Plate tectonics, plate moving mechanisms and rifting

Tectonophysics, 215 (1992) 9-34 Elsevier

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9

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Plate tectonics, plate moving mechanisms and rifting Peter A. Ziegler Geological-Palaeontological (Received

Institute, Vnicersity of Basel, Bemoullistr. 32, CH-4065 Basel, Switzerland

March 5, 1991; revised version

accepted

November

28, 1991)

ABSTRACT Ziegler, P.A., 1992. Plate tectonics, plate moving mechanisms and rifting. Volume III. Thematic Discussions. Tectonophysics, 215: 9-34.

In: P.A. Ziegler

(Editor),

Geodynamics

of Rifting,

Analysis of a sequence of palaeo-reconstructions of the distribution of continents in the Western Hemisphere suggests that frictional forces exerted on the base of the lithosphere by the slowly convecting sub-lithospheric upper mantle play an important role as a driving mechanism of plate movements. With such a scenario. slab-pull and roll-back, ridge-push and deviatoric tensional stresses, related to upwellings of the asthenosphere as well as to lithospheric over-thickening in erogenic belts, can be considered as secondary, albeit important, plate moving forces. The present stress state of the globe and circumstantial geological evidence suggest that major compressional stresses can be transmitted over great distances through continental and oceanic lithosphere. Assembly of major continental masses (like Pangea) probably has an insulating effect on upper-mantle convection systems, causing their decay and reorganization. Development of new asthenospheric upwelling systems under mega-continents causes their break-up by development of deviatoric tensional stresses in the lithosphere and by exerting drag forces on its base. Extension of the lithosphere, culminating in its failure, is followed by passive advection of asthenospheric material into the space opening between the diverging plates to which it is accreted as oceanic lithosphere. At the same time developing ridge-push forces contribute to plate divergence. However, activity along seafloor spreading axes can terminate abruptly if far-field compressional stresses impede further divergence of the respective plates. This may explain the nearly contemporaneous decay of seafloor spreading axes in often distant areas during periods of plate boundary reorganization. Assuming a finite globe, the generation of new oceanic lithosphere at seafloor spreading axes has to be compensated for elsewhere by the subduction of commensurate amounts of oceanic lithosphere and/or shortening of continental crust and subduction of lower crustal material and the subcrustal lithosphere. Plate interaction, driven by mantle convection systems and their changes, ridge-push and slab-pull and roll-back, plays probably an all-important role in the development of intra-continental rift systems, the opening of new oceanic basins and the inception and development of subduction zones. Mantle plumes rising from the deep mantle do not appear to play a significant role in the development of intra-continental rifts though they probably contribute to a weakening of the lithosphere. The relative importance of the various processes contributing to the movement of lithospheric plates differs probably during the assembly and break up of Pangea-type mega-continents.

Introduction

During the last decade, compilation of a global inventory of seafloor magnetic anomalies, palaeomagnetic, palaeo-biogeographical and palaeoclimatological data, combined with the development of computer programs permitting integration of these expansive data sets, has facilitated

Correspondence ningen,

to: P.A. Ziegler,

Kirchweg

41, CH-4102

Bin-

Switzerland.

0040-1951/92/$05.00

0 1992 - Elsevier

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the construction of global palaeogeographical base maps that give the distribution of continents during Phanerozoic times. Such maps have greatly contributed towards the understanding of the drift patterns of cratons and the opening and closing of ocean basins (Smith and Briden, 1977; A.M. Ziegler et al., 1979; A.M. Ziegler, 1981; Scotese, 1986; Scotese and Barrett, 1990; Scotese and McKerrow, 1990; Scotese et al., 1980, 1987, 1988; Zonenshain et al., 1990). A review of Phanerozoic palaeo-reconstructions of the continents in the Western Hemi-

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sphere, with an emphasis on tectonic aspects, such as the evolution and collapse of erogenic belts, the development of sedimentary basins and particularly the evolution of rifts in time and space, indicates that plate interaction is a major factor controlling the development of rifts (Ziegler, 1989, 1990). The observed drift patterns of major continental blocks are difficult to explain purely in terms of ridge push, slab-forces and deviatoric tensional stresses developing over asthenospheric upwellings, and also as a cons
spreading axes are not shown as their locatioil can only be determined from xeafloor magnetic anomalies from Mid-Mesozoic times onward. Ordovician to Late Carboniferous rcconstructiom arc based mainly on palaeomagnztic data and on palaeo-climatological and palaeo-biogcographicai criteria (Ziegler. 1989: Scotesc and McKerrow 1989; Zonenshain et al.. 1990). The configuratiott of Pangea is based on the fit ot continents after closure of Mesozoic and Cenc.rzoic oceanic domains, taking into account &I_ position <,t ihr: geophysically defined ocean-cckntinent transition zones as mapped by Shell Intrrntrtu~mdc~Perroltwt~ Mij.‘s Global Geology Study Group. The latitudenal position of Pangea is given by palaeomagnctlc and palaeo-ciimatological data. The Mesozoic anli Cenozoic reconstructions are based on a global inventory of seafloor magnetic anomalies (Candc et al., 19891 and on palaeomagnetic data. The palaeotectonic reconstructions presented here arc based on the assumptions that seafloor magnetic anomalies provide a record of oceanic lithosphere generation at spreading ridges f Arkani-Hamed, 1991), and that during Phanerozoic times the magnetic poles did not undergo major long-term excursions (lack of true polar wandering, long-term constant spin-axis of the globe: f<,r an alternate view see Storetvcdt. 1990, 19921 These assumptions entail the risk of making mistakes in the assessment of plate motions and also of the geodynamic processes that govern them. Nevertheless, and in spite of some nagging doubt?, about certain aspects of the generally accepted global tectonic concepts. these form the basis fctr the discussion of the plate boundary rcorganizutions given below. These paiacc,-r~constructjoni describe a sequence of fundamental changes in plate movements and interaction that demand a review of the currently favourcd driving mechanisms of plate motions.

reconstructions Main phases of plate boundary reorganization

Figures 1 to 4 give reconstructions of the distribution of continents in the Western Hemisphere during Early Ordovician to Pliocene times. On these reconstructions the patterns of active and inactive fold belts, as well as of rifts and major wrench fault systems is indicated. Seafloor

This summary account of plate motions during Phanerozoic times is essentially an excerpt from Ziegler (1988, 1989, 1990), to which the reader is referred for further details and a comprehensive reference list. For the sake of brevity, the folluw-

PLATE

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ing discussion desists from giving detaiIed references. for each of the reconstructions which are, to a large extent, similar to those given by Scotese and McKerrow (1989) and Scotese et al. (1987). The reconstructions shown in Figs. l-4 were prepared with the aid of an Evans and Sutherland PS 300 vectorgraphics terminal. An orthographic projection was chosen to simulate a readily readable view of the globe; the maps are centred on Northwest Europe. Following the Late Precambrian GrenviilianDalslandian orogeny, during which LaurentiaGreenland and Fennoscandia-Baltica became sutured along the Arctic-North Atlantic Grenvillian fold belt, wrench and rifting activity culminated during the latest Vendian in crustal separation between Greenland and Fennoscandia and the rapid opening of the Iapetus Ocean. During the latest Precambrian and early Cambrian PanAfrican (Cadomian, Baikalian) erogenic cycle Gondwana became assembled.

Already during the Middle to Late Cambrian, opening of the Iapetus Ocean had ceased and Fennoscandia-Baltica began to converge with Laurentia-Greenland. This marked the beginning of the Caledonian erogenic cycle, spanning Mid-Cambrian to Late Silurian-earliest Devonian times (Fig. 1). Ordovician and Silurian convergence and collision of Laurentia-Greenland and Fennoscandia-Bahica resulted in their suturing along the Arctic-North Atlantic Caledonides and the creation of the mega-continent Laurussia. This was paralleled by rifting activity along the northern margin of Gondwana and the successive detachment from it of continental terranes. At the same time, northward subduction of the Proto-Atlantic-Proto-Tethys Ocean along an arc-trench system, paralleling the southern margin of the newly forming Laurussian mega-continent, was accompanied by the transfer of a number of Gondwana-derived terranes across this ocean and their accretion to the southern margin of Laurussia. During the early phases of the Caledonian erogenic cycle, the Barentsia craton was accreted to the northwestern margin of

Fennoscandia and by the Late Ordovician-Early SiIurian, the Arctic craton collided with the northern margin of Laurentia-Greenland. In Ordovician and Silurian times, Laurentia-frreenland underwent only minor latitudinal motions, whereas Gondwana drifted over the South Pole in a sinistral, rotational fashion. Rifting activity along its northern margin presumably involved important shear movements. During the late phases of the Caledonian orogeny the Laurentian craton was subjected to compressional intra-plate stresses, causing broad lithospheric deflections and local upthrusting of basement blocks, for instance in the Hudson Bay.

