Csontos 2004-Mesozoic Plate Tectonic Reconstruction of the Carpathian Region

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Palaeogeography, Palaeoclimatology, Palaeoecology 210 (2004) 1–56

Mesozoic plate tectonic reconstruction of the Carpathian region

Laszlo Csontosa,*, Attila Vorosb

aGeological Department, Eotvos University of Sciences, Budapest 1117 Pazmany P: setany 1/a, HungarybGeological and Palaeontological Department and HAS-HNHM working group for palaeontology, Hungarian Natural History Museum,

Budapest, Muzeum krt. 14-16, Hungary

Received 4 April 2003; received in revised form 28 January 2004; accepted 20 February 2004

Abstract

Palaeomagnetic, palaeobiogeographic and structural comparisons of different parts of the Alpine–Carpathian region suggest

that four terranes comprise this area: the Alcapa, Tisza, Dacia and Adria terranes. These terranes are composed of different

Mesozoic continental and oceanic fragments that were each assembled during a complex Late Jurassic–Cretaceous–

Palaeogene history. Palaeomagnetic and tectonic data suggest that the Carpathians are built up by two major oroclinal bends.

The Alcapa bend has the Meliata oceanic unit, correlated with the Dinaric Vardar ophiolite, in its core. It is composed of the

Western Carpathians, Eastern Alps and Southern Alcapa units (Transdanubian Range, Bukk). This terrane finds its continuation

in the High Karst margin of the Dinarides. Further elements of the Alcapa terrane are thought to be derived from collided

microcontinents: Czorsztyn in the N and a carbonate unit (Tisza?) in the SE. The Tisza–Dacia bend has the Vardar oceanic unit

in its core. It is composed of the Bihor and Getic microcontinents. This terrane finds its continuation in the Serbo–Macedonian

Massif of the Balkans.

The Bihor–Getic microcontinent originally laid east of the Western Carpathians and filled the present Carpathian

embayment in the Late Palaeozoic–Early Mesozoic. The Vardar ocean occupied an intermediate position between the Western

Carpathian–Austroalpine–Transdanubian–High Karst margin and the Bihor–Getic–Serbo–Macedonian microcontinent. The

Vardar and Pindos oceans were opened in the heart of the Mediterranean–Adriatic microcontinent in the Late Permian–Middle

Triassic. Vardar subducted by the end of Jurassic, causing the Bihor–Getic–Serbo–Macedonian microcontinent to collide with

the internal Dinaric–Western Carpathian margin.

An external Penninic–Vahic ocean tract began opening in the Early Jurassic, separating the Austroalpine–Western

Carpathian microcontinent (and its fauna) from the European shelf. Further east, the Severin–Ceahlau–Magura also began

opening in the Early Jurassic, but final separation of the Bihor–Getic ribbon (and its fauna) from the European shelf did not

take place until the late Middle Jurassic.

The Alcapa and the Tisza–Dacia were bending during the Albian–Maastrichtian. The two oroclinal bends were finally

opposed and pushed into the gates of the Carpathian embayment during the Palaeogene and Neogene. At that time, the main N–

S shortening in distant Alpine and Hellenic sectors was linked by a broader right-lateral shear zone along the former Vardar

suture.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Plate-tectonics; Carpathians; Oroclinal bending; Tectonic transports; Palaeomagnetic data

0031-0182/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.palaeo.2004.02.033

* Corresponding author.

E-mail address: [email protected] (L. Csontos).

L. Csontos, A. Voros / Palaeogeography, Palaeoclimatology, Palaeoecology 210 (2004) 1–562

1. Introduction

1.1. Aims, structure of the study

In recent years, our systematic review of the

Mesozoic formations and structures of the Carpa-

thian–Pannonian region have yielded new insights

to the plate-tectonic evolution of the area that enable a

reevaluation of the palaeogeographic evolution of this

region during the Mesozoic era (Fig. 1). Other studies

have already dealt with the Cenozoic development of

the area in more detail (e.g. Balla, 1984; Csontos,

1995; Csontos et al., 1992; 2002; Fodor et al., 1999;

Haan and Arnott, 1991; Kovac et al., 1994, 1998),

therefore this study focuses on a Mesozoic plate-

tectonic reconstruction.

This study is composed of six main parts. The aim

of the first part is to briefly introduce the basic

geological features and tectonic events of the Carpa-

thian area. The different Cenozoic events and related

Fig. 1. Geography of the Carpathian area. Grey shaded digital terrain mod

30-arc-second Digital Elevation Model. Major geographic units and impo

deformations are discussed in the second part. The

third part deals with the nappes formed in the Meso-

zoic. The fourth part attempts a correlation between

the different structural units to arrive at the key

intervening oceans and continents. The fifth part lists

the geologic evidence for the timing of the main plate

tectonic events of the area. Finally, the sixth part

concentrates on the Mesozoic reconstruction.

1.2. Geographic–geologic outline of the Carpathian

area

Medium high mountains (1500–2500 m above sea

level) encircle an Intra-Carpathian basin system (ca.

100 m above sea level; Fig. 1). Geographically, it

appears that the Alpine chain is split into two east-

wards: one branch continues to form the Carpathian

arc in the north and the other the Dinaric chain in the

south. Then the two chains are reunited in Serbia

(Mahel’, 1973). To a first approximation, the geolog-

el from National Oceanic Atmospheric Administration (USA) global

rtant mountains are marked.

L. Csontos, A. Voros / Palaeogeography, Palaeoclimatology, Palaeoecology 210 (2004) 1–56 3

ical structure of the two mountain branches is sym-

metrical. Both branches form outward verging nappes.

Investigation of the basin floor revealed that there is

no oceanic crust beneath the Cenozoic infill of the

Intra Carpathian basin system, but according to its

Mesozoic–Palaeogene composition it can be subdi-

vided into two distinct parts.

1.3. Structural events

The Mesozoic–Cenozoic sedimentary and struc-

tural evolution is summarised in a generalised and a

Fig. 2. Simplified stratigraphic diagram showing the main nappe units and

of the terranes within the Carpathian area. Numbers correspond to tectoni

cover a roughly NW–SE section across the area. EC=External Carpathian

much simplified terrane analysis diagram (Fig. 2).

Five main structural phases can be recognised in the

Mesozoic to Cenozoic evolution of the Carpathian

area, from young to old:

1. Middle Miocene large-scale back-arc extension in

internal zones, coeval with subduction in the

external zones, interrupted by smaller amplitude

strike slip and positive inversion episodes.

2. Palaeogene amalgamation of two composite ter-

ranes, Late Palaeogene–Early Miocene right-lateral

shear along the Mid-Hungarian zone (Periadriatic

structural phases recognised in the Mesozoic to Cenozoic evolution

c phases described in the text. The schematic stratigraphic columns

s.


lineament system), followed by major rotations of

the terranes.

3. Late Cretaceous oroclinal bending of the two

composite terranes, i.e. development of a system

of lateral shears coupled with major thrust and

normal faults.

4. Mesozoic nappe emplacement with a Late Jurassic,

an Early Cretaceous and a mid-Cretaceous (Albian)

peak; collision of microcontinents.

5. Middle Triassic to Late Jurassic rifting in several

distinct zones resulting in oceanic troughs or large

oceans; drifting of microcontinents.

In the following, the main structural units and their

lithologic content corresponding to each event are

described in more detail. Then the structural events

responsible for the particular tectonic situation will be

described. The first event is responsible for the

present-day geology, so this will be detailed first.

1.4. Middle–Late Miocene tectonic events

During this event, the whole Carpathian area can

be subdivided into three major domains (Fig. 1): the

External Carpathians, composed mainly of Late Cre-

taceous–Cenozoic turbidites; the Internal Carpathians

and the Intra-Carpathian basin. All the internal moun-

tain areas, as well as the Dinaric chain contain a more

or less continuous exposure of Mesozoic rocks and in

some cases their Palaeozoic or crystalline basement.

The present geologic pattern is the result of Mio-

cene subduction-docking in the external parts of the

Carpathian arc (Fig. 3A) (Lillie and Bielik, 1992). The

roll back of the subducted European lithosphere

created the Intra-Carpathian (=Pannonian) back-arc

basin, the opening of which was synchronous with

thrusting of the External Carpathian foredeep–fore-

land basin (Balla, 1984; Horvath and Royden, 1981;

Linzer et al., 1998). A Middle–Late Miocene calcal-

kaline volcanic arc parallel to the outlines of the chain

borders the Intra-Carpathian basin (Balla, 1984; Szabo

et al., 1992). During this tectonic episode the whole

Intra-Carpathian area behaved as a uniform, but not

rigid upper plate against the subducting European

plate. The Intra-Carpathian basin is underlain by a

thin continental crust (Meissner and Stegena, 1988),

which is variable in thickness and composition. There

are isolated internal mountains (inselbergs) within the

Intra-Carpathian basin, like the Transdanubian Range

or the Mecsek Mountains (Fig. 1). The crustal thin-

ning of the upper plate due to the roll-back effect

(Horvath and Royden, 1981), and the geochemistry of

the Miocene volcanic arc rocks (Szabo et al., 1992)

strongly suggest that at least part of the subducting

European lithosphere was of oceanic nature. This

statement remains valid in spite of missing evidence

of the oceanic crust itself (Winkler and Slaczka,

1992). The subducted European margin can be seen

on seismic reflection sections (Tomek et al., 1987,

1989) and colder, denser detached material is visible

on seismic tomography (Spakman, 1990; Sperner et

al., 2001). The most plausible location to place this

oceanic basin is the now detached, subducted, original

substratum of the Alpine Flysch belt, and the External

Carpathian flysch nappes (Fig. 2).

Rocks underwent different styles of deformation in

the Neogene. In the Intra-Carpathian area the dominant

style was an intra-continental stretching concentrated

either along low-angle normal fault zones or distrib-

uted to zones of wide rifting (Fig. 3A) (Dunkl and

Demeny, 1997; Horvath, 1993; Tari et al., 1992, 1999;

Fodor et al., 1999). As a consequence, pre-Neogene

basement is commonly broken up and tilted in smaller

blocks. Strike-slip faulting seems to form local basins

in the external part of the Intra-Carpathian Basin, e.g.

Vienna Basin (Fodor, 1995; Royden, 1988) or along

Late Neogene internal shear zones (Horvath, 1993;

Csontos, 1995; Prelogovic et al., 1998). During Late

Neogene to Quaternary, especially in the SW part of

the Intra-Carpathian Basin, but also along NE–SW

striking deformation belts, the basins were inverted

and sometimes formed narrow pop-ups or transpres-

sional–compressional belts, like the Balaton fault area

or the Sava Folds (Fig. 3A) (Balla et al., 1987;

Csontos, 1995; Csontos and Nagymarosy, 1998; Fodor

et al., 1998, 1999; Tomljenovic and Csontos, 2001).

Palaeomagnetic data acquired in the Dinaric chain

suggest that the whole chain and the southernmost

inselbergs of the Pannonian basin show 35j counter-

clockwise rotation during Late Miocene–Pliocene

(Fig. 3) (Marton, 1987, 1993b; Marton et al., 1999,

2002). Meanwhile, the southern part of the Dinaric

chain (=Hellenic arc) suffered similar clockwise rota-

tions (Kissel et al., 1985; Marton, 1987). These

opposite rotations can be compensated in the Scu-

tari–Pec deformation zone (Fig. 1) (Aubouin et al.,

Fig. 3. Middle–Late Miocene structural features of the Carpathian area. (A) Palaeomagnetic data after Marton and Marton (1989, 1996, 1999),

Marton et al. (1992, 1999, 2002), Mauritsch and Marton (1995), Panaiotu (1998) and main structural elements after Csontos (1995), Csontos et

al. (2002), Tomljenovic and Csontos (2001). (B) Reconstruction of the Middle Miocene situation modified after Csontos et al. (2002). (C)

Reconstruction after Late Miocene–Pliocene inversions and rotations.


Fig. 4. Late Palaeogene–Early Miocene structural features of the Carpathian area. (A) Palaeomagnetic data after Bazhenov et al. (1993), Krs et

al. (1982, 1991), Marton and Marton (1989, 1996, 1999), Marton et al. (1992, 1999), Mauritsch and Marton (1995), Panaiotu (1998), Patrascu et

al. (1990, 1992, 1994) and main structural elements after Csontos (1995), Csontos and Nagymarosy (1998), Fodor et al. (1992, 1998, 1999,

2002), Tari et al. (1999). (B) Reconstruction of the right lateral shear along the Periadriatic lineament system. (C) Reconstruction of the opposite

major rotations.



1970), where oblique thrusts and a slight bending are

suggested. This bending did not concentrate on the

coastal part of the two chains, but possibly affected

the internal parts as well. A reconstruction of the

original Middle Miocene positions (Csontos et al.,

2003) suggests that all the now dogleg shaped internal

Dinaric–Hellenic units and structural zones were

straight (Fig. 3B,C). The Late Miocene–Pliocene bulk

rotation of the Dinaric chain might be one of the

driving forces of inversion and uplift in the Carpathian

area at the same time.

Apart from low-angle normal fault-bound core-

complexes and localised fold-thrust belts, this last

tectonic phase did not seriously disturb the Palae-

ogene–Early Miocene structural pattern. Only some

attempts to estimate the amount of stretching do exist.

Horvath and Royden (1981), Tari (1994) and Tari et

al. (1999) proposed a stretching factor of h=1.5, basedon estimates of crustal (1.2–1.8) and mantle litho-

spheric (2–2.4) thinning. Higher values were estimat-

ed near the core complexes (Tari et al., 1999).

Because reconstruction of individual structural ele-

ments is often impossible, we applied these rough

estimates in our reconstruction (Fig. 4C).

The Late Tertiary formations are of continental or

shallow marine facies in the Intra-Carpathian area.

The varied topography also resulted in a variety of

heteropic facies rocks, like shallow marine limestone

and basinal clay. A deep lake prograding delta

sequence of Late Miocene–Pliocene age is the

thickest formation of the basin fill. In the foredeep

a typical shallowing-up stratigraphic sequence was

deposited. The deposits are dominated by siliciclastic

rocks at all places.

