Shallow-water deposition took place onthis part of the arc until the end of the Moscovianeroded from this region, 50–70 km wide, the level

of erosion reaching volcano-sedimentary rocks and (Chuvashov et al., 1984); however, no Kasimoviandeposits have been found in the region. This mostophiolites of the Sakmara–Magnitogorsk nappe.

The eroded surface did not reach these rocks in likely indicates that superposition of the arc ontothe microcontinent slope occurred around thethe frontal part of the nappe, which was at a lower

altitude. About 2 km of rocks have been eroded transition from the Moscovian to the Kasimovianevent of collision C 2 (Fig. 8a) (Seliverstov andfrom this region.

Deposition of conglomerates and turbidites in Denisov, 1982). During ~4 m.y., the arc overrodethe western slope of the microcontinent for ~50–the Eastern flysch basin had ceased by the late

Moscovian when shallow-water carbonates and 70 km in the south and ~100 km in the north(Yazeva and Bochkarev, 1993). The arc wasevaporites began to form. This indicates planation

of the topography at the end of collisional event broken into tectonic slices and shortened by abouta factor of two. On its way, this nappe (EastC 1 and a low altitude of the nappe surface above

sea level (most probably, ≤0.5 km). Magnitogorsk nappe) pushed out Dhs=2–3 km ofthe deposits and incorporated them into its frontalThe Sakmara–Magnitogorsk nappe overrode a


Fig. 7. Tentative scheme of the crustal structure along the URSEIS’95 profile; see Fig. 1 for location (modified after Skrypyi andYunusov, 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 theMagnitogorsk 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 basinin the present Uralian foredeep in the late Gzhelian. (b) Shortening of the Sakmara–Magnitogorsk nappe and the Bashkir shelf inthe late Early Permian collision C 4; collision in the central and eastern parts of EUM (event C 5). (c) Planation of the topographyin the Uralian thrust belt by the end of the Early Permian.


part. Dher=5–7 km of rocks have been eroded 6. Evolution of the Uralian foredeep in the LatePrecambrian–Carboniferousfrom the nappe surface later — in the Permian

and Triassic (Sigov and Romashova, 1984). Thepresent thickness of the nappe h1n reaches the 6.1. Basin structure and deposition in the Late

Proterozoicmaximum value above the western edge of EUMwhere h1n=hmaxn ~27 km (Fig. 7). After collisionC 2, an increase in the crustal thickness due The thrust belt of the Southern Urals is bounded

by the Uralian foredeep, 40–130 km wide (Fig. 1).to convergence was reached in this regionDhc=h1n+Dher−Dhs~30 km. The depth to the crystalline basement in this basin,

on the eastern margin of the East European plat-The East Magnitogorsk nappe was superim-posed onto a shallow-water shelf. Without density form and in the Peri-Caspian basin is shown in

Fig. 9. Over most of the Southern Urals, thechanges in the lithosphere, the surface of thenappe, up to ~30 km thick, would reach a very foredeep is up to 12–15 km deep. The age of the

crystalline basement varies over the basin. In mosthigh altitude. No coarse clastic material has beenbrought from the East Magnitogorsk nappe into cases, this boundary coincides with the base of the

Riphean (Fig. 10). In some regions, the basement,the Western flysch basin in the Kasimovian andearly Gzhelian (C2k+g1 in column III in Fig. 5). shown in Fig. 9, refers to the top of metamor-

phosed deposits of the Lower Riphean (Fig. 11).Fine-grained turbidites with minor carbonateswere predominantly deposited on its eastern slope. The thickness of the Phanerozoic deposits varies

along the foredeep, and the deposits are consider-Only fine-grained sediments and carbonates wereformed in shallow marine basins that bounded ably thicker in the south. Thus, in the profile of

Fig. 10, the thickness of the Paleozoic and Triassicthe Urals in the south (Chuvashov et al., 1984).This shows that after collision C 2, the East reaches 10 km, while in the profile of Fig. 11, the

Phanerozoic deposits are only 2–3 km thick. WeMagnitogorsk nappe was hardly higher than~0.5 km. Hence, a density increase occurred in will describe the evolution of the foredeep on the

example of the southern part of Bel’sk depression,the lithosphere below the nappe synchronously tosuperposition. using the profile of Fig. 10 and the corresponding

stratigraphic columns V–VII of Fig. 5. Their posi-In the early Gzhelian, a layer of large blocks,olistoliths, 100–150 m thick, was formed on the tion is shown above the profile. Upper Permian

and Triassic deposits in the foredeep are stronglyeastern slope of the Western flysch basin. Thislayer was composed of Visean and Moscovian deformed by salt-dome tectonics. Their average

thicknesses in columns V–VII are given accordinglimestones derived from the Bashkir shelf ( legend6, C2g1 in columns III, IV, Fig. 5) (Misens, 1995b). to Unified Stratigraphic and Correlational

Schemes of the Urals (1980). In the west, theAt the same time conglomerates, up to ~100 mthick, were formed on the eastern basin slope in foredeep is bounded by the East European plat-

form. Column VIII refers to the eastern platformthe northern part of the Southern Urals(Chuvashov, 1975). These data indicate a short margin. In the east, the foredeep is bounded by

the nappe, which includes shortened deposits ofimpulse of convergence C 3 on the Bashkir shelfwith erosion of several hundred metres of carbon- the Western flysch basin (right-hand side of

Fig. 10).ates (see the hiatus between C2k and C2g2 incolumn I and between C2m and C2g2 in column Slow deposition took place on the eastern

margin of the East European continent in theII ). According to the small amount of erosion onthe Bashkir shelf and small volume of synchronous Riphean (R) and Vendian (V ) in the Late

Proterozoic. In the Uralian thrust belt, deposits ofcoarse deposits in the adjacent flysch basin, colli-sion was weak, and the altitude of eroded topogra- this age are exposed to the surface in many places

(Figs. 1 and 11). The thickness of R+V variesphy was low. The total shortening of the regionduring this period of time could not exceed along the foredeep and increases towards the thrust

belt (Figs. 10 and 11). In the profile of Fig. 11, at~10 km.


of superposition of the Phanerozoic Uralian nappeover the Proterozoic sedimentary basin, whosedepth increased eastwards.

6.2. Slow subsidence in the Devonian–Carboniferous and the crustal uplift in the middle ofthe Gzhelian age in the latest Carboniferous

Almost no deposition took place in the Uralianforedeep and on the eastern platform margin inthe Cambrian-Silurian (see the hiatus between Vand S, D in columns V–VIII ). A slow crustalsubsidence resulting in deposition of 1–2 km ofshallow-water limestones with minor sandsoccurred under cratonic conditions in theDevonian–Carboniferous. Near to the end of theLate Carboniferous, in the middle of the Gzhelian,a short regression occurred for ~1 m.y. on ashallow-water shelf in the Uralian foredeep to thewest of the Western flysch basin (Melamud, 1981).In a band, ~50–100 km wide, this resulted inerosion and hiatus increasing eastwards towardsthe Urals. Up to 500 m of limestones had beeneroded in the eastern part of the foredeep (columnV in Fig. 5). The lower Gzhelian, Kasimovian andupper part of the Moscovian are missing. Fiftykilometres further to the west, in columns VI andVII, only 50–100 m of the deposits had beenwashed out. No regression or erosion took placeon the present platform margin (column VIII ).Strong lateral non-uniformity of erosion precludesa eustatic origin of regression and indicates that itresulted from a tectonic crustal uplift. This upliftcannot be recognized in the western part of theflysch basin where deep-water deposition pro-ceeded until the very end of the Carboniferous(C2g1−C2g2 in column IV ).

Fig. 9. Depth to the crystalline basement in the foredeep of theSouthern Urals and on the adjacent East European platform 6.3. Rapid crustal subsidence in the late Gzhelianand Peri-Caspian basin (km) (modified after Semenovichet al., 1982).

The crustal uplift in the middle of the Gzhelianwas followed by rapid subsidence in the lateGzhelian. A deep-water basin, 40–100 km widethe nappe front, these deposits are 12–13 km thick,

and maybe even 15–17 km thick, if layer R1 (?) is and up to ~1 km deep, was formed on the easternplatform margin to the west of the flysch basinincluded. Similar thicknesses of R+V can be seen

in the thrust belt in column II in Fig. 5. Hence during ~1 m.y. (Figs. 4b and 8a) ( Khvorova,1961; Chuvashov, 1975; Melamud, 1981;variations in the thickness of R+V deposits at the

nappe front can mostly reflect a variable magnitude Chuvashov et al., 1990). This was the initial fore-


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

deep of the Southern Urals. In its deepest eastern Gzhelian (conglomerates 1, sands 2 and limestones3 of C2g2 in columns I, II ) and on the easternand middle parts, eroded shallow-water limestones

were abruptly overlain by deep-water marls, black slope of the flysch basin at the beginning of theEarly Permian (P1a — Asselian in column III ).limestones and argillites ( legend 8 in columns V

and VI in Fig. 5; see also Figs. 10 and 12). On the After the rapid subsidence in the Gzhelian, acontinuity has been preserved of the Upperouter western foredeep margin, the subsidence was

gradual. This basin part was separated by a normal Carboniferous shallow-water carbonates in theforedeep (9 in Fig. 12), which underlie the upperfault from the adjacent East European platform.