Late Silurian -Early

Carboniferous

During the Late Silurian-EarIy Devonian, an important reorganization plate boundaries occurred. It involved abandonment of the ArcticNorth Atlantic Caledonides subduction system, inception of the intra-oceanic Sakmarian arctrench system (to the east of Fennoscandia-Baitica), and development of a sinistral mega-shear system transsecting the Arctic-North Atlantic Caledonides along their axis. At the same time, the drift pattern of Gondwana changed and it began to converge in a clockwise rotational fashion with Laurussia (Fig. 1). During the Devonian and earliest Carboniferous, continued northward subduction of the Proto-Atlantic-Proto-Tethys plate, possibly at variable rates, was associated with intermittent back-arc extension in the domain of the Variscan geosynclinal system and the opening of limited oceanic basins. During the Middle Devonian Acado-Ligerian diastrophism, additional Gondwana-derived continental terranes were accreted to the southern margin of Laurussia. Major sinistral translations between Laurentia-Greenland and Fennoscandia-Baltica accompanied the Devonian-earliest Carboniferous suturing of Laurussia and the Arctic craton along the Inuitian-~monosov fold belt. In the domain of the oceanic Sakmarian back-arc basin, a phase of back-arc compression, at the transition from the Early to the Middle Devonian, was followed by back-arc extension during the Give-

Laur

1

ACTIVE FOLDBELTS

1 OCEANICCRUST CONTINENTAL

Fig. 1. Tentative

reconstruction

-

CRATONS of continents

in Western

Hemisphere

INACTIVE FOLD BELTS during Early Ordovician

to Mtddlr

lkvonian.

PLATE

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PLA’l‘t

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AND

Klf”r‘lNG

ACTIVE FOLDBELTS -

INACTIVE FOLD BELTS

Fig. 2. Tentative reconstruction of continents in Western Hemisphere during Late Devonian to Eady Permian.

a

,

1 OCEANICCRUST CONTINENTAL

ACTIVE FOLDBELTS =

CRATONS

Fig. 3. ‘Tentative reconstruction

of continents

in Western

INACTIVE

Hemisphere

FOLD BELTS

durmg Late Permiau to tariv

( r~t~wous

r--:1

OCEANIC CRUST CONT~N~~JAL CRATONS

Fig. 4. Tentative reconstruction of continents in Western Hemisphere during Middle Cretaceous to Pliocene.

tian and a resumption of back-arc compression during the earliest Carboniferous (Fig. 2). During Gedi~n~an to Frasnian times, Laurentia-Greenland underwent only minor latitudinal displacements. However, as a consequence of collision of Gondwana with the southern margin of Laurussia in the area of Iberia during the Famennian, Laurussia (including the Laurentia-Greenland craton) began to move northward in a clockwise rotational fashion. From the Middle Devonian to the Early Carboniferous, Laurussia was rimmed by active erogenic belts and showed no evidence of internal extension (Fig. 21. Carboptiferous-Early

Permian

The latest Devonian-earliest Carboniferous collision of Gondwana with Laurussia marked the beginning of the Herdsman erogenic cycle, during which the western parts of the Proto-Tethys and the Proto-Atlantic oceans became closed. Increasing collisional coupling between Gondwana and Laurussia was accompanied by their joint clockwise rotation and northward drift. The resulting reorganization of plate boundaries is expressed in Western and Central Europe by the Early Carboniferous compressional overpowering of the long-standing Variscan geosynclinal backarc extension systems and by Late Westphalian consolidation of the Variscan fold belt (Fig. 2). The latest Carboniferous-Early Permian Alleghenian consolidation of the AppalachianMauretanides fold belt was accompanied by a modification of the convergence direction between Gondwana and Laurussia, giving rise to the development of an E-W trending dextral shear system transsecting the newly consolidated Variscan fold belt; this shear system represented a diffuse plate boundary between Africa and the Fennoscandia-Baltica subcontinents. Lithospheric thickening in the Appalachian-Mauretanides fold belt was coupled with major intraplate cumpressionai deformation of the Sahara Platform and of the southern parts of Laurentia. Along the western margin of Lauren&, the Middle Devonian-Early Carboniferous Antler orogen became inactive at the end of the Mississippian and the Havalla intra-oceanic arc-trench

system began to evolve: to the north it probably was linked with the Uralian subduction system. In the Arctic domain, the ~~rer~ynia1~plate reorganization is expressed by early Visean termination of erogenic activity in the lnuitian fold belt and the onset of tensional subsidence of the Sverdrup Basin, as well as by Late Carboniferous tensional reactivation of the Arctic-North Atlantic mega-shear, causing development ot the Norwegian-Greeniand Sea rift system: its car& rifting phases may be related to the build-up oi intra-plate compressional stresses emanating from the Variscan orogen. The Carboniferous clockwise rota&i(~~~ of Larirussia was paralleled by the convergence of the Siberian and Kazakhstan cratons with each othct and with the eastern margin of the Fennoscandia-Baltica subcontinent (Fig. -71.‘Their Late Carboniferous-Early Permian ~o~~~sionwith the passive margin of the Mos~w-Barents Sea Platform marked the onset of the Uralian erogenic cycle, which culminated during the Late Permian- Early Triassic in the con~lidation of the ~~ral-N[JYa~~ Zemlya-Taim~r and the Kol~~-Tom--Ku%~ss fold belts (Fig. 2 and 3). With the Late Carboniferous and Early Per. mian consolidation of the Variscan and Gppalachian fold belts, which form the magasuturu between Gondwana and Laurussia, their underlying long-standing subduction systems were abandoned. At the same time, activity along the subduction system rimming the Pacific (Panthalassal margin of the newly assembled Pangea supercontinent increased (Fig. 3; e.g., Sonoma orogeny of Cordillera, Cape fofd belt orogeny). This represents a first-order plate-boundary reorganization.

The ~st-Her~n~an plate reor~n~~atio~ was coupled with a Late Permian and Triassic 4” counter-clockwise rotation of Pangea around ;I pivot located in the Gulf of Mexico. This was followed by the Early Jurassic Rorthward drift of Pangea. During the Late Permian, Triassic and Early Jurassic, rifting along the northeastern passive margin of Gondwana resulted in the separa-

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tion of a number of terranes, referred to as Cimmeria (such as Central and East Iranian and Tibet blocks). At the same time, rifts propagated southward into areas between Africa and Madagascar and India and Australia (Fig. 51. Continued northward subduction of the Proto-Tethys along the southern, Asian margin of Pangea, and development of the Mesozoic Tethys seafloor spreading axis, was coupled with the northward transfer of the Cimmeria terranes and their Late Triassic-Early Jurassic accretion to Asia. In the Black Sea area, Late Permian-Early Triassic back-arc extension, presumably controlled by the decay of the Variscan subduction system, caused opening of an oceanic basin, which closed again during the Early Jurassic. During the Late Permian and Triassic, the interior of Pangea was broken up by the gradually westward propagating Tethys rift system and the southward propagating Norwegian-Greenland Sea rift system. By the Late Triassic, these rifts had propagated through the North and Central Atlantic into the area of the Gulf of Mexico. Particularly during the Triassic and Early Jurassic, large areas around these future divergent plate boundaries were affected by tensional stresses (Fig. 31. In the Uralian domain, erogenic activity continued into the Early Triassic with a last pulse occurring in its northern parts at the transition from the Triassic to the Jurassic. Back-arc extension in the area of the West Siberian Basin commenced during the Late Permian and persisted into the Early Jurassic, when abandonment of the Uralian subduction system completed the assembly of a new mega-continent, referred to as Laurasia. This was paralleled by the development of a new arc-trench system along the eastern and northern margins of Siberia, which presumably was connected to the west with the Cordilleran subduction system. As such, Pangea was encircled by active erogenic belts, with exception of the tension-dominated northeastern margin of the Gondwana sub-continent (Fig. 3). Late Triassic and Early Jurassic accelerated crustal distension in the Gulf of Mexico-Central Atlantic-Tethys rift zone, accompanied by the extrusion of flood basalts in the Central Atlantic

domain, culminated in the development of a new, discrete divergent/transform plate boundary between Gondwana and Laurasia during the Middle Jurassic; this boundary is largely superimposed on the axes of the Hercynian fold belts, along which Laurussia and Gondwana were welded together (Fig. 3). With this Mid-Mesozoic reorganization of plate, boundaries a new kinematic regime was established, which governed the further break-up of Gondwana and Laurasia. Late Jurassic-Cretaceous