2. Palaeogene events

2.1. Terrane nomenclature

Terrane for us means a collage of structural units of

different geodynamic origins (e.g. Hamilton, 1990;

Voros, 1988), which, however, behaves as a main

and more or less rigid structure during a particular

tectonic event. This concept will be used for the pre-

Middle Miocene period, since after that the whole area

could be considered a single terrane. Naturally, ter-

ranes evolve through time: get new amalgamated

material or lose some by rifting. To avoid confusion,

the terranes we use stand for the Late Cretaceous–

Paleogene situation. The precursors will be called

differently. Since palaeo-plates are very hard to define,

we rather use the term geodynamic unit to designate

once (micro)continents and oceans, bearing in mind

that the former plate distribution could comprise both

oceanic and continental material. All this material is

now found in individual thrust slices, nappes. These

nappes form in fact groups of tectonic slices with

similar stratigraphic/facies content.

In the last years a succession of papers appeared on

the terrane analysis of the European Alpides from the

Western Alps (Neubauer et al., 1997) and the Intra-

Carpathian basin (Kovacs et al., 1997, 2000; Vozarova

and Vozar, 1997) to the Dinaric chain (Karamata and

Krstic, 1996). In the present paper we use a simplified

terrane classification, following the above definitions.

2.2. Palaeogene terranes of the Carpathian area

The Carpathian area, as an Alpine Cretaceous to

Cenozoic orogenic collage, can be subdivided into

three major composite terranes (Fig. 5) (Balla, 1984;

Csontos, 1995; Kovacs et al., 2000). The terranes are

defined by contrasting Triassic and Jurassic sedimen-

tary facies, most spectacularly demonstrated between

the now neighbouring Transdanubian Range and

Mecsek Mts. in Hungary (Fig. 1) (Kovacs, 1982;

Kovacs et al., 2000; Voros, 1977, 1984, 1988). Even

Palaeogene–Early Miocene stratigraphy is different

(Csontos et al., 1992). Based on palaeobiogeographic

work in the Transdanubian Range and the Mecsek and

Villany Mountains, a distinction between the Meso-

zoic fossil assemblages of the Intra-Carpathian area

could be made (Geczy, 1973, 1984). The Transdanu-

bian Range, and Mecsek-Villany Mountains belonged

to two different faunal provinces in the Jurassic (Fig.

5). Since this work, there is now a wealth of palae-

obiogeographic data especially for the Jurassic period.

Diagnostic fossils include brachiopods (Dulai, 1990;

Voros, 1977, 1984, 1988, 1993) ammonites (Geczy,

1973, 1984; Meister and Stampfli, 2000), bivalves

(Szente, 1990), gastropods (Szabo, 1988, 1990), and

palynomorphs (Lachkar et al., 1984). All the Austro-

alpine nappes and the nappes of the Inner Western

Carpathians belong to the southern, Mediterranean

faunal province, just as the Transdanubian Range,

Fig. 5. Major terranes of the study area shown on a palaeobiogeographic map for the lower half of the Jurassic (Sinemurian–Bathonian).

Modified from Voros (1992, 2001). Contours of main terranes marked as thick dashed lines after Csontos (1995), modified. TR: Transdanubian

Range, SMM: Serbo–Macedonian Massif. (For color see online version).


the Eastern Carpathian Persani unit and the Dinaric

Mountains (Fig. 5). On the other hand, the Lower

Jurassic formations of the Helvetic zone of the Alps

and the Mecsek, Villany and Apuseni inselbergs,

together with the Eastern and Southern Carpathians

have a stable European continental margin fauna,

indicating a close palaeogeographic contact between

them (Voros, 1993, 2001).

2.2.1. Alcapa terrane

This is an elongate and structurally complex ter-

rane, extending from the Alps to the Western Carpa-

thians (Fig. 5). Its northern limits are the Gresten–St.

Veit Klippen, the Pieniny Klippenbelt (and the grad-

ually accreted External Carpathian flysch nappes,

Oszczypko, 1992), the southern limit is the Mid-

Hungarian zone (Fig. 4). An important internal strike

slip zone: the Balaton–Periadriatic line runs north of

this limit. There is a Late Palaeogene basin along and

north of this structural zone, which is dissected by

later movements. The fill is a deepening upward series

of Late Eocene limestones, marls, grading to Oligo-

cene clays, tuffitic shales. In the Early Miocene

shallowing upwards clastic rocks fill up the basin.

Another Palaeogene basin is found in the northern

part of the Western Carpathians and is called Intra-

Carpathian flysch basin. Its fill is an Eocene–Oligo-

cene deep marine siliciclastic turbidite.

2.2.2. Tisza–Dacia terrane

This terrane occupies the central and eastern part of

the Intra-Carpathian area (Fig. 5). Its Mesozoic rocks

appear on the surface only near the eastern (Eastern

and Southern Carpathians) and western terminations

(Slavonian inselbergs, Mecsek, Villany) and in the

Apuseni Mountains; the intervening parts form the

basement of the Great Hungarian Plain and the

Transylvanian Basin and are covered by thick Ceno-

zoic sediments. The northern boundary of this terrane

is the Mid-Hungarian zone, whereas in the south the

boundary is formed by the Sava fault (Fig. 4). The

terrane can be subdivided into a Tisza (northwestern)

and a Dacia (curvilinear part in the Eastern and

Southern Carpathians) part. Their internal limit

marked by ophiolites is beneath the Transylvanian

Basin and is sealed by Palaeogene continental to

marine beds. On the northern periphery of the

Tisza–Dacia terrane a Late Cretaceous–Early Mio-

cene turbidite basin is found. This is called Szolnok

flysch beneath the Great Hungarian plain, and Borsa


basin in exposures of northern Transylvania (Nagy-

marosy and Baldi-Beke, 1993; Szepeshazy, 1973).

2.2.3. Adria terrane

This is the largest terrane of the area commonly

called the peri-Adriatic region, or Apulia, or Adriatic

promontory (Fig. 5). The eastern, Dinaric and west-

ern, Apenninic margins are strongly compressed into

huge nappe systems and the southern margin is

concealed under the modern Mediterranean Sea. Only

the northeastern margin of Adria, the Dinaric chain

belongs to the main area of the present study. The

limit of the Dinarides with adjoining units is formed

by the Sava fault zone (Fig. 4). This must have been a

mobile zone, where Late Cretaceous–Eocene turbi-

dites were deposited and a series of Oligocene gran-

ites were emplaced (Pamic, 1998b, 2002). Another

mobile zone is found on the Adria margin, along the

Budva–Pindos zone, where Late Eocene–Early Mio-

cene turbidites are found.

2.3. Palaeogene–Early Miocene tectonic events

Systematic palaeomagnetic study of Cenozoic to

Late Cretaceous rocks corroborated the distinction

between the terranes: in a first approach Alcapa is

characterised by Cenozoic counter-clockwise rota-

tions, while Tisza and Dacia are characterized by

Cenozoic clockwise rotations (Fig. 4A) (e.g. Balla,

1987a; Bazhenov et al., 1993; Krs et al., 1982, 1991;

Marton, 1987, 1990; Marton and Marton, 1978, 1989,

1996, 1999; Marton et al., 1992, 1999, 2002; Marton

and Fodor, 1995; Patrascu et al., 1990, 1992, 1994;

Surmont et al., 1990).

These terranes moved as major uniform blocks, but

were not rigid. This is also best shown by palae-

omagnetic data from Upper Cretaceous–Lower Mio-

cene rocks. Detailed studies of some inselbergs

showed that there exist differences in the angle of

rotation between members of the same terrane (e.g.

Transdanubian Range vs. the Bukk Mts. area; Fig.

4A) (Grabowski and Nemcok, 1999; Marton and

Fodor, 1995; Marton and Marton, 1996). This may

be explained by the detachment of some elements on

low angle normal faults or thrusts. On the other hand,

similar total amount of rotation of two parts within the

same terrane has a different timing (e.g. Early Mio-

cene in the Mecsek vs. Early and Middle Miocene in

the Apuseni–Transylvanian basin; Figs. 3A and 4A)

(Csontos et al., 2002; Marton and Marton, 1999;

Panaiotu, 1998). This suggests major deformation

belts across the terranes. Local differences in the

rotation sense (e.g. clockwise and counter-clockwise

rotations within the Mecsek Mts; Fig. 4A) can be

explained by local shear and rotation (Marton and

Marton, 1999; Csontos et al., 2002).

Three major tectonic events happened during the

Palaeogene–Early Miocene. Starting in Late Eocene

(Fodor et al., 1992), but best developed in Oligocene, a

continental escape of the Alcapa terrane from the

Alpine sector took place (Fig. 4B) (Csontos et al.,

1992; Fodor et al., 1992, 1998; Kazmer and Kovacs,

1985). The amount of this escape is estimated to 60–

100 km in the Alps (Schmid et al., 1989) and it is very

likely to have the same original displacement values in

the Carpathian sector. The present-day larger offsets are

the result of subsequent large opposite rotation and

Middle–Late Miocene extension, lateral extrusion

(Balla, 1984; Tari, 1994; Sperner et al., 2002). The

continental escape took place along the Periadriatic–

Balaton line, which is the limit of Alpine vs. Dinaric

elements in the Alcapa terrane (Balla, 1984; Csontos et

al., 1992; Csontos and Nagymarosy, 1998; Fodor et al.,

1998; Wein, 1969). A major right lateral shear took

place in the Vardar zone of the Dinarides during Late

Paleogene (Grubic, 2002; Gerzina and Csontos, 2003).

The amount and exact timing of this shear is not yet

given.

The second major event (possibly partly synchro-

nous with the first one) was a Late Eocene thrusting

and out of sequence nappe stacking in the Dinarides.

This thrusting occurred along the western part of the

Vardar belt, along the Sarajevo sigmoid and along the

Budva–Pindos zone. Thrusting was southwest ver-

gent and created blueschist metamorphism in the

Hellenic Olympus window (Ricou et al., 1998, and

references therein). Therefore, an oceanic subduction

along the Budva–Pindos zone is inferred, although

there is no direct evidence of Pindos oceanic litho-

sphere in the Dinarides. In the Hellenides, smaller

fragments were detected (Bellini, 2002; Champod et

al., 2003). Ricou et al. (1998) thinks that the Vardar

ocean remained open during Mesozoic to close in

Eocene. We rather speculate that Vardar was closed in

mid-Cretaceous and a southern, Budva–Pindos ocean

was closed during the Eocene event.


The third major event in this period was the

opposite rotation of the Alcapa and Tisza–Dacia

terranes (Fig. 4C). Timing of this major rotation is

given by detailed palaeomagnetic studies in these

terranes (Marton et al., 1992; Marton and Fodor,

1995; Marton and Marton, 1999; Panaiotu, 1998), as

Early Miocene (19 Ma). Before that time, until Late

Cretaceous, the two terranes seem to have suffered no

rotation (Fig. 4A). The opposite rotation is thought to

be a consequence of the eastward escape of Alcapa

(Balla, 1984; Csontos et al., 2002), but eventually the

slab subduction roll back in the External Carpathians

might have pulled the terranes northeastward, too

(Csontos, 1995; Ratschbacher et al., 1991).

Deformation due to these tectonic phases is either

concentrated along narrow lateral shear belts, like the

Periadriatic lineament, the Pieniny Klippenbelt (Fodor

et al., 1998; Ratschbacher et al., 1991, 1993a; Fig.

4A), or is consumed in the Flysch nappe overthrusts

of the External Carpathians, or along the Mid-Hun-

garian zone. The Alcapa and Tisza–Dacia terranes

have a common rotation pole (Fig. 4C) documented

by the exposures of Poiana Botizii, Transylvania

(Gyorfi et al., 1999). As a consequence, the Mid-

Hungarian zone along the contact of the terranes

experienced NW–SE shortening deformation (Balla

et al., 1987; Csontos and Nagymarosy, 1998). Syn-

chronous NE–SW elongation should have occurred

because of geometric constraints (Fig. 4B,C). This

elongation is also inferred from the lack of crustal

thickening in the Mid-Hungarian zone, in spite of the

strong across-strike shortening. The along-strike ex-

tension might also be indicated by the presence of

large amounts of Miocene volcanic material (e.g. Tari,

1994; Csontos, 1995).

The reconstruction of the Palaeogene–Early Mio-

cene tectonic phase is best done by rotating back-

wards the two terranes (Alcapa and Tisza–Dacia) by

the amount indicated by palaeomagnetic measure-

ments (Fig. 4C). Although the details of the precise

rotation history are interesting (Csontos et al., 2002),

the simplest way is to rotate back the Late Cretaceous

palaeomagnetic directions to the north. Because of the

uncertainties in the amount of shortening and stretch-

ing, this operation probably contains the least error as

well. This Late Cretaceous–Early Palaeogene posi-

tion of the two terranes, first suggested by Balla

(1984, 1987a), is now accepted with some modifica-

tions (e.g. Csontos, 1995; Csontos et al., 2002; Fodor

et al., 1999; Kovac et al., 1994). Since the Late

Palaeogene right lateral shear along the Mid-Hungar-

ian zone could not have taken place along the present,

curvilinear trends, this motion is to be restored after

the reconstruction of the pre-rotation situation (Fig.

4B) (e.g. Fodor et al., 1998). This reconstruction

brings the Periadriatic line, the Mid-Hungarian belt

(northern margin of Tisza) and the Dinaric Sava–

Vardar belt on one trend. This latter is also considered

to be a major right lateral shear belt (Gerzina and

Csontos, 2003; Mercier, 1968; Ricou et al., 1998).

This structural zone is marked by syn-kinematic

granites all along its length.

Rotating the intra-Carpathian terranes into their

original (pre-Cenozoic) position leaves an open space

within the Carpathian arch, between the terranes and

the European margin. This suggests that there was

consumable oceanic crust in this embayment even in

the Palaeogene (Fig. 4C) (see also Csontos et al.,

1992; Ratschbacher et al., 1991).

3. Mesozoic structures and events

3.1. Mesozoic structural units

The subdivision of individual structural units pre-

sented here (Fig. 6) is generally accepted and is the

result of more than 100 years of Carpathian–Dinaric

geologic knowledge (for a review, see Plasienka,

1999, 2002). The tectonic transport direction, where

known, is indicated according to present coordinates

(Fig. 7). On the other hand, Mesozoic nappe transport

directions compiled and shown in their present posi-

tion should be re-located and reoriented due to the

Cenozoic rotations.