During the period of time, ~30 m.y. long, until Gzhelian deep-water marls and argillites (8).Hence, the subsidence occurred without significantthe Artinskian age of the Early Permian, 0.7–

1.5 km high reefs were formed in this region lithospheric stretching. In Fig. 11, except at someminor thrust faults, Riphean and Vendian strata(C2g−P1ar in column VII ) (Geology of the USSR,

v. XIII, 1964; Melamud, 1981). A marginal reef are continuous in the Uralian foredeep, 15–20 kmdeep, which also indicates no stretching. In thecan also be seen in the left-hand side of profile

of Fig. 12. profile of Fig. 10, a continuity of the strata inthe foredeep can be reliably traced at the base ofRapid subsidence probably also occurred in the

deepest western part of the flysch basin. It cannot the Devonian. Several minor thrust faults exist inits inner part. The only normal fault occursbe recognized in this region, however, since a deep-

water basin had already existed there before the between the foredeep and platform margin in thewest. This ensures a negligible mean lithosphericGzhelian subsidence. No intensive subsidence

occurred further to the east — on the eastern slope extension in the foredeep.At the end of the Carboniferous, collision C 2of the Western flysch basin and on the Bashkir

shelf. After a short hiatus, shallow-water depos- took place only in the eastern Urals, far from theforedeep (Fig. 8a). No strong subsidence occurredition resumed on the Bashkir shelf in the late


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, P1s — Sakmarian, P1k — Kungurian, P2u — Ufimain, P2k —Kazanian, P2t — Tatarian.

Fig. 12. Upper part of the sedimentary cover on the outer margin of the Uralian foredeep; see map in Fig. 6 for location (modifiedafter Kamaletdinov, 1974). The transition from Upper Carboniferous shallow-water limestones (9) to upper Gzhelian deep-watermarls, 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.


in a wide region between the foredeep and the indications of intensive erosion all over westernUrals. Folds and cordilleras, ≥1 km high, werearea of collision. The Bashkir shelf to the east of

the foredeep remained near sea level or slightly probably formed in this region during the colli-sional phases. Piles of conglomerates are, however,above it (C2g2 and P1a in columns I, II of Fig. 5).

These features show that the rapid crustal subsi- separated by layers of sands, carbonates and evap-orites formed during the epochs of stability 5–dence in the foredeep in the late Gzhelian did not

result from collision. In the absence of stretching, 6 m.y. long between the collisional phases. Theycan be seen, for example, in the interval P1s−P1kthis could be caused only by a density increase in

the lithosphere. in columns I–III. This shows that a rugged topog-raphy formed at the phases of collision becamerapidly eroded, and between them, the nappe was≤0.5 km above sea level.7. Evolution of the Uralian foredeep and thrust belt

Near to the end of the Artinskian, the inner-in the Early Permianeastern part of the Uralian foredeep became filledwith ~3.5 km of turbidites with minor pelagic7.1. Collision in western Urals in the late Earlymarls and limestones (column V ). The top of thePermian and deposition in the foreland basinsArtinskian (P1ar) in this column is represented byshallow-water limestones. The central part of theThree collisional phases took place in the secondforedeep (column VI) remained at deep water untilhalf of the Early Permian. They occurred in thethe Kungurian age of the Early Permian. Onlymiddle of the Sakmarian (C 4a), Artinskian (C150–200 m of black pelagic limestones and marls4b) and Kungurian (C 4c) ages (Chuvashov, 1975;were formed in this region during ~30 m.y. afterChuvashov and Misens, 1980; Misens, 1995a,b).the rapid subsidence in the Gzhelian. The heightThese phases were each 2–3 m.y. long and sepa-of the reefs on the western basin margin (0.7–rated by longer time intervals ~5–6 m.y. Under1.5 km) is a good constraint for a minimum waterthe pressure of the advancing Sakmara–depth in the deep-water central part of the fore-Magnitogorsk nappe, the Bashkir shelf was short-deep. In the Kungurian, the Uralian foredeep andened by 10–20 km (Figs. 4b and 8b). This resultedthe Peri-Caspian basin became semi-isolated fromin strong folding and slicing with a completethe ocean, and during the following ~7 m.y. theoverriding of the shelf by the nappe in some places.central and outer parts of the foredeep have beenMost likely, the Sakmara–Magnitogorsk nappefilled with salt (P1kg in columns VI, VII )was also shortened at that time with thrusting in( Khain, 1977).the crystalline basement (Fig. 7).

No significant additional subsidence occurredEach phase of collision was accompanied byin the foreland region at the collisional phasesdeposition of conglomerates on the western marginC 4a–C 4c. A slow shallow-water deposition pro-of the Bashkir shelf and on the eastern slope ofceeded on the eastern slope of the flysch basin inthe Western flysch basin (columns I–III in Fig. 5).the Sakmarian–Artinskian (P1s–P1ar in columnOlistostromes were formed further to the west —III ), and only 1.5 km of the deposits were formedin the deep-water western part of the flysch basinduring this period of time, 20 m.y. long. The(column IV ) and in the adjacent eastern part ofadjacent western deep-water part of the flyschthe foredeep (column V ). Coarse material is mostlybasin had been filled with turbidites by the laterepresented by cherts, mafic volcanites and ultra-Artinskian and became shallow water after phasemafic rocks that were derived from the Sakmara–C 4b (conglomerates at the top of P1ar in columnMagnitogorsk nappe (Geology of the USSR, v.IV ). Only about 300 m of shallow-water salts andXIII, 1964; Chuvashov, 1975; Melamud, 1981).sands were formed in this region in the KungurianConglomerates deposited at each phase of colli-age (P1k), which includes collisional phase C 4c.sion C 4 were 200–500 m thick. In addition, atThus, the collision in the late Early Permian didthat time, 1.5–3 km of turbidites had been formednot result in the formation of high mountains andin the outer western part of the flysch basin and

in the eastern part of the foredeep. These are strong crustal subsidence in the Uralian foredeep.


7.2. Continental collision on the East Uralian collision C 5. However, at the end of the EarlyPermian, salts with a small amount of fine-grainedmicrocontinent in the late Early Permianterrigenous material were forming on the Uralianforeland in the west (P1k in columns I–VII, Fig. 5).Shallow-water deposition proceeded in theAt the start of the Late Permian, thin layers ofcentral and eastern parts of EUM until the end ofcarbonates were formed between the Urals andthe Kasimovian age of the Late CarboniferousKazakhstan. This shows that the topography pro-(Chuvashov et al., 1984). After a hiatus, ~40 m.y.duced by collision C 4, C 5 in western and easternlong, marine transgression occurred in the easternUrals was rapidly eroded, and the thrust beltpart of the microcontinent in the Kazanian age ofbecame smooth, remaining at a moderate altitudethe Late Permian. Between these two epochs,≤0.5 km (Chuvashov et al., 1984).continental collision took place in the eastern and

central parts of EUM (event C 5) ( Kamaletdinovand Kazantseva, 1983; Puchkov and Svetlakova,

8. Deposition in the Uralian foredeep, mountain1993). The shortened region includes numerousbuilding and collision in the Late Permian–Triassicgranitic plutons ( Keilman, 1974; Perfiljev, 1979)

that were emplaced into the upper part of the8.1. Rapid crustal subsidence in the Uraliancrystalline crust. Many of them were formed alongforedeep and mountain building in thrust belt in thethrust faults, most probably, synchronously withLate Permianthe deformations (Popov and Rapoport, 1996).

The plutons are dated as the second half of theSince the start of Late Permian and until theEarly Permian (Smirnov, 1997). Hence, shortening

Middle Triassic, terrestrial deposition prevailed inof EUM could probably occur at the time ofthe Uralian foredeep. Sands, clays and coastalcollision C 4 in western Urals (Fig. 8b).marine carbonates were forming during theBefore collision C 5, the central part of EUMUfimian and Kazanian (P2u and P2kz in columnswas approximately at sea level, and it can beV–VII ). The sediment loaded subsidence occurredsupposed that the crustal thickness at that timeat a rate ≤300 m m.y.−1, and minor conglomerateswas h0c~40 km. The mean present crustal thicknessappeared in the deposits. This indicates the onsetin this region is h1c ~50 km. No further collisionof crustal uplift in the Uralian thrust belt in the

took place on EUM after collisional event C 5, Ufimian (Geology of the USSR, v. XIII, 1964;and ~5 km of rocks have been eroded from this Khain, 1977). At that time, the Western flyschregion in the Late Permian and Triassic (Sigov basin and the shortened Bashkir shelf slightlyand Romashova, 1984). Hence, it is probable that emerged above sea level, and no deposits youngercollision C 5 resulted in an increase in the crustal than the Kungurian deposits occur in these regionsthickness in the central part of EUM by (P1k in columns I–IV ).Dhc~15 km. Taking this value in Eq. (3) with the A strong acceleration of the subsidence tookmean density of the crust rc=2830 km m−3, we place in the inner-eastern and central parts of theobtain an average altitude of the mountains that foredeep in the Tatarian age (3 km of P2t inwould be formed in the absence of a density column V and 2 km in column VI). Rapid depos-increase in the lithosphere as Df~2.3 km. This is ition of coarse molasse took place in these regionscomparable with the average altitude of the Alps. in the first half of the Tatarian. The cumulate

