Dynamics of intra-plate compressional deformation: the Alpine foreland and other examples

TECTONOPHYSICS ELSEVIER

Tectonophysics 252 (I 995) 7-59

Dynamics of intra-plate compressional deformation: the Alpine foreland and other examples Peter A. Ziegler a7*,Sierd Cloetingh b, Jan-Diederik van Wees ’ b Department

a Geological-Palaeontological of Sedimentary Geology. Faculty ’ Rijks Geologische

Institute.

of Earth Dienst,

University of Bad, Bernoullistr. 32, 4065 Bawl, Switzerland Sciences, Free University, de Boelelaan 1085, 1081 HV Amsterdam, Richard Holkade IO,2000 AD Haarlem, Netherlands

Netherlands

Received 8 November 1994; accepted 18 July 1995

Abstract Intra-plate compressional structures, such as inverted extensional basins and upthrusted basement blocks, play an important role in the tectonic framework of the European Alpine foreland. Similar structures are observed on many continental cratons but occur also in oceanic basins and more rarely along passive continental margins. The World Stress Map shows that horizontal compressional stresses can be transmitted over great distances through continental and oceanic lithosphere. Although a number of geodynamic processes contribute to the build-up of intra-plate horizontal compressional stresses, forces related to collisional plate interaction appear to be responsible for the most important intra-plate compressional deformations. Such deformations can involve whole-lithosphere buckling and folding, crustal folding and, by reactivation of pre-existing crustal discontinuities, upthrusting of basement blocks and inversion of tensional hanging-wall basins. Mechanical aspects of basin inversion depend on the interplay of stresses and rheology of the lithosphere. Pre-existing crustal discontinuities weaken the lithosphere and play a crucial role in localizing intra-plate compressional deformations. Reactivation of relatively steeply dipping normal faults occurs when the angle between their sbike and the compressional stress trajectory is smaller than 45”. Compressional deformations restricted to crustal levels involve ‘simple-shear’-type detachment of the crust at the level of the rheologically weak lower crust from the mantle-lithosphere; whole-lithospheric ‘pure-shear’-type compressional deformation is indicated for certain inverted basins. A distinction must be made between collision-related and anorogenic compressional/transpressional intra-plate deformations. The hypothesis is advanced that the stratigraphic record of collision-related intra-plate compressional deformations can contribute to the dating of erogenic events affecting the margin of the respective craton.

1. Introduction Western post-Variscan

and Central structuring

* Corresponding author. CWO-I 95 I /95/$09.50 SSDl

0040-1951(95)00102-6

0

I995

Europe owes its intense to multiple deformation

phases ranging in age from Permo-Carboniferous to Cenozoic. During this time span regional stress regimes changed repeatedly, causing the development of new crustal discontinuities and time and again reactivation of pre-existing ones. This applies specifically to Mesozoic basins which developed un-

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der a tensional to transtensional setting and became inverted during the Late Cretaceous and/or Cenozoic in response to the build-up of intra-plate horizontal compressional stresses. Inversion of these basins was accompanied by upthrusting of major basement blocks, for instance in the Bohemian Massif and along the Fennoscandian border zone (Ziegler, 1987, Ziegler, 1990). Irma-plate compressional structures play an important role in the tectonic framework of Western and Central Europe. Some of these tectonic features are exposed and accessible to surface geological analyses (e.g., Osning zone, Harz, Pays-de-Bray anticline, Scania, Bohemian Massif, Provence and Languedoc), whereas others are concealed, for instance by the thick Cenozoic sediments of the North Sea basin or those of the Alpine-Carpathian foredeep, or were overprinted by Cenozoic rift systems. The evolution of inverted basins, which are now recognized as representing compressionally/transpressionally deformed extensional and/or transtensional hanging-wall basins, has been the subject of debate since a long time, particularly in terms of the origin of the stresses which controlled their post-rift deformation. In Europe, inverted basins were recognized as such already in the 1920’s and 1930’s (Lamplugh, 1920; Stille, 1924; Pruvost, 1930). The mechanism of basin inversion was variably attributed to local isostatic movements compensating for anomalies generated during the initial rapid basin subsidence (Weald basin: Lamplugh, 1920) or to halokinesis (Lower Saxony basin: Lackmann, 1910; Boigk, 1968). Voigt (1963) presented a first regional synopsis of inverted Mesozoic basins in Western and Central Europe, to which he referred, in the sense of Stille (1940), as ‘ parageosynclines’ or ‘marginal troughs’ (Randtriige) and stressed their association with the margins of stable cratonic blocks. He noted that these basins were all inverted during the latest Cretaceous and earliest Tertiary, and that a sharp distinction between ‘germanotype’ (intra-plate) and ‘alpinotype’ (plate-margin) fold belts is not always possible, thus hinting at the compressional nature of basin inversion, without stating this explicitly. In Russia, the concept of aulacogens and aulacogeosynclines, corresponding to rifts of variable dimensions which transect Precambrian platforms, was devel-

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oped by Shatsky (1964) and modified by Milanovsky (1983) who referred to inverted aulacogens and aulacogeosynclines as intra-cratonic fold belts (see Milanovsky, 1992). During the last decades the understanding of the structural configuration and evolution of Western and Central Europe has greatly advanced. In this the results of intense exploration activity carried out by the petroleum industry played a very important role. The acquisition of extensive reflection-seismic data bases and the drilling of countless wells during the search for hydrocarbons in basins located in on-shore as well as off-shore areas has permitted to establish an inventory of compressional intra-plate deformations and to calibrate the timing of these deformations. On the basis of these sub-surface data the structural style and the regional pattern of these deformations could be analyzed and related to the megatectonic evolution of Europe. Results of these studies are documented in a literature that is too extensive to summarize here (see e.g., Ziegler, 1982, Ziegler, 1987, Ziegler, 1988, Ziegler, 1990; Brooks and Glennie, 1987; Cooper and Williams, 1989; Parker, 1993; Chadwick, 1993; Buchanan and Buchanan, 1995). At the same time deep reflection and refraction-seismic surveys, carried out by national and international scientific consortia (e.g., BIRPS, DEKORP, ECORS, EGT, BABEL), paralleled by analogue and theoretical basin modelling research, has provided a better insight into geodynamic processes which govern the deformation of the lithosphere (Kusznir and Park, 1987; McClay, 1989; Balling, 1992; Kusznir and Ziegler, 1992; Cloetingh and Kooi, 1992; Cloetingh and Banda, 1992; Nalpas et al., 1995; Brun and Nalpas, 1995). Moreover, analysis of the present stress distribution in the lithosphere, resulting in the compilation of the World Stress Map, has shown that regionally consistent horizontal compressional stresses can be projected from plate boundaries deep into continental as well as oceanic intra-plate domains (Zoback, 1992), where they can give rise, depending on their magnitude and the elastic strength of the crust and mantlelithosphere, to a broad spectrum of intra-plate deformations. These include broad-scale positive and negative deflections of the lithosphere, lithospheric and/or crustal folding and, by reactivation of pre-existing crustal discontinuities, basin inversion and the

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upthrusting of basement blocks (Cloetingh et al., 1993). Irma-plate compressional structures occur on many continents, are recognized in oceanic domains, but are only rarely observed on today’s passive margins. Intra-continental upthrusted basement blocks and graben structures that were inverted after an initial rifting stage, range in age from Proterozoic to Cenozoic. In most cases development of these intra-plate compressional features appears to coincide with major erogenic events affecting one or more of the margins of the respective craton. However, transpressionally deformed grabens are also associated with multi-directional rift systems and zones of major wrench faulting. The first author has proposed in his publications that the bulk of the Late Cretaceous and younger intra-plate compressional deformations observed in Western and Central Europe developed in response to stresses that developed as a consequence of collisional coupling of the Alpine and Pyrenean orogens with their forelands, and that the minor inversion structures evident along the Norwegian and Scottish continental margins are related to differential movements across intra-oceanic transform faults which project under the continent (Figs. 10-14). This view was challenged, for instance by Cartwright (1989) who visualizes a much stronger contribution of ridge-push forces associated with the sea-floor spreading axes of the North Atlantic and the Norwegian-Greenland Sea. On the other hand, Gillcrist et al. (1987) proposed a genetic relationship between the timing of basin inversion in the Channel-Westem Approaches-Celtic Sea area and the development of the Rhine-Rh6ne rift system. A major concern of Dewey and Windley (1988) was that the intense Paleocene compressional deformation of the Alpine foreland apparently does not correlate with major erogenic activity in the Alps (‘Paleocene restoration’ of Trtimpy, 1960). Baldschuhn et al. (1991) questioned whether the inversion of Mesozoic troughs in Northern Germany can indeed be related to collisional coupling of the foreland and the Alpine orogen and suggest that mantle processes are the cause of these deformations, without specifying these. Lastly, Drozdzewski (1988) Chadwick (1993) and Nalpas et al. (1995) support the first author’s hypothesis and conclude that Late Cretaceous and Cenozoic

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transpressional deformation of the heavily fractured northern Alpine foreland must be related to its collisional coupling with the evolving Alpine orogen. In this paper we review the megatectonic setting of selected examples of intra-plate compressional deformations observed on the Russian Platform, in North America and Africa, on passive margins and in oceanic domains, as well as the setting of the Late Cretaceous and Cenozoic compressional and extensional deformations occurring in the European Alpine foreland. Possible mechanisms governing the development of intra-plate compressional deformations are discussed in terms of lithospheric rheology and stress. We discuss the evolution of the Alpine Foreland lithosphere and draw the conclusions that collisional coupling between orogens and forelands is responsible for many, tough not for all of the observed intra-plate compressional deformations.