3.1.1. Structural units of Alcapa terrane

Mesozoic nappes are exposed in the Eastern Alps,

Western Carpathians, Bukk Mts. and in the Trans-

danubian Range (Figs. 6 and 8). The structural build-

up is described in two parts, because the Alpine and

Western Carpathian sectors are geographically sepa-

rated by large basins (Fig. 1). In the Western Car-

pathian sector, from Cracow to the Bukk and

Transdanubian Range the structural edifice is similar

to that of the Eastern and Southern Alps (for a review,

Fig. 6. Mesozoic tectonic units of theCarpathian areawithCenozoic formations removed (outcrop and subcrop).Modified fromCsontos et al. (1992) after Arkai (1990), Arkai andBalogh

(1989), Beck-Managetta andMatura (1980), Canovic andKemenci (1988), Dicea et al. (1980), Ebner et al. (1998), Flugel (1988), Fulop andDank (1987), Fusan et al. (1987), Glushko and

Kruglov(1986),GreculaandEgyud(1989),Gnojeketal. (1991),Haasetal. (1988,2000),Hovorka(1985),Mahel’ (1973),Nastaseanu(1975),Pamic (1998a),Pap(1990),Proticetal. (2000),

Sandulescu (1975a,1976,1980b,1988),SandulescuandVisarion (1978), Simunic et al. (1979),Sotaket al. (1993)Tari (1994),Wessely (1988), andownwork. (Forcolor seeonlineversion).

L.Csontos,A.Voros/Palaeogeography,Palaeoclim

atology,Palaeoeco

logy210(2004)1–56

11

Fig. 7. Tectonic transport directions in present coordinates. Compiled from Csontos (1999), Dallmeyer et al. (1996, 1999), Faryad and Henjes-

Kunst (1997), Frank et al. (1987), Fritz et al. (1991), Grill (1989), Hok and Hrasko (1990), Hok et al. (1993, 1994, 1995), Linzer et al. (1995),

Maluski et al. (1993), Marko (1993), Matenco and Schmid (1999), Neubauer et al. (1995), Pana and Erdmer (1994), Pana (1998), Plasienka

(1991), Plasienka et al. (1991), Putis (1991), Ratschbacher and Neubauer (1989), Ratschbacher et al. (1991, 1993a,b), Ring et al. (1989),

Schweigl and Neubauer (1997a), Tari (1994), Tomljenovic (2002), Willingshofer and Neubauer (2002), Koroknay (personal communication,

2001), own works. Numbers indicate the ages of radiometrically dated shear zones.


see e.g. Andrusov et al., 1973; Fuchs, 1984; Kovacs et

al., 2000; Mahel’, 1974; Plasienka, 1999, 2002). The

Transdanubian Range is described within the Alpine

chapter, which is justified by the Late Cretaceous

reconstruction.

3.1.1.1. Structural units of the Eastern Alps. Three

major geodynamic units are exposed in the Alps. The

lowermost Helvetic nappes are related to the European

margin (Figs. 8A and 9). The overlying units represent

the Penninic=Ligurian–Piemontais ocean. The Flysch

nappes detached from their substratum are exposed in

the External Alps, while a metamorphosed sequence: a

thinned continental fragment of the European margin is

exposed in tectonic windows, beneath ophiolites. The

third geodynamic unit above the Penninic/Flysch

nappes is the Austroalpine nappe complex, which is a

set of nappes mostly with a Variscan crystalline base-

ment and a Permian–Mesozoic cover. Lower and

Middle Austroalpine nappes contain abundant crystal-

line basement and a more or less metamorphosed cover

succession. The Upper Austroalpine nappes have a

weakly metamorphosed Variscan basement overlain

by thick, non-metamorphosed, mainly carbonate Me-

sozoic complexes. The lower (Bajuvaric) and the

higher (Tirolic) nappes differ in their Mesozoic facies.

The latter is composed of rocks of continental margin

origin overlain by turbidites with ophiolite clasts of a

Late Jurassic–Early Cretaceous foredeep (Faupl and

Wagreich, 1992; Gawlick et al., 1999; Mandl, 1999).

Fig. 8. Schematic cross sections of the Alcapa terrane. All sections are strongly simplified. (A) After Mandl (1999) and Neubauer et al. (1999),

(B) partly after Plasienka (1998), (C) after own work, (D) after Aubouin et al. (1970), Csontos et al. (2003) modified.


The Tirolic nappe is tectonically overridden by the

Juvavic nappe with mafic fragments in its evaporitic

sole. The ophiolite-derived fragments are thought to

have originated from a fourth, poorly represented geo-

dynamic unit, the Meliata ocean. The Juvavic nappe

might have originated from the other margin of the

Meliata (Schweigl and Neubauer, 1997a,b) or from a

more distal part of the Tirolic margin (Mandl, 1999).

Nappe stacking in the Eastern Alps is thought to begin

in Late Jurassic–Early Cretaceous. These early nappes

probably arrived from the south (Gawlick et al., 1999;

Mandl, 1999; Schweigl and Neubauer, 1997a,b). Later

mid-Cretaceous nappe formation started in the SE and

propagated toward the W–NW (Linzer et al., 1995).

Terminal Cretaceous–Eocene–Early Miocene nappe

transport was directed more to the north.

The Northern and Southern Alps are divided by the

Periadriatic lineament, a shear zone along which the

Transdanubian Range was displaced (Haas et al., 1995;

Kazmer and Kovacs, 1985). The latter is a thick nappe

(Adam et al., 1985; Horvath et al., 1987; Tari, 1994)

with its Variscan weakly metamorphosed basement

(Arkai and Balogh, 1989; Dudko and Lelkes-Felvari,

1992) which overlies Middle Austroalpine nappes

(Figs. 6, 8B and 10) (Tari, 1994). The sedimentary

facies of the Transdanubian Range are very similar to

those of the Southern Alps (Kazmer and Kovacs,

1985). On its NE periphery, the Transdanubian Range

has an Early Cretaceous foreland basin turbiditic suc-

cession with ophiolite clasts (Argyelan, 1996; Csaszar

and Bagoly-Argyelan, 1994; Tari, 1994; Vasko-David,

1991). The source area of the sedimentary infill of this

Fig. 9. Simplified tectono-stratigraphic diagram of the western part of the Alcapa terrane (Eastern Alps). Structural units are arranged in a

palinspastic order. Data taken from: Mandl (1999), Neubauer et al. (1999). Palaeozoic rocks (metamorphic and sedimentary) indicated by simple

boxes. Thick lines indicate nappe contacts. Upper, less inclined portions of the line suggest the emplacement time of this nappe. Lower, less

inclined portions indicate the detachment. Lines with several less inclined portions indicate reactivation and further transport.


foreland basin was located to the present-day north of

the Transdanubian Range (Csaszar and Bagoly-Argye-

lan, 1994; Sztano, 1990). It is assumed that there

was an obducted ophiolitic nappe, Meliata (Balla,

1987b), on the northern, now eroded periphery of the

Transdanubian Range in the latest Jurassic–Early

Cretaceous. Bada et al. (1996) demonstrated N–S

convergence in the Late Jurassic–Early Cretaceous.

According to Tari (1994), the Transdanubian Range

unit was first transported to SW in the Aptian–Albian,

then to NW (or to SE) in the Albian and Turonian onto

lower Alpine units. A Santonian–Campanian cover

postdates most major tectonic transport.

3.1.1.2. Structural units of the Western Carpa-

thians. Six geodynamic units are exposed in the

Western Carpathians (Figs. 6, 8C and 11). The Carpa-

thian foredeep represents the lowest, subducted Euro-

Fig. 10. Simplified tectono-stratigraphic diagram of the southern part of the Alcapa terrane (Transdanubian Range and Bukk areas). Data taken

from Balogh (1981), Csontos (1988, 2000), Galacz et al. (1984), Voros et al. (1990), Voros and Galacz (1998). Same legend as for Fig. 9.


pean margin. This is overridden by the second unit: a

strongly folded sedimentary complex, thought to have

been accumulated in several branches of an ocean,

mainly exposed in the External Carpathian Flysch

nappes, in the Pieniny Klippenbelt and in tectonic

windows in the western part of the chain (Fig. 6)

(Birkenmajer, 1965, 1985; Plasienka, 1987; Plasienka

and Marko, 1993). The northern branch of this ocean

is called Magura and is represented solely by turbi-

dites sheared off their substratum. A central, conti-

nental fragment called Czorsztyn or Oravicum unit

(Birkenmajer, 1985; Plasienka, 1987) may have sep-

arated the northern, Magura from the southern, Pie-

niny–Vahic oceanic branch (Fig. 8C) (Birkenmajer,

1986). In the east, a window covered by thick Neo-

gene sediments exposes weakly metamorphosed Eo-

cene shales and clastic deposits of the Inacovce–

Krichevo unit beneath metamorphic mafic rocks and

a Mesozoic succession (Sotak et al., 1993, 1994).

These metamorphic rocks are interpreted as remnants

of the Pieniny Klippenbelt by Kovac et al. (1994) or

of the Magura.

The bulk of the Alcapa terrane is built by the third

geodynamic unit, the Austroalpine nappe complex

with the following nappes from bottom to top: Tatric,

Fatric/Veporic, Hronic (=Choc) and Gemeric (Figs. 2,

8C and 11) (Plasienka, 1998). The Tatric and Fatric

nappes have mostly Variscan granitic basement with

non-metamorphic Mesozoic cover. The Variscan

gneissic basement and the Mesozoic cover of the

Veporic nappe suffered eo-alpine metamorphism with

Albian–Late Cretaceous formation resp. cooling ages

(Cambel and Kral’, 1989; Dallmeyer et al., 1996;

Maluski et al., 1993; Plasienka, 1991; Putis, 1991).

The Hronic nappe contains an Upper Palaeozoic

volcanic-sedimentary succession and a mostly Trias-

Fig. 11. Simplified tectono-stratigraphic diagram of the northern part of the Alcapa terrane (Western Carpathians). Lithology is figured by main

facies. Data taken from Andrusov et al. (1973), Birkenmajer (1977), Bujnovsky and Polak (1979), Kovacs (1984), Kovacs et al. (1988), Kozur

and Mock (1973, 1985), Lefeld et al. (1985), Less et al. (1988), Mello et al. (1983, 1996), Michalik (1977), Plasienka (1987, 1998), Rakus et al.

(1990), Vozarova and Vozar (1992). Same legend as for Fig. 9.


sic Mesozoic sedimentary cover. A characteristic

Lower Cretaceous turbiditic sequence with ophiolite

material derived from the south is also found here

(Plasienka, 1998).

The Gemeric nappe lies above the Veporic one. It

is composed of Variscan polymetamorphic rocks

which have a weakly metamorphosed Carboniferous

and a non-metamorphic Permian volcanic-sedimenta-

ry cover. The southern margin of the Gemeric nappe

has an Alpine metamorphic overprint. A group of

small nappes in S Slovakia–N Hungary: Borka,

Torna, Szendr}o, represent the metamorphic Mesozoic

cover of the Gemeric (Kovacs et al., 1988; Mello et

al., 1983, 1996; Plasienka et al., 1997). At some

places, they were subjected to blueschist facies, at

some others medium-high pressure, low-temperature

metamorphism (Arkai, 1983; Dallmeyer et al., 1996;

Faryad and Henjes-Kunst, 1997; Ivan and Kronome,

1996; Maluski et al., 1993).

Above the Gemeric and Borka–Torna nappes, the

fourth geodynamic unit, dominated by dark shales and

redeposited material occurs. This very variable se-

quence is called Meliata nappe in Southern Slovakia–

Northern Hungary (Kozur and Mock, 1973, 1985).

Remnants of Triassic mid-oceanic ridge basalts and

serpentinized gabbros (Kozur and Reti, 1986; Faryad

et al., 2002), Jurassic mafic flows and related shallow

intrusives, Jurassic acidic volcanic rocks, Triassic

carbonates and Triassic–Jurassic radiolarites (DeW-

ever, 1984; Dosztaly and Jozsa, 1992; Harangi et al.,

1996; Szakmany et al., 1989; Vozarova and Vozar,

1992) are found within the shaly succession. Most

rocks suffered diagenetic–anchimetamorphic trans-

formation (Arkai, 1983).

Emplacement of this dissected ophiolitic and me-

lange material probably occurred in Late Jurassic–

Early Cretaceous (Balla, 1987b). The syn-kinematic

blueschist metamorphism in the underlying Borka


nappe is dated as Kimmeridgian (150 Ma) (Maluski et

al., 1993), while more marginal units suffered a

Barremian metamorphic overprint (120 Ma) (Arkai

et al., 1995). Shear direction from blueschists indi-

cates northward transport (Fig. 7) (Faryad and Henjes-

Kunst, 1997). Ductile structures in the Szendr}o nappe

are also north vergent (Arkai et al., 1995). Nappe

emplacement in the Austroalpine nappes was sup-

posed to be Turonian in age (Mahel’, 1974), but new

data suggest a longer, Early to latest Cretaceous

stacking period (e.g. Maluski et al., 1993; Dallmeyer

et al., 1996; Plasienka, 1998). Nappe emplacement is

younging towards the NNW, as also indicated by

propagation of foreland basins (Plasienka, 1998,

2002). Measured tectonic transport directions are also

to the N–NNW (Fig. 7) (Hok and Hrasko, 1990; Hok

et al., 1993, 1994, 1995; Plasienka, 1991; Putis, 1991;

Ratschbacher et al., 1993a). These transport directions

refer to the Albian–Maastrichtian period (Dallmeyer

et al., 1996).

The Gemeric and all other, Austroalpine Western

Carpathian nappes are overthrust by the fifth unit, the

uppermost, non-metamorphic Szilice–Straov nappe

(Figs. 6, 8B and 11) (Kovacs, 1984; Kovacs et al.,

1988; Plasienka, 1999). It consists of shallow to deep

marine Triassic margin succession, topped by pelagic

and condensed Jurassic succession. The Permian

evaporitic sole of Szilice nappe comprises tectonically

incorporated Mesozoic ophiolitic fragments. This

implies that during its emplacement, the nappe tore

off some parts of the Meliata nappe and transported

them above other units. Tithonian shallow water lime-

stones present as reworked clasts and assumed to be

the cover of the Szilice nappe might indicate a Late

Jurassic nappe formation (e.g. Plasienka, 1998). Final

emplacement of the Szilice–Straov nappes is Late

Cretaceous in the northern Western Carpathians. In

Hungary, the Szilice unit was transported towards the

south (Pero et al., 2003).