No deposition took place near to EUM on the thickness of conglomerates is 0.5–1 km. Dep-western margin of the Kazakhstan continent in osition of coarse molasse was caused by erosionthe Early Permian. Clastic material that had been of the rapidly rising Uralian thrust belt (Fig. 13a).brought into this region from the Urals became The material had been transported to the foredeeperoded and removed by rivers into the northern by rivers that were flowing across the slightlypart of the West Siberian basin (Bush et al., 1995). elevated region of the former Bashkir shelf andHence, the data on the Kazakhstan cannot be used Western flysch basin. Conglomerates were depos-

ited in piedmont fluvial fans in a band, severalfor estimates of the altitude of eastern Urals after


Fig. 13. Evolution of the Southern Urals in the Late Permian to the Middle Triassic. (a) Mountain building in thrust belt and rapidsediment 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. boulders (up to ~50 cm), at that time, the Uralianmountains probably reached the highest altitude.XIII, 1964). The thickness of the Lower Tatarian

conglomerates rapidly decreases westwards. Coarse molasse includes the material derived fromthe area, 200–300 km wide, namely carbonatesSands, clays and lacustrian carbonates are pre-

dominant in the inner part of the foredeep in the from the shortened Bashkir shelf, cherts, diabasesand ultramafic rocks from the Sakmara–upper Tatarian (the upper part of P2t in column

V ). No significant changes in the climatic condi- Magnitogorsk nappe, metamorphic rocks from theUraltau, juspers from all the Magnitogorsk zone,tions took place in the Southern Urals during the

Late Permian and Early Triassic. Hence, it is and granitoids from its eastern part.In the Early Triassic, a slight crustal upliftprobable that at the Late Permian time, the highest

altitude of the Uralian mountains was reached in occurred in the inner part of the Uralian foredeep,and the deposition ceased in this region (columnthe early Tatarian. At the end of the Late Permian,

weathered mantles were forming in eastern Urals V ). The central and outer parts of the foredeepproceeded to subside until the Late Triassic (T1(Bochkarev, 1973). This indicates the end of the

first epoch of mountain building and planation of and T2–3 in columns VI, VII ). The rate of subsi-dence in the central basin part decreased several-the topography.fold as compared to the Late Permian. Since thelate Late Carboniferous and until the Triassic,8.2. Second impulse of mountain building in the

Early Triassic 1 km of sediments had been deposited on theeastern margin of the East European platform.During the same epoch, up to 7 km of depositsA new impulse of intensive uplift occurred in

the Uralian thrust belt in the Early Triassic. This had been formed in the Uralian foredeep (Fig. 5).This led to the formation of a deep foreland basin.gave rise to deposition of up to 1 km of conglomer-

ates in the central part of the foredeep (molasse No extensive compression occurred in theUralian foredeep at the epoch of mountainT1 in column VI ). According to the large size of


building in the Late Permian–Early Triassic. Fig. 10). The nappe pushed out Permian depositsfrom the latter region, and most of them had beenPreservation of a thin layer of Kungurian deposits

on the former Bashkir shelf and in the Western eroded. A minor part of these deposits becameincorporated into the frontal part of the nappe.flysch basin (P1k in columns I–IV, Fig. 5) shows

that no shortening occurred in these regions, which Nappe superposition was associated withextensive erosion on its surface. Sands and clayswere at a low altitude above sea level. The front

of the Uralian nappe remained stable (Fig. 13a) with minor conglomerates were formed in theMiddle Triassic in the adjacent Uralian foredeep(Kamaletdinov, 1974; Melamud, 1981). By the

Late Permian, the nappes in eastern Urals had (T2 in columns VI, VII, Fig. 5). A small volumeof coarse clastics indicates that rapid erosion ofalready been cut by granitic intrusions (Popov and

Rapoport, 1996; Smirnov, 1997). Furthermore, the nappe surface limited its average altitude at alevel of ~0.5 km. No more than ~0.5 km ofduring mountain building, since the start of the

Late Permian slight extension with normal faulting sediments was deposited in the central part of theforedeep in the Middle and Late Triassic (T2–3 intook place in the Uralian thrust belt (Bochkarev,

1973; Chuvashov et al., 1984; Popov and column VI ). Thus, the last collisional event in theUrals resulted in no mountain building, and noRapoport, 1996; Ivanov, 1997). The intensity of

extension increased in the Early Triassic, which significant subsidence in the foreland region.By the end of the Triassic, a planation surfaceresulted in graben formation and basaltic magma-

tism (Bochkarev and Nesterov, 1987; Surkov et al., was formed over most of the Southern Urals atan altitude of several hundred metres (Sigov and1987). Thus, mountain building in the Urals could

not be produced by shortening of the crust in the Romashova, 1984). At that time, the thrust beltwas surrounded by land areas. Since the Jurassicthrust belt or underthrusting under it of the litho-

sphere from the foreland basin. This required a and until the Eocene, the Southern Urals remainedat an altitude 200–300 m above sea level. Harddensity decrease in the lithosphere.

Conglomerates, up to several hundred metres Precambrian blocks of the Uraltau and BashkirAnticlinorium were staying 0.5 km above the pla-thick, occur in the lower part of the sedimentarynation surface.pile in Early Triassic to Jurassic grabens in eastern

Since the Oligocene, the crustal uplift has takenUrals (Bochkarev, 1973; Ivanov, 1997; Puchkov,place in the east of the East European platform1997). Since the end of the Early Triassic, onlyand in West Siberia. Its magnitude is ~200 m insands, clays and coals were forming in the grabensthe former region and 100–150 in the latter region.(Bochkarev, 1973; Ivanov, 1997; Puchkov, 1997),The Southern Urals were also involved in a dome-which shows that the Uralian mountains becamelike uplift with magnitudes ~400–500 m ( Khain,eroded.1977). All over these regions, the crustal upliftoccurred without any significant compressivedeformations.9. Collision in the Middle Triassic and post-

Using the above description of data, we cancollisional vertical crustal movementsformulate the answers to questions 1–9 inSection 1.At the start of the Middle Triassic, over a few

million years, shortened deposits of the Bashkirshelf were pushed westwards for 15–30 km

10. Basic regularities of vertical crustal movements:and overrode the inner-eastern part of theno direct correlation with collisionWestern flysch basin, collision C 6 (Fig. 13b)

(Kamaletdinov, 1974; Melamud, 1981; Zhivkovich10.1. Slow and rapid subsidences: their role in theand Chekhovich, 1985; Brown et al., 1997). Theformation of foreland basinssedimentary cover of the flysch basin was short-

ened by 20–50% and superimposed as a nappeonto the innermost part of the Uralian foredeep, To discuss the main regularities of the crustal

subsidence on the Uralian foreland, we use the15–20 km wide (see Fig. 6, and right-hand side of


subsidence curves of Fig. 3. These are compiled impulses of rapid tectonic subsidence ensured themajor input into the formation of the forelandbased on the stratigraphic columns of Fig. 5. The

numbers of the curves correspond to those of the basins. In the east, a slow sediment loaded subsi-dence was predominant (curves I–III ). However,columns. For the periods of shallow-water depos-

ition, the amount of subsidence is taken as the the subsidence rate strongly varied in time andsometimes reached 100–200 m m.y.−1 in the Latesediment thickness. The depth of water loaded

basins is usually rather uncertain. For the rapid Carboniferous and Permian.water-loaded subsidences, we conventionally taketheir magnitude as 80% of the sediment thickness 10.2. Absence of strong crustal subsidence on the

Uralian foreland at the epochs of collision in thrustthat filled the deep-water basins. The remainder of20% is attributed to a slow subsidence during the beltfollowing period of time until the recovery ofshallow-water deposition. We presume that the As can be seen from Figs. 3 and 5, the rapid

crustal subsidences on the Uralian forelandduration of the rapid subsidences was 1 m.y. Thisapproximation is very schematic. However, it occurred when there was no strong collision in the

adjacent western part of thrust belt. Rapid subsi-allows us to distinguish between the epochs ofslow and rapid subsidence and to compare the dence in the Western flysch basin in the early Late

Carboniferous preceded superposition of theepochs of rapid subsidence with the main colli-sional events. Sakmara–Magnitogorsk nappe (curve IV at the

epoch of collision C 1). The nappe began to moveAs can be seen from Figs. 3 and 5, a slow sedi-ment loaded subsidence at a rate ~10–30 m 2 m.y. later, and its front was initially ~150 km

from the basin (see Sections 4.1 and 4.2). The firstm.y.−1 took place on the Uralian foreland overmost of time. In the Silurian–Carboniferous, rapid subsidence in the Uralian foredeep at the

end of the Carboniferous took place at the end ofduring 100–140 m.y., it resulted in the form-ation of 0.5–2.5 km of predominantly shallow- collisional event C 3 (see curves V and VI in the

Gzhelian). This event was very weak (shorteningwater deposits. At the beginning of the LateCarboniferous, the deep-water Western flysch by ≤10 km) and could not result in the formation

of the deep-water basin, ≥50 km wide. The secondbasin was formed by rapid crustal subsidence(curve IV ). Another deep-water basin was rapidly rapid subsidence in the foredeep (same curves)

took place in the Late Permian. It was synchronousformed in the Uralian foredeep at the end of theLate Carboniferous (curves V, VI ). The depth of to mountain building in the thrust belt; however,

no collisional events occurred at that time.water in these two basins was ~1 km. In curvesIV–VI, the magnitude of the rapid subsidence that However, collision in the western Urals did not

significantly influence the subsidence on theformed the basins is shown as ~2–3 km. Thiscorresponds to 80% of the sediment thickness that Uralian foreland. Strong collisional events C 1, C

4 and C 6 were accompanied by only minor verticalfilled the basins. It is very likely that at the end ofthe Late Carboniferous, rapid subsidence also displacements in curves I–VI.

