carpathian foreland fold and thrust belt

8/7/2019 Carpathian Foreland Fold and Thrust Belt

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The American Association of Petroleum Geologists BulletinV 66, N o. 9 (SEPTEMBE R 1982), P. 1179-1195, 12 Figs.

Carpathian Foreland Fold and Thrust Belt andIts Relation to Pannonian and Other Basins^

B. C . BURCHFIEL and L. ROYDEN^

ABSTRACTThe foreland fold and thrust belt of the Carpathians canbe traced from Austria through the western and easternCarpathians to the south Carpathian bend in Romaniawhere most of its structural units plunge beneath youngerPUocene-Pleistocene cover.Folds in the youngest rocks at the south Carpathianbend continue westward, until all surface expression disappears before reaching the Danube R iver. The fold and

thrust belt is flanked by elements of the European, Russian, and Moesian cratonal areas which are overlain by avariable width and thickness of foredeep deposits that arepartly involved within the external fold and thrust belt.Internal to the fold and thrust belt are older parts of theCarpathian orogene formed mainly on continental crust.These older structures have significant differences in evolution between the western and eastern Carpathians.The foreland fold and thrust belt consists predominantly of flysch in its inner parts and molasse in its outerparts. These sedimentary rocks are cut into a series ofthrust sheets that have moved relatively toward the craton.Oldest flysch units are Late Jurassic and the youngestunits are Miocene. Some of the flysch units of the innerpart of the belt may have been underlain by oceanic crust,whereas the younger, more external molasse units wereunderlain by continental crust. Timing of thrusting is onlywell-constrained locally, but suggests that deformation inthe western Carpathians developed during Oligocene toMiocene time. In the eastern Carpathians, structuralactivity continued into the Pliocene , and deep earthquakeactivity at Vrancea suggests subduction. Deformation isstill locally active. A volcanic arc of Miocene to H oloceneage lies internal to the fold and thrust belt. Magmaticactivity appears to have migrated externally and eastwardin time.Contemporaneous with thrusting was the development

of Neogene basins in the intra-Carpathian region. Thesebasins exhibit a two-phase subsidence history (except for1982. The American Association of Petroleum Geologists. All rightsreserved.'Manjscript received, June 9, 1981; accepted, September 11, 1981. Thispaper was an invited paper presented at the symposium on overthrust beltsheld at the AAPG annual meeting in Denver, Colorado, 1960.^Department of Earth and Planetary Sciences, Massachusetts Institute ofTechnology, Cambridge, Massachusetts 02139.The sections of this paper on the overthrust belt w ere written by Burchfiel,and the sections on the basins by Royden.Preliminary studies on the Carpathians were financed by the InternationalBranch of the National Science Foundation, and a grant from Shell Oil Co, Current studies on the intra-Carpathian basins are financed by the N ational Science Foundation Grant No, INT 7910275 and the Hungarian Academy of

Sciences.

the Pannonian and Transylvanian basins)fast initialsubsidence, followed by a period of slower linear subsidence. The Pannonian basin shows only a reasonably fastlinear subsidence. Structural and thermal models indicatethe basins were formed by about 100


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1180 Carpathian Foreland Fold and Thrust Beltthe outer Carpathians were deformed during the Cenozoicand consist of a sequence of predominantly flysch andflysch-hke terrigenous sandstone, shale, and local conglomerate that ranges from latest Jurassic to Miocene inage. Rocks in the outer Carpathians grade both structurally and stratigraphically into the foreland foredeep. Theforedeep rocks are dominately Neogene in age and form aterrane of varying width around the Carpathian fold belt.The focus of this article is the outer Carpathian tectonicelement that forms the foreland fold an d thrust belt of themountain chain, and the intra-Carpathian basins. Theevolution of these basins is partly contemporaneous andclosely related to the structural development of the foldand thru st be lt. However, it is necessary to examine brieflythe other tectonic elements to understan d the regional setting of the fold and thrust belt.

INNER CARPATHIANRocks within the inner Carpathians consist of deformedPaleozoic and perhaps older rocks metamorphosed tovarying degrees and overlain by essentially nonmeta-morphosed Mesozoic rocks. These rocks were thrust

northward in the western Carpathians and ApuseniMountains during the early Late Cretaceous (Fig. 3). Inthe eastern Carpathians pre-Mesozoic crystalline rocksform a series of east-directed thrust plates (Fig. 3). Eachplate carries small remnants of an unmetamorphosedMesozoic shelf sequence; the highest plate contains onlyMesozoic sedimentary rocks. Thrusting b egan in the EarlyCretaceous and was completed by the Cenom anian.The southern Carpathians consist of two thrust sheetsof crystalline rocks tha t were thrust south and east over the

European Platform

FIG. 1Tectonic sketch map showing position of Carpathians within alpine mountain belts of eastern Europe. V, Vienna basin; WD,West Danube basin; To, Transcarpathian basin; B, B akony Mountains; Bu , Bukk Mountains; A, Apuseni Mou ntains. Horizontallyhachured areas are underlain by oph iolite.


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B. C. Burchfiel and L. Royden 1183thrust belt (Fig. 4). The majority of the folds and thrustsheets were displaced toward the European and Russianforeland, although thrusting in the reverse direction islocally important (Figs. 5, 6). The thrust sheets form aneastward convex loop that ranges from about 30 to 62 mi(50 to 100 km) in width and is about 750 mi (1,200 km) inlength. Westward, the fold and thrust belt can be tracedinto the eastern Alps of Austria, and southward it can betraced to the south Carpathian bend in Romania where itis covered by you nger sedim ents.