The sixth geodynamic unit of the Alcapa terrane is

exposed in the Bukk Mts., N. Hungary (Figs. 6, 8B

and 10). Here the Bukk parauthochthonous unit is

tectonically overlain by the Meliata nappe (Csontos,

1988, 1999, 2000). The Bukk parauthochthonous unit

has a Dinaric provenance (Balogh, 1964) and consists

of a south-vergent imbricate system. The Palaeozoic–

Mesozoic rocks including Upper Jurassic foreland

deposits are strongly folded and suffered anchizonal

metamorphism (Csontos, 1999). Two distinct episodes

of metamorphism occurred at 120 and 90 Ma (Arkai

et al., 1995; Arva-Sos et al., 1986). The Meliata nappe

above the Bukk parautochthonous comprises a Juras-

sic succession which locally contains large slivers or

olistoliths of Triassic basalts, gabbros and Jurassic

basalts (Dosztaly and Jozsa, 1992). Chemistry of all

mafic rocks indicates an oceanic origin (Harangi et al.,

1996; Faryad et al., 2002).

Meliata nappe emplacement occurred after the

extrusion of mafic flows (160 Ma) (Arkai et al.,

1995; Arva-Sos et al., 1986), but before peak meta-

morphism at 120 Ma, probably in latest Jurassic–

earliest Cretaceous (Balla, 1987b; Csontos, 2000).

Preliminary work on structural transport suggests a

top to the west shear, followed by top to the south

shear and asymmetric folding. The latter deformation

is thought to be synchronous with Early Cretaceous

peak metamorphism. Transport directions suggest that

the Meliata unit was located to the E or NE of the

Bukk parautochthonous unit.

The contacts of the Bukk parautochthonous unit

and the overlying Meliata nappe with other struc-

tural units are not exposed in northern Hungary.

From seismic sections, however, it seems that Bukk

and Meliata as a whole are part of the general

north-vergent Gemeric nappe structure, as an upper

nappe (Tomek, personal communication, 1997). If

this is correct, then this contact is probably a Late

Cretaceous feature.

3.1.2. Structural units of the Dinaric chain

For a general description, the reader is referred to

Aubouin et al. (1970), Dimitrijevic (1982), Herak

(1986), Jacobshagen (1979), Pamic (1982), Tari and

Pamic (1998). The general structural concept of the

Dinaric chain is revised based on recent visits in

some key parts in Croatia, Bosnia, Serbia (Csontos et

al., 2003; Gerzina and Csontos, 2003). The Dinaric

nappes (Figs. 6, 8D and 12) expose four geodynamic

units, bound to the east by the overthrust of the

Serbo–Macedonian (Supragetic) massif and to the

southwest by the Ionian (Hellenic) subduction (Med-

iterranean Ridge). The lowest unit is the Adriatic–

Ionian Platform (parautochthonous) with its Palae-

ogene turbidite cover. It is overridden by the Budva

nappe with a Ladinian mafic basement and a Meso-

zoic–Palaeogene trough slope fill. The third unit is

Fig. 12. Simplified tectono-stratigraphic diagram of the Adria terrane (Dinarides). Structural units are arranged in a palinspastic order. Data

taken from: Aubouin et al. (1970), Blanchet (1970), Canovic and Kemenci (1988) Cousin (1970), Dimitrijevic and Dimitrijevic (1973, 1991),

Obradovic and Gorican (1988), Pamic (1982, 1984, 1998b), Rampnoux (1970), Simunic et al. (1979). Same legend as for Fig. 9.


the High Karst nappe with a thick carbonate platform

of Mesozoic age. In some southern localities, there is

an Anisian turbidite at the base of the Triassic

carbonate platform. This might indicate Palaeotheys

subduction.

The High Karst platform has a NE-facing margin

which is locally marked by an Early Cretaceous

metamorphic event. A now dislocated Early Creta-

ceous turbiditic foredeep basin called Bosnian flysch

is found on the High Karst margin (Aubouin et al.,

1970; Dimitrijevic, 1982). This turbidite contains both

continent- and ocean-derived clasts.

The Drina Ivanjica and the Golija–Pelagonian

metamorphic nappes are also affiliated to the

underthrust High Karst margin and are supposed

to be Palaeogene out of sequence nappes. More-

over, anchi to epimetamorphic Triassic–Jurassic (?)

successions found adjacent or within the Vardar

belt, like Kopaonik, Jadar, Medvednica are also in

the same structural position as the High Karst

nappes (Csontos et al., 2003). These are thought

to be the deepest underthrust elements of this

margin.

The fourth geodynamic unit is the Vardar, which is

now a thrust and laterally sheared belt containing

lenses of different mafic rocks, Early to Late Creta-

ceous–Paleocene turbidites, metamorphic Palaeozoic

and Mesozoic rocks and even Jurassic granites (Dimi-

trijevic, 1982; Pamic, 2002). The complexity is

explained by the strong shearing and lateral displace-

ment of an earlier nappe pile: ophiolites above High

Karst margin and below Serbo–Macedonian crystal-

line. The Vardar nappe proper contains Jurassic shaly

melange and remains of Triassic and Jurassic ophio-


lites. This is interpreted as an ocean containing a

Jurassic accretionary prism and arc–back-arc basin

complex, obducted in Late Jurassic and covered later

by different turbidites (Zachariadou and Dimitriadis,

1995). There are two parallel belts of ophiolites: the

Dinaric Ophiolite Belt and Vardar. In spite of the

differences in chemical composition, the western and

eastern ophiolitic belts of the Dinarides seem to derive

from the same Vardar ocean. Huge olistoliths and

gravity nappes of Mesozoic carbonates are found in

the ophiolitic melange (Dimitrijevic and Dimitrijevic,

1973, 1991).

In general, all the Dinaric nappes are at present

southwest vergent. Recent field visits indicate a four-

step tectonic evolution (Csontos et al., 2003). Thrust-

ing started as early as Late Jurassic (Dimitrijevic,

1982). In the innermost zones, Tithonian shallow

water limestones seal the tectonic contact of ophiolites

and their foreland. It was suggested that Dinaric

Ophiolite Belt and the Vardar oceanic crust obducted

to the east and west, respectively (Robertson and

Karamata, 1994). However, obduction was most prob-

ably oblique or almost parallel to the Dinaric margin,

since strong stretching lineation with top to NW fabric

(i.e. parallel to present structural trends) was found in

many places, including a metamorphic ophiolite sole.

A section in Serbia suggests that Vardar ophiolites

were also obducted on the Serbo–Macedonian Mass

prior to earliest Cretaceous.

The Tithonian obduction was followed by south-

westward propagating thrusting onto the Bosnian

flysch foredeep. This event created syn-cleavage tight

folds and nappes in the Dinarides, as well as a weak

metamorphism in the more southern and eastern units.

An Early Cretaceous Aptian–Albian emplacement

event (possibly a collision) is marked by the age of

low-grade metamorphic overprint of the underthrust

High Karst nappe at 120 Ma (Belak et al., 1995;

Milovanovic, 1984) and Lower-mid-Cretaceous

coarse grained conglomerates containing granite

boulders (Neubauer et al., 2003; Pamic and Tomlje-

novic, 2000). Albian and later deposits cover the

eroded ramp-anticlines above the nappe thrust surfa-

ces (Dimitrijevic, 1982).

Shortening seems to have continued during the

Late Cretaceous in parts of Vardar, in the Bosnian

flysch and in Budva with the onset of sedimentation in

turbiditic basins (Dimitrijevic, 1982; Pamic, 2002).

This event ended by the third major event: a Late

Eocene folding and nappe emplacement, involving

many out of sequence nappes. This main shortening

phase is also top SW. The fourth main structural event

(eventually synchronous with Palaeogene thrusting) is

a pervasive right lateral shear along the Vardar belt

(Grubic, 2002).

The Hellenides has a very similar evolution (e.g.

Ricou et al., 1998). The first obduction of the single

Vardar ocean is of non-precised direction, although

NE and SW directions were both proposed (Jones and

Robertson, 1990; Ricou et al., 1998). The second,

Albian event created high-grade metamorphism of the

lower Rhodope (Drama unit) (Ricou et al., 1998),

which is considered here as the deepest involved

Dinaric High Karst margin. The third, Paleogene

event is either characterised by right-lateral slip along

the Vardar belt (Mercier, 1968), or a SW-vergent

thrusting, which created blueschist metamorphism in

the Olympus window.

3.1.3. Structural units of the Tisza terrane

The Tisza terrane has scattered surface exposures

such as the Papuk–Krndija, Moslovacka Gora (NE

Croatia), the Mecsek, Villany (S. Hungary) and Apu-

seni Mountains (Romania) (Fig. 1). It is now gener-

ally accepted that the Apuseni and Great Hungarian

Plain pre-Palaeogene structures can be correlated

(Berczi-Makk, 1986; Bleahu, 1976; Kovacs, 1982).

Tisza is a composite terrane, consisting of three geo-

dynamic units.

The lowermost Bihor nappe system is built (from

bottom to top) of north-vergent Mecsek, Villany,

Lower Codru nappes (Figs. 6, 13E and 14) (Bleahu

et al., 1996). All these nappes comprise a polymeta-

morphic Variscan basement and a Permian–Mesozoic

cover (Dallmeyer et al., 1999; Lelkes-Felvari et al.,

1996). Early Cretaceous (and possibly Late Jurassic)

mafic rift-related volcanic rocks are present at the

northern margin of the lowermost nappe (Berczi-

Makk, 1986; Harangi et al., 1996). The small Urmat

nappe composed of Toarcian to Lower Cretaceous

turbidites may indicate a deep basin or an ocean at the

southern boundary of the Lower Codru nappe (Bleahu

et al., 1981, 1996). This would be an additional

geodynamic unit.

Nappe emplacement directions are generally to-

wards the NNW (Fig. 7) (Dallmeyer et al., 1999;

Fig. 13. Schematic cross sections of the Tisza–Dacia terrane. All sections are strongly simplified. (E) after own work, (F) after Sandulescu et al.

(1981a), modified, (G) Sandulescu et al. (1981b), modified.


Pana, 1998; Pana and Erdmer, 1994) and constitute a

northward propagating sequence of Barremo–Aptian

(115–120 Ma), Albian (100 Ma), and possibly Turo-

nian thrusting events (Bleahu, 1976). This nappe

emplacement is synchronous with mid-Cretaceous

sedimentation. Later, Turonian–Campanian sedi-

ments seal nappe contacts in the Apuseni (Bleahu et

al., 1981) and under the Great Hungarian Plain

(Szentgyorgyi, 1989).

In vast areas, the Lower Codru nappes are cov-

ered by the second geodynamic unit called Biharia

(Figs. 6, 13E and 14). The Biharia nappes s.str. are

mostly composed of polymetamorphic Variscan base-

ment (Dallmeyer et al., 1999), but some weakly

metamorphic post-Variscan cover rocks (Bleahu et

al., 1981, 1996) and Mesozoic sediments (Balintoni,

personal communication, 1992) are also locally pres-

ent. Nappe transport is towards the NNW (Fig. 7)

(Pana, 1998) and dated as 115–120 Ma. The over-

lying Baia de Aries nappe contains also polymeta-

morphic Variscan rocks, but of different composition

and grade (Balintoni, 1994). Some ENE–WSW

trending lineations have the same or an earlier age

(150 Ma) (Dallmeyer et al., 1999; Pana, 1998). The

southern portions of the nappe are covered by

Tithonian limestone and by Santonian–Maastrichtian

foredeep deposits.

The Lower Codru nappes are partly overlain by the

Upper Codru nappes (Figs. 2, 6, 13E and 14). They

are detached Mesozoic cover nappes composed main-

ly of thick Triassic carbonates and eventually their

Permian basement (Bleahu et al., 1981). The proposed

original basement of the Upper Codru nappes is the

Biharia crystalline basement (Patrulius, 1971), but the

original contact is nowhere observed. The origin of

the Upper Codru nappes is not yet known. Their

emplacement seems to be towards the NNW, but the

emplacement age is not known.

Fig. 14. Simplified tectono-stratigraphic diagram of the Tisza terrane. Data taken from Balogh (1981), Berczi-Makk (1986), Bleahu et al. (1981,

1996), Bordea et al. (1975), Fulop (1966), Ianovici et al. (1976), Lupu (1976), Nagy (1968), Nagy and Nagy (1976), Voros (1972) and the

corresponding sheets of the 1:50.000 map series of the Roumanian Geological and Geophysical Institute. Same legend as for Fig. 9.


The uppermost, third geodynamic unit of the Tisza

terrane is the Mures nappe system (Figs. 6, 13E and

14). These nappes contain large amounts of magmatic

rocks and turbidites (Lupu, 1976). The structural

build-up of the Mures belt appears to be symmetric

towards north and south (Lupu, 1983). The external

Late Cretaceous–Paleocene turbiditic foredeep basin

nappes are overlain by remnants of a Late Jurassic–

Early Cretaceous oceanic island arc with shallow and

deep water carbonate cover (Bortolotti et al., 2002;

Cioflica and Nicolae, 1981; Stefan, 1986), while in

the internal part Early Jurassic sheeted dyke ophiolites

(Cioflica et al., 1981; Savu and Stoian, 1988) are

found. In the Mures belt, a Late Jurassic?–Early

Cretaceous tectonic phase was followed by another

shortening marked by Albian conglomerates and olis-

tostromes, then by a latest Cretaceous nappe forma-

tion. The emplacement of thrust sheets related to the

last, Maastrichtian phase is towards the north in the

Southern Apuseni (Lupu, 1983).

Late Cretaceous to Palaeogene deposits, not only

in the Apuseni sector, but also in the southern part of

the Tisza terrane, are marked and plugged by huge

amounts of calcalkaline volcanic and plutonic material

(Canovic and Kemenci, 1988; Pamic, 1998b; Stefan et

al., 1988). These rocks called Banatites are subduc-

tion-related, but their source is debated. The banatitic

belt continues through the Balkan peninsula in the

Sredno–Gorje belt of Bulgaria (e.g. Balla, 1984).

3.1.4. Structural units of the Dacia terrane

The Dacia terrane can be subdivided into the more

or less distinct Eastern and Southern Carpathians. The

nappe structures are concentric but because of sub-


stantial differences in stratigraphy and structural evo-

lution, they will be described separately.

3.1.4.1. Eastern Carpathians. In the Eastern Carpa-

thian section, the eastwards thrust nappes expose four

geodynamic units. The European margin in lowermost

position is a Cenozoic foredeep (Figs. 2, 6, 13F and

15) (Sandulescu et al., 1981a,b). The Flysch nappes

sheared off this substratum are overridden by a second

geodynamic unit, which is exposed in the Ceahlau

nappes. Mid-Late Jurassic rifting and Tithonian oce-

anic basaltic crust is documented here. A Late Juras-

sic–Early Cretaceous turbiditic sequence indicates

early thrusting/subduction.