Crustal separation between Gondwana and Laurasia was achieved during the Late Jurassic. Late Jurassic-Early Cretaceous rapid opening of the Central Atlantic Ocean was accompanied by an 8” clockwise rotation of Laurasia, major sinistral translations between Africa and Europe, and the transtensional opening of the oceanic Piedmont-Penninic basin in the Alpine domain. At the same time, space constraints developed in the central Mediterranean area as a consequence of the translation of the Italo-Dinarid block relative to Europe; these plate movements underlie the development of a new subduction system and the closure of the Jurassic Hellenic-Dinarid ocean (Figs. 3 and 4). Moreover, stress regimes governing the evolution of Western and Central European rift systems changed as a result of crustal separation between the Italo-Dinarid Block and Europe. Late Jurassic and Early Cretaceous seafloor spreading in the Central Atlantic was only partly compensated by opening of the oceanic Piedmont-Penninic basin of the Alpine domain; much of the sinistral motion between Africa-Arabia and Europe was taken up along an intra-continental transform fault zone, partially separating the Italo-Dinarid block from Africa. During the Cretaceous collision of the Italo-Dinarid Block with Europe, the former rotated counter-clockwise and was internally deformed. Increasing Late Jurassic-Early Cretaceous collisional coupling of Europe and Africa in the Mediterranean domain impeded further clock-wise rotation of Eurasia (see latitudinal position of Caspian Sea; Figs. 3 and 41, whereas sinistral motions between Africa

ix and Europe persisted. However, clock-wise rotation of the North American craton, paralleled by rapid opening of the Central Atiantic, continued during the Cretaceous. This was accompanied by northward rift and later seafloor spreading propagation into the North Atlantic domain, the Norwegian Greenland Sea and the Labrador SeaBaffin Bay. Crustal distension across the Norwegian-Greenland Sea rift was compensated by dextral movements along the trans-Arctic shear zone and rifting along the western margin of the Lomonosov Ridge. In the Arctic domain, rifting culminated during the Valanginian in crustal scparation between the Laurentia-Greenland subcontinent and the Alaska North Slope-Chukotka and New Siberian Island blocks, their counterclockwise rotation away from the Canadian Arctic shelf, and the opening of the oceanic Canada Basin; these rotating blocks collided with the Pacific arc-trench system (Figs. 3 and 4). During the Late Jurassic and Early Cretaceous, Go~dwana remained more or less stationary but was internally deformed by new rift systerns, such as those paving the way for the opening of the South Atlantic and the Indian Ocean. At the same time, Africa, and to a lesser degree also South America, became dissected by wrench faults and rift systems (e.g. Central African rifts, Amazon fracture system; Niirnberger and Miiller, 1991; Binks and Fairhead, 1992; Guiraud and Maurin, 1992). The Late Jurassic-Early Cretaceous plate bounda~ reorganization underlies the MidCretaceous to Early Cenozoic break-up phases of Laurasia and Gondwana and the early phases of Alpine suturing of Africa-Arabia and Eurasia. In the Arctic-North Atlantic rift system. crustal separation was achieved during the Early Cretaceous in the North Atlantic area, during the Senonian in the Labrador Sea, and during the earliest Eocene in the Baffin Bay, between Greenland arid Eurasia and between the Lomonosov Ridge and the Barents Shelf. The Senonian onset of seafloor spreading in the Labrador Sea coincides with the termination of seafloor spreading in the Canada Basin, the Bay of Biscay and the Rockall Trough. The first step in the break-up of Gondwana

was its Late Jurassic split into a western half. consisting of Africa and South America, and an eastern half comprising Madagascar, India, Antarctica, Australia and New Xealand. Rifting, preceding crustal separation, was accompanied by the extrusion of the massif Karoo and Parana-Etendeka flood basal& this suggests that rifting took place in an area underlain by anomalously hot, upwelling asthenospherc, !?r that mantle plumes existed beneath the evoiving rifts fWhitc and McKenzie, 1989; Campbell and Griffiths. 19901. In the southern South Atlantic. seafloor spreading commenced during the Barremian and propagated northward during the Aptian. Crustal separation between Africa and South America. along the Equatorial transform t’ault system was achieved shortly thereafter; with this, the seafloor spreading axes of the Central and the South Atlantic linked up. Some of the Central African rifts remained active into Early Cenozoic times. however. Following Mid-Cretacc~~us separation of Madagascar and India from Eastern Gondwana, rapid opening of the South Atlantic-lndia~~ Ocean was accompanied by a counter-clockwise rotational northward movement i,f Africa. This motion of the Africa-Arabia craton underlies the Alpine plate reorganization of the Tethys domain. during which collisional plate boundaries propagated rapidly into the Western Mediterranean area. The Campanian abandonment of the Bay of Biscay seafloor spreading axis coincides with the onset of the Pyrenean progeny during which oceanic l~th~~sphere of the Bay of Biscay was subducted. Activation of the North Pyrenean-Cantabrian subduction zone can be rclated to the build-up of a regional compressional stress field as a consequence of the northward drift of Africa (,Fig. 4).

Cretaceous and Cenozoic opening of the South, Central and North Atlantic was paralleled by major erogenic activity in the Cordilleran and Andean fold belts and the development of the Caribbean arc-trench systems. By mid-Eocene time the western margin of the North American craton apparently approached the East Pacific

I’LATt

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Rise and overrode it during the Neogene. This resulted in the development of the Basin-andRange extensional system that does not find its equivalents in the Canadian and Alaskan Cordillera nor in the Andes; these, in part still active fold belts are separated from the Pacific seafloor spreading axis by a variable width of oceanic lithosphere, ranging in age from Eocene to Recent (Cande et al., 1989). During the Late Cretaceous to Paleocene Early Alpine erogenic cycle, progressive closure of oceanic domains in the Western and Central Mediterranean area, followed by the full-scale collision of Africa-Arabia and Europe, was accompanied by important intra-plate compressional deformations both in Europe and in North Africa. The Eocene-Oligocene Main Alpine orogenie phases coincide with a further phase of pIate-boundary reorganization in the ArcticNorth Atlantic domain. Earliest Eocene onset of seafloor spreading in the Baffin Bay, the Norwegian-Greenland Sea and in the Eurasian Basin was accompanied by continued clock-wise rotation of North America and Greenland relative to Eurasia; this was paralleled by compressional deformation of the eastern Sverdrup Basin, northwestern Greenland and the western margin of the Barents Shelf. The resulting Eurekan and Spitsbergen orogens are of a passive collisional nature and are not associated with long-standing subduction zones. Upon early Oligocene abandonment of the Labrador Sea-Baffin Bay seafloor spreading axis, these orogens became inactive. With the late Oligocene stabilization of spreading systems in the Norwegian-Greenland Sea, crustal separation was achieved between northeastern Greenland and the Barents Shelf (Fig. 4). The Late Palaeogene and Neogene plate reorganization in the Tethys domain can be reIated to gradually increasing dextrai translations between Europe and Africa during the Main and Late Alpine erogenic phases; these were paralleled by differences in the rates of seafloor spreading in the Central and Arctic-North Atlantic. This plate reorganization was accompanied by the Miocene and Pliocene development of important intramontane wrench systems, the subsidence of the extensional Tyrrhenian, Aegean and Pannonian