2. Examples of intra-plate compressional mations outside the Alpine foreland

defor-

The Russian Plagorm underwent major rifting cycles during Riphean and Vendian times. Some of these aulacogens were inverted during the Baikalian orogeny, at the time Pangea II was assembled (Fig. 1; Milanovsky, 1987, Milanovsky, 1992). During the late Middle Devonian to Early Carboniferous a new rifting cycle affected the Russian Platform, giving rise to the subsidence of, for instance, grabens in the Timan-Pechora area (foreland of the Polar Urals) and the Dniepr-Donets rift system. The grabens of the Timan-Pechora area were inverted during the Permian main phases of the Uralian orogeny; therefore, these deformations can be directly related to collisional coupling between the evolving orogen and its foreland (Ulmishek, 1982; Matviyevskaya et al., 1986); the most distal compressional/transpressional foreland structures are located 500 km to the west of the Uralian thrust front. The southeastern parts of the Dniepr-Donets rift, corresponding to the Donbass and Karpinsky Swell, were also inverted during the Permian Uralian orogeny (Milanovsky, 1992; Chekunov et al., 1992; Stephenson et al., 1993); however, subsequent intraMesozoic pulses of inversion movements must be related to erogenic activity in the Crimean-Caucasus

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domain. This structure, the greater parts of which are deeply buried beneath Mesozoic sediments, strikes in a northeasterly direction away from the southern termination of the Uralian fold belt; partial inversion of the Dniepr-Karpinsky rift presumably entailed differential movements between the Scythian and the Russian platforms, possibly involving an intra-lithospheric detachment of the former (Konashov, 1980; Sobomov, 1995). In the North American cruton, the Midcontinent

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rift (U.S.A.), which is located in the foreland of the Grenville orogen, is a classical example of a compressionally deformed extensional intra-plate structure (Fig. 2); this rift evolved during the Middle Proterozoic and was inverted during the Late Proterozoic Grenvillian orogeny (van Schmus, 1992; Cannon, 1992). During the latest Silurian-earliest Devonian, the North American craton was bordered to the north and east by the active Pearya (early Inuitian) and

CARPATHIANS

-

-

_ Riphean

rifts

Fig. 1. Megatectonic map of the East European Platform (after Milanovsky, 1987; Zonenshain et al., 1990; A.N. Nikishin, pers. commun., 1995). Precambrian rifts: widely spaced barbs; Devonian rifts: closely spaced barbs. D&KS = Donbass-Karpinsky swell, PDD = PrypiatDniepr-Donets Graben, SP = Scythian Platform, TP = Timan-Pechora area, C/S = Ukrainian Shield, VZ = Voronesh Massif.

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Caledonian orogens, respectively, and to the southeast by the Appalachian arc-trench system (Fig. 3; Ziegler, 1989). Evidence for contemporaneous compressional deformation of the craton is provided by upthrusting of the Boothia arch-Comwallis fold belt in the Canadian Arctic (Kerr, 1977; Trettin, 1989) and by upthrusting of basement blocks in the Hudson Bay (Sanford and Norris, 1973; Sanford, 1987; Thorpe, 1988); the latter are located some 1600 km to the northwest of the contemporaneous Salinian (Appalachian) deformation front. At the same time the broad Trans-Continental arch, the Ozark uplift and the Nashville dome were upwarped, giving rise to the partial destruction of the Silurian-Ordovician carbonate platform which had covered much of the North American craton (Cook and Bally, 1975; Bally, 1989). As the resulting Tippecanoe-Kaskaskia sequence boundary is not evident on the Sahara Platform, it is not of a eustatic nature but probably is the result of broad lithospheric buckling in response to the build-up of horizontal compressional stresses in

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the lithosphere during the late phases of the Caledonian orogeny (Bally, 1989; Ziegler, 1989). During the Late Devonian-Early Carboniferous terminal phases of the Inuitian orogeny, the over 1100 km long and some 12.5 km wide, basement-involving and fault-bounded Boothia arch, which trends at a right angle to the Inuitian thrust front, was repeatedly reactivated (Kerr, 1977, Kerr, 1981; Ziegler, 1989). During the latest Carboniferous to Early Permian Alleghenian suturing of Gondwana and Laurentia, the cratonic part of the United States was again affected by major intra-plate compressional stresses. Collisional coupling between the evolving Ouachita-Marathon orogen and its foreland gave rise to the development of the Ancestral Rocky Mountains (Fig. 4); these are characterized by a system of upthrusted basement blocks and intervening transtensional, fault-bounded basins (Ross and Ross, 1985, Ross and Ross, 1986; Kluth, 1986; Stevenson and Baars, 1986; Oldow et al., 1989). The

MOHO

I

I

Rift sediments ~ “i’:,“:, .,:gxfjy ‘izp:‘.:Rift volcanics Rift intrusives mj

Basement rocks Fig. 2. Schematic

cross-section

through

Precambrian

Midcontinent

rift, U.S.A.

(after

Cannon,

1992).

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most distal transpressional structures forming part of the Ancestral Rocky Mountains are located some 1500 km to the north of the contemporaneous Marathon-Ouachita thrust front (Ziegler, 1989); moreover, compressional microtectonic deformations are evident at distances of up to 1700 km from this erogenic front (Craddock et al., 1993).

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During the latest Cretaceous-Paleogene Laramide orogeny, the Cordilleran foreland was affected by an important phase of intra-plate compressional deformation. Major features are the inverted early Palaeozoic Richardson Mountain trough, which projects from the Mackenzie Mountains into their northern foreland (Fig. 5; Ziegler, 19691, the basement blocks

Fig. 3. Geological map of North America with Devonian and younger strata removed, illustrating subcrop pattern beneath Kaskaskia sequence (Devonian-Early Carboniferous) (after Cook and Bally, 197.5; Bally, 1989). BA = Boothia arch, HB = Hundson Bay, NV = Nashville Dome, OZ = Ozark uplift, TCA = Transcontinental arch.

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of the Central Rocky Mountains of Wyoming, Colorado, Utah and New Mexico (Bally and Snelson, 1980; Oldow et al., 1989) and possibly also the broad Sabine and San Marcos arches, located in the northwestern parts of the Gulf of Mexico basins (Laubach and Jackson, 1990). The Rocky Mountains consist of an array of upthrusted basement blocks which are apparently detached from the mantle-lithosphere at a lower crustal level (Fig. 6; Brewer et al., 1980; Egan and Urquhart, 1993); crustal shortening achieved in this thick-skinned foreland thrustbelt amounts to some 35 km and was presumably compensated at mantle-levels at the Cordilleran A-subduction zone. The most distal compressional foreland structures of the Rocky Mountains are located 750 km to the east of the Cordilleran Sevier trust front. The sedimentary record of basins separating the major upthrusted blocks of the Rocky Mountains shows that these evolved during the latest CretaceousPaleogene, presumably in response to compressional

Fig. 4. Tectonic elements of Late Carboniferous-Early U.S.A. (after Ross and Ross, 1985).

Permian

Ancestral

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13

reactivation of pre-existing crustal discontinuities (Heller et al., 1993); this is indicative for strong coupling between the Laramide orogen and its foreland (Oldow et al., 1989). Similarly, the Paleocene uplift of the broad Sabine and San Marcos arches may be attributed to buckling of the entire lithosphere in response to the build-up of collision-related compressional stresses within the Cordilleran foreland. During the Late Cretaceous and Paleogene, opening phases of the North Atlantic and the Norwegian-Greenland Sea, Greenland played the role of a semi-independent plate that shifted in a rotational mode northward relative to the Canadian Shield in conjunction with progressive opening of the Labrador Sea and the Baffin Bay. This motion of Greenland was accompanied by major transpressional deformation of the western margin of the Barents shelf (Senja ridge-West Spitsbergen Alpine fold belt: Lyberis and Manby, 1993; Lepvrier, 1994)

Rocky

Mountains

in foreland

of the Marathon-Ouachita

orogen,

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and of the eastern parts of the Sverdrup basin, resulting in its inversion and the development of the Eurekan fold belt and, at the same time, reactivation of the Boothia arch (Ziegler, 1988; Trettin, 1989; Ricketts and Stephenson, 1994). Development of the Eurekan fold belt entailed major crustal shortening and possibly folding of the entire lithosphere (Stephenson et al., 1990). The Sahara Platform of northwestern Africa experienced major intra-plate compressional deformations during the Variscan erogenic cycle and again during the Alpine orogeny. During the Late Carboniferous-Early Permian the broad Reguibat and Anti-Atlas arches were uplifted, resulting in the individualization of the Tindouf and Taoudeni basins;

Fig. 5. Richardson

Mountains,

northwestern

Canada:

(a) location

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this was paralleled by inversion of the CambroOrdovician Saoura-Ougarta rift, which cross-cuts the Sahara Platform in an southeasterly direction, and by the reactivation of the northerly striking Panafrican El-Biot axis (Fig. 7; Ziegler, 1988; van de Weerd and Ware, 1994). The Saoura-Ougarta anticlinorium, which has a length of some 800 km and strikes oblique to the North African Alpine chains, was affected by a second inversion phase during the Eocene. This deformation phase was contemporaneous with the inversion of the Triassic-Jurassic Atlas troughs and the emplacement of the internal Kabylian and ultra-Tellian nappes (Vially et al., 1994). Both phases of compressional intra-plate deformation can be clearly related to collisional events affecting the

map, (b) stmctiural

cross-section,

(cl schematic

palinspastic

reconstruction.