Fig. 15. Simplified tectono-stratigraphic diagram of the eastern part of th

(1975b), Sandulescu and Tomescu (1978), Sandulescu et al. (1981a,b) a

Roumanian Geological and Geophysical Institute. Same legend as for Fig

The third geodynamic unit, the Bucovinian nappe

system (Figs. 2, 5, 13F and 15), can be subdivided

from bottom to top into Infrabucovinian, Subbucovi-

nian and Bucovinian nappes (Sandulescu, 1975b;

Sandulescu et al., 1981a,b). The upper nappes cover

the lower ones almost completely and the whole

nappe pile is deformed into large folds. The eastern-

most synform is exposed in the Eastern Carpathians.

The other folds are mostly hidden by Cenozoic

deposits of the Transylvanian basin. All the Infrabu-

covinian–Bucovinian nappes are characterised by

Variscan crystalline basement and a relatively thin

Mesozoic succession. The fragmentary record of Me-

sozoic sediments in these nappes may be partly due to

e Dacia terrane (Eastern Carpathians). Data taken from Sandulescu

nd on the corresponding sheets of the 1:50.000 map series of the

. 9.


the strong tectonic elongation and chain-parallel shear

along the Eastern Carpathian range (Pana and Erdmer,

1994). Upper Albian and younger deposits form the

post-tectonic cover.

The fourth unit, the Transylvanides (Figs. 6, 13F

and 15) are sedimentary debris and gravity nappes

in and on top of Lower Cretaceous olistostrome of

the Bucovinian nappe (Sandulescu et al., 1981b).

The Transylvanide successions can be grouped into

three geodynamic units: a continental margin series

(Persani), an ophiolitic series (Olt), and an oceanic

island arc series (Haghimas). It is debated, whether

the Persani series represents a marginal realm of

the Bucovinian continent, or whether it forms a

different continental fragment. Faunal data (Voros,

1993, 2001) suggest that it belongs to the Medi-

terranean faunal province. In this case, the Olt

series (Russo-Sandulescu et al., 1981) may repre-

sent an oceanic realm between the Bucovinian and

Persani margins.

In the Eastern Carpathians, shortening began in the

Late Jurassic?–Early Cretaceous and continued until

the Albian. Nappe emplacement apparently propagat-

ed towards the east. Continent-derived pebbles in the

lower, eastern nappes suggest a collision with a

hypothetical, ‘‘Coumanian’’ cordillera during the

Albian (Sandulescu et al., 1981b). Recently, along-

strike elongation or transpression at 120 Ma or be-

tween 115 and 80 Ma was suggested (Fig. 7) (Pana

and Erdmer, 1994; Dallmeyer et al., 1996). There is a

renewed phase of shortening in the Late Cenozoic

(Sandulescu et al., 1981a,b).

3.1.4.2. Southern Carpathians. In the Southern Car-

pathians, three geodynamic units are exposed (Figs. 6,

13G and 16). All upper nappes cover the lower ones

almost completely and the lower units outcrop in large

windows. The Late Cenozoic deposits of the Europe-

an foreland are overthrust by crystalline-Mesozoic

nappes called Danubian (Nastaseanu, 1975; Nasta-

seanu et al., 1981). This is a sheared-off fragment of

the European–Moesian platform. Some parts of the

Danubian suffered low-grade Alpine metamorphism

(Berza et al., 1988a,b).

The overlying Severin unit is represented by two

nappes. One nappe contains voluminous Middle Ju-

rassic rift-related basalts (Iancu, 1986). The second

nappe is reconstructed from clasts of a Danubian Upper

Cretaceous olistostrome (Cioflica et al., 1981; Savu,

1985). The reconstructed lithologic composition of

Severin (including Tithonian mid-ocean ridge basalts

and Lower Cretaceous turbidite) is identical to that of

the Ceahlau nappe of the Eastern Carpathians.

The Severin nappe is overlain by the third geo-

dynamic unit, the Getic nappe system (Nastaseanu,

1975; Sandulescu, 1975a). Crystalline basement and

the rarely preserved Triassic carbonates of the Getic

and Supragetic nappes are covered by Jurassic and

Cretaceous deposits. There is a well-documented Early

Jurassic rifting event with intra-plate basalts in all

Getic nappes.

The earliest shortening is Early Albian in age,

followed by major shortening episodes of Turonian

and Maastrichtian age and other, Cenozoic deforma-

tions (Nastaseanu, 1975; Sandulescu, 1975a). All

shortening episodes are synchronous with turbiditic

sedimentation. Tectonic transport directions oblique to

the chain point to the ESE (Fig. 7) (Ratschbacher et

al., 1993b), and are partly due to later exhumation

(Matenco and Schmid, 1999). Mid-Late Cretaceous

ages were measured for these tectonothermal over-

prints (Dallmeyer et al., 1996; Ratschbacher et al.,

1993b).

Thick Late Cretaceous–Palaeogene calc-alkaline

magmatic rocks (banatites) plug all northern units

(Nastaseanu, 1975; Berza et al., 1998). This is attrib-

uted to subduction of unknown oceanic crust.

3.2. Mesozoic tectonic problems in the Carpathians

There are three major tectonic problems in the

Carpathians (Fig. 7): (1) the arcuate shape of the

Western Carpathian structural units; (2) the non-con-

formable tectonic transport directions in the Alcapa

and in the Tisza–Dacia terranes, e.g. the southern

versus northern structural vergencies in adjacent Bukk

and Szendr}o units; or the arcuate shape and centripetalthrust directions of the Tisza–Eastern Carpathian–

Southern Carpathian Dacide units; and (3) the prov-

enance of large amounts of siliciclastic material in

Late Jurassic–Early Cretaceous turbidites. In the

following, these will be discussed.

3.2.1. Late Cretaceous arching and related structures

The Late Cretaceous tectonic phase is usually

considered as the terminal nappe emplacement event

Fig. 16. Simplified tectono-stratigraphic diagram of the southern part of the Dacia terrane (Southern Carpathians). Data taken from Berza et al.

(1988a,b), Nastaseanu (1975), Nastaseanu et al. (1981), Nastaseanu and Maksimovic (1983), Sandulescu (1976, 1988), Savu (1985) and the

corresponding sheets of the 1:50.000 map series of the Roumanian Geological and Geophysical Institute. Same legend as for Fig. 9.


in the Inner Carpathians. However, a Late Creta-

ceous ‘‘Gosau’’ extension with basin formation,

synchronous with lateral shears and perpendicular

shortening can be separated from the general nappe

stacking events (Willingshofer et al., 1999). This

Gosau event may also be recognised in the Alps

(Froitzheim et al., 1997; Neubauer et al., 1995;

Wagreich and Faupl, 1994), and in all other major

terranes. In Alcapa, the biggest Gosau basin is

exposed in the Transdanubian Range. A similar but

smaller basin is found near Kainach, Austria (Neu-

bauer et al., 1995). Both are controlled by major

normal faults. A set of smaller basins is found in the

Northern Calcareous Alps (Wagreich and Faupl,

1994) and in the Inner Western Carpathians (Pla-

sienka, 1998). In the northern part, a bigger Puchov

basin hosts pelagic sediments. The Tisza–Dacia

terrane has a deep-sea trough succession at its

northern rim: the Szolnok Flysch basin. In the Great

Hungarian Plain area and in the central Transylva-

nian Basin, a bigger Gosau basin is found (Szent-

gyorgyi, 1989; Ciulavu et al., 1994). A series of

main deep-sea troughs is located in the Vardar–

Mures zone. In the Adria terrane, Late Cretaceous

forms generally broad basins above earlier nappes.

These basins host mostly marly to calcareous sedi-


ments. In the Sava–western Vardar zone, a turbidite

trough is found (Pamic, 2002).

The Western Carpathians are dominated by an

arcuate structure. This is manifested by the arrange-

ment of the nappe units bound by arched major

tectonic surfaces (Fig. 7). These surfaces are arranged

in an onion-shell-like concentric pattern, the centre of

which seems to be near the Bukk Mountains, North-

ern Hungary. All these surfaces are composed of NE–

SW striking portions with left-lateral, NW–SE strik-

ing portions with right-lateral, and E–W striking

portions with thrust offsets (Csontos, 1999; Grecula

et al., 1990; Hok et al., 1995; Putis, 1991). Major

shears also bend previous structures like nappe bound-

aries and lineations (Balogh, 1964; Balla, 1984;

Csontos, 1988; Grecula et al., 1990).

Arching was considered to be Cenozoic (Balla,

1984), mainly because the Cenozoic orogen of the

External Carpathians is of similar shape. There are

several arguments, though, that this structural fea-

ture is of Late Cretaceous age. Big offsets in

crystalline basement and Mesozoic strata are found

across one of the curved tectonic surfaces, Muran,

but overlying Palaeogene strata are not displaced

(Marko, 1993). Ar/Ar and K/Ar data on sheared

metamorphic rocks adjacent to the same surfaces

indicate an 88–90 Ma tectonic event (Arkai et al.,

1995; Dallmeyer et al., 1996; Maluski et al., 1993).

Palaeogene–Miocene strata covering differently bent

portions of the Bukk Mountains have the same

palaeomagnetic rotation (Marton and Fodor, 1995;

Marton and Marton, 1996), implying that ductile

shear and related arching must be pre-Palaeogene

(Csontos, 1999).

Smaller Western Carpathian basins may have

opened synchronously and adjacent to the mentioned

shear zones (Brezsnyanszky and Haas, 1984). The

opening directions of these basins are not known

because of later overprint, but ductile extension in

their basement is oriented E–W (Hok et al., 1993).

In the Alps, near Graz, a well-documented E–W

left-lateral shear was dated as 88 Ma old (Neubauer et

al., 1995). This ductile shear was synchronous with

the uplift of an adjacent crystalline dome and the

opening of a basin. Ductile extension directions are

towards the ENE. The major Late Cretaceous basin of

the Transdanubian Range has a synchronous opening

and is supposed to be part of the same wrench-normal

fault system (Tari, 1994). Fodor et al. (2002) suggest a

Late Cretaceous ENE-oriented extension of the Trans-

danubian Range basin along flat normal faults, reac-

tivated in similar extension direction in the Middle

Miocene.

When viewed in the reconstructed Late Creta-

ceous position, the arching and extension forms a

logical system throughout the Alps–Western Carpa-

thians (Fig. 17). Ongoing convergence creates east-

vergent shortening while the same compression gen-

erates conjugate strike-slip shear zones. All the

measured extension directions suggest along-chain

extension. This situation resembles much the model

of Neubauer and Genser (1990) proposed to explain

the Cenozoic structures (N–S shortening, E–W

extension; conjugate strike-slip faults) for the Eastern

Alps.

It seems certain that the wings of the onion-shell

structure did rotate to some extent, but in some cases

this rotation is formed by the drag effect along semi-

ductile or ductile strike-slip shear zones like in the

Bukk Mountains (Csontos, 1999). The concentric

arching could be explained by some southern indent-

er, but no such body is known so far. Another

possible explanation would be to suggest that the

two strike-slip branches form distributed transfer

fault zones, which accommodate thrusts of different

direction. We speculate that at least the present

eastern (in reconstructed directions southern) branch

could act as a wide transfer zone, linking the north

(east)-vergent nappes of the Western Carpathians to

the southwest (west)-vergent nappes of the Dinaric

chain (Fig. 17).

3.2.2. Discrepant tectonic transport directions

At a first glance, it seems that tectonic transport

directions in the Alps and Western Carpathians are in

prefect harmony. If the Tertiary palaeomagnetic results

are taken into account, however, the transport direc-

tions do diverge (Fig. 18). Naturally, it is not meant

that tectonic transport directions should be absolutely

parallel, but divergence should be explained. Even if

Late Cretaceous deformation is taken into account, the

shear directions of similar age do diverge. This can

either mean that the Eastern Alps and the Western

Carpathians were submitted to different stresses dur-

ing Cretaceous deformation, or that the two parts were

not rigidly coupled: rotation was possible between

Fig. 17. Late Cretaceous structural elements of the Carpathian area. Thick lines indicate arched shear zones in the Western Carpathians.


them. In the plate tectonic reconstruction, we use both

hypotheses.

The Late Jurassic–Early Cretaceous tectonic

transport directions show an even wider scatter. In

the southern termination of the Alcapa terrane

(Figs. 6 and 8C), similar structural units occur, in

similar general order, but with opposite vergence.

Late Jurassic–Early Cretaceous tectonic transport

directions reconstructed for Late Cretaceous are

markedly different in the same area (Fig. 18). At

the Bukk Mountains–Szendr}o Mountains interface,

they are 180j apart and the mid-Cretaceous shear

directions also differ by the same amount. Since the

lithologic content, deformation history are very

similar, it is thought that the angular difference is

due to major rotation. This symmetric situation

suggests a late folding of a more linear, uniform

margin (Fig. 19A). This rotation can be indirectly

supported by palaeomagnetic data. A set of palae-

omagnetic measurements made on pre-Late Creta-

ceous rocks indicates that parts of Alcapa had a

complicated rotational history (Fig. 20) (Grabowski

and Nemcok, 1999; Haubold et al., 1999; Marton,

1993a, 1998, 2000; Marton and Marton, 1978;

Mauritsch and Frisch, 1980; Mauritsch and Marton,

1995). As seen later, the Western Carpathians, the

Fig. 18. Late Jurassic–mid-Cretaceous tectonic transport directions in reconstructed Late Cretaceous position. Transport data from Fig. 7,

palaeomagnetic declination data from Fig. 4A.


Northern Calcareous Alps and the Transdanubian

Range should have been located on the same shelf

(Fig. 19A) (Kovacs, 1982; Haas, 1987). Palaeomag-

netic data suggest that the Northern Calcareous

Alps and the Transdanubian Range were at different

respective positions during Mesozoic (making

angles from acute to 180j), to become roughly

parallel by Late Cretaceous. Oroclinal bending of

the same shelf can eventually explain the symmet-

rical structural positions and vergencies of the

nappe units at the southern termination of the

Western Carpathians (Fig. 20). Mesozoic data for

the Western Carpathians (Grabowski and Nemcok,

1999) are not clear and numerous enough to draw

major conclusions.

Similar to the Alcapa case, the Tisza–Dacia

terrane also shows widely diverging tectonic trans-

port directions (Fig. 18). The structural situation is

symmetrical with respect to the Mures–Vardar zone.