Each thru st sheet within the belt is underlain by a thrustor thrust zone with a displacement of a few tens of kilometers and can be subdivided into numerous smaller subunitsbound ed by faults of lesser displacemen t. No single thrustsheet can be traced around the eastward convex loop of the

Carpathians, as the bounding thrust faults form a complex anastomosing pattern of intersecting faults; only themajor thrust sheets of the belt are shown in Figure 4. Th ebelt ends in the subsurface somewhere between the southCarpathian bend and the concave eastward loop of thesouth Ca rpath ians . Folds in the Pliocene-Pleistocene sedimentary rocks overlap the main structures of the forelandbelt but con tinue the belt west about 62 mi (100 km), an dthe structures of the main fold and thrust belt are deeplyburied . D rilling in the foreland area of the south Carpathians near the Danube River has reached the basement rocksof the Moesian platform without evidence of the continuation of the foreland fold and thrust belt, but the positionand nature of the termination of the belt are unknow n.The external boundary of the Carpathian foreland fold

/

FIG. 4Major thrust sheets of foreland fold and thrust belt of outer Carpathians. West Carpathians (starting at left of map): Z,Zdanlce sheet; S, SUesian sheet; sS, Subsilesian sheet; M, Magura sheet; Sk, Skiba-Skole sheet; pD , pre-DukIa sheet; D , Dukla sheet;IF, uiner foredeep unit. Tc is Transcarpathian area. Eastern Carpathians: A, Audia sheet; T, Tarcau sheet; sC, Subcarpathian unit;Mf, Marginal Folds unit; C, Ceahleau sheet (belongs to inner Carpathian element); Cu,Curbicortical flysch sheet. Pieniny klippenzone is shown in black. Locations of cross sections for Figure 5 (through eastern Carpathians at right side of figure) and Figure 6(through western Carpathians at left of figure) are indicated.


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1184 Carpathian Fo reland Fold and Thrust Belt

and thrust belt is marked by an undeformed sequence oflargely Neogene sedimentary rocks forming a molassicforedeep. The most external structures ofthe foreland beltcontain rocks that belong to the inner parts o fthe molassicforedeep caught in the thrusting. The internal boundaryof the belt is more difficult to define. In the western Carpathians it occurs structurally along the Pieniny klippenzone (Fig. 4), which was deformed with the inner Carpathians during the Cretaceous and with the outer C arpathians during the Cenozoic. In the Transcarpathian region ofthe USSR and n orthern Rom ania, the innermost units ofthe outer Carpathians and the Pieniny klippen zoneoccupy a more internal position. They strike south of thewestern termination of structural and stratigraphic unitsofthe inner west Carpathian of northern Romania. Fromthe Transcarpathian region similar rocks can be tracedsouthwestward in the subsurface into a zone of deformedLate Cretaceous and early Cenozoic rocks in the Szolnoktrough of eastern Hungary (Fig. 3) . The relations betweenthis northeast-trending structural zone of Szolnok and theforeland fold and thrust belt are unclear, but the Szolnokzone of rocks was deformed in the Paleogene and represents the structural break between the west Carpathianand Apuseni Mountain inner Carpathian units. Eastwardfrom the Transcarpathian area, the internal bounda ry ofthe foreland belt is marked by the contact with thrustsheets containing mainly Cretaceous flysch. The mainthrust sheet, the Ceahleau napp e, contains rocks similar tothose in the foreland belt (Figs. 4, 5); however, it was

deformed during the Cretaceous and is regarded by mostworkers as part oft he internal eastern Carpathian s. (Theterms "Dacides" and "Moldavides" are commonly usedfor the internal and external parts of the foreland beltunits of the eastern Carpathians.)Rocks in the foreland fold a nd thru st belt consist mainlyof various types of flysch and range in age from LateJurassic to M iocene and reach several miles (kilometers) inthickness (Figs. 7, 8). Most of the rock sequence is LateCretaceous to Oligocene in age. In the western Carpathians, flysch units of the innermost Magura thrust sheetgrade south into flysch units that are tectonically part ofthe Pieniny klippen belt. Similarly, these flysch units gra desouthward into the relatively undeformed flysch units(Podhale flysch) that lie unconformably on the structuresofthe inner Carpathians. Likewise, in the external part ofwestern foreland belt, flysch un its grade northwa rd acrossthe most external nappe boundaries into molassic rocksthat represent the most southerly pa rt ofth e molassic foredeep basin. Within the foreland belt of the western C arpathians, some major changes appear in rock units andfacies across nappe con tacts, such as between the Mag ura,Duk la, and Silesian napp es (Fig. 8) and within the nappesboth across and parallel with structural strike. In the eastern Carpathians, there are important changes of rockunits and facies between napp es such as between the Audiaand Tarcau nappes and between the Curbicortical andAudia nappes (Fig. 7). Eastward, the external nappes ofthe eastern Carpathians grade into molassic rocks of the

FIG. 5Cross section through foreland fold and thrust belt of the eastern Carpathians. Location of section shown in Figure 4 (fromBurchfiel, 1976).

K L I P P E N B E L T M A G U R AS U B S I L E S I A N

F O R E M A G U R A

l O k m

FIG. 6Cross section through foreland fold and thrust belt of western Carpathians. Location of section shown on Figure 4. Crossesare pre-Cenozoic basement rocks. Dashes are Oligocene-Miocene rocks of the autochthon (section redrawn from PesI et al, 1968).


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B. C. Burchfiel and L. Royden 1185

foredeep. Similar to the western Carpathians, the innermost parts of the foredeep were involved in the deformation of the foreland fold and thrust belt.Many studies have been conducted on the sedimenttransport within the flysch formations of the forelandbelt. Sediment transport from the internal zones of theCarpathians is very prono unce d, with major dispersal centers in the Transcarpathians and south Carpathian bendregions (Contescu, 1974). Sediment transport along theaxis of the belt is perhaps the dominant direction withinmuch of the belt. In the west Carp athians, there is a majorsediment source in the west, perhaps coming from theeastern Alps where a continental collision between Europeand an Apulian fragment took place in the Paleogene. Inthe eastern Carpathians and Transcarpathian areas, thesediment trans por t directions are fanlike, reflecting m ajorlocal sources from the inner zones; however, there is adominant northward transport direction from the southCarpathian bend. Along the external part of the belt, sedi

ment sources in the foreland are comm on, although theyprobably did not supply the great bulk of detritus for theflysch belt as a whole. The foreland sources are most significant in the most external thrust she ets.One of the surprising features of the sediment transp ortstudies in the Carpathians has been the evidence of localpaleocurren t directions , which suggests sources within theflysch terrane itself. Additionally, rare units within theflysch sequences contain con glomerates and oUstostromesof apparently exotic ma terial. For example, the Pasierbiecsandstone of the Magura thrust sheet contains cobbles ofgranite, slate, amphibolite, and limestone (Gszcypko,1975). These rocks were ap parently derived from an areabetween the Magura and underlying Silesian nappes.These data have led to the interpretation that short tolong-lived cordilleras were present, some perhaps underlain bycontinental crust, which served as source terranes(e.g., see Ksiazkiewicz, 1956, 1963). Similar suggestionshave been m ade for the eastern Carpathians (e.g., P atru-

INTERNALUNITS

EXTERNALUNITS

-**

Ceahleau Curb icortical AudioNappe Nappe Nappe

TarcauNappe

MarginalFolds SubcorpothianNappe Nappe

Clue Bodoc inner outer

FIG. 7Stratigraphic sections from east Carpathians fold and thrust belt showing some of the major flysch units. Flysch units areshown in dash-dot pattern. Dot patterns are molassic rocks. Lines are more pelagic rocks; open circles are conglomeratic beds. Blackis Vraconian variegated shale, th, Tithonian; ne, Neocomian; al , Albian; vr, Vraconian; Eo, Eocene; pr, Priabonian; ch, Chattian; h e,Helvetian; and sm, Sarmatian. Vertical scale only approximate.