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et al. / Tectonophysics

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15

(b) W EAGLE

E PLAINS

RICHARDSON

0

PEEL PLATEAU

MOUNTAINS

50

looKhI

80

200

400

600

KM

Fig. 5 (continued).

northwestern and northern margin of the African craton, respectively (Ziegler, 1988 Ziegler, 1989). In the European, northern foreland of the Variscan orogen, a first phase of compressional deformation occurred during the late Visean-early Namurian, resulting in the disruption of the Kohlenkalk platform which occupied the northern shelf of the Rhenohercynian basin (Ziegler, 1990); this defortna-

tion phase probably coincided with the collision of the Variscan orogen with the passive northern margin of the Rhenohercynian basin. During the subsequent thrust-loaded subsidence of the Rhenohercynian shelf, governing the evolution of the Variscan foreland basin, there is little evidence for compressional foreland deformation. Only during the Westphalian, terminal phase of the Variscan orogeny were

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100

150

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200

DISTANCE

250

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350

300

m

Tertiarysediments I:;l!‘ili”‘:‘:l Paleozoic

I

Mesozoic

Fig. 6. Tectonic elements of Laramide Central Rocky crustal section (after Egan and Urquhart, 1993).

400

450

500

(KM)

Precambrian Mountains,

compressional stresses exerted again on the foreland, presumably as a consequence of erogenic over-thickening of the lithosphere in the Variscan fold belt,

Wyoming,

western

U.S.A.

(after

Bally

and Snelson,

1980) and balanced

causing inversion of the Carboniferous rifts of the British Isles (Fig. 8; Midland Valley, Northumberland-Solway, Craven, Dublin, Munster basins) at

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252 (1995)

SW

Reggan Basin

7-59

17 Ougarta Chain

Timimoun Basin

N

Atlantic

0

MZ km

3

660 km

Fig. 7. Permo-Carboniferous intra-plate deformation of the Sahara Platform forming the foreland Arrows indicate axes of major arches. Schematic evolutionary diagram of Ougarta trough (after Ta = Taoudeni basin, Tm = Timimoun basin, Gh = Cihadames basin, MO = Mourzouk basin.

distances up to 550 km to the north of the Variscan thrust front (Ziegler, 1989, Ziegler, 1990). The complex Cretaceous rift system of West and

-j/J

rkHB

Fig. 8. Westphalian tectonic framework of Northwest Europe (after Ziegler, 1990). CRB = Craven basin, DBT = Doublin trough, MSB = Munster basin, MVG = Midland Valley Graben, NHB = Northumberland basin, SE = Slyne-Erris trough, SWB = Solvay basin.

of the Mauretanides-Atlas orogen. Ziegler, 1988). Ti = Tinduff basin,

Central Africa includes several sub-basins that were inverted during the Santonian, such as the Benue, Bongor and Doba-Doseo basins (Fig. 9; Genik, 1992, Genik, 1993). Although this transpressional deformation phase correlates with a change in the drift pattern of Africa (Guiraud et al., 1992; Maurin and Guiraud, 1993), it is thought to reflect a sequential reaction of the lithosphere to progressive extension that was controlled by the same, persistent large-scale divergent deviatoric stress regime that had also governed the preceding subsidence phases of these rifted basins (Pavoni, 1992). As such, the inversion of basins forming part of the West and Central African rift system cannot be related to a collisional event, but resulted apparently from progressive changes in strain distribution in a multi-directional rift system. Further examples of wrench-induced, frunspressional basin inversion come from Californian basins which are associated with the San Andreas fault system (Oldow et al., 1989), from the Dead Sea wrench system (Palmyra trough: McBriden et al., 1990; Chaimov et al., 1993; Negev fold belt: Quennell, 1984; Moustafa and Khalil, 1995) and from the Old Red basins which evolved during the Devonian sinistral translation of Baltica relative to Laurentia-

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Greenland (Ziegler, 1988, Ziegler, 1990; Coward et al., 1989). Such deformations are associated with the development and activity along transform intra-con-

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tinental plate boundaries. Examples of transpressionally deformed plate boundaries, which evolved into passive margins, are the western margin of the Ba-

-\\.%\I

,/

5

Ethiopia

\--Cameroon

Atlantic Ocean

‘, j 0 I

Iceous-Tertiary nentary Rocks

I

Et;

I

Pre-Cretaceous Metamorphic

IgneousRocks

Volcanics D

+

~$.~~~~~

\

Cretaceous Extension Normal

Rift

Fault

I

N DOSE0

BASIN

BOROGOP

FAULT

ZONE

3.0

Fig. 9. Central African Cretaceous rift system: (a) tectonic crossing Doseo and Doba basin (after Genik, 1992).

DOBA

BASIN

PRECAMBRIAN

map (after

McHargue

et al., 1992);

(b) line drawing

of reflection-seismic

line

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rents shelf (Ziegler, 1988; Faleide et al., 1993) and the Central Atlantic margin of Brazil and the conjugate African margin of Ghana and Ivory Coast (Kumar, 1981; Blarez and Mascle, 1988; Basile et al., 1992; Guiraud and Maurin, 1992). Examples of compressional and transpressional structures

occurring

on passive

continental

margins,

which clearly post-date the onset of sea-floor spreading, come mainly from the continental slopes of the British Isles (Porcupine Bank, Faeroe-Rockall area: Roberts, 1989; Boldreel and Andersen, 1993) and the Mid-Norway shelf (Fig. lOa). Whereas the broad anticlinal feature of the Molde high is interpreted as a partly inverted basin (Fig. lob; Bukovics and Ziegler, 19851, wrench faults, small-scale folds and thrust-faulted structures are evident on the WyvilleThomson ridge, which separates the Rockall from the Faeroe-Shetland trough, and within the Faeroe basin (Fig. 10~; Earle et al., 1989; Boldreel and Andersen, 1993). These structures developed mainly during the late Eocene to Oligocene opening phases of the northern North Atlantic and NorwegianGreenland Sea. During this time, sea-floor spreading axes had not yet stabilized to their present arrangement; gradual abandonment of the Aegir ridge at the expense of the newly developing Kolbeinsey ridge was accompanied by transform motions along the Jan Mayen and the Iceland fracture zones. In areas where these fracture zones project into the continental margins, these were apparently tectonically destabilized as evident by their deformation. That such deformations can reach deep into a continent is illustrated by the late Eocene-Oligocene activity along a zone of wrench-faulting which extends over a distance of 1000 km from the Hebrides shelf through the Irish Sea to Cornwall, as indicated by subsidence of a chain of pull-apart basins (Ziegler, 1988, Ziegler, 1990; Roberts, 1989). Late Cretaceous and Paleocene compressional deformations in the Rockall-Faeroe area pre-date crustal separation between Greenland and Europe and are probably related to the same stress system which controlled the main inversion of Mesozoic basins in Northwest Europe. However, Miocene and younger deformations, as evident on the RockallHatton Bank and its margins (Boldreel and Andersen, 19931, were possibly induced by a combination of ridge-push forces and collisional foreland stresses.

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At the base of the Armorican and the Porcupine B ank shelf slope m inor Eocene compressional/transpressional structures, including steep reverse faults, are observed (Fig. lOa). These structures, which are superimposed on the continent-ocean transition zone, are thought to have developed in response to the Pyrenean collision of Iberia with the southern margin of Europe (Boillot et al., 1985; Sibuet et al., 1985; Roberts, 1989; Masson et al., 1994). There is also evidence for Mid-Miocene compressional reactivation of the continent-ocean transition zone west of Iberia; this is thought to be related to erogenic activity in the Betic Cordillera (Masson et al., 1994). In oceanic domains, intra-plate compressional deformations are evident, for instance, in the North Atlantic and the Indian Ocean. In the eastern part of the North Atlantic, a zone of compressionally and wrench-deformed oceanic crust extends from the north Spanish trough some 1000 km westward towards the Mid-Atlantic ridge and terminates in the tensional Kings trough and Palmer ridge (Fig. lOa). As sea floor younger than anomaly 13 (35-36 Ma) is not affected, these structures are thought to have developed during Paleocene-Eocene times (Grimaud et al., 1982). As such, these deformations coincide with the main phases of the Pyrenean orogeny during which Iberia was accreted to the southern margin of France and the southern parts of the oceanic floor of the Bay of Biscay were subducted. Senonian to Paleogene convergence of Iberia with Europe was induced by collisional coupling of Iberia with northward-drifting Africa. Thus, initiation of the south-dipping Pyrenean-Cantabrian subduction zone must be related to the build-up of far-field compressional stresses that were large enough to cause failure of the lithosphere along the northern continent-ocean transition zone of Iberia. Eocene minor compressional/transpressional deformation at the foot of the continental slope of the Armorican margin and the Porcupine Bank suggest that stresseswere also transmitted through the oceanic lithosphere of the Bay of Biscay (Boillot et al., 1985; Sibuet et al., 1985; Roberts, 1989). Collisional coupling between Africa, Iberia and Europe is held responsible for severe Senonian to Paleogene intraplate deformations of Iberia (inversion of Celtiberian Range, upthrusting of Sierra Guadarrama basement

20

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block: Vegas et al., 1990; Salas and Casas, 1993; van Wees, 19941, along the southern margin of France and imbrication of the oceanic crust in the southern Bay of Biscay and the eastern North Atlantic (Ziegler, 1988). In the northeastern Indian Ocean, the ocean floor and its sedimentary cover (Bengal fan) are involved in a system of broad folds and smaller faulted anticlinal structures which trend normal to the main axis of compression as inferred from earthquake focal mechanisms and stress field modelling. Sub-parallel whole lithospheric anticlines are spaced 150-250 km, whereas the crust is deformed into smaller folds, complicated by thrusts and reverse faults. Based on

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ODP well results, development of these structures commenced during the late Miocene (Cloetingh and Wortel, 1986; Neprochnov et al., 1988; Bull and Scrutton, 1990). Deformation of oceanic crust, which has an age of 60-80 Ma and therefore is relatively strong, requires the build-up of a high stress level. Such a stress level developed probably as a consequence of collisional coupling of the Eurasian and Indian plates and their differential northward movement which was not fully compensated at the Himalyan subduction zone and by deformations of the Eurasian plate (Nikishin et al., 1993; Beekman, 1994). The examples summarized above indicate that the

GREENLAND

1988): CIR = Celtiberian Fig. IO. (a) Oligocene tectonic framework of Western Europe (after Ziegler, WT = Wyville-Tomson ridge, KT = Kings trough. (b) M@lde high on Mid-Norway shelf (after Bukovics compressional deformationsons on Wyville-Thomson ridge (after Boldreel and Andersen, 1993).

range, MH = Mslde high, and Ziegler, 1985), and (c)

PA. Ziegler

0 w

100

et al./

Tectonophysics

200

252 (1995)

7-59

300

400 Km

sw

Fig.