10

back-arc basins, opening of the oceanic AlgeroProvensal Basin, and concentration of crustal shortening to within the Western Alps, the Apennine-Calabrian arc, the Carpathians and the Hellenic arc. During the late Eocene and Neogene, the East African-Red Sea-Gulf of Suez rift system, which finds its continuation in the rifts of Libya and the Pelagian Shelf, came into evidence. Development of this mega-rift system was paralleled by the evolution of the Rhine-RhBne graben system, which extends southwestwards through the western Mediterranean and crosses the Rif fold belt. Development of these rift systems may herald a post-Alpine plate reorganization that may ultimately lead to the disintegration of the present continent assembly. The Mid- to Late Eocene plate boundary revolution is expressed by a nearly gIoba1 change in lithospheric stress fields, plate motions and plate interaction (e.g., South Atlantic shear zones, La Breque and De Souza, 1991; collision of India and Eurasia, Scotese et al., 1988; reorganization of seafloor spreading axes in Indian Ocean, Patriat and Segoufin, 1988; inception of East African and Rhine rift systems; onset of development of Basin-and-Range extensiona system, Keith and Wilt, 1985, Jones et al., 1992). Underlaying causes may be seen in a partial decoupling of the lithosphere from the asthenosphere and of the latter from the Earth’s core in response to deceleration of the Earth’s rotation during the last 40 Ma; apparently this entailed a westward shift of the lithosphere relative to the Earth’s core, as suggested by the hot-spot reference frame (Doglioni, 1990, 1991; Ricard et al., 1991). This summary illustrates that the Palaeozoic to Recent evolution of the Western Hemisphere was governed by a sequence of plate boundary reorganizations during which a broad spectrum of geodynamic processes was responsible for the development and subsequent partial to total destruction of sedimentary basins. Amongst these, rifts play a major role. Some examples of such rifted basins are discussed in papers contained in the Tectonophysics volumes 208 and 213 which are dedicated to the review on “Case History Studies on Rifts”.

Plate moving forces

The present stress regime of the globe, as documented by the World Stress Map (Zoback et al., 1989), as well as current plate movements can be adequately explained in terms of ridge push. slab forces and collisional resistance and do not require a significant contribution from sublithospheric mantle flow (Richardson, 1992). However, the Phanerozoic motion and interaction of major lithospheric plates, illustrated in Figs. l-4 and summarized above, are difficult to reconcile with the concept that ridge push, slab-forces and deviatoric tensional stresses developing over mantle plumes and associated with zones of orogenie lithospheric over-thickening are the only plate driving forces (Bott and Kusznir. 1984: Dewey, 1988a; Park, 1988; Bott et al., 1990; Bott. 1992). Specific questions are raised, for instance, by the behaviour of Laurentia-Greenland, which stayed almost stationary during the Ordovician to Devonian time span of some 125 Ma, even during its Caledonian suturing with Fennoscandia-Baltica, and then underwent a 55” clockwise rotation during the Carboniferous upon collision of Gondwana with the southern margin of Laurussia. Similarly, the 40” counter-clockwise rotation of Pangea during the Permian and Triassic, and its Early Jurassic northward drift, which was paralleled by the development of its internal rift systems, cannot be explained by these standard plate moving forces. Roll-back of the subducted iithospheric slab. referred to also as trench suction, is thought to give rise to intra-plate tensional stresses, particularly if the respective continent is rimmed by active subduction zones (Bott, 1982; Dewey, 1988b). This model has been applied to the devclopment of the Mesozoic intra-Pangean rift systems, evolution of which was paralleled by major erogenic activity along most of the margins of Pangea. However, the available geological record indicates that during major erogenic cycles compressional stresses are often transmitted over large distances into continental cratons where they can give rise to broad-scale lithospheric deflections (upwarping of arches, stress-induced

subsidence of basins; Cloetingh. 1988) and the reactivation of crustal discontinuities (upthrusting of basement blocks, basin inversion; Ziegler, 1988. 1990). Examples are the Late Paiaeozoic deformation of the North African craton and the dcvelopment of the Ancestral Rocky Mountains of the Southern U.S.A. during the collision tri Gondwana and Laurussia (Kluth, 1986; Ziegler. 1989), the Late Cretaceous-Paleocene intra-platc deformation of Europe during the early phases of the Alpine orogeny, the Laramide deformation of the U.S. Rocky Mountains and the Cenozoic dcformation of the Argentinian Andean foreland (Ziegler, 1990; Peterson, 1986). Jntra-plate compressional deformation of this type is associated with Himalayan-type (continent--continent coilision) and Andean-type (continent-ocean co&sion) orogens (Uyeda, lY82j. The recorded patterns of intra-plate compressional deformation are in keeping with the present-day stress field of the globe, which shows that large parts ot the cratons are affected by consistently oriented corn-pressional stresses (Zoback and Zoback, 1980: Klein and Barr, 1986; Zoback et al., 1989). The present stress fields are generatly related to ridge-push whereas the role of collisional coupling is less clear (Richardson, 1992). Summarizing the above, the occurrence of major intra-plate compressional features during cycles of increased erogenic activity casts doubt on &hevalidity of the slab roll-back concept as the cause of large-scale intra-plate extension. Similarly, slab-forces and ridge push do not seem to offer a satisfactory explanation for the Permo-Triassic rotation of Pangea. The same applies for the clockwise rotation of Laurasia during the Late Jurassic-Early Cretaceous opening phase of the Central Atlantic Ocean, the MidCretaceous termination of this motion and the subsequent northward propagation of rifting and later of seafloor spreading into the Arctic-North Atlantic domain. These events, which were accompanied by continued clockwise rotation ot Laurentia-Greenland and major erogenic activity in the U.S-Canadian-Alaska Cordillera, as well as in the Alpine-Mediterranean domain, demand a unified kinematic explanation.

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RIFTING

Subduction zones

The palaeo-reconstructions, given in Figs. 1 to 4, indicate that subduction zones migrate with the drifting plates. Earthquake data indicate that subducted slabs do not penetrate the f670 km discontinuity, and that they are therefore not rooted in the deep mantle (Thompson, 1991). Slab-forces include the negative buoyancy of the subducting slab and the resistance of the slab entering the mantle. The sum of these two large forces, which counteract each other, is exerted on the colliding plates (Richardson, 1992). Slab-pull and roll-back (trench-suction) of the subducted oceanic lithosphere slab can account for the gradual consumption of oceanic lithosphere along the leading edge of a drifting continent, and can possibly explain its movements (e.g. Mesozoic evolution of U.S. and Canadian Cordillera and opening of the Atlantic). On the other hand, subduction zones located at the trailing edge of a drifting continent (Late Devonian to Carboniferous development of Appalachian orogen, Fig. 2; Carboniferous-Triassic southern margin of Pangea, Fig. 51, cannot account for the motion of that continent. Furthermore, slab-pull cannot account for the Neogene overriding of the active East Pacific Rise seafloor spreading axis by the Laurentian craton (Coney, 1987; Verall, 19891, nor for the projection of the active Carnegie and Chile Rise seafloor spreading axes into the Andean arc-trench system (Cande et al., 1989). The dip angle of subducted lithospheric slabs generally increases with increasing age of the subducted oceanic lithosphere (see e.g. Vlaar and Wortel, 1976). On the other hand, the dip of Benioff zones appears to decrease with increasing convergence rates (Uyeda, 1981, 1982; Uyeda and Kanamori, 1979; Hsui and Toksiiz, 1981; Cross and Pilger, 1982; Zonenshain and Savostin, 1981; Jarrard, 1986; Nakamura et al., 1990; Bott, 1990), whereby the length of subducted slabs (measured in the dip direction) increases with increasing convergence rates (Forsyth and Uyeda, 1975). Furthermore, the steepness of a subducted slab may depend on whether it dips in the direction of general mantle flow (shallow angle) or opposite to it (steep angle; Doglioni, 1990, 1991). Steeply