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1186 Carpathian Foreland Fold and Thrust Beltlius, 1969). Most of these studies were done prior to thedevelopment of depositional m odels for deep-sea fans andthe recognition of the importance of transport throughsubmarine canyons. Although still imperfectly understood, these features may explain many of the paleocur-rent and exotic source terrane data (e.g., see Contescu,1974). It is surprising that none of the Cordilleras has everbeen mapp ed or located within the present f lysch terran e,and that they have been interpreted to be either covered byhigher thrus t sheets or completely subducted.

The time of emplacement of the thrust sheets in the Carpathian foreland fold and thrust belt is not well defined,except in its external part, beneath the V ienna basin and atthe south Carpathian bend where younger rocks uncon-

formably overlie deformed rocks. Beneath the Viennabasin early Miocene rocks (Eggenburgian and Ottnan-gian) overlap thrusts and folds in the inner part of the foreland belt (Brix and Schultz, 1980), but rocks as young asearly middle M iocene (Karpatian) are involved in the mostexternal thrusts. Similar age thrusting can be documentedat the south Carpathian bend, but rocks as young as lateMiocene are thrust in the most external structural units,and the folded molasse of the foredeep involves Pliocene-Pleistocene beds (Burchfiel, 1976). Jiricek (1979) has triedto dem onstrate th at the age of thrusting in the most external part of the foreland belt migrated eastward in time, anidea in harm ony w ith other features of the Carpathian arc.The onset of thrusting is more difficult to dem onstrate .

M a g u r a D u k i a Si les ian Subsi lesian Skolem e d i a l ou te r

TuCen

L _ - _ - l - -

1=^==?^o o e

o o o, '~ZZ.

\ L

Cen

FIG. 8Stratigraphic sections from eastern part of tlie west Carpathians showing some of major flysch sequences in foreland foldand thrust belt. Flysch units are shown by dash-dot pattern. Long dash lines are more pelagic rocks. Dark shading is Menelite shalesand cherts. Mio, Miocene; Olig, Oligocene; Eo, Eocene; Pal, Paleocene; Sn, Senonian; Tu, Turonian; and Cen, Cenomanian. Vertical scale only approximate.


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B. C. Burchfiel and L. Royden 1187Most of the inner nappes contain rocks as young as Eoceneand locally Oligocene (Figs. 7, 8), but unconformablyoverlapping beds are missing. Structural relations betweenthrust sheets in the belt suggest thrusting first began in theinnermost nappes and migrated outward with time.Higher thrusts are often folded with footwall unitsacommon feature of many foreland fold and thrust belts.Only some of the most external thrusts remain relativelyplanar. Althou gh evidence is unclear, most writers suggestthe deformation began sometime in the Oligocene, perhaps in the late Oligocene. Thrusting in the foreland belt isspatially and temporally related to volcanism in the intra-Carpathian region, which began in the early Miocene andmigrated eastward in time. A fairly consistent interpretation is tha t thrusting did begin in the late Oligocene and didmigrate eastward in time, but complete confirming evidence is presently not available. Internal structure withinthe foreland belt is often dominated by the nature of therock sequence. Where massive sandy flysch units arepresent, they often form more open o r less complex foldsand may be locally or regionally detached, forming separate subunits within the major thrust sheets. The moreshale-rich sequences commonly form very complex foldsand imbricate thrust zones; locally they may form thecores of diapiric structures. Most folds and thrusts arevergent toward th e external pa rt of the belt; however, foldsand thrusts in the inner part of the belt are commonlyinternally vergent. Often these internally vergent structures deform earlier externally vergent structures, whichsuggests that the earliest deform ation began in the internalpart of the belt, but that as thrusting migrated outward intime, deformation took place across a broad zone withinthe belt (Kusmierek, 1979). The most internal part of thebelt and the Pieniny klippen zone show significant internally vergent thrusting and folding.

One of the most dramatic changes in structural stylerelated to rock sequence is where evaporite units, comm onin Oligocene and particularly in Miocene rocks, becomeimportant in the external thrust sheets. These evaporitesform zones of high ductility; their presence results in anextremely supple style of folding during deformation. Inthe Marginal Folds nappe in Roman ia, the internal structure consists of strongly overturned and dislocated folds(see Fig. 5). Similar structures are present along strike inthe USSR, where the folds are cored by Upper C retaceousshaly flysch and complicated by evaporite-related structures at shallower s tructural levels. It should be noted thathydrocarbon production is obtained from the MarginalFolds nappe in Rom ania and that at m ost places this nappeis completely covered by the higher and structurally simpler Tarcau napp e (Figs. 4, 5).

Rocks of the foreland fold and thrust belt have beencompletely detached tectonically from their substratum.Only along the most internal part of the belt in the westCarpathians and Transcarpathian area do the flysch unitsoverlie older rocks. H ere they lie on the older rocks of thePieniny klippen belt, which itself consists of allochtho-nous Mesozoic rocks. Locally along some thrust sheetboundaries, pre-Cretaceous rocks are present as tectonicslices mixed with a shaly matrix of sheared rocks (thesetypes of tectonic slices are referred to as klippen, a usage of

the term that is not common outside the region). In thewest Carpathians, the most external thrust zones containklippen of Jurassic and Cretaceous limestones that aresimilar to those that were deposited on the crystalline basement of the European foreland. From the molassic character of these rocks in the external part of the thrust beltand tectonic slivers of their probable basement, it can beinferred th at rocks of the external part of the foreland beltwere deposited on the European continental crust. TheAndruchow klippen of western Poland are more difficultto interpret. They consist of tectonic slices between theSilesian and Subsilesian nappes (Ksiazkiewicz et al, 1972).The khppen contain crystalHne rocks, shallow-water carbonate rocks, and other rocks of Mesozoic and Cenozoicages. They may represent fragments from European basement rocks, but their source is still somewhat u ncertain.