It can be suggested that the collisional eventsoccurred in the flysch basin (curve IV ). The thirdrapid subsidence took place in the Uralian fore- are determined incorrectly, and collision actually

took place at the epochs of rapid subsidence ondeep in the Late Permian (curves V, VI). Thesubsidence by 2.5–3 km was sediment-loaded and the foreland, which are timed precisely from the

stratigraphic records. The magnitude of the subsi-reached a rate of ~0.6 km m.y.−1.The magnitudes of the total subsidence since dence due to deflection of the elastic lithosphere

should increase towards the thrust belt. However,the start of the Carboniferous are rather similarin the west (in the Uralian foredeep and in the the areas of rapid subsidence on the foreland were

separated by stable blocks from the nappe. At thewestern part of the Western flysch basin) and east(in the eastern part of the flysch basin and on the time of rapid subsidence in the Western flysch

basin in the early Bashkirian (curve IV in Fig. 3),Bashkir shelf ). In the west (curves IV–VI), short


the basin was separated from the nappe by the ments also took place during the crustal subsidenceon the Uralian foreland, resulting in large lateralBashkir shelf (Fig. 2b), where no strong subsi-

dence took place (curves I, II ). No strong subsi- variations in thicknesses of synchronous deposits.Thus, in columns II and IV of Fig. 5, shallow-dence occurred on the eastern slope of the flysch

basin and on the western margin of the deformed water Silurian beds were initially horizontal. Atthe end of the Early Permian, the base SilurianBashkir shelf at the end of the Carboniferous when

rapid subsidence took place in the foredeep and in became 3–4 km deeper in column IV than incolumn II. The distance between these two regionsthe western part of the Western flysch basin

(Fig. 8a) (compare curves I–III with curves IV–VI was L0~30–40 km. Column V in Fig. 5 includes3 km of Tatarian deposits (P2t), while no subsi-in the Gzhelian). At the epoch of rapid subsidence

in the foredeep in the Late Permian (Fig. 13a), the dence occurred at L0~20–30 km further to theeast in column IV. In the western part of theflysch basin between the foredeep and thrust belt

was stable and remained above sea level (compare Uralian foredeep (Fig. 10), the initially flat layerof Late Carboniferous shallow-water limestones iscurves V, VI with curves III, IV in the Late

Permian). Deepening of the Paleozoic–Triassic bent down by 2.5 km in a band, L0~25 km wide.The lithosphere of the Uralian foreland is dis-foreland basins away from thrust belt can also be

clearly seen in Fig. 5. Thus, irrespective of timing rupted by numerous faults (Fig. 9). Steep bendingand faulting of the lithospheric layer indicate itsof continental collision in western Urals, it could

not be responsible for the formation of foreland high degree of weakening under the Uralian nappeand on the foreland.basins.

10.3. High degree of weakening of the lithosphere10.4. Absence of significant lithospheric stretching

indicated by its narrow-wavelength deformationson the foreland

The absence of flexural reaction of the litho-The depth of a sedimentary basin formed by

sphere on the Uralian foreland on collision inuniform lithospheric stretching by b times in a

thrust belt indicates either a strong decrease in thestate of isostasy is:

effective elastic thickness of this layer, Te, or itsdisruption by large normal faults. The magnitude hs=[(rm−rc)/(rm−rs)]h0c (1−1/b), (4)of crustal subsidence under the Sakmara–Magnitogorsk nappe increased from zero at its where h0c is the thickness of the prestretched crust.

The thickness of Phanerozoic sediments in thefront to 10–15 km at a distance, L0~20–30 kmfurther to the east (Fig. 4a and b). According to Uralian foredeep reaches ~10 km (Fig. 10). At

the mean value of rs=2500 kg m−3 and h0c=Eq. (1), these values of L0 correspond to very lowTe values of ~3–5 km. In fact, at a dip angle of 40 km, the value of b necessary to form this basin

by stretching would reach ~1.7. Lithosphericthe basement ~25–30°, the deformations werebeyond the limit of elasticity. A system of normal stretching of such an intensity should result in

extensive normal faulting with a formation offaults could arise in this region.Large narrow-wavelength variations of the strongly tilted fault blocks. Except at the steep

boundary fault and minor thrust faults, thenappe thickness have been preserved until present.Thus, in the Kimpersay massif, the nappe thickness Riphean strata are continuous in the foredeep

(Figs. 10 and 11). This means that the basin wasincreases from hn~3–5 km to hn~10–15 km at aninterval, L0~15 km wide (Saveljev and Saveljeva, formed without any significant stretching in the

upper crust at least. Stretching of only the lower1991). The thickness of the East Magnitogorsknappe on the western slope of the East Uralian crust can be presumed, however, to be a cause of

the subsidence. Designate the density of the lowermicrocontinent reaches 25–28 km at a distance,L0~75 km from the nappe front (Fig. 7). crust by rlc, and its thinning from stretching by

Dhlc. In a state of local isostasy and after aStrongly differentiated vertical crustal move-


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, to the East European platform actually occursbeneath the Uralian foredeep (Figs. 7 and 14).

Dhlc=[(rm−rs)/(rm−rlc)]hs . (5)This means that the subsidence in the foredeeptook place without any significant stretching ofhs#7 km of the deposits were formed in the fore-

deep after the rapid subsidence at the end of the lower crust.As shown in Section 4.1, no significant stretch-the Carboniferous. Taking rs#2500 kg m−3and

rlc=2930 kg m−3, which is typical of gabbro in ing took place in the Western flysch basin at thetime of initial rapid subsidence between thethe lower crust, we have: Dhlc#14 km. The occur-

rence of garnet granulites in the lower crust Serpukhovian and Bashkirian. There are no datathat could reliably resolve the presence or absenceincreases its mean density, rlc. Then, according to

Eq. (5), a higher degree of thinning of the crust of stretching on the Bashkir shelf where only aslow subsidence took place. However, it iswill be necessary to ensure a subsidence of the

same magnitude. extremely unlikely that a strong extension couldhave occurred in this region near to the front ofThinning of the crust by 14 km with an addition

of 7 km of deposits will raise the base of the crust the nappe and in the absence of stretching on theUralian foreland further to the west.by 7 km. Downwarping of the Moho with respect


10.5. Density increase in the lithosphere as a cause thick nappes superimposed onto shallow-waterregions would be eroded almost completely. Inof crustal subsidence on the Uralian forelandfact, after the end of collisional events, the crustalsurface remained at an altitude ≤0.5 km, withThe crustal subsidence in the Uralian forelandonly a partial erosion of the nappes. In somebasins occurred without any significant lithosphe-regions, thick nappes have been preserved untilric stretching and was independent of collision inthe present.the thrust belt (see Sections 10.2 and 10.4). Under

Collisional event C 1 (Fig. 2c) occurred soonsuch circumstances, a density increase in the under-after closure of the Kazakhstan oceanic basin tolying lithosphere was necessary to ensurethe east of EUM in the middle of the Serpukhoviansubsidence.(Fig. 2b), which could be associated with theappearance of slab pull. At that time, however, no10.6. Continental collision without a synchronousstrong subsidence occurred on this microcontinentformation of high mountainsnear the collisional front. This indicates delamina-tion of the subducted slab synchronously withAll the events of continental collision in theplate collision. A density increase in the underlyingSouthern Urals (C 1–C 6) were accompanied bylithosphere was necessary to maintain thedeposition of moderate volumes of coarse clasticSakmara–Magnitogorsk nappe at a low altitude.material in the adjacent basins. Deposition of onlySimilarly, a density increase should have occurredsands, clays, carbonates and evaporites took placein the lithosphere during the following epochs ofin these basins after the end of each collisionalstrong collision that took place in the absence ofevent. This shows that after rapid erosion of localslab pull.folds and cordilleras formed at the epochs of

Shortening of the crust in the Southern Uralscollision, the areas remained at a low altitudehad been predominantly completed before the≤0.5 km above sea level. The occurrence of severalonset of mountain building in the Late Permian.

epochs of strong thrusting in the Southern Urals Only ~15–30 km of shortening took place in thelong before the onset of mountain building has Middle Triassic, after the end of mountain build-been described by many authors (Khain, 1977; ing. Hence, the Uralian crust had been almostPerfiljev, 1979; Chuvashov et al., 1984; Zonenshain completely formed by the Late Permian. The pre-et al., 1990; Artyushkov, 1993; Khain and Lomize, sent crustal thickness in the Urals reaches 50–1995; Puchkov, 1997). Continental collision orog- 60 km (Figs. 7 and 14), and up to ~5–10 km ofeny is, however, commonly considered as a syn- rocks have been eroded since the Late Permianonym of mountain building. We considered the (Sigov and Romashova, 1984). At the end of theabove stratigraphic data in detail to confirm that Early Permian, the crustal thickness reachedcollision in the Urals did not result in the formation hc≥60–65 km. These values of hc are typical ofof high mountains. high mountains. However, deposition of salts with

minor sands and shales on the Uralian foreland at10.7. Density increase in the lithosphere at the the end of the Early Permian indicates that theepochs of collision Uralian thrust belt was no higher than ~0.5 km

(Section 8.1). Therefore, at that time, the litho-Events of collision in the Southern Urals began sphere in the Urals was considerably denser than

when the crust was near, or ≤0.5 km above, sea in low cratonic areas where the mean crustallevel. Collision produced a large increase in the thickness was hc~40 km.crustal thickness, which, at a constant lithospheredensity, must result in the formation of high 10.8. Mountain building without synchronousmountains. Then, during the following ~10– shortening of the crust20 m.y. a thick layer of rocks would be erodedfrom the mountains with a recovery of approxi- According to the appearance of large volumes

of coarse molasse in the Uralian foredeep, moun-mately the initial crustal thickness. In particular,


tain building in thrust belt took place in the Late us take, for example, d2−d1~50 km. Then, at amean sediment density, rs=2650 kg m−3, it fol-Permian–Early Triassic (Fig. 13a) (Khain, 1977).