10 (continued).

most important compressional intra-plate deformations are the result of collisional and transpressional plate interaction. For instance, Permo-Carboniferous suturing of Gondwana and Laurussia was accompanied by spectacular compressional deformation of the North American craton as well as of northwestern Europe and northern Africa. Large-scale compressional deformation of these intra-plate domains at distances of over 1500 km from the thrust front of the Variscan-Appalachian-Ouachita-Marathon fold belt ceased with the consolidation of this orogen. Similarly, deformation of the Uralian foreland can be closely related to the suturing of Siberia and Kazakhstan to the eastern margin of Baltica. Laramide compressional deformation of the North American craton must be related to its collisional coupling with the Cordilleran orogen. Collisional resistance between the Eurasian and Indian plate is thought to underlay compressional deformation of oceanic crust

in the northeastern parts of the Indian Ocean. Transpressional interaction of the Greenland sub-plate with the Laurentian and European cratons during the progressive opening of the Arctic-North Atlantic resulted in deformation of the western Barents shelf and the Sverdrup basin. A collisional plate-interaction model can also be applied to the Late Cretaceous and Cenozoic deformation of the European Alpine foreland.

3. Compressional Alpine foreland

deformation

of the European

The inventory of latest Cretaceous and Cenozoic compressional/transpressional it-ma-plate deformations in the Alpine foreland of Europe permits to map their temporal and spatial distribution, to ana-

22

P.A. Ziegler

et al./

Tectomphysics

lyze their intensity and to relate them to the megatectonic evolution of Europe. For descriptions of individual inverted basins and basin complexes the reader is referred to Ziegler (1987), Brooks and Glennie (1987), Drozdzewski (1988) Cooper and Williams (1989) Baldschuhn et al. (1991), Parker (1993) and Chadwick (1993). Regional syntheses are provided by Ziegler (1988, , 1990 and in the maps given in Figs. 11-14. In Western and Central Europe, structures which evolved in response to Late Cretaceous and Cenozoic intra-plate compression include an array of inverted Mesozoic tensional and transtensional basins as well as the upthrusted basement blocks of the Bohemian Massif and Scania. Zones affected by intra-plate transpressional deformation are separated by such little deformed stable blocks as, for instance, the Franconian Platform and the Paris basin which were possibly affected by crustal buckling. In general, the intensity of basin inversion decreases with increasing distance from the Alpine thrust front. However, small-scale compressional structures occur also along the continental margins of the British Isles and Norway. The age of these intra-plate deformations ranges from intra-Senonian to Pliocene. Through time, a systematic westward shift of compressional tectonic activity is evident. To the northeast, the distribution of intra-plate compressional deformations is limited by the Tomquist-Teisseyre and Fennoscandian border zone (Sorgenfrei line), corresponding to the margin of the stable EastEuropean Platform During the Middle and Late Cretaceous, rifting activity abated rapidly in Europe in conjunction with progressive strain concentration on the Arctic-North Atlantic rift system. Consequently, the tensional basins of Western and Central Europe entered a post-rifting stage that was dominated by their regional subsidence in response to thermal relaxation of the lithosphere. At the same time, rising sea-levels accounted for a regional transgression and rapid overstepping of basins margins. However, Late Cretaceous and Cenozoic convergence of Africa-Arabia with Eurasia, involving deformation of such intervening microplates as the Italo-Dinarides block, resulted in the gradual closure of oceanic basins in the Mediterranean domain and ultimately the collision of the evolving Alpine orogen with the southern, pas-

252 (1995)

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sive margin of Europe. Collisional coupling of the Alpine orogen with its foreland is thought to be responsible for the onset of intra-plate compressional deformation of the highly fractured metastable platform of Western and Central Europe (Ziegler, 1988, Ziegler, 1990; Dercourt et al., 1993). In the area of the Bohemian Massif, earliest compressional reactivation of basement blocks, delimited by faults that were active during Permo-Carboniferous and Late Jurassic-Early Cretaceous times, is dated as late Turonian (Malkovsky, 1987). Similarly, inversion of the Polish trough started during the late Turonian (Stephenson et al., 1993; Dadlez et al., 1995). Moreover, in the Sole Pit and Broad Fourteens basins earliest inversion movements are also dated as intra-Turonian (van Hoom, 1987; Nalpas et al., 1995). During the Senonian deformation of the Bohemian Massif, the Polish trough and the Sole Pit basin intensified and inversion of the North Danish, Altmark-Brandenburg, Lower Saxony, West Netherlands and the southern parts of the North Sea rift commenced. Most distal structures related to this so-called ‘Sub-Hercynian’ phase of foreland compression are located 1200 km to the northwest of the present Carpathian deformation front (Fig. 11). Following Early Cretaceous closure of the oceanic Pieniny basin and collision of the internal Carpathian Tatra block with the Czorsztyn ridge, the North Carpathian Magura basin, which was floored either by strongly attenuated continental crust or at least partly by oceanic crust, began to close during the Turonian and Senonian, as indicated by uplift of the Silesian Cordillera (Kovac et al., 1993; Winkler and Slaczka, 1994). At the same time horizontal compressional stresses were transmitted into the Carpathian foreland. It is likely that compressional deformation of the Alpine North Penninic-Rhenodanubian trough had also commenced during the Senonian (Triimpy, 1980); however, as only minor intra-Senonian deformations are evident on western parts of the Helvetic shelf, coupling of the nascent Central Alpine orogen with its foreland was apparently less intense than in the Eastern Alpine and Carpathian domains. Despite these intra-Senonian foreland deformations the Chalk basin of Northwest Europe did not experience a major disruption and persisted until the end of the Danian during which time its compres-

PA

Ziegler

et al./

Tectottophysics

EXTENSION

loo

252 (199.5)

39

7-59

INVERSION

200

300 400 dlstance(km)

INITIAL

!500

loo

200

CENTER EXTENSION

300 400 dlstance@m)

500

CENTER INVERSION

----

----a

8

3

p Z %ii

8

0 strength

1000 @Pa)

0 strength

1000 (MPa)

0 strength

1ooo (MPa)

Fig. 20. Thermal and rheological evolution for a IO-Ma phase of pure-shear extension (lithospheric stretching factor 1.5), followed by a lO@Ma period of tectonic quiescence which precedes a IO-Ma phase of compressional basin inversion. Model and figure conventions as in Fig. 19. (A) Thermo-tectonic configuration at end of the period of tectonic quiescence for Q, = 70 mWm-a. (B) 2-D rheological reconstruction corresponding to (A). (C) Thermo-tectonic configuration at end of basin inversion. (D) 2-D rheological reconstruction corresponding to (C). (E-G) Strength profiles for Q, = 60, 70 and 80 mW m- *, indicated by light, medium and dark shading, respectively. (E) Initial, unstretched stage. (F) Basin centre after stretching, at end of tectonic quiet phase. (G) Basin centre at end of inversion phase.

40

P.A. Ziegler

Table 1 Timing of rifting Rifted

and inversion

basin

stage of extensional

et al./

Pechora-Kolva (Russia) Kirov-Via&a aulacogen (Russia) Dniepr-Donbass (Ukraine) Midland Valley (Scotland) Maestrat Basin (Spain)

stage

1109-1087 580-380(?) 530-340(?)

Atlas Troughs

(Morocco)

390-350 380-365 380-340 370-310 260-205 154-104 255- 152 250- 100 250-80 245- 157 135-85 24& 180

Lower

Basin (Germany)

146-100

Polish Trough (Poland) Southern North Sea graben Pyrenees (Spain, France) Celtiberian Range (Spain)

Saxony

252 (1995)

7-59

basins Rifting (Ma)

Midcontinent rift (USA) Richardson Mms (Canada) Ougarta trough (N.Africa)

Tectonophysics

extended basin. The indicated rheological evolution would inhibit extensional as well as compressional basin reactivation after long, tectonically inactive periods. Such a model, which assumes standard rheologies and indicates that basin inversion is unlikely to occur, is contradicted by the observation of repeated basin deformation on time scales of 300 Ma (see Table 1; Ziegler, 1989, Ziegler, 1990). In view of this, it is likely that pre-existing faults and shear zones are most likely marked by a weaker rheology than their surroundings. Therefore, it must be assumed that in the basin centres the strength of the lithosphere is considerably weakened by faulting. 6.3. Mechanisms of lithospheric weakening and localization of deformation A number of possible effects of lithospheric weakening can be described in terms of events intrinsically accommodated in standard models of extensional basin evolution. These include episodic, multiple rifting phases and thermal blanketing which are commonly incorporated in stretching models. The role of pre-existing faults has been analyzed in detail in a modelling study of the Celtiberian Range,