Except for these rare klippen, there are no certain representatives of basement for most of the flysch terrane.Studies of similar flysch in Aus tria have suggested th at atleast part of it was deposited at oceanic depths (Butt andHerm, 1978). In Austria, fragments of basalt, diabaseultramafic rocks, and serpentinite are present as tectonicblocks in scattered localities. Such rocks are rare in theCarpathians, except in the flysch nap pes that belong to theinternal element in the eastern Carpathians (Burchfiel,1976). Although the flysch nappe s probably lie above continental crust, as suggested by geophysical data (Roth andLesko, 1974; Radulescu et al, 1976), they are entirelyallochthonous. Sparse data suggest the flysch was at leastpartly deposited in deep water, locally of oceanic depths; itcould be inferred deposition occurred partly on very thincontinental or even oceanic crust. Deep earthquakesbeneath Romania suggest a downgoing slab has reacheddepths of about 93 mi (150 km) and subduction beneaththe south Carpathian bend is still active, although it maybe in its final stages (Marza, 1979). These da ta suggest tous that part of the flysch terrane was floored by oceanicand thinned continental crust that has been wholly subducted, but much further study will be necessary to substantiate this conclusion.

The amount of shortening within the foreland fold beltis difficult to calculate. In the western Carpathians, simpletectonic overlaps accoun t for a minimum of 37 mi (60 km)of shortening, and some writers have estimated 62 to 75 mi(100 to 120 km) of northward translation of the thrustsheets (Roth, 1974). Others, based on less constraineddata, have suggested as much as 310 to 375 mi (500 to 600km) of shortening in this part of the belt (Unrug, 1979). Inthe eastern Carp athia ns, palinspastic reconstructions yielda minimum Cenozoic shortening of about 62 mi (100 km )for the foreland belt (Burchfiel, 1976), but rock unitsbetween nappes are missing in many places, thus the realvalue must be much greater. The relation of volcanism inthe intra-Carpathian region to shortening (and subductionbeneath the internal elements) suggests shortening of 93 to124 mi (150 to 200 km) may n ot be unreasonable.PIENINY KLIPPEN BELT

The Pieniny klippen belt is one of the m ost remarkable tectonic units in the Carpathians. It can be traced for more


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1188 Carpathian Foreland Fold and Thrust Beltthan 375 mi (600 km) from the Vienna basin in the west tothe Transc arpathian area in the east (Figs. 3, 4). Generallyit is only a few kilometers wide, reaching about 6 mi (10km) at its widest point. It consists of small and large (up toseveral kilometers long) masses of Jurassic, Cretaceous,and rare Triassic rocks with a variety of facies types. M ostof the rocks are carbonates, but cherts, shales, and sandstones are also present. They are surrounded tectonicallyby shales of Cretaceous and Paleocene age and Eoceneflysch. Study of contemporaneous rock sequences withinthe Mesozoic rocks of these allochthonous masses (calledklippen) has shown they can be divided into several different paleogeographic elements consisting of basin and swellfacies (Andrusov, 1968).

The structural evolution of this belt is indeed complexand interpretations of its history vary greatly. Most w ritersagree that some or all of the paleogeographic units wereinvolved in northward thrusting during the middle to LateCretaceous along with the internal structural elements.Rocks of the klippen belt were then redeform ed, and theiryounger cover deformed for the first time during one ormo re events in the Paleogene and early Neog ene. This history of mu ltiple deforma tions has resulted in a steeply dipping zone which has the aspect of a tectonic melange.Structural features w ithin the zone are sub vertical, dipping both south and no rth. In many places the klippen beltrocks are thrust south against rocks of the internal structural element and its relatively undeformed flysch cover.Even where the structure of the klippen belt dips north , asin Poland , the belt dips to the south at depth beneath theinternal elements (Sikora, 1971). Estimates of shorteningwithin the klippen belt are of the order of 62 mi (100 km;Andrusov, 1968).

FOREDEEPThe Carpathian foredeep is characterized by predominantly Neogene sedimentary rocks largely derived fromthe Carpathians but also with a significant co ntribution ofdetritus from the more external European craton. Theforedeep range s in width from only a few kilome ters in thewestern part of the west Carpathians where it is adjacentto the Bohemian massif, to m ore than 125 m i (200 km) inthe eastern west Carpathians and at the south Carpathianbend (Fig. 2). Neogene rocks, mainly middle Miocene andyounger, lie unconformably on many different rock units,

indicating that the present structure of the foredeep islargely Neogene. Paleogene rocks are present, p articularlyin the inner par t of the foredeep; locally their distributionsuggests that they filled topo graph ic depressions and faulttroughs not associated w ith the present trend of the Carpathians. In the western west Carpathians, Paleogene,Mesozoic, Paleozoic, and even Precambrian rocks under-He the Neogene along northwest-trending fault blocks andsedimentary troughs that reflect fauhs in the Bohemianmassif of the European craton. Similarly, Neogene rockseast of the south Carp athian bend lie on a variety of olderrocks with northwest trends that reflect the structure ofDobrogea in eastern Romania (Fig. 1).In the west Carpathians the Neogene sediments areabout 1.2 mi (2 km) thick. They thicken eastward reaching

more than 3 mi (5 km) locally in Poland and the USSR(Ukranian part of the Carpathians). The greatest thickness of foredeep rocks is east of the south Carpathian bendwhere Neogene rocks are probably 5 to 6 mi (8 to 10 km)thick. Subsidence is still active in this latter sector whereabout 0.6 mi (1 km) of Q uaternary deposits are present.Rock types vary greatly within the foredeep. Conglom erate, sands tone, and shale of molassic character are common as are locally important units of evaporites.Conglomerate tends to be most abundant along the internal side of the foredeep, but in some places like the westernpart of the west Carpathians, only fine grained sedimentsmake up the bulk of the thin foredeep sequence. Coaldeposits are present but rare. The older and more internalunits in the foredeep are locally flyschlike, ch aracterizingthe transition into the more internal flysch units of theforeland fold and thrust belt.Structurally, the foredeep forms a wedge of sedimentary rocks that thickens toward the Carpathians, and the