21

dipping subducted lithospheric slabs are generally associated with large slab pull-down forces and back-arc extension (decoupling between the subducted and overriding plate), whereas back-arc compression is associated with shallow dipping slabs and smaller pull-down forces (coupling between subducting and overriding plate; Bott, 1982; Bott and Kusznir, 1984; Wortel, 1986; Bott et al., 1990). The Cenozoic evolution of the Sunda arc and the Sea of Japan (Uyeda and McCabe, 1983; Jolivet et al., 1989; Letouzey, 1990), as well as of the Late Palaeozoic Variscan geosynclinal system (Ziegler, 1990), shows that back-arc compression and back-arc extension can alternate, presumably in response to changes in the convergence rate between the colliding plates. The same may also apply to intermittent activity along such longstanding subduction zones as the one controlling the evolution of the Cordilleran system, in which the Devono-Carboniferous Antler orogeny is separated from the Permo-Triassic Sonoma orogeny by a phase of back-arc extension (Figs. 2 and 3; Burchfiel and Davis, 1975; Frazier and Schwimmer, 1987; Ziegler, 1989). Roll-back and detachment of the subducted slab from the lithosphere may be responsible for termination of the subduction process. Convergence rates between plates are thought to be a function of plate interaction. Plate interaction is presumably governed by a combination of slab-forces, ridge-push and drag forces exerted on the base of the lithosphere by the convecting sub-lithospheric mantle. A further question mark is raised on the concept of slab-pull and roll-back as a major plate driving force by the fact that subduction zones, located at the trailing edge of a drifting continent, can also remain active for a long period of time and can be associated with the development of major fold belts. Examples in point are, the Appalachian orogen, which developed along the trailing edge of Laurussia during its Carboniferous clockwise northward rotation, and the contemporaneous closure of the Proto-Atlantic Ocean (Fig. 2; Ziegler, 1989), as well as the erogenic system which evolved at the southern margin of Pangea during its Late Carboniferous-

290 MA -

250 MA

OCEANIG CRUST

CONTINENTAL

CRATONS

ACTIVE FOLDBELTS

INAGTlVE FOLDBELTS

Fig. 5. Tentative reconstr~~tj~n of w~tin~nts

’‘W in Southern Hemisphere during Late ~a~~~jf~r~?us to Middie Triassic.

1’LA’I’lC ‘I’L:(‘I’C)NlC‘S. PLAI‘L

MOVING

ME(‘HANlSMS

AND

Triassic counter-clockwise northward rotational movement (Fig. 5). The evolution of such fold belts is not compatible with the slab-pull and roll-back concepts and must be related to other plate moving mechanisms. Abandonment of subduction systems underlying Himalayan-type megasutures is generally accompanied with the inception of new subduction zones or an increase in activity along pre-existing subduction systems fringing the margins of the colliding continents. Examples are, for instance, the Late Silurian-Early Devonian locking of the Caledonian Arctic-North Atlantic subduction system and the development of the Sakmarian arc/trench system in the Proto-Uralian Ocean (Fig. 1; Zonenshain et al., 19901, and also the Pcrmo-Carbonifer~us locking of the Hercynian subduction system and increased activity along the arc/trench system marking the southern margin of Pangea (Fig. 5; Ziegler, 1989). Such firstorder suture progradations can apparently entail the development of new subduction zones in oceanic basins. This suggests that convergence and collision of major plates is not solely governed

by slab-forces

23

KIFI‘ING

and ridge-push.

The inception of new subduction zones, particularly in young oceanic basins, is probably related to the build-up of regional compressional stress fields as a consequence of plate interaction (Cloetingh et al., 1989) and not so much to gravity driven, passive sinking of oceanic lithosphere into the asthenosphere. For instance, closure of the Hellenic-Dinarid ocean during the Late JurassicEarly Cretaceous is directly related to the opening of the Central Atlantic and the clock-wise rotation of Laurasia (Fig. 3). Stress-induced subduction systems become inactive upon relaxation of their controlling stress fields. For example, the Pyrenean-Cantabrian subduction system was activated during the Senonian convergence of Africa and Europe; with the development of increasing space constraints, this subduction zone propagated during the Oligocene deep into the Atlantic Ocean and became inactive during the Mid-Miocene when lithospheric shortening concentrated on the Betic Cordillera (Fig. 4). An other example is the collisional deformation of the eastern Sverdrup Basin (Eurekan fold belt)

and the western margin of the Barents Shelf (Spitsbergen Alpine orogen) during the Palaeogene phases of simultaneous seafloor spreading in the Labrador Sea-Baffin Bay and the Norwegian Greenland Sea. Development of the Eurekan fold belt was associated with the subduction of 300-350 km of lithosphere; upon early Oligocene abandonment of the Labrador SeaBaffin Bay seafloor spreading axes, this subduction system locked and the Eurekan and Spitsbergen Alpine fold belts became inactive (Fig. 4; Ziegler, 1988). From the above it is concluded that slab-pull and roll-back do not exclusively govern activity along subduction zones; additional forces are necessary to explain the observed drift patterns of continents, their collisional interaction, first-order suture progradations and changing levels of activity along subduction zones. Seafloor

spreading

axes

The mayor axes of the globe-encircling system of seafloor spreading ridges appear to be rather long-lived. Ridge-push forces associated with such mature seafloor spreading axes result from lateral density changes involving the progressive cooling and thickening of the oceanic lithosphere away from the spreading axis where hot asthenospheric melts wells up to the seafloor. Ridge-push forces are amplified in the case hot-spot activity is centred on a spreading axis (Lister, 1975: Dahlen, 1981; Bott, 1982, 1991; Bott and Kusznir, 1984; Dewey, 1988a). Ridge-push forces appear to account for much of the present global stress distribution (Richardson, 19921. During the early opening phases of oceanic basins, ridge-jumps occur rather frequently, entailing abandonment of earlier axes at the expense of new ones (e.g_? Paleocene Labrador Sea, Oligocene NorwegianGreenland Sea; Cande et al., 1989). Plate reconstructions indicate that seafloor spreading axes migrate relative to the Earth’s core and deep mantle during the opening of Atlantic-type oceans. This is exemplified by the Mesozoic and Cenozoic palinspastic reconstructions of the continents, which are based on seafloor magnetic anomalies, tlow-lines of the

24

ocean floor and palaeomagnetic data (Scotese et al., 1988; Gahagan et al., 1988). These indicate that during the Jurassic and Early Cretaceous opening of the North, Central and South Atlantic and the Indian Ocean, Africa remained nearly stationary (Figs. 3 and 4). Approximately radial growth of the African plate during the opening of the Atlantic and Indian oceans implies that active seafloor spreading axes moved away from Africa at the half-rate of new seafloor generation (Pavoni, 1985, 1988), and that the upwelling asthenospheric systems, which are associated with seafloor spreading axes, are not rooted in the deep mantle. Furthermore, as seafloor spreading axes can be offset at transform faults, it is hard to see how they could be related to majar upweiling convection cells of the deep mantle. Moreover. most modern seafloor spreading axes do not coincide with first order gedid gravity and deep man-

tle low-velocities anomalies (Bowq 1983, 199 I ; Dziewonski, 1984; Vigny et al., 1991). Geochemical criteria indicate that bulk of magmas extruded at seafloor spreading axes is extracted from relatively shallow parts of the depleted upper mantle in response to its t~nsj~~n~~ decompression, causing partial melting to depths of 60-80 km; melts segregating at depth ranges (ti 30-40 km rise diapirically to the seafloor or surface. These melts differ from Hawaiian-type hot.spot magmas which are formed by decomprcssional melting of deep mantle material, rising from the boundary between the upper and lowe mantle ( & 670 km seismic discontinuity) and/or the core-mantle boundary. However. there are many examples of near-ridge or ridge-ccntred hot-spots (c.g., ‘Tristan da Cur&a, St. Helena, Iceland), which influence the c~)rt~~)siti~n of tile Mid-Ocean Ridge Basalts (MOKBI McKenzie and

Fig. h. Time table of seafloor spreading in Ar~tic-~(~rth

Atlantic

and Western

Tethys domains

PLAT1: TEtTONlfS.