Inversion (Ma) 1039-970 75-45(?) 320-270 50-30 265-248 loo-? 275-260 300-290

stage

Time (Ma)

lapse

50-l 17 305-335 20-70 290-310 85-102 265-? 65-80 IO-30

35-20 85-55 85-57 80-20

69-84 67-97 15-43 O-60

35-20 65-55 40-35 85-55

50-65 115-125 140-145 15-45

which developed by inversion of the Central Iberian basin, by van Wees and Stephenson (1995). In the following we concentrate on the main features of the interplay of processes and strength evolution. Crustal faults and shear zones in the subcrustal mantle are thought to contribute materially to the weakening of the lithosphere. In order to quantify the relative importance of this weakening mechanism, detailed investigations were carried out on the evolution of the Central Iberian basin, a well constrained intra-continental area that is characterized by repeated, superimposed intra-plate deformations, as shown by back-stripped subsidence curves (Fig. 21). These curves, in combination with forward modelling, establish a quantitative framework for the understanding of the evolution of the lithosphere during the Mesozoic extensional subsidence of this basin. During the subsequent Paleogene inversion phase, deformation is largely restricted to the basin. It is therefore expected, that throughout the rifting and inversion phases, the Central Iberian basin was underlain by a relatively weak lithosphere as compared to surrounding areas. To test this assumption, the strength evolution of the lithosphere during the development of the Central Iberian basin was calculated as constrained by the thermal and kinematic

PA

Ziegler

et al. / Tectonophysics

reconstructions of forward modelling of tectonic subsidence and uplift. Fig. 22 gives the results of these calculations for extensional and compressional strength. It demonstrates that the observed localization of Cretaceous basin extension and Paleogene inversion cannot be explained by the thermal evolution of the lithosphere. Therefore, mechanical weakening of the lithosphere through faulting is required to explain the timing and nature of basin inversion during the Pyrenean orogeny. Permo-Carboniferous (late Variscan) faults, probably characterized by a significant reduction of the frictional angle, provide a mechanism for a significant decrease of the lithospheric strength. This is in agreement with the observed localization of the Central Iberian basin and its inversion axis. Lithospheric strength reduction, due to pre-existing faults, probably controlled localized deformation of a large number of intra-plate basins in Europe (Ziegler, 1990). Long-term and intermittent episodes of transtensional and transpressional fault reactivation strongly suggest that sliding on pre-existing faults is favoured over formation of new ones. In the previous section, models were presented which focused on thermal perturbation of the lithosphere, resulting from its extension. This assumption, which is intrinsic to these models, cannot fully address other thermal effects, such as dyke swarms, slab detachment and changes in material composition associated with moving phase boundaries. Such pro-

0

252 (1995)

41

7-59

cesses could have large effects on the strength evolution of the lithosphere, but are difficult to quantify. Although first steps have been made to quantify the role of stresses on melting (Pedersen, 1994) and diapiric processes (DaudrC and Cloetingh, 1994), future work has to integrate these effects with mechanical weakening effects in order to obtain a quantitative understanding of the spatial and temporal evolution of lithospheric strength under such complex, coupled conditions. 6.4. Folding and warping of the lithosphere Although absolute values of horizontal intra-plate stresses are not known by direct information, estimates based on brittle/ductile rheology of the lithosphere suggest that stresses of 75-200 MPa are already sufficient to activate folding of young, crust-controlled lithosphere. An example is the Eurekan folding of the Sverdrup basin in Arctic Canada (Stephenson et al., 1990). Whole-lithosphere folding of older lithosphere requires stress levels in the order of 500 MPa. For areas where prominent folding is not observed, the level of vertically averaged compressional horizontal stress will be typically in the order of a few tens to a few hundreds of MPa, although localized stress fluctuations, such as superimposed bending stresses, may exceed average strength values at certain depth intervals. The presence of horizontal stress will weaken the lithosphere;

TECTONIC SUBSIDENCE CENTER IBERIAN BASIN

I t

, I

II

Ill

I H

Fig. 21. Celtiberian range (for location see Fig. 10): back-stripped and modeled Thin line connecting squares: back-stripped tectonic subsidence curve; intervals breaks in sedimentation. Heavy black line: modeled subsidence curve.

tectonic subsidence curves between white and black

for the Central Iberian basin. squares correspond to major

42

P.A. Ziegler

et al./

Tectonophysics

252 (1995)

7-59

: COMPRESSION

1 EXTENSION w--e y--\ \ ttD

---/ /snlndard

0

HC HIc /’

0

\

-\

0

--e-

0 0 0 0

-\ 0 --_4,:. :. .:.:.:.:, .:.:.:.:, ........... ............ .,.,...,, .,,..: ‘.. . . . ..::. ...:. ... .: ..:. ..:, .: .:.. .. .: .: ............ .. ,.:... .................... .,...,.....,...,.,. ,...,..... ...... . ....,..,., .,.,....,.. .:.::.:.: .:.::.:.: :’:’ .:. .......... ::~~~~!~~~~~~~~~~~~~~~~i ::~~~~!~~~~~~~~~~~~~~~~i ........... . ... .. ... ..... ... :. :./:./::::..: ..: ...... ..\. ..:.::.:.:::. .y.::.:.:::. ::...: ... . ;; ......... .

Fig. 22. Integrated lithospheric strength evolution 21 (Mesozoic rifting event, followed by tectonic

F--

---,@-@@@

rsscovated teulta .. ,.,::_:.:_: ..... ... ....:::: ,., .... .. ::::, ::::, :.: :.: ,::_:.:_: ,., :j: :j: ....... ...., ..:. ..:‘, ..: .. ,,:.:,., ,,:.:,., ,..: ... . ‘y::‘.:. .:::..:. :. :. ;::: ;::: :.:. :.:. .;::.. .;:: :; :; ::.:7: ::.:7: ... _Ii _Ii .~~~:i:si:‘:i:-:i-iii”::~,.i!-ii,~ ~;~~i.i:i:..~~~~~~~~~~.~:~::~.~~~ .::i jNj&@ON jNj&@ON ................... . .. . . . .:.: ...........

of the Central Iberian basin, based on forward quiescence and Paleogene basin inversion).

this can strongly affect effective elastic thickness (EET) values. As demonstrated by Cloetingh and Burov (19951, a stress field of 100 MPa can lead to a 10% reduction of the EET of stable continental lithosphere, with values being much larger for young lithosphere. For example, tectonic stresses of 100500 MPa may decrease the EET of 400-Ma-old, relatively cool lithosphere by 20%, but for 200 Ma old lithosphere this decrease may be as large as 50%. If stresses exceed a limiting value, unstable conditions develop and folding or buckling of the lithosphere can occur. For a constant strain rate, the critical stress threshold generally depends on the thickness-to-length ratio of the competent lithospheric layers, and partly on the distribution of surface loads. The predominant wavelength of folding is

modeling

of subsidence

curve given

j

in Fig.

roughly proportional to a factor of 4-6 of the thickness of the strong layer subjected to folding. Weak layers in the lower crust can permit detachment folding of the upper crust and of the lithospheric mantle. In the case of detachment folding, two dominant wavelengths of deformation can be observed. As upper crustal folding is not associated with the deformation of major density boundaries (assuming absence of strong density contrasts between the upper and lower crust), whereas mantle folding causes deflections of the Moho and produces gravity anomalies (Martinod and Davy, 1992; Burov et al., 1993), disharmonic folding of the lithosphere, giving rise to structures with different wavelengths, can be recognized as such. Fig. 23 displays estimates for wavelengths of lithospheric folding inferred for Arc-

0

200

400

600

800

1000

lzoc

1400

1600

1800

2000

2200

2400

Articcanada

Central Australia whole

lithosphere

folding -600

-700.

-800

-800

-900.

-900

-1000

-1000

Time [Ma] Fig. 23. Diagram depicting dependence between thermo-tectonic age and wavelength of bi-harmonic folding of lithosphere. MSC = mechanically strong part of crust. MSL = mechanically strong part of mantle-lithosphere. Examples: Arctic Canada (young), Central Asia (medium), Central Australia (old) and Transcontinental Arch of North America (very old lithosphere). Gray fields: wavelengths predicted for folding of decoupled crust and mantle lithosphere and for whole-lithosphere folding. Crosses represent observation error bars.

tic Canada (Stephenson et al., 1990) Central Australia (Stephenson and Lambeck, 1985), Central Asia (Martinod and Davy, 1992; Nikishin et al., 1993; Burov et al., 1993; Cobbold et al., 1993) and the Transcontinental arch of North America (Ziegler, 1989). Also shown in Fig. 23 are theoretical relationships between lithospheric age and predominant wavelength of lithospheric folding derived from analytical expressions which relate the wavelength of folding to the thickness of the competent layer(s) (Burov et al., 1993). Fig. 23 illustrates that the wavelength of lithospheric folding is directly related to the thickness of the competent crustal and mantle layers and the EET of the lithosphere. Whereas the thickness of the competent sub-crustal mantle is generally controlled by the age of the lithosphere, the thickness of the rheologically strong parts of the

crust depends on their mineralogical composition, the content of fluids and melts and the concentration of heat producing elements.