inner part of the foredeep is caught up in the folding andthrusting of the foreland belt. These deformed foredeeprocks form a transition to the more highly deformed anddislocated rocks of the main fold and thrust belt. Somethrust sheets carrying molassic rocks have been transported a considerable distance, and autochthonousmolasse has been reached by drilling 12 to 15 mi (20 to 25km) south of the front of the foreland belt in Poland(Wdow iarz, 1974).In the region of the south C arpathian b end and extending westward beyond the southern limit of the forelandfold and thrust belt are folds and small thrusts within theNeogene rocks (Fig. 4). Beds as young as Pliocene, andlocally even Pleistocene, are folded. Many of these foldscontain important oil fields such as the famous Ploiestifield. Many of these folds are complicated by salt diapir-ism which takes a variety of complex structural forms(Patru t et al, 1973). These diapiric structures are the type-diapirs first described and n amed by Mrazec (1907).Beneath the youngest molassic rocks of the foredeep arefaults that cut pre-Neogene rocks and the older part of theforedeep sequence. These faults are normal faults withtheir hangwalls dominan tly down thrown toward the internal part of the foredeep. Similar fauhs, with larger displacements up to 0.6 mi (1 km) or more, are presentbeneath the outermost thrust sheets of the foreland foldand thrust belt. Most of these faults do not cut the overriding thrust sheets. The normal faults characteristicallytrend more or less parallel with the Carpathians althoughsome are more oblique and influenced by preexistingfaults in the European craton. These faults can easily beexplained by crustal flexing and loading in response to theweight of the advancing thrust sheets, which has beenexplained by Price (1973) in his model of a migrating foredeep.

INTRA-CARPATHIAN BASINSWithin the Carpathian arc, a series of discrete basinsdeveloped contemporaneo usly w ith the Neogene thrusting

in the foreland fold and thrust belt (Fig. 9). These basinswere superposed on the complex structural units of the


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r+ 1 L a t e Cenoz o i c' * * ' v o l c a n i c r o c k s

I S O P A C H I N T E R V A L2 k m

'- 2 - 3 k m

> 3 km

p00o3-3Q.J3O>


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1190 Carpathian Foreland Fold and Thrust Beltinternal structural element (and locally on the externalstructural element) of the foreland fold and thrust belt(Vienna basin). Basin development and thrusting wereaccompanied by a magmatic arc that extended from theGraz basin in Austria to the south Carpathian bend (Lexaand Konecny, 1974). The rocks are primarily rhyolite,andesite, and dacite. Lherzolite inclusions within theandesitic rocks suggest a mantle source. Eruption ofmagm a was episodic (Fig. 9); there is a clear migration ofvolcanic activity along the arc becoming younger to theeast, and perhaps from the center of the intra-Carpathianregion outward toward the foreland fold and thrust belt.The internal position of the magmatic arc and its spatialand temporal relation to the foreland fold and thrust beltsuggests it is related to south-dipping subduction beneaththe foreland belt (Burchfiel, 1980).

Neogene BasinsDuring and after the last stages of thrusting in the outerCarpathians, a series of discrete basins opened within theintra-Carpathian region. The basins are characterized byhigh heat flow a nd thin crust, and exhibit a two-phase subsidence history (Sclater et al, 1980). The first phase consists of (usually) rapid, K arpatian-Baden ian (17. 5 to 13.0Ma) subsidence that corresponds to a m aximum sedimentthickness (in the Vienna and Transcarpathian basins) of2.5 mi (4 km). These mostly shallow-water sediments arewell localized within distinct fault-bounded basins, aretypically cut by synsedimentary normal faults and exhibitrotated bedding.The second phase consists of a slow, long-term subsid

ence which began at the end of the Badenian and has continued to the present. Sedimentary rocks deposited d uringthis phase are generally flat-lying and unfaulted. Theirhorizontal extent is much greater than th at for first-phasesediments. These sediments onlap the pre-Neogene basemen t. The onset of both phases is approximately th e samefor all basins, although subsidence began slightly earlier(18.5 Ma) in the Vienna and Transcarpathian basins, andthe first phase is poorly developed in the Pannonian basin.This pattern of developm ent does not seem to fit the Tran-sylvanian basin, which shows normal crustal thickness,low heat flow, and little or no signs of extension or faulting during the Miocene (Ciupagea et al, 1970). Furthermore, this basin has undergone recent uplift and erosion,although there has been little or no folding and deformation of the basin fill. Surface elevation of the Transylva-nian basin is approxima tely 1,930 to 2,300 ft (600 to 700m) whereas the other basins have elevations of ab out 325 ft(100 m) above sea level. Basin developmen t w as accompa nied by folding in the southwest Pannonian region (Savafoldbelt). These folds are arranged en echelon, and thefold axes trend approxim ately east-west (Fig. 9).The generalized development of an extensional basinmay be divided into two stages (McKenzie, 1978; Roydenand Keen, 1980). During and immediately after extensionthere is a rapid change in elevation (usually subsidence).This occurs in isostatic response to net density changes

resuhing from c rustal thinning and from heating and thermal expansion. The second stage of subsidence is a rela

tively long-term process caused by cooling andcontraction of the lithosphere following the extensionphase. The overall subsidence is generally amplified by theeffects of sediment loading. If original crustal thickness,elevation, and temperature structure are known, a detailedanalysis of subsidence history can be used to determ ine themagnitude of extension.The subsidence of the intra-Carpathian basins may beinterpreted as the result of Badenian extension thataffected the entire intra-Carpathian region, but was inho-mogeneous and left some blocks emergent and relativelyundeformed (Sclater et al, 1980). The two types or phasesof subsidence can be summarized as follows: the firstrapid, fault-boun ded, and extremely localized phase is dueto crustal extension; the second, long-term subsidence ofgreater area extent is thermally controlled by decay of athermal anomaly produced by extension of the crust andunderlying lithosphere. The first phase is a response toactive processes in the crust and lithosphere, which arereflected by synsedimentary faulting and rotation of bedding. The second is the result of the passive cooling of thelithosphere toward thermal equilibrium, and sedimentsare correspondingly flat-lying and undisturbed.

Sclater et al (1980) gave a detailed analysis of subsidence, heat flow, and crustal thickness of the intra-Carpathian basins. They conclude that these basinsformed by 100% extension of the crust and underlyinglithosphere (although the Pannonian and East Danubebasins seem to require an additional heat source whichmay be related to subduction along the arc). On the average, these estimates of extension are probably too high,since they were generally determined from subsidence ofthe deepest part of the basins.Although the major phase of tectonic activity andextension was completed in the middle to late Miocene,minor tectonic activity continues through much of theCarpathian region. Strike-slip and normal fauhs of smalldisplacement cut Pliocene sediments in the Pannonianbasin and along the arc. Folding tmd deep earthquakes inRomania suggest continued subduction (Fuchs et al,1979). Weak seismic events have also been detected in theintra-Carpathian area (magnitude less than 5.5), and whileit is not clear that present seismic activity is a direct reflection or con tinuation of earlier events, the sense of m otionand local stress pattern inferred from fault plane solutionsis compatible with tha t inferred for Miocene events (Gut-deutsch and A ric, 1976).