PLATE

MOVING

MECHANISMS

AND

25

RIFTING

Bickie, 1988; White and McKenzie, 1989; Christensen, 1989; Wilson, 1989). It cannot be excluded that the magmas of such hot-spots are entrained by the general flow of the convecting asthenosphere. In this context it should be noted that most hot-spot tracks crosscut ocean floor flow-lines, suggesting that hot-spots have little bearing on the motion of the lithospheric plates and that the present association of some hot-spots with active seafloor spreading axes is perhaps coincidental (Morgan, 1981; Cande et al., 1989; Wilson, 1989; Zonenshain et al., 1991). Activity along seafloor spreading axes can apparently terminate almost abruptly, as for instance in the case of the Canada Basin, the Bay of Biscay and Rockall Trough during the Senonian, and also in the Labrador Sea-Baffin Bay during the early Oligocene, at the time of stabilization of spreading axes in the Arctic-North Atlantic to the present system of mid-oceanic ridges (Fig. 6; Rowley and Lottes, 1988; Ziegler, 1988). Furthermore, repeated reorganizations of spreading ridges are evident during the opening of the Indian Ocean, entailing the abandonment of earlier active spreading axes (Scotese et al., 1988; Patriat and Segoufin, 1988; Cande et al., 1989). The nearly contemporaneous and often rapid termination of activity along often distant seafloor spreading axes could be explained by plate interaction. Most seafloor spreading axes are probably the loci of passive advection of asthenospheric material into the space opening between diverging plates (Pavoni, 1992, this volume). Such passive advection can terminate abruptly if far-field stresses impede further divergence of the respective plates, This is exemplified, for instance, by the latest Jurassic-earliest Cretaceous compressional overpowering of the Hellenic-Dinarid ocean seafloor spreading axis (Vardar suture) during the early opening phases of the Central Atlantic. Similarly the Senonian termination of seafloor spreading in the Bay of Biscay is probably a consequence of the build-up of regional compressional stresses in conjunction with the convergence of Africa and Europe (Ziegler, 1988, 1990). During the early stages of seafloor spreading,

modifications in the stress pattern controIling the divergence of the respective plates can entail a rearrangement of the seafloor spreading axes (e.g. Arctic-North Atlantic, Indian Ocean). With progressive opening of an oceanic basin, ridge-push forces gain in importance and may become a dominant plate moving mechanism. However, if plate interaction impedes further divergence of the respective plates, seafloor spreading terminates and ridge-push forces decay rapidly. ~eLlia~ori~tensional stresses

Deviatoric tensional stresses develop in the lithosphere above upwelling asthenospheric convection cells and diapiric intrusions of the asthenosphere into the lithosphere (Bott, 1992, this volume), and also in conjunction with over-thickening of the lithosphere in erogenic belts (Bott, 1990; Dewey, 1981(b). Modefling indicates that u~weI~in~ asthenospheric conuection cells exert a loading stress on the overlaying lithosphere, causing the development of broad-scale, low relief positive lithospheric deflections and the development of deviatoric tension in the lithosphere; at the same time density driven advection and lateral outflow of the anomalous asthenosphere exerts frictional forces (shear-traction) on the base of the Iithosphere (Bott, 1992, this voiume). Permo-Triassic and Jurassic regional uplift of Africa, coupled with hot-spot type intra-plate magmatic activity, can be related to the gradual development of upwelling convection cells under Pangea. Also the present topographic elevation of Africa, which is considered by some authors to be underlain by a single or by several upwelling asthenospheric convection cells (Pavoni, 1985, 1988, 1992, this volume; Bowin, 1991) could be explained in these terms. In this context it should be noted that the intra-continental Mesozoic ultra-alkaline magmatic centres of Africa and the Cenozoic Ahaggar and Tibesti hot-spots are not associated with major crustal extension (Burke and Whitemann, 1973; Le Bas, 1987; Wilson and Guiraud, 1992; Guiraud et al., 1992). On the other hand, there is only limited geological evidence for regional lithospheric doming during the

Triassic ~~v~~~p~~nt of the Tethys and ArcticNorth and Central Atlantic rift systems which transsected Pangea (Ziegler, 19X8). Stress-induced thinning of the litbospherc causes decompression of the asthenosphcrc and lower lithosphere, their partial melting and the diapiric rise of melts into the lithoLy?here (McKenzie and Bickle, 1988; White and Mciienzie, 1989; Wilson, 1989; Latin et al., 1990; Latin and Waters, 1991, 1992). Geochemical criteria indicate that magmas extruded in rift zones iirt’ intialfy derived by partial melting from the Iower lithosphere and the upper asthenospherc at depths ranging between l(H) and 200 km and later also at shallower as well as at greater depths; melts originating from the deeper mantle, which are typical for intra-plate mantle plumes, appear to lack or to play only a sub(~rdinate role in most Phanerozoic rift volcanic suites (Wilson. 1989; Christensen, 1989; Loper, 1991). Progressive thinning of the subcrustal l~thosphcre, as well as the intrusion of melts at the crust/~dntle boundary, can cause thermal doming of a rift zone (Ziegler, 1992, this volume). The cv~~luti(~nof most rifts indicates that major thermal. doming rarely occurs during the initial rifting phases and generally post-dates the onset of rifting by tens of millions of years (LS-20 Ma Gulf of Suez, 20-25 Ma East African rift, 20-30 Ma Baikal rift, 25-35 Ma Rhine Graben, 30-40 Ma Central Atlantic rift, HO Ma North Sea, respectively). Moreover, the ~~bse~ati~~n that also voicanic rifts can become abruptly tectonically inactive, as for instance the Dniepr-Donetz graben during the Early Carbon~ferous ~Chckunov et al,, 1992) and the Oslo Graben at the end of the Earty Permian (Neumann et al., 1992), suggests that deviatoric tensional stresses related to mantle diapirisms are, on their own, not able to drive apart major lithospheric plates. This suggests that deviatoric tensional stresses associated with diapiric intrusion of melts into the lithosphere are probably not the primary driving mechanism of rifting, but are a consequence of lithospheric stretching. However, deviatoric tensional stresses may contrjbute towards the break-up of Pangea-type mega-continents, particularly if they act in constructive interference with

far-field stresses. On the other hand, they UC’ unlikely to be the primary driving force of this process. Deviatoric tensional stresses, dcve~(~p~~lgin responsc to olw-thickening of !hr lithosphtw m orogertic bdts, particularly of the continenr-tocontinent collisional type. arc thought to cause their p~)st-~.)r~~geniccollapse. and may enhance tensional piate-boundary forces associated with slab roll-back (Dewey, 198&b: Bott, 1990). Such tensiona body-forces arc LikcIy to be targcst shortly after the consolidation iit’ it fold belt, and thus should cause its early colfapsc. However, in the Arct~c-~or~~~ Atlantic C.:alcdonides, regional crustal extension cornminced only during the Namurian, soott: X-HO Ma alttcr consolidation of this fold belt. Ijuring Devonian and Early C’arhoniferous times, the post-erogenic development of the Arctic-North Atlantic C’afcdonides was governed by major s-inisrrai shear rn(~ve~lenfs along a complex fracture system subparalleling the axis of this f<>idheit; related puflapart basins give evidence for major extensional strain along the axis of the C’aiedonides and not across them as would be expected if subsidence of these basins were controiled by the body-forces inherent to this fold hclt. Simltarly, tho Stephan&n-Autunian initial phase nf disintegration oi the Variscan fold belt was governed by dextral translations between Africa and .Europe, giving rist: to the deve~opme~t of a complex pattern of shears and wrenc~~-il~ducc~~ dull-apart basins: their evolution reflects a progressive rotation rti the controlling stress field that rctJeccs a gradual change in the convergence d~rect~ol~ between Gondwana and Laurussia during the ~le~benian consolidation of the Appalachian-Mauretanidcs fold belt. A period of reiativbt tectonic yuiesccnce, spanning some 20 Ma, separates the tcrmination of wrench-fault activity in the domain of the Variscan fold belt from the Triassic onset of regional crustal extension. Also in the hppa~achian domaj~, the Mod-.~riass~c onset of PZgional extensional tectonics post-dates the termination of the ~leghenian orogcny by some 211-25 Ma (Ziegler, 1988, 1989, 1990). Moreover, the Permo-Triassic counter-ciockwisc northward rotation of Pangea can neither be explained by

PLATE

TEC-I-ONICS,

PLATE

MOVING

MECHANISMS

AND

RIFTING

deviatoric tensional stresses related to the collapse of the Hercynian megasuture along which Gondwana and Laurussia were welded together, nor by subduction related forces governing the consolidation of the Uralian fold belt. For these reasons, it is questioned whether deviatoric tensional forces developing in response to lithospheric over-thickening in erogenic belts are indeed able to cause, on their own, the disintegration of mega-continents such as Pangea and Laurasia, particularly in the face of counter acting far-field stresses. In constructive interference with far-field stresses, however, deviatoric tensional body-forces inherent to erogenic belts may contribute towards their post-erogenic collapse. Drag forces exerted on the base of the lithosphere