7. Lithosphere

of the Alpine foreland

Large parts of the metastable Palaeozoic crustal domain of Western and Central Europe are characterized by crustal and lithospheric thicknesses of less than 35 and 90 km, respectively, an elevated mantle heat flow and an upper asthenosphere displaying anomalously low shear-wave velocities. In contrast, the stable Precambrian East European Platform and the Fennoscandian Shield are characterized by crustal thicknesses of 40-55 km and lithosphere thicknesses of 100-190 km, a low heat flow and an upper

44

P.A. Ziegler

rt al./

Tectonophysics

asthenosphere having high shear-wave velocities. The boundary between these two lithospheric domains coincides with the Tomquist-Teisseyre zone which delimits areas of Palaeozoic crustal consolidation towards the northeast (Cermak, 1982; Suhadolc et al., 1990; Ansorge et al., 1992; Zielhuis, 1992; Zielhuis and Nolet, 1994). The area affected by latest Cretaceous and Cenozoic intra-plate compressional deformations is also delimited to the northeast by the Tomquist-Sorgenfrei line (Ziegler, 1990). The upper asthenospheric low-velocity anomaly, which is observed beneath Western and Central Europe, is restricted to depths of less than 200 km. The magnitude of this anomaly cannot be explained alone by an increased temperature of the upper asthenosphere but requires the presence of partial melts in order to be compatible with the observed surface heat flow (Zielhuis and Nolet, 1994; Ansorge et al., 1992). This suggests that partial melting of volatileenriched mantle source regions, having lower than normal solidus temperatures, has occurred (Wilson, 1993). Therefore, the observed anomaly must be interpreted in terms of either the head of a large, not very energetic mantle-plume, or partial melting of volatile-enriched layers within the mantle-lithosphere which is underlain by an asthenosphere having a somewhat higher than ambient temperature (> 13OO”C), possibly forming part of an upwelling and out-flowing branch of the mantle convection system. The geological evolution of Western and Central Europe indicates that its lithosphere/asthenosphere system was repeatedly disturbed during Palaeozoic, Mesozoic and Cenozoic times. During the Caledonian and Variscan erogenic cycles of crustal consolidation, a number of Gondwana-derived Precambrian continental terranes was accreted to the southern margin of the Fennoscandian-East European craton. This involved the closure of pre-existing oceanic basins, the opening and closure of limited oceanic back-arc basins, and the subduction of supra-crustal rocks, continental and oceanic crust, as well as of substantial volumes of mantle-lithosphere (Ziegler, 1989, Ziegler, 1990). However, neither Caledonian nor Variscan erogenic roots are preserved (Ansorge et al., 1992). Present crustal thicknesses bear little relationship to these fold belts but appear to be controlled by the pattern of Mesozoic and Cenozoic rifts and, in the case of the Tomquist line, by

252 (1995)

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compressional foreland tectonics (Ziegler, 1990). This indicates that in post-erogenic periods the Moho discontinuity was repeatedly destabilized by processes that affected also the mantle part of the lithosphere. Following the late Westphalian consolidation of the Variscan orogen, its deep lithospheric roots were destroyed during the Stephanian-Autunian phase of wrench faulting and widespread magmatism that must be related to a re-orientation of the convergence direction between Gondwana and Laurussia during the terminal phases of Pangea suturing (Ziegler, 1989, Ziegler, 1990). These wrench deformations, which affected the entire Variscan orogen and its northern foreland, presumably were accompanied by the detachment of the subducted lithospheric slab(s) and at least partial delamination of its deep lithospheric roots. Resulting uplift of the orogen and upwelling of the asthenosphere was coupled with the ascent of mafic melts to the base of the crust, causing crustal anatexis and intrusion of granitic melts into the crust, and, in some areas, with magmatic underplating and basification of the lower crust (Dowries, 1993). This, in combination with transtensional uplift of core complexes and erosional unroofing of the crust (e.g., Massif Central: Malavieille et al., 1990; Burg et al., 1994; Black Forest: Eisbacher et al., 19891, resulted in a regional re-equilibration of the Moho discontinuity at depths of about 35-40 km. During the Late Permian and Triassic, the lithosphere of Western and Central Europe relaxed thermally. Consequently the surface of the crust subsided below the erosional base-level and sedimentation resumed in progressively expanding areas. The evolution of Mesozoic basins was governed by rift- and wrench-tectonics, entailing a renewed destabilization of the lithosphere, which culminated in Mid-Jurassic opening of the Western Tethys and in Cretaceous opening of the North Atlantic Ocean. In Western and Central Europe, Mesozoic rifting activity was accompanied by relatively minor magmatic activity only; the volcanic Mid-Jurassic central North Sea arch is an exception (Hendrie et al., 1993). The Late Cretaceous and Paleocene phases of major intra-plate compressional deformation, which must be related to the collision of Iberia and of the Alpine orogen with the metastable platform of Western and Central Eu-

P.A. Ziegler

et al./

Tectonophysics

rope, preceded the Mid-Eocene onset of a renewed cycle of rifting and lithosphere destabilization (Ziegler, 1988, Ziegler, 1990). In areas which were not affected by major Mesozoic rifting, such as the North German Platform, the Bohemian Massif, the Rhenish Shield, the Franconian Platform, the Paris basin and the Massif Central, the lithosphere had apparently stabilized towards the end of the Cretaceous to thicknesses of some loo-120 km. Consequently, the strength of the lithosphere had increased considerably. During and after the Caledonian and Variscan orogenies significant amounts of crustal material were probably incorporated into the mantle-lithosphere by means of syn-erogenic ultra high-pressure metamorphism (phase transitions) and by post-erogenic interaction of basaltic melts with the felsic parts of the lower crust and by delamination of the lower crust in response to the intrusion of melts and their subsequent cooling into the granulite and ultimately the eclogite stability field (Austrheim, 1987, Austrheim, 1990; Downes, 1993; Kern, 1993). The occurrence of felsic lower crustal xenoliths (Downes, 19931, as well as reflection seismic data indicating that in many areas the erogenic fabric of the crust is truncated by the Moho discontinuity (Ziegler, 1995 ‘), suggests that the lower crust can contain a significant amount of felsic material and that it is not necessarily exclusively composed of mafic granulites (Dawson et al., 1986). Moreover, there are reasons for doubting whether on a regional basis the geophysitally and the petrologically defined crust-mantle boundaries always coincide (Mengel, 1992; Kern, 1993). For instance, deep reflection seismic data from the southern North Sea, the Skagerrak (Lie and Husebye, 1994) and the Hebrides shelf (Flack and Warner, 1990; Klemperer and Hobbs, 1991; Morgsan et al., 1994) show reflectors extending deep into the mantle-lithosphere. In analogy with the deep structure of the Pyrenees (Servey Geologic de Catalunya, 1993) and of the Alps (Guellec et al., 1990; Pfiffner, 19921, such reflectors are generally interpreted as

’ For example, Variscan fold belt ( DEKORP, 1988; Vollbrecht et al., 1989; Meissner and Bortfeld, 199O;Bois et al., 1990; Bankwitz and Bankwitz, 19941, Western Baltic (Baldschuhn, 1992) and British Caledonides (Klemperer and Hobbs, 1991).

252 (1995)

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45

being related to shear zones and relict subduction zones along which lower crustal material was inserted into the mantle. Moreover, the refractionseismically defined velocity layering of the lithosphere (Ansorge et al., 19921, as well as the composition of mantle xenoliths, indicate that the mantlelithosphere is characterized by considerable chemical heterogeneities and stratification and apparently includes volatile-enriched zones (Menzies and Bodinier, 1993). Correspondingly, it is questioned whether the sub-continental mantle-lithosphere is exclusively composed of frozen asthenospheric material, the thermal boundary layer of which was metasomatically enriched through time by very-low-fraction melts leaking up from the asthenosphere (McKenzie, 1978; Smith, 1993; Thompson and Gibson, 1994). Only in areas where during an earlier rifting event the mantle-lithosphere was completely removed and replaced by upwelling asthenosphere is it likely that the mantle-lithosphere is exclusively composed of solidified asthenospheric material. In the area of Western and Central Europe, mantle convection patterns presumably changed repeatedly through time, particularly at the end of the Caledonian and Variscan orogenies, during the opening of the Tethys and Atlantic oceans and again in conjunction with the Cretaceous development of subduction zones in the Tethys domain and the evolution of the deep-reaching Alpine roots (Brousse and Bellon, 1983; Ziegler, 1988, Ziegler, 1990; Ansorge et al., 1992). In view of the generally low level of volcanic activity associated with Mesozoic extensional basins, it is speculated that the upper asthenospheric S-wave anomaly of Western and Central Europe came only into evidence toward the end of the Cretaceous and evolved further during the Cenozoic. This notion is in keeping with the occurrence of Paleocene pre-rift dykes (alkali basalts, nephelinites, melilites) in the area of the Massif Central, the Rhine Graben and the Bohemian Massif (Horn et al., 1972; Maury and Varet, 1980; Brousse and Bellon, 1983; Suk et al., 1984; Ziegler, 1994). These magmas were derived either by partial melting of the upper asthenosphere and volatile-enriched zones in the thermal boundary layer of the lithosphere, or from a mantle plume impinging on the base of the lithosphere (Wilson et al., 1995). Emplacement of these dykes coincides with the Laramide phases of compressional intra-plate

46

PA.