A MEC HANICAL MODEL FOR MIOCENE BASINFORMATIONAlthough Sclater et al (1980) demonstrated that the intra-Carp athian b asins were formed by lithospheric and crustalextension, the exact relationship of these extensional processes to the geometry and structural evolution of thebasins is unclear. T he small size, limited areal extent, a ndapparent isolation of the intra-Carpathian basins presentcertain c onceptual difficulties when this local extension isviewed in light of some consistent pattern of regional

extension and deformation. Small, isolated basins of similar character occur along the San Andreas fault (Crowell,


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B. C. Burchfiel and L. Royden 11911974) and in other regions where they appea r to be associated with major strike-slip faults (Q uennell, 1959). In m apview, these "pull-apart" basins are typically rhombohe-dral and are bounded by steep normal faults. This phenomenon has been demonstrated theoretically by finitedifference calculations (Segall and Pollar d, 1980).

In map view, the Vienna basin, which was superimposed on early Miocene and older thrust nappes, has arhombohedral shape, suggesting that it may have formedas a pull-apart bas in along strike-shp fault segments (Fig.10). Large normal faults that cut the basement and oldersediments indicate that the major phase of extension wasearly and middle Badenian and that strike-slip faultingmay have been most active at this time. The orientatio n ofthe rhombohedron-shaped sedimentary basin fill and thelocation of known faults suggest that the sense of lateralplacement was left-sHp. Minor seismic activity along anortheast-trending fault at the southeastern end of thebasin produce fault plane solutions that indicate east-westextension and steeply dipping or vertical nodal planes(Gutdeutsch and Aric, 1976). These data are also consistent with left-slip. The strike-slip components of motionalong this and other faults that strike northeast are difficult to show directly by surface and subsurface mappingbecause these faults strike parallel with the thrust fronts.Because the deepest part of the V ienna basin appears to bethe resuk of 50 to 100% extension of the crust, probably afew tens of kilometers of strike-slip displacement are suf-ficient to produce the Vienna basin. This general style ofbasin formation can be extended to other intra-Carpathian basins.

Figure 11 shows one system of strike-slip faults tha t canexplain the extension of the intra-Carp athian basins, theirindividual geometries, and their relationships to eachother. Some faults are known from geologic and subsurface mapping. Other faults were identified by steep iso-pach gradients in the sediments or from air photos, andsome of these may have large components of lateralmotion. Direct geologic corroboration is somewhat difficult to obtain for many of these fauUs. Much of the intra-Carpathian region is covered by post-extensional(Pliocene and Q uaternary) sediments. Drill hole data andreflection seismic profiles show norm al faults and growthfaults bounding many of these basins, but the ability todetect horizontal displacements from these data in the subsurface is limited. Some faults in the emergent blocks andin the Carpathians have surface traces that disappearunder the Neogene basins. Many of these have been interpreted as thrust and normal faults, as in the Vienna basin,but because they strike parallel with the thrust fro nts, theymay have a significant component of lateral offset. Theage of many of these faults is also uncertain, althoughsome are known to be late Miocene. Up to 19 mi (30 km) ofBadenian (16.5 to 13.0 Ma) displacement can be shown onsome of these faults (Z. Balla, personal commun.).

There is some latitude in connecting the va rious faults inFigure 11, but the variations appear to be of only minorimportance. Regional extension occurred along mostlynorth east- a nd no rth west-trending sets of conjugateshears, and the sense of motion is consistent w ith east-westextension (northeast sinistral shear, northwest dextral

shear). Not all the basins are directly analogous to simplerhombic pull-aparts, Uke the Vienna basin, and manyformed at the intersection of conjugate faults (the Grazand Pannonian basins). However, all (with the probableexception of the Transylvanian basin) are regions of localextension related to motion along strike-slip faults, andthe extended regions form a series of eastward-openingarcs. This style of extensional faulting is similar to recentextension in the Basin and Range province of the UnitedStates, where areas of significant extension are separatedfrom relatively undefo rmed regions by a complex systemI S O P A C H I N T E R V A L

50 km

1 0 5 - 0 ma ( Ponnonion - Presen

3 -10 5 ma (S ormQ Iian)

16 5- 13 ma ( Badenian )

2 2 - l 6 5 m a [ E g g e n b u rg i o n -K a rp l i a n )

l l \ Pre-Neogene basement

FIG. 10Upper diagram: isopach map to base of Miocene forVienna basin. C.I. = 1,000 m. Dots indicate depth of 1 to 2 km;shaded areas, 3 to 4 km; black, more than 5 km. Heavy lines arelocations of major faults (modified from Brix and Schultz,1980). NW-SE line is location of cross section of lower diagram.Lower diagram: cross section through northeast corner ofVienna basin. X's indicate pre-Neogene basement; dashes,Eggenburgian-Karpatian (22 to 16 Ma); shaded area, Badenian(16 to 13 Ma); dots, Sarmatian (13 to 10.5 Ma); white area, Pannonian to present (10.5 Ma to present). After Steininger et al(1975).


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1192 Carpathian Foreland Fold and Thrust Beltof shear zones (Davis and B urchfiel, 1973).The synchronous development of en echelon folds in thewestern part of the intra-Carpathian region (Sava foldbelt) is probably related to this extension and, locally, todextral shear along west-northwest-trending strike-slipfaults. The east-west trend of the fold axes suggests relative north-south compression and is consistent with thestress field inferred from strike-slip faults. Similarly, enechelon folding associated with large shears has been demonstrated experimentally in clays (Wilcox et al, 1973).Regional seismic activity is quite weak, but the few shockslarge enough to produce reliable fault plane solutions arealso consistent with east-west extension and north-southcompression. Most of these Miocene faults appear to bereactivated parts of older Mesozoic tectonic lines. In thenorthwest, left-slip faults parallel the northeast strike ofthe thrust fronts in that part of the west Carpathians (Brixand Schultz, 1980). Similarly, in the region of the T ranscar-pathian depression, dextral shear probably occurred alonga northwest-trending fault system parallel to the thrustfronts. The large shear zone north of the Pannon ian basinappears to have been a major Mesozoic tectonic line ofuncertain significance.