Drift patterns of major continents, such as of Gondwana during the Devonian and Carboniferous, Pangea during the Permian and Triassic and Eurasia during the Late Jurassic and Cretaceous, are difficult to explain in terms of the conventional plate moving forces such as ridge-push, slab-pull and roll-back, deviatoric tensional body-forces of orogens and deviatoric tensional stresses developing over mantle plumes. In view of the above, it is suspected that convection currents in the asthenosphere and upper mantel play an important role as a plate driving mechanism, primarily by exerting drag forces (shear-traction) on the base of the lithosphere, and also by causing, over upwelling cells, the development of deviatoric tension in the lithosphere (Forsyth and Uyeda, 1975; McKenzie, 1977; Richter and McKenzie, 1978; Christensen, 1983; Wortel, 1986; Zoback et al., 1989; Irvine, 1989; Loper, 1985; Forte and Peltier, 1991; Vigny et al., 1991; Bott, 1992, this volume). Geochemical criteria are strongly in favour of a separation of the convection cells in the depleted sub-iithospheric upper mantle from the convection systems of the more primordial, less depleted, lower mantle at the +_670 km seismic discontinuity (Richter and McKenzie, 1981; Wilson, 1989; Christensen, 1991; Thompson, 1991). However, convection cells in the lower and upper mantle may be partially coupled during the

LI

break-up of a Pangea (Mesozoic African hot-spot activity), and largely decoupled at times of dispersed continents, as appears to be the case today (Le Pichon and Huchon, 1984; Le Pichon and Gamier, 1986). On the other hand, seismic mantle-tomography favours a whole mantle convection system (Forte and Peltier, 1991; Vigny et al., 1991). Shear-traction at the base of the lithosphere is thought to be relatively small per unit area; however, considering the size of lithospheric plates, cumulative stresses exerted on a plate by the convective asthenosphere can be very large. Moreover, deflection of asthenospheric convection currents around the deep lithosphere keels of old cratons, which reach down to 200 km as evident, for instance, along the European Geotraverse (Blundell et al., 1992) may cause an increase in frictional forces exerted on the latter. Drag forces are likely to be proportional to the horizontal component of asthenospheric flow velocity relative to the lithosphere. In the direction of increasing horizontal sub-lithospheric mantle flow velocities, corresponding to the peripheries of an upwelling cell, the overlying lithosphere experiences deviatoric tensional stresses, whereas in the direction of decreasing flow velocities, corresponding to areas of downwelling cells, it experiences compressional deviatoric stresses (Pavoni, 1981, 1992, this volume). Furthermore, the contribution of shear-traction to plate motion may depend on whether sub-lithospheric mantle flow is essentially unidirectional beneath the respective plate or whether mantle flow converges or diverges. For instance, during the Mesozoic and Cenozoic opening of the Atlantic and Indian oceans, Africa, which is thought to overlay an upwelling and diverging branch of the mantle convection system, underwent only minor displacements and was subjected to almost continuous multi-directional crustal extension (Wilson and Guiraud, 1992); on the other hand, South America drifted rapidly westward and straddles now a downwelling branch of the mantle convection system (Pavoni, 1991, 1992, this volume). With such a scenario, ridge-push, slab-forces and deviatoric tensional stresses related to mantle plumes and orogens played probably a sec-

ondary, albeit im~rtant role as pfate moving forces during the break-up of Pangea and the radial dispersal of its components. As the Pangea fragments moved away from the African asthenosphere upwelling centre and mature spreading ridges developed, the contribution of shear-traction to plate movements decreased and ridge-push and slab-forces gained in importance as plate moving mechanisms. Their constructive interference with shear-traction can account for high drift rates, pa~icularIy of relatively small plates, as observed, for instance, during the CretaceousPalaeogene northward drift of India. In time and space, the relative importance of the different plate moving forces is likely to change, dependent on plate interaction. Moreover, variable degrees of coupling between the lithosphere and asthenosphere may account for differences in the rate of lithospheric plate motions (Sabadini and Yuen, 1889; Doghoni, 1990). In this context it should be noted that a correlation exists between plate velocities and the proportion of plate boundaries occupied by subduction zones, whereas no correlation can be established between plate velocities and the surface area of the respective plate (see Forsyth and Uyeda, 1975). Major reorganizations of lithospheric plate boundaries probably result from plate interactions that are governed largely by the convection systems of the sub-Iithospheric mantle. The presence, or absence, of low velocity zones at the base of the Iithosphere may indicate whether or not a lithospheric plate is coupied or decoupled from the underlying asthenosphere. Furthermore, changes of convection-induced mantle inhomogeneities and the distribution of major continental masses can apparently cause a shift of the lithosphere relative to the Earth’s core and mantle, as mirrored by the hot-spot tracks (Scotese and Denham, 1988), and consequentiy changes in plate interaction (Neogene westward shift; Doghoni, 1990, 1991; Ricard et al., 1991). AssembIy of a major continental mass, such as Pangea, presumably has an insulating effect on mantle convection systems that were active during its assembly and may cause their decay; this is followed by a reorganization of the global convection system (Le Pichon and Huchon, 1984). De-

cay of old downweIi~ng and evolution of new upweliing convective cefls under a super-continent, such as the Permo-Triassic Pangea, gives rise to changes in the stress regime in the lithosphere; regional compression terminates and gives gradually way to a new tensional stress system which will govern the break-up of the respective mega-continent. Broad-scale thermal doming over new upwelling asthenospheric cells is accompanied by the development of regional deviatoric tensional stresses, which are enhanced by sheartraction of the diverging mantle flow at the base of the lithosphere. On the other hand, rift-teiated lithospheric thinning causes development of local deviatoric tensional stresses (Bott, 1992; Pavoni, 1992, this volume). In the case of Permo-Triassic Pangea, the process of convection cell reorganization spanned some 40 Ma from the moment of consolidation of the Hercynian megasuture until the onset of regional crustai extension, as refIected by the development of the Triassic rift systems transsecting Pangea, and about 80 and 12tl Ma until the opening of the Central and South Atlantic. I’Cspectively. Geophysical data indicate that at present, the Earth’s mantle convection system is organized in an approximately symmetrical, bipolar configuration with upwelling cells underlying Africa and the Pacific Ocean (Pavoni. 198.5, 1988. 1991; Irvine, 1989; Vigny et al., 199i); evolution of this modern convection celi system commenced presumabIy during the Permian and governed the break-up of Pangea. Interaction

of lithospheric

plates

As we have to deal in all likelihood with a finite globe, the generation of new oceanic lithosphere has to be compensated for elsewhere by the subduction of a commensurate amount of oceanic lithosphere and/or shortening of continental crust and subduction of tower crustal material and subcrusta1 Iithosphere fForsyth and Uyeda, 1975; McKenzie, f977; Bott, 1982; CIoetingh and Wortei, 1985). This relationship is evident, for instance, in the coincidence of opening of the Atlantic Ocean with the main erogenic activity in the Andean-Cordilleran erogenic sys-

PLATE

TECTONICS,

PLATE

MOVING

MECHANISMS

AND

RIFIING

tem, and also in the timing of opening of the South Atlantic-Indian Ocean and the onset of the Alpine Orogeny. Thus plate interaction, largely driven by upper-mantle convection-systems and their changes through time, plays probably an all-important role in the development of intra-continental rift systems, the opening of new oceanic basins and the inception of subduction zones. Moreover, circumstantial evidence suggests also that compressional stresses can be transmitted over great distances through continental and oceanic lithosphere (Letouzey and Tremoliere, 1980; Cloetingh and Wortel, 1985); this concept is supported by the World Stress Map compilation (Zoback et al., 1989). Plate interaction can have a profound effect on drift patterns of major cratonic blocks. For example, Laurentia-Greenland underwent only minor latitudinal and longitudinal displacements during the Ordovician to Devonian time span, and was presumably located over converging, downwelling, deep asthenospheric convection systems that keep it in place. However, upon impact with Gondwana, or perhaps as a consequence of the reorganization of these convective systems, Laurentia-Greenland became decoupled from the sub-lithospheric mantle during the Famennian and started to drift rapidly over great distances, essentially in unison with Gondwana (Fig. 3; Ziegler, 1989). As the Late DevonianCarboniferous drift pattern of Laurussia cannot be explained by subduction mechanisms along its northern margin, it is likely that motion of Gondwana, driven mainly by traction of the convecting asthenosphere and ridge-push, was imparted on Laurussia upon their collision. Similarly, the locking of subduction systems underlying intra-continental collision zones, such as the Arctic-North Atlantic Caledonides and the Hercynian megasuture of Gondwana and Laurussia, is likely to be a consequence of lithospheric thickening having reached the critical point at which crustal shortening is no longer possible. Locking of these subduction systems was apparently compensated for by the transfer of subduction activity into the Ural Ocean and to the southern margin of Pangea, respectively (Figs.