These are the most internal and highest nappes in

both Tisza and Dacia (Fig. 2). The ophiolitic

material apparently separates Tisza and Dacia, but

Fig. 19. Schematic position of facies belts in Alcapa (A) and Tisza–Dacia (B) terranes, respectively, for the present situation (top) and simplified

reconstructed early Mesozoic situation (bottom). Further explanations in the text.


it is no more found in the basement of Northern

Transylvania (Sandulescu and Visarion, 1978):

Bihor and Bucovinian nappes are in direct contact

(Fig. 6). Beneath the ophiolitic nappe of common

origin, all the lower nappes with similar stratigraphy

and facies polarity are thrust centripetally towards

the external parts, in a present-day radial pattern

(Fig. 19B). As in the former case, either different

stresses, or major terrane-rotation are supposed to

explain the situation. Mesozoic palaeomagnetic data

suggest that at least the Tisza terrane underwent an

important rotation prior to the equally important

(and contrary) rotation in Cenozoic (Figs. 4A and

20). We consider the very rapid back and forth

rotations suggested by the diagram in Fig. 20 for

the earliest Cretaceous unrealistic and to be

explained by some local factor as the data come

from redeposited sediments (Marton, 2000). Unfor-

tunately, no pre-Late Cretaceous palaeomagnetic

data exists for Dacia, because of a strong Late

Cretaceous remagnetisation. Still, combining Late

Cretaceous counterclockwise rotation of Tisza with

radial structural vergencies, a major oroclinal bend-

ing of a formerly more linear ribbon-continent can

be proposed. The shape of the Vardar–Mures belt

was interpreted as a triple junction, or a side-branch

of a main oceanic trend (Sandulescu and Visarion,

1978), but we suggest that the present form is

rather due to late bending around a subvertical axis

(Fig. 19B).

When all the Cretaceous and post-Cretaceous

bends are restored (Csontos et al., 2003; Tomlje-

novic, 2002), the Early Cretaceous shear directions

are all aligned and parallel to the Dinaric main

structural strike. Even some early shear directions

at the southern part of Tisza show the same direc-

tions. It is therefore proposed, that this first major

structural event was characterized by a transpressive

Fig. 20. Palaeomagnetic data for Mesozoic rocks of the Northern Calcareous Alps (NCA), Transdanubian Range (TR), Tisza (TI) and Western

Carpathians (WCA). Accepted rotation path marked by thin lines. The small number of data points for WCA yet inhibits to define a rotation

path. Double data set at TR Triassic interval indicates two, slightly different parts of that chain. Double data set at NCA Triassic interval and

WCA Cretaceous interval indicates different structural zones and ambiguous data. Numbers at data points indicate averaged inclination data.

Data from Grabowski and Nemcok (1999), Haubold et al. (1999), Mauritsch and Frisch (1980), Mauritsch and Marton (1995), Marton (1993a,

1998, 2000). Little boxes at right represent schematic positions of the respective terrane elements. ?=no or doubtful data. (For color see online

version).


shear along the Dinaric High Karst margin. Since

ophiolites and at least Tisza (but probably Serbo–

Macedonian, i.e. Dacia) are apparently involved in

this major left-lateral shear, we speculate that the

first, Early Cretaceous collision, or docking was the

result of lateral shear, rather than head-on collision.

The shear directions apparently changed during Early

Cretaceous to be perpendicular to the Dinaric mar-

gin. This might have occurred either in the Aptian,

or Early Albian. This could have been a more head-

on collisional stage of the Austroalpine–Dinaric

margin on one side, and of the Tisza–Dacia on the

other. The two oroclinal bends described formerly

were possibly formed later, during a complex Creta-

ceous tectonic evolution.

3.2.3. Provenance of siliciclastic material in Late

Jurassic–Early Cretaceous turbidites

The Late Jurassic–Early Cretaceous turbiditic

foreland basins in the Austroalpine–Dinaric High


Karst margin (Hronic, Tirolic, Transdanubian Range,

Bukk, Bosnian flysch) bring up a major problem

(Figs. 9–12). The clastic material is derived from

an ophiolite sequence and a siliciclastic source

(Argyelan, 1995; Csontos et al., 1991; Dimitrijevic,

1982; Faupl and Wagreich, 1992). The ophiolitic

source is not a problem, since we know of ophiolites

obducted by Late Jurassic onto the Dinaric margin.

On the other hand, siliciclastic material has to be

explained, even if oceanic island arc volcanic rocks

can be an effective source. Detrital material such as

muscovite–chlorite schists (Csontos et al., 1991),

metamorphic rocks (Argyelan, 1995) cannot come

from these island arcs, neither huge Early Mesozoic

granitic boulders in latest Jurassic–Early Cretaceous

coarse conglomerate in Bosnia (Neubauer et al.,

2003; Pamic and Tomljenovic, 2000). Siliciclastics

cannot come from a local, Dinaric source for two

reasons: with one small exception the whole shelf

was covered by a thick platform-to margin carbonate

succession; Palaeozoic massifs now exposed in the

Dinarides were also possibly covered, but anyway,

with minor exceptions they were also metamorphosed

during the Early Cretaceous tectonic event, and not

during the Variscan event (Pamic and Tomljenovic,

2000). Granites of the given age are lacking in the

whole Dinarides.

If the source area of siliciclastic material is not

found on the lower plate, it should be located on the

upper plate. We speculate that this upper plate was

Tisza–Dacia. This microcontinent has thin Mesozoic

cover above the widely exposed Variscan basement. It

also shows major unconformities cutting down to

basement during Mesozoic (e.g. Barremian directly

upon crystalline in the Southern Carpathians; Titho-

nian above crystalline in the Southern Apuseni Mts.)

(Bleahu et al., 1981; Nastaseanu et al., 1981). There-

fore, it is logical to think that the denudated south-

western margin of Tisza–Dacia could provide the

needed siliciclastic material. If this is true, then the

Tisza–Dacia and Alcapa–High Karst margins should

have been in a more or less close contact from the

Late Jurassic on. This means that by the Tithonian

there should have been at least some kind of docking

between the two continental margins. In other words,

the intervening Vardar ocean should have been

obducted, and almost entirely subducted by the latest

Jurassic–earliest Cretaceous.

4. Correlation

4.1. Correlation of oceans: how many of them?

Since oceanic remnants are important for the geo-

dynamic reconstruction, the similarities and prove-

nances of different oceanic fragments are discussed.

The backbone of correlation is also given by ophio-

litic units (Fig. 21). In the following, we try to

minimize the number of oceans (as also suggested

by the review of Prof. Stampfli).

In the Alcapa terrane, the Southern Penninic and

Vahic/Pieniny units are considered equivalents by

Fuchs (1984), and Plasienka (1999), although no

ophiolitic rocks are preserved in the Western Carpa-

thians. It is debated, whether Penninic–Vah and

Magura were two or one ocean(s). In our opinion,

this question loses importance, because the two might

have been separated by one or several minor conti-

nental fragments, like the Czorsztyn ridge of the

Pieniny Klippenbelt, (Birkenmajer, 1998), but where

there was no such a continental fragment, they formed

one ocean (Fig. 19).

The mafic fragments and related metamorphic

rocks found at the southern margin of the Austroalpine

nappes are correlated to the mafic rocks of the Meliata

unit, found in similar position in the Western Carpa-

thians (Hallstatt–Meliata; Kazmer and Kovacs, 1989;

Mandl, 1999; Schweigl and Neubauer, 1997a,b). This

Meliata ocean can be correlated to the Dinaric Vardar

ocean.

It may be proposed that Meliata and Penninic–Vah

ophiolitic and related units are in fact the remnants of

the same ocean and they acquired their structural

position due to out of sequence thrusting or some

other tectonic process. If the whole Western Carpa-

thians are taken into consideration, such a process can

be excluded, since all available evidence suggests an

early (Late Jurassic–Early Cretaceous) consumption

of Meliata in the southern portions of the chain, while

at the same time Penninic–Vah was still open and

received sediments (Argyelan, 1996; Balla, 1987b).

Moreover, a large thrust of Meliata towards the north,

followed by some in sequence or out of sequence

thrust can be excluded because of active Western

Carpathian sedimentary basins between the two

ophiolitic units, where such a tectonic event with

related debris is not detected (Plasienka, 1998). An-

Fig. 21. Remnants of Mesozoic oceanic troughs (ophiolite belts, suture zones) and microcontinents in the central Mediterranean area.


other argument is that Meliata started to form in

Triassic, while Penninic opened only during Jurassic.

The Meliata and Penninic–Vah should thus be two

different oceans in the southern and northern periph-

ery of the Western Carpathians.

The Dinaric–Hellenic sector comprises three

oceans: the present-day Ionian; the Budva–Pindos

and the Vardar (Fig. 21). The Ionian might be a

remnant of the Palaeotethys ocean; this solution

reduces the number of needed oceans. It is, however,

equally possible that these were two different oceans.

This question falls out of the main scope of the paper,

therefore it will not be further discussed.

As discussed above, we accept the general view

(Papanikolaou, 1985) that most ophiolitic material of

the Dinarides–Hellenides was derived from the Var-

dar ocean. This also includes the so-called Pindos

ophiolites, which are in fact nappe outliers above the

Pindos succession (e.g. Ricou et al., 1998). However,

there are indications of basaltic crust beneath the

Pindos, and the Budva successions, which are gener-

ally correlated (Bellini, 2002; Champod et al., 2003;

Dimitrijevic, 1982). Because of blueschist metamor-

phism of the Olympus window, we think that Pindos

was a well-developed ocean. This metamorphism

cannot be caused by the Vardar, since it overlies first

the Pelagonian microcontinent, which is on its turn

above the Olympus blueschists.

In the Tisza and Dacia terranes, the innermost units

are all formed by the topmost Mures and Olt nappes.


These fragments are continued in the Dinaric Vardar

zone as demonstrated by geophysical and borehole

data (Figs. 6 and 21) (Canovic and Kemenci, 1988;

Lupu, 1976; Sandulescu and Visarion, 1978).

The external Ceahlau and Severin (Troyan in

Bulgaria) oceans are also well correlated (Grubic,

1983; Nastaseanu and Maksimovic, 1983; Sandu-

lescu, 1976, 1980a) because of their similar structural

position, time of opening/closure and stratigraphic

content. This oceanic branch may be continued north

of Tisza, where a Late Jurassic–Early Cretaceous rift

was postulated (Harangi et al., 1996). This hypothetic

ocean certainly existed, because it separated fauna of

the Bihor microcontinent from the European mainland

from Late Jurassic onwards (Voros, 1993, 2001). The

continuation of this ocean is proposed in the Magura

ocean of the Western Carpathians. The main argument

for this is the presence of Early Cretaceous rift-related

volcanic rocks in an External Western Carpathian

Silesian succession very similar to that of the north-

ernmost Bihor nappes.

The innermost Vardar–Mures related ophiolitic

units cannot be derived from the same ocean as the

outer Ceahlau–Severin ophiolites and their different

positions cannot be explained by complicated out-of-

sequence tectonics. The sedimentary record in the

Bucovinian and Getic nappes as well as the different

times of opening/closure exclude this possibility (e.g.

Sandulescu et al., 1981a,b). Moreover, sedimentary

transport directions and provenance studies indicate

that the Ceahlau trough was adjacent to the Infrabu-

covinian nappe. Albian conglomerates seal the nappe

contacts and the tectonic situation did not drastically

change from the Albian on.

4.2. Correlation of continental units between the

external and internal oceans

The most diagnostic stratigraphic differences can

be found in Triassic and Lower-Middle Jurassic rocks.

According to the occurrence or absence of Upper

Triassic continental to shallow water, variegated red-

beds, Upper Triassic neritic limestones, Middle-Upper

Triassic pelagic limestones, Lower-Middle Jurassic

ammonitico rosso-type limestones, and Lower Juras-

sic coal-bearing succession, two different facies

domains may be separated (Figs. 8–16 and 19)

(Kovacs, 1982). The first one, characterized mainly

by calcareous sediments, is restricted to the southern,

inner parts of each nappe pile, while the second,

characterized by the abundance of siliciclastic depos-

its and coal, is located in lower, external positions,

like the Helvetic (Gresten), Mecsek, Infrabucovinian

and Getic nappes. Sedimentary transport directions in

these units point to an external, i.e. northern, eastern

(European) continental provenance of the clastic ma-

terial (Nagy, 1968, 1969; Sandulescu et al., 1981a,b).

The distribution of Middle Triassic calcalkaline

volcanic rocks is also remarkable. These occur in

the Southern Alps, the Bukk, a part of the Trans-

danubian Range and a large part of the Dinarides,

while they are thin or lacking in the Austroalpine and

Bihor–Getic. On the other hand, the Late Cretaceous

calcalkaline Banatites are found in the southern parts

of the Tisza and Dacia terranes and the related Serbo–

Macedonian massif (Fig. 6).

4.2.1. Correlation of the Alcapa terrane units

A number of studies compared the units of the

Eastern Alps and the Western Carpathians. Fuchs

(1984), Kovacs et al. (2000), Plasienka (1999), Voros

(2000) and Wessely (1988) successfully correlated the

Lower Austroalpine with the Tatric, the Middle Aus-

troalpine with the Fatric and the Upper Austroalpine

with the Hronic nappe systems (Fig. 19A). Based on

the occurrence of Lower Cretaceous turbidites with

ophiolite-derived clasts in both Tirolic and Hronic

nappes (Figs. 9–12 and 19), these are correlated.

It is difficult to correlate and place the Trans-

danubian Range and Bukk parautochthonous in the

Alpine–West Carpathian edifice. Transdanubian

Range facies zones were correlated with those of the

Southern Alps (Galacz et al., 1984; Haas and Budai,

1995; Haas et al., 1995, 2000; Kazmer and Kovacs,

1985; Majoros, 1980). The equivalents of the Bukk

parautochthonous have not been found in the Alps.