DISCUSSIONExtension in the intra-Carpathian basins, whichoccurred mostly during the Badenian, can be correlated

with the timing of thrusting around the Carpathian loop.At this time north-directed thrusting was essentially completed in the west Carpathians, but major thrusting andcrustal shortening continued in the east Carpathians until13 Ma (Bu rchfiel, 1976; Jiricek, 1979). Hence , it appearsthat east-west crustal shortening in the east Carpathianswas compensated by east-west extension of the intra-Carpathian region. We cannot exclude the possibility thatsome north- south extension also occurred. The migrationof thrusting around the Carpathian belt probably corresponds t o the migration of A-type subduction (Bally, 1975;Burchfiel, 1980), also suggested by the eastward m igrationof calc-alkaline volcanics. In plate tectonic terms, back-arc-type extension of the upper plate seems to haveoccurred while continued subduction along the plateboundary in the east Carpathians was not matched by therate of plate convergence. A similar relation was proposedby Molnar and Atwater (1978) for oceanic back-arcbasins.

Rigid plate convergence during the Miocene was proba bly limited by plate geometry, since Moesia, the easternAlps and the Dinarides have been locked approximatelyinto their present position relative to Europe since earlyMiocene time. Motion between Adria and Europe mayhave been taken up by dextral shear along the peri-Adriatic-Vardar fauh system, thus isolating the Carpathian system from th e rest of the Mediterranean. Miocenesplays off this fault system into the Carpathian region

4 0 0 kmFIG. 11Generalized map of intra-Carpathian basins. Stippled pattern covers areas that have undergone significant extension.Some basins are interpreted as "pull-apart" basins, whereas others are regions of extension bounded by zones of differential shear.Arrows show direction of overall extension.


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B. C. Burchfiel and L. Royden 1193

indicate that isolation was not total, but it is the opinion ofthe writers that the Carpathian region was not stronglycoupled to Adria during the middle and late Miocene.Therefore, if extension of the intra-Carpathian basins didoccur to accommodate crustal shortening in the east Carpathians, the net east-west extension across the entireintra-Carpathian region should be roughly equal to thetotal Badenian and post-Badenian shortening in the EastCarpathians. Such a reconstruction depends on theassumption that in the intra-Carpathian region the crustmay be treated as a system of more or less rigid blocks tha tare bounded by zones of deformation. These zones may begrouped as extensional (basins), transform boundaries

(strike-slip faults and more diffuse shear zones), and com-pressional (overthrusting and folding only along the Carpathian arc). This system is very roughly analog ous to thelarge-scale tectonics of ridges, subduction zones, andtransform faults. The major difficulties in attempting tomake "tectonic reconstructions" at this scale are: (1) thedivision between a "rigid block" and a zone of deformation is somewhat ambiguous, and all the blocks may besubject to some internal deformation, and (2) the magnitude and direction of crustal extension are not alwaysclear. Hence, rigorous reconstruction of pre-Miocenegeometries is not feasible for the intra-C arpathia n region.However, this approach may be useful as a first-order

6 Ma

0 MaMm p p a p p n v n f ^ M' i i " ^ ' * * ! '

0r 400 kmFIG. 12Diagram showing proposed downbending of subducted plate under east Carpathians. Flow is induced in overriding plate tofill space formerly occupied by downbending plate, resulting in extension of the Pannonian basin and temporary suction on the Tran-sylvanian basin. P, Pannonian basin; A, Apuseni Mountains; Ts, Transylvanian basin; EC, east Carpathians; V, Vrancea zone; E,Europe. Dip of plate at 0 Ma constrained by epicenter locations of recent seismic events (Fuchs et al, 1979).


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1194 Carpathian Foreland Fold and Thrust Beltapproximation in determining the relative motions andstyle of deformation within the Carpathian system.Additional information is supplied by estimates ofcrustal shortening in the Carpathians. This constraint islacking for large-scale tectonic reconstruction s, w here therates of subduction of oceanic plates must be inferredfrom the plate motions. By further analogy with the consistency condition for relative motion of large-scale plates,the motions of all the blocks within the system should beinternally consistent. This also applies to the zone ofshortening along the Carpathian loop and to the European foreland.

Assuming 50 to 100% extension in the basins, totaleast-west extension is estimated as 45 to 62 mi (75 to 100km). Com paring this to abo ut 62 mi (100 km) of Mioceneshortening estimated for the east Carpathians by Burch-fiel (1980), the magnitudes are seen to be in fair agreement . However , i t should be remembered thatpalinspastic restoration of thrusts and folds within erogenic belts provides only a minimum estimate of crustalshor tening. This temporal and spat ia l connect ionbetween thrusting, volcanism, and basin extension is notunique to the history of the Carpath ian reg ion. A similarsituation seems to have prevailed during the Eocene,when north-south crustal shortening in the west Carpathians and to the south (Zagreb line) was accompanied bybasin formation and eruption of andesitic magmas,apparently from a deep source. Because the orientationof the regional stress field for this Eocene event musthave been totally different from that for later Mioceneextension, it is not surprising that the orientation andlocation of the Paleogene basin appears totally unrelatedto that of the Neogene basins.

The development of extensional basins adjacent to azone of subduction and apparent compression suggeststhat this subduction cannot be driven by rigid plate convergence, as it is difficult to understand how compres-sional stress can be transmitted across an extendingterrane. Furthermore, motion between Europe andAdria appears to be taken up along a transform boundary parallel with the strike of the D inaric Alps and unrelated to Miocene shortening in the east Carpathians.Hence, in the east Carpathians subduction must bedriven from the subduction zone itself and might beaccomplished by gravitational forces acting on the down-going slab. One possibility is that shortening in the eastCarpathians and corresponding basin extension are theresult of bending of an initially shallow-dipping subducted plate into a vertical position and retrogrademotion of the position of the trench (Fig. 12). This mechanism is supported by intermediate-depth earthquakes inRomania that indicate the existence of a vertical slab todepths of 93 mi (150 km) with downdip component ofextension (Fuchs et al, 1979). Furthermore, this mechanism can explain the migration of the andesitic volcanicarc toward the suture. In the east Carpathians, some ofthese volcanic rocks are very near to and almost superimposed on the suture zone. This mechanism m ay be a common process fo l lowing the terminat ion of t ruesubduction and may help to explain the spatial and temporal distribution of volcanism in some Mediterraneanback-arc basins.