At that time, the nappe front remained stable. lows from Eq. (6) that hths ~4.7 km. The sedimentthickness on the Uralian foreland reaches ≥15 km.Slight extension with the formation of grabens and

basaltic volcanism occurred in the rising thrust Hence, thermal relaxation could be responsible foronly about a third of the total subsidence.belt. Thus, mountain building in the Urals was

associated with no synchronous shortening of the Furthermore, the characteristic cooling time of thelithosphere, 150–200 km thick, is ~90–170 m.y.crust.This phenomenon could not influence the subsi-dence of the crust, ~1000 m.y. old, in the10.9. Mountain building as a result of a density

decrease in the lithosphere Paleozoic and Early Mesozoic when the Uralianthrust belt was evolving.

Erosion of the lower crust by convective flowsThe Southern Urals are wide, and their crustwas (and still is) close to the isostatic equilibrium in the mantle can occur when the asthenosphere

wells up to the base of the crust. In cratonic areas(Artemjev et al., 1994; Doring et al., 1997).Mountain building in the Late Permian–Early with a thick mantle lithosphere, such an upwelling

must be associated with a crustal uplift by ≥1 km.Triassic took place without any additional thicken-ing of the crust. Under such circumstances, a Only slight uplifts took place on the Uralian

foreland in the Phanerozoic. Furthermore, as inconsiderable density decrease in the lithospherewas necessary to ensure the formation of the the case of stretching of the lower crust, subcrustal

erosion should result in a large uplift of the Mohomountain range.boundary. Under the Uralian foredeep, the Mohois warped down with respect to the adjacent plat-form margin (Figs. 7 and 14). These features pre-11. Mechanisms of the crustal subsidence and

lithospheric weakening clude the subsidence from subcrustal erosion.

11.2. Crustal subsidence from metamorphism in the11.1. Possible role of thermal relaxation andsubcrustal erosion lower crust

In the absence of strong thermal relaxation andA density increase in the lithosphere can resultfrom thermal relaxation (Sleep, 1971; McKenzie, subcrustal erosion, a large density increase beneath

the Uralian foreland could be produced only by1978), subcrustal erosion ( Keen, 1985), and meta-morphism in mafic rocks in the lower crust (Haxby metamorphism in mafic rocks in the lower crust.

In the case of local isostasy, the formation of aet al., 1976; Artyushkov and Baer, 1983;Artyushkov et al., 1991; Baird et al., 1995). layer of garnet granulites with density, rgg, and

thickness, hgg, from a gabbro with density, rgb,Cooling of the lithosphere, which was initially atsea level, produces a sedimentary basin with a results in the subsidence, hs, expressed by:depth:

hs=(rm/rgb)[(rgg−rgb)/(rm−rs)]hgg . (7)hths =a(Ta/2)[rm/(rm−rs)](d2−d

1). (6)

In the profiles of Fig. 14, a layer, ~8–13 km thick,with P-wave velocities, VP~7.4–7.5 km s−1, isHere, a=3×10−5 K−1 is the thermal expansivity,

Ta=1300°C is the temperature of the asthenosph- located in the lowermost crust beneath the fore-deep. At such a velocity, garnet granulites canere, and d1 and d2 are the initial and final lithosphe-

ric thickness, respectively. The subsidence on the have a density, rgg~3250–3300 kg m−3 (Sobolevand Babeiko, 1994). Similar P-wave velocities areUralian foreland began in the Riphean on the

lithosphere that was already rather cool and thick. typical of the lowermost crust below the Uralianthrust belt, where the density has been estimatedHence, the lithospheric thickness could not be

considerably increased during the subsidence. Let as ~3240 kg m−3, according to the gravity data


(Doring et al., 1997). Taking rs=2400 kg m−3, dence under a sediment load by ~3 km at a rateof ~3 km m.y.−1, which is 1.5–2.5 orders of mag-rgb=2930 kg m−3, rgg~3250–3300 kg m−3, hgg=

8–13 km in Eq. (7), we obtain: hs~4–5 km. The nitude higher than that of slow sediment loadedsubsidence. The increase in the subsidence ratesediment thickness that has been formed in the

foredeep after the end of thermal relaxation should be associated with a drastic increase in therate of gabbro–garnet granulites–eclogite trans-(~10 km) is ~5 km larger, and an additional

density increase was necessary to ensure this subsi- formation. Consider possible causes of thisphenomenon.dence. P-wave velocities under the Moho beneath

the foredeep are VP~8.0–8.1 km s−1. Such veloci- Typical diagrams of the gabbro–garnet granu-lite–eclogite transformation are shown in Fig. 15.ties can be typical of both the mantle and denser

eclogites with re~3500 kg m−3 (ibid., Manghnani Fig. 15b is more representative of the continentalcrust. Due to a very low reaction rate, the experi-et al., 1974; Christensen and Mooney, 1995).

Hence, it cannot be precluded that the Moho ments were carried out at high temperature, andonly extrapolations can be used for a mediumbeneath the foredeep in this profile is underlain by

eclogites. Let us take the density of sediments and temperature typical of the lower crust in cratonicareas. If these extrapolations are correct, at pres-metasediments in the lower part of the sedimentary

cover as rs=2700 kg m−3. Then, according to Eq. sure, p≥0.5–1 GPa and temperature, T ≤400–600°C, the gabbro in the lower crust is metastable(7), to ensure an additional subsidence of ~5 km

at rgg=re~3500 kg m−3, the thickness of eclogitic and should transform into dense eclogites. In cooland dry mafic rocks, the transformation is verylayer below the Moho should also be

he=hgg~5 km. slow and could be responsible for only a slowsediment loaded subsidence on the Uralian fore-In the profile of Fig. 14a, the Moho boundary

beneath the foredeep is at a depth of ~45–48 km. land. An increase in the rate of the transformationby 1.5–2.5 orders of magnitude requires heatingAccording to their composition, mafic eclogites

pertain to the crust. If we add ~5 km of eclogites of the rocks by several hundred degrees or thepresence of a small amount of water-containingto the crust, its thickness will be 50–53 km. Then,

the thickness of the crystalline crust that is overlain fluid (CO2+H2O) to catalyse the reaction (Ahrensand Schubert, 1975). Several tens of millions ofby ~15 km of sediments will be ~35–38 km, as

on the adjacent platform margin. In this case, the years are necessary to heat the lower crust in athick cratonic lithosphere after an increase in theformation of the Uralian foredeep can be explained

by thermal relaxation in the lithosphere and meta- heat flow from the asthenosphere. This is incom-patible with a short duration, ~1 m.y., of rapidmorphism in the lower crust with a preservation

of the material in the crystalline crust. crustal subsidences.Numerous examples are known for rapid meta-The high value of VP~8.5 km s−1 occurs under

the Moho in the Uralian foredeep in profile of morphism in mafic rocks in the presence of vola-tiles (Austrheim, 1987, 1998; Wayte et al., 1989;Fig. 14b. The composition of rocks with such

velocities is a special problem. These might be Rubie, 1990; Walther, 1994; Austrheim et al.,1997). The lower crust includes hydrous mineralseclogites with a low mean atomic number.such as biotite, muscovite, epidot and amphibol(Spear, 1993). Segregation of water occurs at11.3. Rapid metamorphism and crustal subsidence

due to infiltration of volatiles from the T~700–950°C, and it is commonly supposed thatvolatiles appear in the lower crust due to theasthenosphereheating of rocks. Several tens of million years arealso necessary to cause dehydration reactions in aAt the epochs of slow crustal subsidence on the

Uralian foreland, the subsidence evolved at a rate cratonic lower crust by heating from the asthen-osphere. This mechanism cannot produce strong~10–100 m m.y.−1. The rapid subsidences pro-

duced water-loaded basins ~1 km deep during crustal subsidences with a duration of ~1 m.y.To explain this, infiltration of volatiles into the~1 m.y. This corresponds to an isostatic subsi-


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 ofthe curves.

lower crust from the asthenosphere has been sug- concentrations of these elements occur in layerswith both high and low contents of organic matter.gested (Artyushkov et al., 1991; Artyushkov,

1993). Rapid subsidence with a formation of deep- This precludes a significant influence of organicsubstances on the concentration of the microele-water basins occurred in many areas, e.g. in West

Siberia in the Late Jurassic and in the Peri- ments. Infiltration of volatiles from the mantle hasbeen suggested as a mechanism of enrichment. OilsCaspian, Volga–Urals and Timan–Pechora basins

in the Late Devonian. Volatiles can be brought to and bitumens in these regions include Nd and Sr,with the isotopic characteristics typical of lam-the base of the lithosphere by small fluid-contain-

ing plumes (Artyushkov and Hofmann, 1997, proites in Australia and Spain derived from themantle: eNd=−(9−12), and 87Sr/86Sr=0.708–1998). Their emergence to the lithospheric base is

manifested by slight crustal uplifts, ~100 m, of a 0.719. In the Peri-Caspian basin, carbonates witha low content of organic matter, which wereshort duration, ~1 m.y., that preceded rapid

crustal subsidences in the above regions, as well formed at the time of rapid subsidence in theVisean age, have a content of uranium (Pisotsky,as in many others (Artyushkov and Baer, 1986a;

Artyushkov, 1993). A slight uplift also occurred 1999) that exceeds that typical of marine carbon-ates by one order of magnitude (Taylor andbefore the rapid subsidence on the Uralian fore-

land at the end of the Late Carboniferous McLennon, 1985). This has been interpreted as aresult of transportation of uranium with fluids(Section 6.2).