29

1 and 5; Ziegler, 1988; Zonenshain et al., 1990). Such first-order suture progradations to the distal margins of colliding continents indicate that their motion was not exclusively governed by subduction processes associated which the evolution of their common megasuture. A further example of plate interaction is the impeded rotation of Laurasia during the Early Cretaceous phases of opening of the Central Atlantic, which resulted from increased collisional coupling between Eurasia and Africa (Fig. 3). However, continued rotation of the North American craton caused northward rift propagation and later seafloor spreading into the Arctic-North Atlantic domain (Fig. 4). It is likely that sheartraction of the convecting sub-lithospheric mantle played an important role in this rotational movement of the North American craton which, coupled with progressive opening of the Central and North Atlantic and subduction of Pacific oceanic lithosphere, culminated in the Neogene overriding of the East Pacific Rise. Ridge-push in the Atlantic and subduction-related forces at the Cordilleran arc-trench system may have contributed toward the post Mid-Jurassic drift pattern of North America. The Late Cretaceous and Palaeogene Alpine plate-boundary reorganization in the Mediterranean domain can be related to convergence of Africa and Europe, which caused, by the build-up of regional compressional stresses, the initiation of subduction zones in the young oceanic basins of the Alboran, Piedmont and Penninic seas, as well as in the Bay of Biscay. These stresses are thought to be responsible for large scale intraplate compressional deformation in Europe and in North Africa, and, particularly during the Paleocene, for impeded opening of the Norwegian Greenland Sea (Fig. 4; Ziegler, 1988, 1990). Interaction of lithospheric plates, driven by a combination of shear traction of the convecting sub-lithospheric mantle, ridge-push and slabforces, and possibly also by wholesale motion of the lithosphere relative to the asthenosphere and core (in response to changes in the rotation rate of the Earth), may be responsible for the nearly contemporaneous cessation of activity along often distant seafloor spreading axes (e.g., intra-

30

Senonian plate-boundary reorganization in Arctic-North Atlantic domain; Fig. 6; Ziegler, 1988). Plate interaction may also play a role in changes in the convergence rate of colliding plates that cause alternations between back-arc compression and extension (e.g., Sunda arc, Uyeda and McCabe, 1983; Jolivet et al., 1989; Variscan geosyncline, Ziegler, 1990). Conclusions The relative importance of the different mechanisms contributing to the movement of lithospheric plates is likely to change in time and space. During the break-up of Pangea-type mega-continents, shear-traction and deviatoric tension developing over upwelling and diverging asthenospheric convection cells probably play a dominant role. During the dispersal of continents, ridge-push forces can act in constructive or destructive interference with shear-traction, depending on their orientation relative to the asthenospheric flow pattern. Assembly of megacontinents occurs probably in areas of downwelling convection cells; under these conditions, slab-forces, ridge-push and shear-traction may act in constructive interference. The development of rift systems of the Atlantic type, which evolve during the break-up of major continental masses, and lead to the opening of large oceanic basins, is probably governed by the reorganization of the sub-lithospheric convection cells after the assembly of mega-continents. Evolution of such rift systems can be related to the build-up of regional tensional stress regimes .in the lithosphere in response to drag forces exerted on its base by the convecting astheno- and mesosphere, and also to the development of deviatoric tension over upwelling asthenospheric convection cells and in individual rift zones. Magmas extruded in intra-continental rift zones are mainly derived by partial melting from the lower lithosphere and the upper asthenosphere. These melts are generated as a consequence of stress-induced extension of the lithosphere and resulting decompression of the lower lithosphere and the upper asthenosphere.

Seafloor spreading is here interpreted as passive advection of upper asthenospheric material into the space opening between diverging plates. Mantle plumes of the Hawaiian-type, rising from the lower mantle, do not appear to play an important role in the evolution of most Phanerozoic rifts, though they may have contributed by weakening of the lithosphere to the localization of rifts (Wilson and Guiraud, 1992). Whether they played a more important role during Precambrian rifting (for instance in the Midcontincnt rift of the U.S.A.; Cannon and Hinze, 1992). remains to he seen. Interaction of lithospheric plates, driven by 2~ combination of shear traction of the convecting sub-Iithospheric mantle slab-forces. ridge push and deviatoric tensional stresses. plays a major role in the development of rifts, the opening of new oceanic basins, the inception of subduction zones and the rates of plate convergence. Activity along subduction zones appears to he largely controlled by the rate of plate convergence and not so much by gravitational sinking of the subducted lithospheric slab. Back-arc tectonics, which can alternate between compression and extension, appear to be controlled also by the convergence rate of the colliding plates; in view of changing convergence rates, back-arc extensional systems are generally short-lived (for instance. Sunda arc, Variscan geosynclinal system). Collisional coupling of major continents can lead to their internal deformation and under given circumstances to foreland splitting (Baikal rift; ‘I’apponier et al., 1986). Stresses developing in rcsponse to plate interaction can counteract deviatoric tensional body-forces related to lithospherrc over-thickening in erogenic belts, to the degree that their early tensional collapse is impeded by tens of millions of years. Overriding of the East Pacific Rise by the North American craton, contributing to the development of the Basin-andRange extensional system, is attributed to plate interaction driven mainly by ridge-push and shear traction of the convecting sub-lithospheric mantle. Thoughts advanced here on the relative importance of plate moving forces are but hypotheses that require careful testing. The thermally driven.

PLATE

TECTONICS,

PLATE

MOVING

MECHANISMS

AND

31

RIFTING

primary mantle convection system envisaged here involves either the entire sub-lithospheric mantle, as suggested by gravity and seismic tomography data (Bowin, 1991; Forte and Peltier, 1991; Vigny et al., 1991) or is two-layered, whereby coupling between lower and upper mantle convection cells may be subject to changes (Le Pichon and Huchon, 1984). However, geochemical considerations speak strongly in favour of a two-layer convection system which is partly or completely separated at the +670 km seismic discontinuity (Richter and McKenzie, 1981; Wilson, 1989; Thompson, 1991; Christensen, 1991). The motion of lithospheric plates and their interaction is thought to be driven by mechanisms facilitating the escape of thermal energy from the interior of the Earth and by motions of the lithosphere relative to the asthenosphere in response to lateral mantle variations (Ricard et al., 1991). The pattern of sub-lithospheric mantle flow and the constructive or destructive interference of resulting shear traction forces with slab-forces, ridgepush and deviatoric tensional stresses probably controls the rate of plate motion. Plate interaction controls secondary, shallow mantle flow at seafloor spreading axes and subduction zones. Physical models of mantle convection must take into account, apart from geochemical, gravity and seismic tomography data and information of the present state of stress of the lithosphere, also the motion of plates, their interaction and changing configuration through time. As geologists, we can only describe the symptoms of the patient, our Earth. During the last decades, considerable advances have been made in deciphering the motion of continents during Phanerozoic times. It is now time to submit these data sets to the physical modellers of mantle convection and lithosphere dynamics, and to await their feedback. Acknowledgments

The author is indebted to Shell Internationale Petroleum Mij. B.V. for releasing for publication the text figures given in this paper. Special thanks go to Mrs. Josje Kriest who prepared the palaeoreconstructions with the aide of an Evans and

Sutherland PS 300 vectorgraphics terminal and to Mr. Ruud van Aarle for drawing the text figures. The author thanks S. Cloetingh, C. Doglioni, X. Le Pichon, M. Wilson and D. Latin for critical and constructive comments to an earlier version of this manuscript.

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