On the other hand, very similar facies and structural

settings exist in areas near Zagreb, Croatia (Medvedn-

ica Mts.) and in the Jadar Mts., Northern Serbia

(Balogh, 1964; Balla, 1987b; Csontos, 1988; Haas

et al., 2000; Pamic and Tomljenovic, 1998; Protic et

al., 2000). Based on the presence of Upper Jurassic–

Lower Cretaceous foredeep sediments commonly

with ophiolite-derived clasts (e.g. Csaszar and Bag-

oly-Argyelan, 1994; Faupl and Wagreich, 1992) and

their structural position, both the Transdanubian

Fig. 22. Proposed position of units in Late Permian and Carnian times. Contours and main latitudes after Stampfli et al. (1998b). Continent arrangement and nomenclature differ from

their construction. Partly inspired by Ziegler and Stampfli (2001). Thin curves indicate present geographic contours in stable Europe and Africa, eventually the contours of the

Adriatic sea are marked. Arrow at the Tunis promontory indicates movement of Africa relative to Europe since the previous stage. Europe is kept fixed for convenience.


atology,Palaeoeco

logy210(2004)1–56

33


Range and the Bukk parautochthonous occupy a

similar position as the Tirolic nappe of the Eastern

Alps and the Hronic nappe of the Western Carpathians

(Figs. 8 and 19A). As shown above, all these nappes

are immediately below or in the direct foreland of

Meliata oceanic slivers. The only difference between

Tirolic, Hron, Transdanubian Range on one hand and

Bukk, Medvednica, Jadar on the other is that the latter

suffered Early Cretaceous metamorphism. The differ-

ence might be caused by the ophiolites overriding the

metamorphic sequences.

Reconstruction of Late Cretaceous positions brings

the Transdanubian Range and Bukk adjacent to the

Dinarides, where a direct continuation is supposed.

There, the correlated units (Medvednica, Jadar Mts) lie

beneath the ophiolitic melange of the Vardar–Dinaric

Ophiolite belt. Therefore, we tentatively correlate the

Meliata unit with the Vardar ocean and the Bukk,

Southern Alps, Transdanubian Range margin with

the High Karst margin beneath the obducted nappes.

4.2.2. Correlation of the Tisza–Dacia terrane units

The Eastern and Southern Carpathian elements of

the Dacia terrane were correlated by Sandulescu

(1976, 1988), although the stratigraphic columns are

not very similar. In both sectors, there are three

common elements: an outer oceanic trough (Ceah-

lau–Severin), an inner (Bucovinian and Getic) conti-

nent and an innermost Vardar ocean (Figs. 13–16 and

19B). The same situation can be seen in the Tisza

terrane. The external Bihor (Mecsek), Bucovinian and

Getic nappes have similar facies, like the Early

Jurassic coal-bearing beds. Furthermore, some other

events, such as Middle Jurassic transgression, Late

Jurassic pelagic sedimentation, Early Cretaceous car-

bonate platform development are also common ele-

ments in the Tisza and at least part of Dacia.

Furthermore, faunal assemblages in critical time peri-

ods agree very well (Voros, 1993). Therefore, we

believe that the Bihor, Bucovinian and Getic units

formed a coherent microcontinent on the northern

margin of the Vardar ocean in the Middle and Late

Jurassic (Fig. 19B).

4.2.3. Correlation of exotic units

Exotic units, such as the Szilice, Juvavic, Upper

Codru and Persani units all lie in a detached, but

apparently uppermost position in the nappe pile (Figs.

6, 8 and 13). Moreover, Szilice and Juvavic nappes

have ophiolite fragments in their evaporitic sole

thrust, indicating that they once overrode an ophiolite,

in both cases Meliata (Csontos, 1988; Gawlick et al.,

1999; Kovacs et al., 1988; Mandl, 1999; Schweigl and

Neubauer, 1997a,b). Szilice has been considered to

represent the opposite shore of the Bukk margin,

because of fragments of Triassic carbonate platform

in the olistostrome and opposite polarities of the

margins (Kovacs, 1984; Csontos, 1988, 2000). The

Persani unit is embedded in Early Cretaceous olistos-

trome with ophiolite debris (Sandulescu et al., 1981b).

Based on the stratigraphy of these units similar to that

of the olistostrome in the Dinaric ophiolite belt (Di-

mitrijevic and Dimitrijevic, 1991), it is suggested that

the Szilice–Juvavic–Upper Codru–Persani units all

correlate with an enigmatic microcontinent that lay to

the NE of the Dinaric margin, across the Meliata–

Vardar ocean. It was formerly proposed (Dimitrijevic

and Dimitrijevic, 1991; Robertson and Karamata,

1994) that the sediments, olistoliths in the ophiolite

melange were accreted from below, but this situation

is unlikely because most of the sedimentary material

and the most complete series are found above the

melange or serpentinites. One of the possibilities is to

propose a Tisza–Getic–Serbo–Macedonian origin to

these rocks, which were on the northeastern margin of

the Vardar ocean (Fig. 22). This possibility is sup-

ported by the fact that denudated basement is present

in the southern part of Tisza: the Biharia–Baia de

Aries crystalline is overlain by Tithonian reefs. It is

therefore proposed that in Jurassic time the Szilice–

Juvavic–Upper Codru–Persani units were detached

from their original Tisza–Getic–Serbo–Macedonian

crystalline basement to glide into the Vardar melange

and then eventually towards their more internal (e.g.

Bucovinian) troughs.

5. Timing of main plate tectonic events

Major plate tectonic events like rifting, opening of

an ocean and collision are best identified by the

stratigraphic content and facies of different nappes.

These data are supported by palaeobiographic, sedi-

mentologic or magmatic–petrologic ones, when avail-

able and needed. In our model, we envisage five

oceanic troughs in the western termination of the


Palaeo/Neotethys. These existed in different time

periods and separated Europe and Africa and different

microcontinents, respectively. After their closure, the

remnants of these oceanic troughs can be recognised

along suture zones and in tectonic windows (Fig. 21).

The opening of oceans is often a long process,

starting with continental rifting and ending in oceanic

spreading. This process is best marked by syn-rift

sediments and volcanic rocks, as well as post-rift

sediments. Complete subsidence histories are rarely

preserved in the Carpathian area, so emphasis is given

to stratigraphy and magmatism (Figs. 9–16). The

closure of oceans is best marked by the presence of

obducted ophiolite masses and turbidite-accretionary

prism belts. However, neither obduction nor turbidite

deposition is necessarily linked to the collision time

itself. Obduction is commonly an intra-oceanic event

well before collision, and turbidite deposition can last

long before or after collision. High-pressure metamor-

phism is another good indicator of subduction.

5.1. Palaeotethys–Ionian–Eastern Mediterranean

ocean

As most or the whole of this oceanic lithosphere

was subducted and no ophiolitic remains exist, the time

of opening is debated. It was estimated from continen-

tal margin development of north Africa by Stampfli et

al. (1991) who, based on transition of syn- to post-rift

sediments, proposed a Permian opening for the Eastern

Mediterranean. There are however, several other rift-

ing events (Early and Late Jurassic, Early–Late Cre-

taceous) with normal faulting and basin formation in

this area (e.g. Dlala, 2002). These deposits suggest that

there was a long-lasting rifting. The Eastern Mediter-

ranean could be continued in the northeastern margin

of the Arabian peninsula and a Permian formation of

this margin has been suggested (Ziegler and Stampfli,

2001). From a palaeobiogeographic point of view, an

early rifting or opening of the Eastern Mediterranean is

preferred, because faunas of Adria have to be separated

from Africa by the earliest Jurassic.

For sake of simplicity, the Eastern Mediterranean

ocean can be also considered a successor of the

Paleotethys. Palaeotethys separated Gondwana from

the European margin. Traces of it are supposed to be

found in the islands of Sicily, Chios and Crete

(Catalano et al., 1991; Champod et al., 2003; Stampfli

et al., 1998b; Ziegler and Stampfli, 2001). There is a

suspect occurrence of early, Anisian turbidite in the

external southern part of the High Karst platform (Fig.

21) (Aubouin et al., 1970). Taken these turbidite and

blueschist units as an indication of the Palaeotethys, a

trace oblique to known facies and structural zones can

be drawn. Permian–Anisian calc-alkaline volcanism

is widespread in the Dinaric and Hellenic chains (Fig.

22) (Pamic, 1984). This volcanism can be attributed to

the northward subduction of Palaeotethys (Ziegler and

Stampfli, 2001). Based on well-dated post-tectonic

sections in Sicily and Crete (Catalano et al., 1991;

Champod et al., 2003), the closure or docking of

Palaeotethys happened by Carnian or slightly later.

The Hellenic subduction of the modern Ionian–

Eastern Mediterranean ocean started in the Oligocene

or Middle Miocene and is still going on (e.g. Angelier,

1979). Back-arc rifting of the Aegean basin and

voluminous Miocene to Recent calcalkaline volca-

nism accompanies this subduction.

5.2. Budva–Pindos ocean

Ladinian rift-related basalts occur in the Budva

sequence (Dimitrijevic, 1982). These are overlain by

Triassic to Upper Cretaceous slope deposits (Fig. 12),

taken here as indicative of ocean margins, though

many authors doubt the existence of oceanic crust

(Aubouin et al., 1970). Pindos ocean is considered

opened by Middle–Late Triassic (Stampfli and Borel,

2002; Stampfli et al., 1991).

The closure of this ocean is documented by volu-

minous Palaeogene turbidite bodies (Fig. 12). How-

ever, the onset of subduction is not clear. The latest

Cretaceous–Palaeogene banatitic volcanism in the

Tisza–North Dinaric–Balkan area is proposed here

to have been generated by this subduction. Therefore,

subduction should have started in mid-Late Creta-

ceous. Closure is suggested to have taken place in

Oligocene (Richter et al., 1995), but blueschist meta-

morphism in the Olympus margin suggests that con-

tinental units entered the subduction zone by the

Eocene (Ricou et al., 1998).

5.3. Meliata–Vardar–Mures ocean

In the Dinaric sector, huge masses of ophiolites are

exposed. Radiolarites intimately associated with


basalts were either Middle-Upper Triassic or Middle-

Upper Jurassic (Fig. 12) (Obradovic and Gorican,

1988). Immediately adjacent, submerged continental

margins (High Karst) also show deepening and onset

of pelagic deposits from the Anisian–Ladinian (Dimi-

trijevic, 1982; Dimitrijevic and Dimitrijevic, 1991).

Upper Permian marine sediments are present on the

High Karst margin. All this suggests a Late Permian

onset of rifting and Anisian–Ladinian opening of this

oceanic trough.

Meliata mafic rocks in N Hungary–S Slovakia are

associated with Middle and Upper Triassic and upper

Middle Jurassic radiolarites (Figs. 10 and 11) (Dos-

ztaly and Jozsa, 1992). Directly adjacent palaeogeo-

graphic units (Bukk and Szilice) have both pelagic

sediments from the Anisian. In some fragments

(Torna), even Anisian–Carnian rift-related mafic

rocks are exposed (Mello et al., 1983). All these facts

point to a Middle Triassic opening of the Meliata.

Late Permian evaporitic and marine sediments in

Bukk suggest that the onset of rifting happened in

Permian.

In the Apuseni sector, there is a preserved Lower

Jurassic (180 Ma) oceanic crust (Fig. 14) (Savu and

Stoian, 1988). Consequently, rifting must have hap-

pened earlier. In the Eastern Carpathians, the Transyl-

vanides contain oceanic rocks of unknown, but

certainly pre-Barremian age (Fig. 15) (Sandulescu et

al., 1981a). They occur together with remains of a

submerged continental margin (Persani), which con-

tains Anisian rift-related volcanic rocks and Middle–

Upper Triassic pelagic sediments. Assuming that this

margin was facing the Vardar ocean, a Middle Triassic

opening seems probable.

The two characteristic ages of rifting coupled with

the demonstrated Guevgueli back-arc and oceanic

island arc sequences in Greece (Ricou et al., 1998)

suggest that this ocean consists of two plates: an

almost completely subducted Triassic–Jurassic plate

and a Middle–Late Jurassic back arc basin, large

masses of which were obducted. These can be taken

as separate oceans (as in Stampfli and Borel, 2002),

but for sake of simplicity we consider them as one,

coupled by an intra-oceanic subduction.

In the Dinaric–Hellenic sector, Vardar ophiolites

were obducted in Late Jurassic time and were overlain

by Tithonian reefs (Dimitrijevic, 1982; Ricou et al.,

1998; Zachariadou and Dimitriadis, 1995). Slightly

later, in Tithonian–earliest Cretaceous time, the Bos-

nian foredeep was formed (Fig. 12) (Aubouin et al.,

1970). This is thought to mark the closure and

collision of the southwestwards advancing nappe

complex of Tisza–Serbo–Macedonian units and the

ophiolites. An Albian shallow water event is thought

to mark collision (Dimitrijevic, 1982). In the Rho-

dope, a collision event is marked by syn-thrusting

metamorphism, with an age span of 140–80 Ma

(Ricou et al., 1998). Albian granite plugs the

Serbo–Macedonian nappe pile above an ophiolite

and above the underthrust Drama unit. In our opinion,

the ophiolite corresponds to the Vardar suture, the

underthrust Drama unit is the equivalent of High

Karst–Pelagonian unit. Later, Palaeogene ages are

interpreted as cooling ages linked with exhumation

(Ricou et al., 1998).

Upper Cretaceous and Palaeogene turbidite and

calc-alkaline magmatic bodies are present in and near

the Vardar belt (Canovic and Kemenci, 1988; Pamic,

2002). This is the main reason why a Maastrichtian–

Palaeogene closure of Vardar was suggested (Pamic,

1998b, 2002). The calc-alkaline Banatites can also be

found in other units outside the Vardar and continue to

the east, in the Sredno–Gorje Mts. of Bulgaria. They

roughly draw a curvilinear pattern oblique to, and

overlapping the Vardar (Fig. 8) (Balla, 1984). Mag-

matic bodies plug nappes north of the Vardar belt

(Ricou et al., 1998 and references therein). We think

that formation of island-arc volcanic rocks at the site

of the accretionary prism and in under- and overlying

nappes cannot be explained by a normal subduction

zone, as the volcanic belt should be located at ca. 150

km from the trench. Therefore, (1) in the Dinaric

sector the Upper Cretaceous–Palaeogene turbidite

units of the Vardar are not issued from an oceanic

trench but from a continental foredeep, activated by

thrust renewal; (2) the calcalkaline Banatitic belt is not

the result of Vardar subduction, but of an ocean more

to the SW, where Pindos is a likely alternative.

In the N Hungarian–S Slovakian sector, the

remains of Meliata are metamorphosed together with

their tectonic substratum in the Early Cretaceous

(Figs. 10 and 11) (Arkai, 1983; Csontos, 2000). By

the Albian, the whole nappe edifice already falls apart

due to unroofing of lower nappes (Plasienka, 1998).