In many cases, slight uplifts that preceded rapid from the mantle.We suggest that the rapid crustal subsidencescrustal subsidences without lithospheric stretching

and collision nearby were associated with slight on the Uralian foreland, as well as rapid additionalsubsidences in the areas of collision in thrust belt,alkaline basaltic volcanism (e.g. the Volga–Urals,

Timan–Pechora and Peri-Caspian basins) and plu- which maintained it at a low altitude, resultedfrom a gabbro–garnet granulites–eclogite trans-tonism (e.g. West Siberia) (ibid.). This can be

another indication of infiltration of volatiles into formation catalysed by infiltration of volatiles fromthe asthenosphere into the lower crust.the lithosphere.

Infiltration of volatiles from the asthenospherealso follows from the geochemical data on the 11.4. Weakening of the lithosphere as a result of

infiltration of volatiles and rapid metamorphismregions of rapid crustal subsidence. Sediments inthe Volga–Urals and Timan–Pechora have veryhigh contents of Se, As, Mo, Hg, U and Re Several mechanisms have been proposed to

explain the high degree of weakening of the litho-(Pushkarev et al., 1994; Pisotsky, 1999). High


sphere. Its extreme bending near thrust belts can much lower than in dry rocks (Poirier, 1985;Strehlau and Meissner, 1987; Kohlstedt et al.,result in yielding and a large decrease in Te

(Washbusch and Royden, 1992; Burov and 1995). Hence, the viscosity drop in the upper crustcan also be caused by infiltration of volatiles. ADiament, 1995). Lithospheric weakening, how-

ever, occurred not only under the Uralian nappe, strong viscosity decrease in the mantle lithospherecan be due to infiltration of volatiles through thisbut also on the foreland at times when there was

no collision in the thrust belt. Lithospheric stretch- layer from the asthenosphere into the crust. Thus,infiltration of volatiles can result in a high degreeing can also thin the elastic lithosphere (Desegaulx

et al., 1991; Stewart and Watts, 1997). No rifting of weakening in the whole lithospheric layer(Artyushkov and Morner, 1997, 1998; Artyushkovoccurred in the Uralian foredeep for ~1000 m.y.

before weakening of the lithosphere in the et al., submitted for publication).Carboniferous. A strong decrease in Te takes placeduring a high heat flow (Burov and Diament,1995). The old cratonic lithosphere of the Uralian 12. Crustal structure in thrust belt and mechanisms

of mountain building in the Late Permian and Earlyforeland was cool.Weakening of the lithosphere on the Uralian Triassic

foreland occurred at the epochs of rapid crustalsubsidence. Steep slopes of the crystalline base- 12.1. Composition of the transitional layer in the

lower crustment, which are a few tens of kilometres wide,were formed during the rapid subsidences in manyother areas (Artyushkov et al., 1996, submitted The crust in the Urals is now up to hc~60 km

thick (Figs. 7 and 14), as in the high mountainsfor publication; Artyushkov and Morner, 1997,1998). This happened, for example, in the of the Tien Shan and Eastern Alps. Since the mean

altitude of the Southern Urals is only 600–800 m,Carpathian and Caucasian foredeeps at times whenthere was no collision in the adjacent thrust belts, their crust should have a high mean density. In

the profiles of Fig. 14, the transitional layer, up toand the Peri-Caspian and Trans-Caspian intraplateareas. Thus, strong lithospheric weakening at times ~20 km thick, with P-wave velocities of

7.4–7.5 km s−1 exists in the lowermost crust. Theseof rapid crustal subsidence is a widely occurringphenomenon. velocities are intermediate between those typical

of gabbro in the basaltic layer (VP=6.8–Rapid subsidences in sedimentary basinsresulted from metamorphism in the lower crust 7.0 km s−1) and mantle peridotites under the

Moho boundary (VP=8.0–8.4 km s−1). In theunder infiltration of volatiles. The same phen-omena ensured an additional rapid subsidence in east, VP in the transitional layer increases to

7.8–8.0 km s−1. The lower crust of the Urals isthe regions of continental collision in the SouthernUrals, which limited their altitude by ≤0.5 km. In characterized by a moderate temperature,

T~600°C (Khachay and Druzhinin, 1993;all cases, lithospheric weakening occurred at timesof infiltration of volatiles. Hence, the latter phe- Kukkonen et al., 1997). At VP=7.4–7.5 km s−1, it

can be composed of (1) a crust and mantle mixture,nomenon can be supposed to be a cause of theformer. and (2) garnet granulites, or their mixture with

eclogites.As evidenced by large viscous deformations inthe lower crust, rapid metamorphism in mafic The formation of the crust and mantle mixture

due to intrusions into mantle peridotites of layersrocks is associated with a strong viscosity decrease(Austrheim, 1991, 1998). At the epochs of rapid of gabbro with a cumulate thickness, hgi, results

in the crustal uplift by:contraction of the lower crust from metamor-phism, large deviatory stresses comparable with

Dfgi=[(rm−rgb)/rm ]hgi , (8)the lithostatic pressure can arise in the upper crust,which can result in a disappearance of the elastic In a volume that includes large blocks of peridotite

and gabbro with P-wave velocities VprP and VgbP ,core in this layer. In wet rocks, the viscosity is


respectively, the average P-wave velocity will be: 12.2. Possible crustal uplift in the future

VP#VprP VgbP /[c(VprP −VgbP )+VgbP ], (9)The lithospheric thickness in the Urals is esti-

mated to be ~200 km (Steer et al., 1998). At awhere c is the volume per cent of gabbro in thecrust and mantle mixture. Taking VgbP = crustal thickness of ~60 km, the thickness of

mantle lithosphere should be hml~140 km.6.8 km s−1, VprP =8.2 km s−1, we find thatVP=7.4–7.5 km s−1 corresponds to a case when Suppose that asthenospheric upwelling with a mag-

nitude Dhml≤hml will occur beneath the Urals ingabbro constitutes c#0.45–0.53 of the totalvolume of rock. In the 20 km thick transitional the future. Assume that the temperature in the

mantle lithosphere increases linearly with depthlayer, the total thickness of gabbroic intrusionswill be hgi#9–10.6 km. Substituting these values from the Moho temperature, TM, at the base of

the crust to the asthenospheric temperature, Ta, atinto Eq. (8), we find: Dfgi#1.1–1.3 km. The totalthickness of the middle and upper crust above the the base of the lithosphere. Then, a rapid replace-

ment by the asthenosphere of the lower part oftransitional layer in the Urals reaches 40 km(Fig. 14). This is similar to the crustal thickness mantle lithosphere, Dhml thick, results in the iso-

static uplift by:on the adjacent East European platform.According to Eq. (8), with ~9–10.6 km of gabbro

Dfuw=a(Ta−TM)Dh2ml/2hml . (10)in the transitional layer, the Urals would standDfgi~1.1–1.3 km above the platform. In fact, over Suppose that, in the future, the asthenosphericmost of the Urals, their elevation with respect to upwelling will reach the base of the crust:this area is only 0.3–0.5 km. Dhml=hml=140 km. Then, at Ta=1300°C and

The most likely time for the formation of the TM=600°C, we have: Dfuw=1.5 km.crust and mantle mixture with the associated Garnet granulites are stable in the lower crustcrustal uplift would be the epoch of mountain only at a moderate temperature. Upwelling of abuilding in the Late Permian–Early Triassic. hot asthenosphere to the base of the crust willIntrusion of a large volume of basaltic magmas result in their heating and retrogression to pyrox-into the mantle under the Moho boundary would ene granulites that are considerably less dense.most likely result in extensive basaltic volcanism. This should produce an uplift at the surfaceIntrusions of minor portions of basaltic magmas (Artyushkov and Baer, 1986b; Artyushkov, 1993;with T~1200°C into the overlying crust should Dewey et al., 1993; Artyushkov et al., 1996; Leresult in granitic magmatism due to melting of Pichon et al., 1997). A transformation of garnetmetamorphosed sediments incorporated into the granulites in a layer, hgg thick, into pyroxenemiddle and lower crust at the preceding stages of granulites with density, rpg, will raise the crustalcollision. Most of the granitic intrusions had been surface by:formed in the Early Carboniferous and EarlyPermian — before the onset of mountain building. Dfggrg=[(rgg−rpg)/rpg]hgg. (11)Slight basaltic volcanism began in the Southern