Ziegler

rt 01. / Tectonophy.sics

deformations of the Alpine-Pyrenean foreland and may have contributed towards a strength reduction of the lithosphere (Menzies and Bodinier, 1993). The observed upper asthenospheric S-wave anomaly can be interpreted as being related to the head of a mantle plume rising from the 670-km discontinuity which developed in conjunction with the development of an upwelling and out-flowing cell of the convecting mantle beneath the Alpine foreland, possibly in response to a progressive rollback of the subducting foreland slab and thrust-loaded downflexing of the foreland lithosphere (see also Laubscher, 1992). The resultant temperature increase of the upper asthenosphere apparently caused partial melting of volatile-enriched phases in the lithosphere/asthenosphere thermal boundary layer and subsequent convective thinning of the lithosphere; this resulted in a gradual upward displacement of the asthenosphere/lithosphere boundary (Latin et al., 1990; Wilson, 1993; Thompson and Gibson, 1994). Thermal thinning of the lithosphere and its infiltration by partial melts was presumably accompanied by a gradual rise in geotherms, causing progressive weakening of the lithosphere. The resulting strength reduction of the mantle-lithosphere over large areas rendered the lithosphere as a whole more prone to deformation in response to far-field stresses (Cloetingh and Banda, 1992). The localization of major magmatic activity in areas affected by Eocene and younger lithospheric extension and doming suggests a rift-related control on further lithospheric thinning, partial melting and melt extraction, the development of melt diapirs and of magmatic pathways to the surface (Wilson and Downes, 1992). The geochemical and Sr-Nd-Pb isotope characteristics of primitive mafic volcanic rocks associated with the Cenozoic rifts of Western and Central Europe suggest that they were derived by mixing of partial melts originating from the convecting asthenosphere and from the thermal boundary layer of the mantle-lithosphere (Wilson and Downes, 1992; Wilson, 1993). Although the geochemical signature of the lithospheric component varies significantly across the Cenozoic European rift system, presumably reflecting compositional differences between the different lithospheric terranes making up the Variscan fold belt (Wilson and Downes, 1992), the asthenospheric component appears to be rather more

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homogeneous and has geochemical similarities to the source of HIMU oceanic island basalts (subducted oceanic lithosphere, see Hart and Zindler, 1989; Wilson, 1989). This may reflect the involvement of an isotopically distinct mantle plume component in the petrogenesis of these magmas (Wilson et al., 1995) that could be derived either from a Neogene mantle plume or a system of plumes, or equally well from a Permo-Carboniferous or Mesozoic plume input into the lithospheric thermal boundary layer. Alternatively, rather than invoking discrete mantle plume activity, latest Cretaceous and Cenozoic development of an upwelling and out-flowing branch of the mantle convection system, causing progressive delamination of the stratified mantle-lithosphere, and/or entraining previously subducted oceanic crust (Hart and Zindler, 1989) or delaminated lithospheric thermal boundary layer material (Smith, 1993), could be envisaged. However, Wilson et al. (1995) suggest that the simplest model is to invoke the impingement of a perhaps not very energetic HIMU mantle plume; this concept is compatible with results of recent tomographic and xenolith studies in the area of the Massif Central (Sobolev et al., 1995). Nevertheless, the occurrence of Paleocene melilitic dykes in the areas of the Massif Central, the Rhine Graben and the Bohemian Massif, suggests that beneath Western and Central Europe the potential temperature of the asthenosphere had increased at the transition from the Cretaceous to the Cenozoic, about at the same time as the Icelandic hot-spot developed (Morton and Parson, 1988; Ziegler, 1990). On the other hand, it is intriguing to note, that the upper asthenospheric S-wave anomaly is essentially restricted to the Alpine foreland and to areas of Caledonian and Variscan crustal consolidation, possibly containing volatile-enriched zones in the mantle-lithosphere. Whatever the cause of this anomaly may be, its development gave apparently rise to a gradual upwards displacement of the asthenosphere-lithosphere boundary and of the lithospheric isotherms, thus causing a progressive weakening of the lithosphere, rendering it more prone to deformation in response to the build-up of compressional and extensional intra-plate stresses. This may have significantly contributed towards the intense Cenozoic intra-plate deformation of the European Alpine foreland.

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8. Discussion The World Stress Map (Zoback, 1992) demonstrates that large parts of the lithosphere are at present affected by regionally persistent compressional horizontal stress systems. Although a number of geodynamic processes appear to contribute to the build-up of intra-plate horizontal compressional stresses, the pattern and timing of corresponding lithospheric deformations suggest that stresses related to collisional coupling of interacting plates play a dominant role in their development. Moreover, palaeo-stress analyses show that stress patterns can change at the scale of a few million years in response to changes in plate interaction. Although the lithosphere can sustain for periods of up to 10 Ma stresses approaching its rheologically controlled strength, it must be realized that neither the structure nor the strength of continental lithosphere is homogeneous and that its weakest parts are prone to fail once the magnitude of stresses exerted on them exceed their strength (Zoback et al., 1993). The examples cited above indicate that intra-plate compressional and transpressional deformations related to plate collision can result in whole lithosphere buckling, crustal folding and the reactivation of crustal discontinuities, entailing uplift of major basement blocks and inversion of rift- and wrenchinduced sedimentary basins (Nikishin et al., 1993). In continental cratons such deformations can occur at distances of up to 1600 km from a collision front. Furthermore, circumstantial and microtectonic evidence suggests that compressional stresses can be transmitted through continental and oceanic lithosphere over even greater distances than the most distally observed megascopic compressional/transpressional deformations; such intra-plate stresses can induce broad-scale positive and negative lithospheric deflections, giving rise to the upwarping of broad arches and accentuated basin subsidence (Cloetingh et al., 1993). Moreover, far-field compressional stresses can cause reactivation of continent-ocean transition zones and potentially the initiation of subduction zones (Masson et al., 1994). Collision-related compressional intra-plate stresses can apparently interfere with the evolution of active rift systems, impeding their development, and can even cause the abrupt termination of activity along

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sea-floor spreading ridges. This is indicated, for instance, by the apparently stress-related Paleocene impedance on crustal separation between Greenland and the European craton, delaying the opening of the northern North Atlantic-Norwegian-Greenland Sea basin, by the Plio-Pleistocene evolution of the Rhane-Rhine graben system and by the Senonian termination of sea-floor spreading in the Bay of Biscay at the onset of the Pyrenean orogeny (Ziegler, 1988, Ziegler, 1990, Ziegler, 1993, Ziegler, 1994). Amongst different intra-plate discontinuities, rifts which are characterized by a strongly faulted crust appear to be prone to early inversion in response to the build-up of collision-related intra-plate horizontal compressional stresses. Major wrench faults, penetrating much of the crust and possibly extending into the upper mantle, are also prone to compressional reactivation. On the other hand, thrust faults and ancient A-subduction zones appear to be less prone to compressional reactivation; this contrasts to their relatively ready tensional reactivation (see Brewer and Smythe, 1984; Lake and Karner, 1987). Analogue models suggest that reactivation of relatively steeply dipping normal and wrench faults occurs only if the angle between their strike and the trajectory of horizontal compressional stress axis is less than 45” (Nalpas et al., 1995; Brun and Nalpas, 1995). This concept is compatible with the observation that during compressional deformation of the Alpine foreland, differential movements between rigid blocks caused predominantly transpressional deformation of the intervening tensional and transtensional basins and the upthrusting of basement blocks. Depending on the pattern of pre-existing intraplate discontinuities, the boundary between an orogen and a compressionally deformed foreland can become diffuse, as seen, for instance, in the Western Alps (Roure and Colletta, 1996). Under such conditions, a distinction between plate-margin (alpinotype) and intra-plate (germanotype) deformations becomes meaningless (Voigt, 1963). The structural style of inverted basins and associated basement uplifts indicates that these deformations were accompanied by a commensurate amount of crustal shortening. Correspondingly, collision-related compressional intra-plate deformations require, on the one hand, a certain amount of coupling be-

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tween the foreland plate and the associated orogen at the A-subduction zone marking their common boundary in order to facilitate stress transmission; on the other hand, within the foreland plate, a decoupling is required, either at the level of the ductile lower crust or at the base of the lithosphere, in order to compensate for crustal shortening achieved during the inversion of basins and the uptilting of basement blocks along steep reverse faults (Fig. 24). The Rocky Mountains style of foreland compressional deformation (Fig. 6), which bears some similarities with the upthrusted basement blocks of the Bohemian Massif, can be explained by ‘simple shear’ detachment of the brittle upper crust from the mantle-lithosphere at the level of the ductile lower crust. Whereas Egan and Urquhart (1993) proposed in their models for the Rocky Mountains that crustal shortening was compensated locally by ductile thickening of the mantle-lithosphere (pure-shear), it is here proposed that during the development of the Rocky Mountains the mantle-lithosphere was subjected to flexural deformation only and that a commensurate amount of shortening of the mantle-lithosphere was probably taken up at the Cordilleran A-subduction zone. Such a crustal detachment (de’collement) model is essentially compatible with the imbrication of the external massifs of the Western and Central Alps, as evident from deep reflection seismic data (Mugnier et al., 1990; Roure et al., 1990; Pfiffner, 1992). A similar model was proposed for the upthrusted basement block of the Sierra Guadarrama (Fig. 10, see Central Iberia) by Banks and Warburton (1991), though Vegas et al. (1990) recognize thickening of the entire crust in the area of this compressional basement uplift. The distance to which compressional (simple-shear) detachment of the crust from the lithosphere at a lower crustal level can propagate into the foreland of an orogen, controlling upthrusting of basement blocks and inversion of extensional basins, is unknown. In the case of the Rocky Mountains a distance of some 750 km appears to be assured. In contrast to the Rocky Mountains, the crustal configuration of the Danish-Polish and the DonetsDonbass troughs (Stephenson et al., 1993; Chekunov et al., 1992) and of the Midcontinent rift (Cannon, 1992) suggests that during their inversion the thinned crust of a rift can be mechanically significantly

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thickened as a result of ‘pure-shear’ crustal shortening. A similar model is invoked for the inversion of the Celtiberian range (Figs. 10 and 13; Salas and Casas, 1993). Deep reflection and refraction seismic data from the Western Baltic Sea (Fig. 18) show that inversion of the Mesozoic Ronne graben involved shortening of the brittle upper crust by reverse faulting and thickening of the lower crust by distributed ductile shear, resulting in depression of the Moho (‘subversion’, Thybo et al., 1994; Thomas and Deeks, 1994; Makris and Wang, 1994). A rift with a pre-inversion thermally still destabilized upper mantle, can attain, depending on the degree of its inversion, isostatic and thermal stability. This is borne out by the lack of post-inversion subsidence of, for instance, the Polish trough and the Lower Saxony basin. Both basins had undergone a relatively short post-rift evolution prior to the onset of their inversion (Table 1). This suggests that pureshear-type basin inversion involves the entire lithosphere, causing its mechanical thickening, a commensurate depression of isotherms and consequently a cooling of the lithosphere (Figs. 19 and 20; van Wees, 1994). However, for many inverted basins insufficient geophysical data are available to test the applicability of their pure-shear whole-lithosphere compressional deformation. Nevertheless, major inversion features, such as the Pyrenees (Servey Geologic de Catalunya, 1993), and perhaps also the Eurekan fold belt of the eastern Canadian Arctic Archipelago, are associated with an imbrication of the entire lithosphere that is rooted locally in the asthenosphere. The evolution of such ‘passive collision zones’ is governed by far-field stresses and not by forces inherent to the gravitational sinking of the subducted lithospheric slab, as visualized for ‘active’ subduction zones (Bott et al., 1990; Ziegler, 1993). Although essentially synchronous intra-plate compressional/transpressional deformations can affect wide areas, sometimes at considerable distances from the respective collision front, the timing of inversion of the different Mesozoic basins in the Alpine foreland of Europe indicates that rifts, characterized by a moderate stretching factor, forming the ‘ weakest’ link in a chain of crustal discontinuities prone to reactivation, will deform first, provided the orientation of their fault systems and of the maximum horizontal compressional stress trajectories permit

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fault reactivation. This can lead either to an in-sequence ‘inversion progradation’ or an out-of-sequence ‘retrogradation’ of inversion movements (Ziegler, 1987, Ziegler, 1990).