This downbending of the subducted plate may alsohelp to explain the existence of the Transylvanian basin.As mentioned, this basin shows little or no evidence ofextension and heating , a ud its 1.9 to 2.5 mi (3 to 4 km) ofMiocene sediment fill a regular, saucerlike depressionthat is not cut by major faults. In addition , the basin hasexperienced recent uplift and erosion, although littlecompression or folding has occurred. The mean surfaceelevation of the Transylvanian basin is 1,970 to 2,300 ft(600 to 700 m) significantly higher than the other intra-Carpathian basins, which are approximately 325 ft (100m) above sea level. If erosion co ntinues until the surfaceof the basin is reduced to sea level, isostasy predicts that1.9 mi (3 km) of sediment will have been eroded, and thatless than 0.6 mi (1 km) of M iocene sediment will remainin the ba sin.This suggests that w hatever the active forces that produced the M iocene subsidence of the Transylvanian basin(amplified by the effect of sediment loading), they weretransient and are no longer active in this region. Oneexplanation may be that this downwarping of the basement was produced by downward pull or suction fromthe downbending plate (Fig. 12). When this action was"turn ed off," the basement would have rebounded to itsoriginal elevation if the weight of the accumulated sediments could have been removed instantly by erosion.Because erosion proceeds slowly, the surface of th e basinhas been uplifted, and is being gradually reduced to sealevel. A similar situation may exist at present in the V ran-cea foredeep region of southeastern Romania, whereextremely rapid subsidence of the area, which is not inisostatic equilibrium, m ay be related to the occurrence ofintermediate depth earthquake on the downgoing plate

(Fuchs etal, 1979).We suggest that, in this region, subduction or downbending of the subducted slab results in an extensionalstress field, and that thrusting and a pparen t compressionalong the mountain belt are only thin-skinned, superficial effects because of the inability to subduct light,upper crustal material and the detachment of the crustfrom the underlying Uthosphere. This may explain thefocal mechanisms for both crustal (thrusting) and m antle(downdip extension) earthquakes in the eastern Carpathians. Extension of the intra-Carpathian basins mayhave occurred because of lithospheric flow to fill thespace left by the retreating or do wnbend ing p late. In thisregion, the dominant forces seem to be those acting onthe downgoing slab, which produce sh ortening along theplate bound ary, and plate geom etry which inhibits rigid-plate convergence.

There is no reason to believe that the processes thatproduced the Carpathian loop and subsided intra-Carpathian region were fundamentally different fromthose operating elsewhere in the Alpine belt. The openingof the intra-Ca rpathia n basins can be explained purely asthe result of local geometric and s tructural c onstrain ts.R E F E R E N C E S C I T E D

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B. C. Burchfiel and L. Royden 1195Brix, F., and O. Schultz, eds., 1980, Erdoel und erdgas in Oesterreich:Verlag Naturhist. Mus. Wien und F. Berger Hoon, 311 p.Burchfiel, B .C ., 1976, Geology of Rom ania: Geol. Soc. America Spec.Paper 158, p. 1-82.1980, Eastern alpine system and the Carpathian orocline as anexample of collision tectonics: Tectonophy sics, v. 63, p. 31-62.Butt, A., and D. Harm, 1978, Paleo-oceanographic aspects of IheUpper Cretaceous geosynclinal sediments of the eastern Alps, in H.Closs, D, Rolder, and K. Schmidt, eds., Alps, Apennines, Hel-lenides: Inter-Union Commission on Geodynamics Sci. Rept. 38, p.87-94.Ciupagea, D., M. Panca, and T. Ichem, 1970, Geologia DepresiungTransilvaniei: Bucharest, Roma nia, Editura A cad., 256 p.Contescu, L. R., 1974, Geologic history and paleogeography of easternCarpathians; example of alpine geosynclinal evolution: AAPCBull., V. 58, p . 2436-2476.Crow ell, J. C , 1974, Origin of late Cenozoic basins in southern C alifornia, in W. R. Dickinson, ed.. Tectonics and sedimentation:SEPM Spec. Pub . 22, p. 190-204.Dany s, V. V., et al, 1974, Flysch Carp athia ns, in M. Mahel, ed .. Tectonics of the Carpathian-Balkan regions: Bratislava, Geol. Inst.Dionyz Stur, p. 210-220.Davis, G. A., and B. C. Burchfiel, 1973, Garlock fault; an intraconti-nental transform structure, southern California: G eol. Soc. America Bull., v. 84, p. 1407-1422.Fuchs K., et al, 1979, The Romanian earthquake of March 4, 1977;aftershocks and migration of seismic activity: Tectonophysics, v.53 ,p. 225-247.Gszczypko, N., 1975, Exotic rocks in the Paleogene of the Maguranappe between Dunajec and P oprod Rivers: Polskie. Towarz. Geol.Rocznik, v. 45, p. 403-431.Gutdeutsch, R., and K. Aric, 1976, Eine diskussion geophysikalischermodelle des Grenzbereiches zwischen Ostalpen und PannonischemBecken: Acad. Sci. HungaricaeA ctaGeol. , v. 21 , p. 287-309.Jiricek, R., 1979, Tectonic development of the Carpathian arc in theOligocene and Neogene, in M. Mahel, ed.. Tectonic profilesthrough the west Carpathians: Bratislava, Geol. Ustav DionyzaStura, p. 205-214.Ksiazkiewicz, M., 1956, Geology of the northern Carpathians: Geol.Rundschau, V. 45, p. 369-411.1963, Evolution structurale des Carpathes Polonaises: LivreMemoire P rofesseur P. Fallot, v. 2, p. 529-562. et al, 1972, On the beds und erlying the An drucho w klippen:Polskie. Towarz. Geol. Rocznik, v. 42, p . 239-269.Kusmierek, J., 1979, Gravitational deform ations and backward over-thrusts with references to deep structures and petroleum prospectsof the Dukla unit foreland in the Bieszczady Mountains: PolskaAkad. Nau k-Pra ce Geo l. 114, 68 p. (in Polish).Lexa, J., and V. Konecny, 1974, The Carpathian volcanic arcdiscussion: Acad. Sci. Hungaricae Acta Geol., v. 18, no. 3-4, p.279-293.Marza, V. 1., 1979, A seismicity pattern of March 4, 1977, Vrancea,Romania earthquake; an earthquake prediction in sight: Tectonophysics, v. 53 , p. 217-222.McKenzie, D., 1978, Some remarks on the development of sedimentarybasins: Earth and Planetary Sci. Letters, v. 40, p. 25-32.

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