Taking hgg=20 km, rgg=3200–3250 kg m−3, rpg=Urals in the Middle Triassic — after the end of2960 kg m−3 (Bousquet et al., 1997), we have:mountain building. At that time, eastern UralsDfggrg=2.0–2.3 km. The present mean altitude ofrepresented only the western margin of a widethe Southern Urals is fp=0.6 km. The mean heightvolcanic province that spread for ~1000 km fur-of the mountains formed due to asthenosphericther to the east into West Siberia.upwelling to the crust will beThus, the occurrence of a large volume of

gabbroic intrusions in the transitional layer is very ff=fp+Dfuw+Dfggrg=4.1–4.4 km. (12)unlikely. This layer would be most likely composedof dense garnet granulites formed from gabbro in Thus, in the future, the present Urals can become

as high as eastern Tien Shan, where the crustalthe lower crust at the epochs of crustal subsidenceand the following collision. thickness is also ~60 km, but where asthenosph-


eric upwelling reaches the base of the crust (Artyushkov, 1993). This situation is supposed forthe present Alps (Laubscher, 1990; Artyushkov,(Yudakhin, 1983; Roecker et al., 1993).1993; Austrheim et al., 1997; Marchant andStampfli, 1997) and Himalayas (Sapin and Hirn,12.3. Possible mechanisms of mountain building1997). Delamination of eclogites of thickness hdeand their replacement by the asthenosphere withThe amount of erosion in the Uralian thrust

belt since the beginning of the Late Permian and density ra produces the uplift:up to the present reaches her=7–9 km on a regional

Dfde=[(re−ra)/ra]hde . (14)scale (Sigov and Romashova, 1984). Let us desig-nate, by f0, the average initial height of the moun- Taking ra=3220 kg m−3, we find that the crustal

uplift by ~1 km in the Southern Urals could betains and, by f1, the average altitude of the areaafter their erosion. Let rer be the density of eroded ensured by delamination of hde~10 km of

eclogites.rocks. Then:Second, as suggested by Richardson and

f0=[(rm−rer)/rm ]her+f

1. (13)

England (1979), at times of collision, eclogitesform from garnet granulites due to a pressureTaking her=7–9 km, rer=2700 kg m−3 and f1=

0.5 km, we find: f0=1.9–2.2 km. The crustal uplift increase in the lower part of a thickened crust.Then, in a course of the following thermal reequili-that formed the Late Hercynian mountains in the

Southern Urals occurred in two impulses: in the bration and heating of the crust, eclogites ret-rogress to less dense granulites, which producesTatarian age and in the Early Triassic. A consider-

able volume of rocks had been eroded during the the crustal uplift (ibid.; Artyushkov and Baer,1986b; Dewey et al., 1993; Le Pichon et al., 1997).first phase before the mountains reached their

maximum height in the Early Triassic. Hence, it is Thermal relaxation evolves gradually during ~40–80 m.y. In the Urals, shortening of the crust beganmost likely that the mean height of the mountains

did not exceed 1.5 km. This is smaller than the in the Devonian, and the shortened regionsremained low for up to 120 m.y. Then, a rapidmean height of the Alps (~2 km).

Asthenospheric upwelling to the base of the uplift began synchronously in all the regions, whichwere shortened at different times. Two peaks ofUralian crust, which was 65–70 km thick in the

Late Paleozoic, would result in a formation of mountain building, each several million years long,occurred in the Late Permian and Early Triassic.mountains, ≥4 km high. The Uralian mountains,

however, had only a moderate height. This indi- This sequence of crustal movements could notresult from retrogression of eclogites due to a slowcates that in the Late Permian and Early Triassic,

the asthenosphere did not reach the base of the thermal relaxation that began at different times indifferent regions.crust. Two mechanisms of mountain building of

moderate intensity can be supposed. Retrogression of eclogites is perhaps possibleonly in the presence of fluid (Heinrich, 1982;The first is convective replacement of the lower

part of mantle lithosphere by a less dense asthen- Austrheim, 1998). In a dry lower crust, they canbe metastable for a very long time. Infiltration ofosphere (Artyushkov, 1983). This could result

from weakening of the lower lithosphere under volatiles will catalyse the reaction. The formationof garnet granulites from eclogite in a layer of theinfiltration of volatiles from the asthenosphere

(Artyushkov and Hofmann, 1997, 1998). initial thickness, he, produces the crustal uplift by:According to Eq. (10), the magnitude of the uplift

Dferg=[(re−rgg)/rgg ]he . (15)is proportional to a square of thickness of thedelaminated layer: Dfuw~Dhml2 . Hence, at Taking re=3550 kg m−3 and rgg=3250–3300 kg

m−3, we find that for an uplift by Dferg~1 km,Dhml≤0.3hml, the uplift is small. However, it couldbe strongly increased if the delaminated volume expansion of eclogites is necessary in a layer of

initial thickness he~11–13 km. Retrogression ofincluded large blocks of dense eclogites that hadsunk into it at the preceding epochs of collision eclogites under infiltration of volatiles from the


asthenosphere can explain why a rapid uplift took type also occurred in many intraplate basins, e.g.the West Siberian, Timan–Pechora and Peri-place synchronously all over the Southern Urals,

1300 km long and several hundreds of kilometres Caspian basins (Artyushkov and Baer, 1986a;Artyushkov, 1993).wide, and occurred concomitantly in regions that

were shortened at different epochs and remainedlow for long periods of time afterwards.Delamination of the lower part of mantle litho- 13.2. Collision without mountain buildingsphere and retrogression of eclogites in thelowermost crust both require infiltration of vola- Strong collision on the continental crust is com-

monly supposed to result in a synchronous moun-tiles. Hence, these mechanisms probably operatedsynchronously. tain building due to the isostatic response to

thickening of the crust (Miyashiro et al., 1982;Zonenshain et al., 1990). In this approach, orogenyis equivalent to mountain building. Several phases13. Discussionsof strong collision took place on the continentalcrust in the Southern Urals in the Late13.1. Formation of foreland basins that was

independent of collision in thrust belt Carboniferous and Early Permian. However, themean altitude of the shortened regions did notexceed ~0.5 km. The Urals are not unique. TheAccording to classic schemes, foreland basins,

including the Uralian basin, were formed due to occurrence of collision and mountain building atdifferent stages of evolution of fold belts was firstdeflection of the elastic lithosphere in response to

collision in thrust belts (Quinlan and Beaumont, suggested by Stille (1920). He pointed out thatcollision orogeny results in the formation of only1984; McNutt et al., 1988; Zonenshain et al., 1990;

Royden, 1993; Stewart and Watts, 1997). The a short-lived relief of a moderate height, whilehigh mountains form later without any strongUralian foreland basin was already up to ~10 km

deep when collision began in the Urals ~410 m.y. collision and exist for a long time. A time lapse of10–100 m.y. between collision and mountainago. Since that time, the subsidence by up to 7–

8 km took place in the foredeep; however, it pre- building occurred in the Tien Shan (Schulz,1948; Sokolov, 1949), Verkhoyansk rangedominantly occurred at times when there was no

collision in the thrust belt. Thus, the foredeep is (Pushcharovsky, 1959), Caucasus (Milanovsky,1968; Lukina, 1990), Alps (Artyushkov, 1993),at the right place, but the subsidence took place

at the wrong times. Furthermore, at the epochs of and Carpathians (Artyushkov et al., 1996). Thisproblem is considered in detail in Artyushkovmajor subsidences, the foreland basins deepened

away from the thrust belt, which precludes their (1993).The formation of only a low topography withformation from collision.

No significant lithospheric stretching took place a mean height ≤0.5 km at each epoch of collisionin the Urals indicates a synchronous densityon the Uralian foreland, and thermal relaxation

could not considerably influence the subsidence of increase in the lithosphere. The epochs of collisionwere short: from a few millions of years tothe lithosphere, which was ~1000 m.y. old. Under

such circumstances, only a density increase due to ~10 m.y. A rapid density increase most likelyresulted from metamorphism in the lower crustmetamorphism in mafic rocks in the lower crust

could produce the subsidence up to 7–8 km. Most under infiltration of volatiles from the asthenosph-ere. This also occurred in the Alps, Carpathiansof the tectonic subsidence in the Uralian foredeep

and Western flysch basin took place in the form and Tien Shan (Artyushkov, 1993; Artyushkovet al., 1996). Rock contraction is supported by aof strong impulses, ~1 m.y. long. This required

infiltration of volatiles from the asthenosphere that low present altitude of a thick Uralian crust(~60 km), and the presence of dense rocks withstrongly increased the rate of metamorphism

(Artyushkov et al., 1991). The subsidences of this high P-wave velocities in its lower part.