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This concept is illustrated, for instance, by the timing of the main inversion phases of the West Netherlands-Broad Fourteens basin and the Sole Pit basin. During the Laramide phase of foreland com-

COLLISION

CONVERGENCE

CRUSTAL

mlIIl El

SYN-RIFT

EXTENSION

pJJ

SEDIMENTS

POST- RIFT AND SYNOROGENIC SEDIMENTS

Fig. 24. Tectonic model illustrating collisional event (not to scale).

CONTINENTAL

OCEANIC

development

of tensional

intra-plate

m

CRUST

CRUST

discontinuities

UPPER

MANTLE

ASTHENOSPHERE

and their

compressional

&,fommtion

during

a

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pression, the bulk of the intra-plate stresses was absorbed by deformation of the West NetherlandsBroad Fourteens basin, thus shielding the Sole Pit basin (Fig. 12). However, the lithosphere of the West Netherlands-Broad Fourteens basin apparently became stabilized to such a degree during the Laramide deformation phase that only minor additional deformations were possible during the subsequent late Eocene-early Oligocene phase of foreland compression; consequently stresses could then be transmitted into areas located to the northwest, causing the main inversion of the Sole Pit basin (Fig. 13). Inversion progradation requires that basins located more proximal to the collision front are characterized by a pre-inversion mechanically weaker lithosphere than more distal basins. Similarly, it appears logical that the ‘weakest’ basins are prone to early inversion, even if they are located at a considerable distance from the collision front. This is exemplified by the Laramide inversion of the Fastnet-Celtic Sea basin and the virtual lack of contemporaneous deformations in the parallel striking Western Approaches basin (Fig. 12). However, under conditions of continued foreland compression, and as a consequence of early deformation of the ‘weakest’ basin, resulting in strengthening of its lithosphere, stresses building up within the foreland plate may reach the magnitude at which they can cause also inversion of more proximally located ‘stronger’ basins. Inversion retrogradation can be inferred from the OligoceneMiocene main inversion of the Western Approaches basin, which clearly post-dates the Laramide main inversion of the Fastnet-Celtic Sea basin (Figs. 13 and 14). On the other hand, stress reorientations can also underlay the locking of certain basins and the reactivation of others, thus mimicking true inversion pro- and retrogradation. In the overall pattern of inverted Mesozoic grabens and troughs in the Alpine foreland, the non-inverted, northeasterly striking Horn and Gllickstadt grabens form an exception in as much as they are not inverted (Fig. 12). These grabens contain 7-8 km of Permo-Triassic sediments; their rifting stage termnated during the Early to Middle Jurassic, some 90-100 Ma before the onset of compressional deformation of the Alpine foreland. These grabens are virtually surrounded by the Laramide-inverted Fennoscandian Border Zone, the southern part of the

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North Sea Central Graben and the Lower Saxony basin. Although the crustal configuration of the Horn Graben is essentially unknown, the Gliickstadt Graben is associated with a distinct shallowing of the Moho discontinuity. Significantly, also the Permian Oslo Graben, which lacks a thick syn- and post-rift sedimentary sequence, was not inverted during the Paleocene phase of foreland compression, despite the very obvious reactivation of the Fennoscandian Border Zone (Sorgenfrei line) (Ziegler, 1990). Keeping in mind, that the compressional yield strength of a palaeo-rift zone is controlled by its crustal and mantle-lithosphere thickness, its thermal regime (Figs. 16 and 17) and the thickness of its post-rift sequence, it should be noted that both the Horn and Gliickstadt Graben had undergone a postrift stage of some 120 Ma and were covered by relatively thin post-rift Cretaceous series at the onset of the Laramide pulse of foreland compression. This suggests that the magnitude of the Laramide intraplate compressional stresses did not exceed the yield strength of the lithosphere of the Horn and Gllickstadt palaeo-rifts but did exceed the yield strength of the North Danish basin, the southern North Sea Central Graben and the Lower Saxony basin, all of which had undergone only a relatively short post-rifting evolution prior to their inversion. If the above-developed concepts of collision-related intra-plate compressional deformation are essentially valid, then it must follow that the stratigraphic record of corresponding foreland deformations can provide valuable information on the dating of erogenic events that affect the collisional margin of the respective craton. As such, the deformation record of a foreland can complement the stratigraphic record which is preserved in the corresponding orogen. In this context, it ought to be kept in mind, that within erogenic systems the stratigraphic record of earlier erogenic events is often cannibalized during subsequent deformation phases, and that the emplacement of major nappes masks the stratigraphic record preserved in their substratum (e.g., beneath the Austro-Alpine nappes of Austria and Germany). Although compressional intra-plate deformations appear to be most commonly associated with continent-continent collisional, Himalayan-type orogens (e.g., forelands of Ural, Variscan, Marathon-

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Ouachita and Alpine fold belts), they can also be associated with continent-ocean collisional, Pacificor accretion-type orogens (e.g., foreland North American Cordillera). The lack of compressional intra-plate deformations in the foreland of an orogen (e.g., Canadian Cordillera of Alberta and British Columbia; Ziegler, 1969) can be related either to the absence of major crustal discontinuities in the respective foreland which can be reactivated under the prevailing stress field and/or to the lack of mechanical coupling between the foreland and the orogen during the evolution of the latter. Mechanisms governing coupling, or the lack thereof, between a foreland plate and its associated orogen at the A-subduction zone, marking their common boundary, are not yet understood. Possible factors could be the lithospheric configuration of the collisionally involved passive margin (upper- versus lower-plate margin, abrupt crustal thinning versus gentle basement ramp), the availability of a thick passive-margin sedimentary prism, the direction (orthogonal or oblique) and rate of convergence, whether or not crustal delamination of the overriding plate facilitates the emplacement of major basement-cored nappes, and erogenic over-thickening of the lithosphere, causing stress propagation into the foreland. For example, the main phases of compressional deformation of the foreland of the Eastern Alps and the Northern Carpathians appear to correlate with the initial collision phases of their nappe complexes with the North European craton; during the EoceneMiocene phases of emplacement of partly basement-cored nappes on the foreland, these orogens were apparently decoupled from their forelands (Fig. 24). On the other hand, in the Western Alps, where basement-cored nappes play a far smaller role (Roure et al., 19901, coupling between the orogen and the foreland progressively increased during late Eocene and Miocene times, as evident by the uplift of the external Alpine massifs and the inversion of Mesozoic basins in the Celtic Sea-Western Approaches area (Figs. 1 I-14). Whereas the development of the Ancestral Rocky Mountains commenced already during the early phases of collision of the South American with the North American cratons and progressed until the Ouachita-Marathon fold belt was consolidated (Ziegler, 19891, the Rocky

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Mountains came into evidence only during the terminal phase of the Cordilleran orogeny (Oldow et al., 1989). Finally, it should be noted that compressional/ transpressional intra-plate deformations, resulting in the inversion of sedimentary basins, can also occur under non-erogenic settings, as seen in the West and Central African rift system (Fig. 9). Similarly, the Late Jurassic-Early Cretaceous polarization of the Northwest European rift systems was associated with wrench-induced partial inversion of, for instance, the West Netherlands and Broad Fourteens basins (Ziegler, 1990). Moreover, the Neogene partial inversion of the Mid-Norway shelf is thought to be related to wrench movements along the continent-ward extension of the Jan Mayen fracture zone (Fig. 10). On the other hand, the general scarcity of compressional deformations along passive margins suggests that ridge-push forces, even if enhanced by a ridgecentred hot spot (e.g., Iceland) (Bott, 19931, are, on their own, not responsible for major intra-plate deformations but may contribute towards the deformation of the lithosphere if they act in constructive interference with other far-field stresses. Overall, circumstantial evidence suggests that the most important intra-plate compressional/transpressional deformations, involving large-scale basin inversion and the upthrusting of major basement blocks, often deep in the interior of continental cratons, are indeed related to collisional plate interaction.

Acknowledgements The authors extend their thanks to Prof. R. Triimpy (E.T.H. Zurich) for his constructive comments on an earlier version of the manuscript. Prof. J-P. Brun (University of Rennes), Dr. L.I. Lobkovsky (Russian Academy of Sciences, Moscow), Prof. A.M. Nikishin (Moscow State University), and Dr. F. Roure (Institut Frangais du P&role) are thanked for their critical reviews and helpful comments to this paper. We are grateful to Shell Intemationale Petroleum Mij. B.V. for releasing text figures 11 to 14 for publication. The strength envelopes given in Figs. 16 and 17 were kindly provided by the Tectonics group of the Vrije Universiteit Amsterdam.

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