13.3. High degree of weakening of the lithosphere thrust belt, high mountains can be formed byinjection of the crustal material into the lowerdue to infiltration of volatiles from the

asthenosphere and rapid metamorphism in the lower crust from the foreland region (Zhao and Morgan,1985; Isacks, 1988). These mechanisms are inappli-crustcable to the Urals where no thrusting occurredduring mountain building in the Late Permian andIn the early Late Carboniferous, a thick nappe

was superimposed onto the eastern margin of the Early Triassic.High mountains can be rapidly formed due toEast European platform, where the crust was

≥1000 m.y. old. The effective elastic thickness of delamination of the mantle lithosphere withasthenospheric upwelling to the crust (Bird, 1979).an old lithosphere is Te≥50 km (Bechtel et al.,

1990; Burov and Diament, 1995; Stewart and Together with retrogression and expansion ofmetamorphosed mafic rocks in the heated lowerWatts, 1997). A similar estimate was obtained for

the present Urals (Kruse and McNutt, 1988; crust, this would raise the Uralian crust, ≥60 kmthick, to ≥4 km. This altitude is typical of theMcNutt et al., 1988). However, according to the

absence of flexural reaction on the foreland and eastern Tien Shan, for example, where the crust,60 km thick, is underlain by the asthenospherenarrow-wavelength deformations under the nappe,

at the time of nappe superposition, Te in these (Roecker et al., 1993). The Late Uralian moun-tains were only ~1.5 km high. Hence, the asthen-regions was no more than ~5 km. A strong

decrease in Te is commonly attributed to steep osphere remained far below the Moho.Mountain building can take place in a coursebending of the lithosphere near to convergent

boundaries (Washbusch and Royden, 1992; Burov of thermal re-equilibration after collision due toretrogression and expansion of eclogites in theand Diament, 1995). However, narrow-wavelength

vertical displacements also occurred on the Uralian lower part of a thick crust underlain by the mantlelithosphere (Richardson and England, 1979).foreland at times of rapid subsidence when there

was no collision in the thrust belt. Similar lith- Thermal relaxation evolves during ~40–80 m.y.(Le Pichon et al., 1997). Hence, this mechanismospheric deformations took place at times of rapid

subsidence in the intraplate Peri-Caspian basin cannot explain a rapid mountain building in theSouthern Urals that occurred synchronously inand in the Transcaspian area far from active plate

boundaries (Artyushkov and Morner, 1997, 1998; regions shortened at different epochs long beforethe uplift. The Uralian mountains are more likelyArtyushkov et al., submitted for publication). They

occurred in the foredeeps of the Caucasus and to have formed due to infiltration of volatiles intothe lithosphere, which triggered two processes.Carpathians when there was no collision in these

thrust belts (ibid.; Artyushkov et al., 1996). Thus, First, the rate of metamorphism in mafic rocksincreased by several orders of magnitude all overlithospheric weakening can occur independently of

plate collision. Its occurrence at times of rapid the thrust belt, thus allowing rapid retrogressionof eclogites in the lowermost crust to less densesubsidences probably indicates that the lithosphere

was weakened by infiltration of volatiles from the garnet granulites. Second, a high degree of weaken-ing of the lower part of the mantle root with denseasthenosphere and rapid metamorphism.block of eclogites ensured its rapid delaminationand replacement by a less dense asthenosphere.13.4. Mountain building without collision

Most mechanisms of mountain building imply 13.5. Lithospheric weakening as a trigger for theformation of thrust belts in the continental crustshortening and thickening of the lithosphere with-

out any significant changes in its density (Molnarand Tapponier, 1975; England and McKenzie, The cratonic lithosphere on the eastern margin

of the East European platform was shortened and1983; Fleitout and Froidevaux, 1983; Royden,1993, 1996; Brown and Beaumont, 1995; Ellis incorporated into the Uralian thrust belt. Strong

shortening of cratonic lithosphere occurred in sev-et al., 1995). In the absence of compression in the


eral other thrust belts, e.g. the Himalayas and fluids. Thus, the major tectonic events in theSouthern Urals were caused by the appearance ofVerkhoyansk range. The integral strength of an

old lithosphere is ≥1014 N m (e.g. Dunbar and volatiles in the lithosphere. In contrast to manyother thrust belts, no postorogenic collapseSawyer, 1989). The forces acting in continental

lithosphere and the forces driving plate motions occurred in the Urals, thus ensuring the preserva-tion of a thick crust. This shows that after moun-are much lower: ≤1013 N m (Artyushkov, 1973;

Harper, 1986). As a result, under normal condi- tain building, the lithosphere became dry and hasremained rigid until present.tions, continental cratons are stable, constituting

rigid parts of drifting lithospheric plates. The presence of volatiles in a deep crust iscommonly attributed to the dehydration of rocksTo allow extensive shortening of an old litho-

sphere under moderate forces acting in this layer, (Heinrich, 1982; Austrheim, 1998). This is a phe-nomenon of a local extent that cannot occurit should be weakened considerably. Narrow-wave-

length deformations occurred during the rapid synchronously in wide areas. Rapid crustal subsi-dences, mountain building, and a strong collisioncrustal subsidence in the Western and Eastern

flysch basins of the Southern Urals in the early in the Southern Urals took place synchronously inregions ≥1000 km long and ≥100 km wide.Late Carboniferous. They indicate lithospheric

weakening that preceded the onset of continental Furthermore, a high degree of weakening occurrednot only in the crust, but also in the mantlecollision by ~2 m.y. A considerable shortening of

the continental crust in the Alpine and lithosphere. These phenomena can be betterexplained by infiltration of volatiles from smallVerkhoyansk belts also occurred only in those

basins where rapid subsidence had taken place plumes emerging to the lithosphere and rapidlyspreading along its base (Artyushkov and(Artyushkov and Baer, 1984, 1986c). Basins where

only a slow subsidence occurred remained as Hofmann, 1997, 1998).A synchronicity of rapid crustal subsidence andslightly deformed inner massifs within the short-

ened areas. In these thrust belts, as in the Urals, uplift, as well as of collision, was typical of manyother areas. For example, in the Late Jurassic,the rapid formation of foreland basins with a high

degree of weakening of the lithosphere was not a within the limits of accuracy of paleontologicaldata ~1 m.y., a deep-water basin, ~1000 km wideconsequence (an egg), but rather a cause (a hen)

of formation of thrust belts. This explains why and ~2000 km long, was concomitantly formedall over West Siberia (Artyushkov and Baer,many thrust belts are adjacent to deep foreland

basins where rapid crustal subsidence of a large 1986a). The crustal uplift by ≥1 km took placeduring the last ~5 m.y. in many cratonic areas,magnitude took place.~1000 km in size, in Africa and north-easternAsia (Artyushkov and Hofmann, 1997, 1998).13.6. Tectonic evolution of the Southern Urals as a

result of infiltration of volatiles from the Several epochs of collision, 1–10 m.y. long, tookplace synchronously in the Alpine belt in theasthenosphereMesozoic and Cenozoic (Schwan, 1980). Thus,infiltration of volatiles from the asthenosphere canRapid crustal subsidence on the Uralian fore-

land resulted from metamorphism in the lower be a common cause of mobilization of continentallithosphere, including both strong vertical crustalcrust catalysed by infiltration of fluids.

Considerable shortening of the continental litho- movements and the formation of thrust belts underconvergent motions of lithospheric plates. This issphere in the thrust belt under convergent plate

motions became possible due to weakening of this why, in the Southern Urals, collision C 1 beganonly 2 m.y. after the rapid crustal subsidence inlayer under infiltration of volatiles. Mountain

building occurred due to rapid retrogression of the flysch basins in the early Late Carboniferous,and rapid subsidence in the foredeep in the Latedense mafic rocks in the lower crust and/or by

delamination of weakened lower part of the mantle Permian was synchronous to mountain building inthrust belt. It is quite probable that lithosphericroot. Both mechanisms required the presence of


stretching under divergent plate motions also mountains were formed only in those regionsoccurs in regions weakened by infiltration of vola- where the crust had been shortened and thickenedtiles from the asthenoshere. at the preceding stages of convergence. Hence,

mountain building occurred as a result of interfer-ence of deep-seated processes and plate motions.Similar regularities have been established for the14. ConclusionsEast Carpathians (Artyushkov et al., 1996). Ananalysis of the development of other thrust beltsVertical crustal movements in thrust belts areand their foreland basins is necessary to under-commonly explained by horizontal plate motions.stand whether or not these regularities are of aAccording to these ideas, foreland basins form duegeneral character.to deflection of the elastic lithosphere caused by

plate collision. The occurrence of collision in thrustbelts is predominantly controlled by the distribu-tion of plate motions. High mountains arise as an Acknowledgementsisostatic response to thickening of the crust fromcollision. We thank I.A. Basov, B.I. Chuvashov, V.S.

As follows from our analysis, this simple scheme Druzhinin, S.N. Kashubin, M. Lindstrom, A.A.is inapplicable to the Southern Urals. The Uralian Mossakovsky, A.S. Perfiljev, V.N. Puchkov, M.S.foreland basins, past and present, were formed Rapoport, A.V. Rybalka, S.G. Samygin, A.A.independently of collision in the thrust belt. In Saveljev and Yu.K. Shchukin for valuable discus-these basins, a predominant part of the crustal sions. We are also very thankful to the referees H.subsidence occurred from the gabbro–garnet gran- Austrheim and J. Touret for their comments whichulites–eclogite transformation in the lower crust. helped to improve the presentation. A large partIntensive convergence took place only at times of of this paper was prepared during the stay of E.V.a high degree of weakening of the lithosphere due Artyushkov at the Stockholm University in 1994–to infiltration of volatiles from the asthenosphere 1996. The support from the Russian Foundationand rapid metamorphism in the lower crust. At all for Fundamental Research and Peri-Tethysthe epochs of collision on the continental crust, Programme is also acknowledged.the crustal surface reached only a low altitude≤0.5 km above sea level. The magnitude of thecrustal uplift was strongly reduced by a synchron- Referencesous density increase from phase transformationsin the shortened regions. The Late Hercynian Ahrens, T.J., Schubert, G., 1975. Gabbro-eclogite reaction rate

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