In the Tisza sector, this ocean was closed probably

by mid-Cretaceous times. A Late Jurassic to Aptian


calcalkaline volcanic arc is documented in the Mures

belt (Fig. 14) (S�tefan, 1986). An Albian olistostrome

covers most nappe units (Figs. 14 and 15) (Sandulescu

et al., 1981a). Turonian shallow water sediments are

also post-tectonic in different parts of the belt. There is

a strong reactivation of nappe movements in the Late

Cretaceous–Paleocene, but this thrusting is not attrib-

uted to collision. Finally, Maastrichtian–Paleocene

calc-alkaline magmatic bodies plug the whole structure

(Figs. 6, 14 and 16) (Berza et al., 1998; Stefan et al.,

1988).

5.4. Penninic–Pieniny–Vah ocean

Late Early Jurassic opening of the Pieniny ocean is

suggested by the drowning of the Czorsztyn ridge and

also by the presence of Lower Jurassic breccias in a

tectonic window of the Western Carpathians (Fig. 11)

(Plasienka, 1987). A Bajocian onset of spreading is

suggested after asymmetric extension. This opening is

almost identical to that of the Alpine Piemont (=Pen-

ninic) ocean dated as Toarcian (Stampfli andMarchant,

1997). There is a controversy in the timing of opening

of this basin, however. Faunal separation of the ‘‘Med-

iterranean’’, i.e. Alpine–Western Carpathian–Adriatic

microcontinent from the European continental area is

well documented for the Pliensbachian onwards. This

separation was most probably due to separation by a

deep-sea barrier at the location of the future Penninic

ocean. It is therefore suggested that rifting with major

spatial separation took place well before spreading. The

only viable way to do that is by low-angle normal

faulting, since symmetric rifting with such an extension

should result in oceanic spreading much earlier (Pla-

sienka, 2002; Stampfli et al., 1991).

A much earlier, Triassic opening of Pieniny was

proposed by Birkenmajer et al. (1990). Their argu-

ments were based on clasts found in an Albian

conglomerate (Misık and Sykora, 1981). Later work

(Plasienka, 1995) suggests that the unit with conglom-

erates was probably formed in the inner parts of the

Western Carpathains and was emplaced in Late Cre-

taceous in its present position. A Triassic oceanic

opening north of the Western Carpathians seems

unlikely from a facies viewpoint, too. Upper Triassic,

and even lowest Jurassic continental, shallow-marine

facies of the northern, Tatric nappes suggest an

exposed land to the north (Fig. 11). Therefore, faunal,

facies relationships described above are considered

much stronger arguments than the clasts in the Albian

conglomerates.

The sedimentary record and a weak metamorphism

suggest a Campanian–Maastrichtian closure of this

trough (Fig. 11) (Birkenmajer, 1986; Plasienka et al.,

1997). A Maastrichtian conglomerate in the Pieniny

Klippenbelt seals earlier nappe structures.

5.5. Ceahlau–Severin–Magura ocean

Data on this rifting come essentially from the

Eastern Carpathians (Sandulescu et al., 1981a,b). Mid-

dle Jurassic rifting is documented in the mafic volcanic

breccias of the Black flysch nappe (Fig. 15). A sliver

of Tithonian ocean floor basalt is preserved in the SE-

bend of the Carpathians, beneath radiolarites and

Lower Cretaceous calcareous turbidites (Sandulescu

et al., 1981a,b). The same rocks and ages were found

in the Southern Carpathians (Fig. 16) (Savu, 1985).

Here marginal units on both sides of Severin show

evidence for alkali-mafic volcanism as early as Sine-

murian. The Arjana succession shows massive mafic

lava flows interlayered with Middle Jurassic sediments

(Iancu, 1986). Therefore, it was proposed that the

Severin portion opened somewhat earlier, in the Mid-

dle Jurassic, and the Ceahlau opened a bit later, in Late

Jurassic times (Sandulescu, 1975a, 1976).

Traces of mafic magmatism are present in the

Hettangian of Mecsek, northern Tisza terrane as well

(Fig. 14) (Nagy, 1969). Voluminous rift-type mafic

volcanism began most probably in the Late Jurassic

and reached a maximum in the earliest Cretaceous

(Harangi et al., 1996). Traces of similar volcanism

were found in the External Western Carpathians as

well (Birkenmajer, 1986). However, pelagic sedi-

ments appear much earlier in both domains (Fig. 11)

(Birkenmajer, 1977; Galacz, 1984). Faunal studies

have shown a separation from the European faunal

realm by the Bathonian (Voros, 2001). Therefore, a

Middle Jurassic onset of rifting and a Late Jurassic

break-up is proposed.

In the Eastern and Southern Carpathian sector, the

closure of this oceanic trough began in the earliest

Cretaceous (Sandulescu et al., 1981a,b). Turbidites

were deposited in this belt until mid-Cretaceous times.

A prograding Albian conglomerate fan covers all

Dacia and Ceahlau nappe contacts (Fig. 15). Appar-


ently, no volcanism accompanied the closure. The

Magura ocean began to close in the Late Cretaceous

(Birkenmajer, 1998). Subduction is accompanied by

deposition of voluminous turbidite bodies (Ksiazkie-

wicz et al., 1968). Closure of the ocean took place in

the Oligocene–Early Miocene. The remains of the

subducted slab produced the calcalkaline volcanic arc

of the Carpathians (Szabo et al., 1992).

6. Plate-tectonic reconstruction

6.1. Basic principles of west-Tethyan plate

reconstruction

Although we have focused on the Carpathian–

Pannonian region, our reconstructions involve a larger

area since the driving forces of plate motion are often

external. The movement of the African and European

plates are particularly important. The palinspastic

maps (Figs. 22–26) are considered a collection of

ideas, rather than true plate-tectonic reconstruction,

since we do not have the means to model all plate

movements in the area. Readers who are familiar with

such reconstructions will recognise the incorporation

of earlier ideas and geometrical solutions (e.g. Der-

court et al., 1986, 1993; Frisch, 1979; Gaetani et al.,

2000; Karamata et al., 1999; Kazmer and Kovacs,

1989; Neugebauer et al., 2001; Plasienka, 1998;

Rakus et al., 1990; Sandulescu, 1980a; Stampfli and

Borel, 2002; Stampfli and Marchant, 1997; Stampfli

et al., 1991, 1998a,b; Ziegler, 1988; Ziegler and

Stampfli, 2001). However, our reconstructions include

many important aspects of Carpathian geology and

palaeomagnetics that have not been fully considered

in previous models. Our model is most applicable to

the Carpathian–Eastern Alpine–Adriatic area. Since

we are less familiar with regions to the west and east,

we are less confident of the suggested geometries and

motions of these areas.

6.1.1. Geometry

The projection, framework and contours of the

European and African plates except for Late Permian,

as well as main latitude lines, geodynamic unit con-

tours in the west and south are taken from Stampfli

and Marchant (1997) and Stampfli et al. (1998b). The

geodynamic unit contours in the Carpathian–Alpine–

Adriatic area were redrawn from a geological map of

the same scale. To remove the effects of Cenozoic

tectonics, retro-deformation was attempted based on

Schmid et al. (1996) in the Western Alpine, Frisch et

al. (1998) in the Eastern Alpine; the estimations of

Roca et al. (1995), Roure et al. (1993), Tari et al.

(1999), the reconstructions of Balla (1984), Csontos et

al. (2002), Fodor et al. (1999), Kovac et al. (1998) in

the Carpathian area, and the works of Fodor et al.

(1998), Schonborn (1992, 1999) and Tomljenovic and

Csontos (2001) in the Dinaric–Southern Alpine sector

(Figs. 3 and 4).

The geometry of the geodynamic units was held

fixed during the tectonically quiet episodes and

attempts were made to account for the transport

directions of particular tectonic events (Fig. 7). Great-

er liberty was taken in the retro-deformation of

Carpathian–Dinaric thrusts and orogens, since these

have not been previously estimated and are necessar-

ily based on incomplete data. During differential

rotations, the shape and area of individual blocks

were held constant, based, in large part, on the palae-

omagnetic constraints of Marton (1988, 1993b) and

others (Fig. 20).

We involved a further palaeomagnetic constraint

by keeping the Adriatic promontory and attached

units as fixed to Africa as possible. Channel and

Horvath (1976), Marton (1993a) and Marton and

Marton (1978) have demonstrated, that the Apparent

Polar Wander curves for these units closely resembles

the curve for Africa, at least until the mid-Creta-

ceous. Most tectonic events can be deduced from the

direct or indirect effects of the Africa vs. Europe

plate movements. This is reflected in the reconstruc-

tions by holding Europe fixed and indicating the

motion of the African plate with respect to its

previous position by an arrow drawn at the Tunis

promontory (Figs. 22–26).

The reconstructions are computer-drawn so that

the geometry of a given reconstruction could be

compared with those immediately preceding and

succeeding. The reconstructions were backward

modelled with each step checked against palaeomag-

netic rotations and consequent deformations. Several

runs were made for each reconstruction. While

backwards modelled, however, the reconstructions

are presented in a forward progression in the follow-

ing discussion.

Fig. 23. Proposed position of units in the Sinemurian and Oxfordian times. Same description as for Fig. 22.


atology,Palaeoeco

logy210(2004)1–56

39

Fig. 24. Proposed position of units in the Tithonian and Aptian times. Same description as for Fig. 22.


atology,Palaeoeco

logy210(2004)1–56

40


The proposed Late Permian starting situation (Fig.

22) differs from most current reconstructions (e.g.

Dercourt et al., 1993; Stampfli and Borel, 2002;

Stampfli et al., 1998b) in the position and size of

the Bihor and Getic microcontinents and the relative

positions of the Eastern Alps and Western Carpathians

within the Austroalpine microcontinent. It is, howev-

er, similar to the pattern proposed by Kovacs (1982).

It is proposed that the northern margin of the Bihor

microcontinent makes an acute angle to its southern

margin, in order to keep this southern margin as linear

as possible. Furthermore, the Szilice–Juvavic nappes

were placed on the southern margin of Tisza. How-

ever, the size and extent of this microcontinent is open

to debate. It is possible for example, that it was only a

discontinuous belt of offshore plateaus similar to the

modern Bahamas.

6.2. Plate tectonic reconstruction of the Alpine–

Carpathian–Pannonian area

6.2.1. Late Permian–Late Triassic

It is generally accepted that the rifting of the

Central Atlantic happened from a classical Pangea

situation (Gaetani et al., 2000; Stampfli et al., 1998b),

which remained fixed from the Permian until the

Early Jurassic (‘‘Pangea A’’). Puzzling palaeomag-

netic data, however, suggest that this fit cannot be

maintained in the Permian–Early-Middle Triassic

period (Irving, 1977; Muttoni et al., 1996; Torcq et

al., 1997). These data indicate that there might have

been a major right-lateral shear and westward dis-

placement of Gondwana relative to Laurasia during

the Permian (Muttoni et al., 1996) or the earlier half of

the Triassic (Torcq et al., 1997) (‘‘Pangea B’’). This

time interval conspicuously coincides with the closure

of Palaeotethys, and the major right lateral shear can

explain many tectonic features involved in our recon-

struction and therefore we preferred the Pangea B

situation (Fig. 22). However, this right lateral shear

does not affect the internal geometry of our Alpine–

Carpathian–Dinaric microplates, only facilitates and

explains better the Paleotethys subduction.

Regardless of the Pangea A or B situation, the

Palaeotethys was subducted obliquely beneath the

northern margin formed by the Southern European

microcontinents (see also Ziegler and Stampfli, 2001).

This subduction created a widespread calcalkaline

volcanic activity in the Dinaric–Hellenic part of this

margin, in some parts as early as the Permian, in

others in the Anisian–Ladinian (e.g. Dimitrijevic,

1982; Karamata et al., 1999; Pamic, 1984). Subduc-

tion roll back created a couple of back-arc basins

along the Pindos (s.str) and the Meliata–Ophiolite

Belt–Vardar oceans in Middle Triassic. The Meliata–

Vardar ocean then reached considerable width. In our

opinion, the southern part of Adria was separated

from Africa by the remains of Palaeotethys or even-

tually by the incipient rift of the Eastern Mediterra-

nean (see Ziegler and Stampfli, 2001).

A third rift was opened between Moesia and the

Ukrainian shield. This rift coincided with the axis of

the Polish Trough, or the Teisseyre–Tornquist Zone.

On land, huge thickness of sediments was accumulat-

ed during the Permian and Mesozoic (Marek and

Pajchlowa, 1997). In the Dobrogean sector of Moesia,

an episode of mafic rift volcanism eventually led to

the opening of an oceanic arm. A branch of this rift

(see also Ionesi, 1994; Tari et al., 1997) might account

for the enigmatic Ladinian pelagic event in the Infra-

bucovinian unit (Fig. 16). Redeposited Middle-Upper

Triassic deep-sea sediments in the Magura turbidite

(Western Carpathians; Sotak, 1985) may similarly

have derived from here. The extent of rifting and

the width of the eventual ocean (Seghedi and Szakacs,

1994; Visarion et al., 1990) are unknown, but must

have been wide enough to produce the later Dobro-

gean orogeny and nappe stacking. Opening of this

trough could be explained by the distant slab roll back

effect of the Palaeotethys (Kure back-arc of Stampfli

and Borel, 2002), or alternatively by rift propagation

from the east.

6.2.2. Early Jurassic–Late Jurassic

From the Pangea A situation reached by the Late

Triassic, Africa began its protracted eastward move-

ment relative to Europe (Fig. 23). The movement

created left-lateral shear between the two main con-

tinents until the Late Cretaceous. This movement,

originating with the opening of the Central Atlantic

(Frisch, 1979; Stampfli et al., 1998b; Ziegler, 1988)

had dramatic consequences for the region under

consideration. With Adria coupled to Africa, the

eastward shift resulted in the gradual opening of the

Penninic–Vahic ocean. Based on faunal differences,

the opening was already significant at the beginning


of Jurassic, although many authors favour an oceanic

spreading at late Early Jurassic (Stampfli et al., 1991;

Plasienka, 2002). If Adria is kept rigidly attached to

Africa, the Penninic realm should have been opened

in the Early Jurassic as a relatively wide ocean. The

same pattern of Penninic spreading persisted during

the Late Jurassic.

Synchronously with these events, the Dobrogea

sphenochasm began to close. Southward-directed sub-

duction is suggested by the vergence of the later

orogenic belt. Closure is completed by Late Jurassic

(Ionesi, 1994).

Eventually, closure of the Dobrogea oceanic arm

may have initiated rifting in the Southern Carpa-

thian Ceahlau–Severin sector where the Early Ju-

rassic is dominated by voluminous volcanic rocks

and thick clastics. However, ocean-formation did

not occur until the late Middle Jurassic. Similarly,

in the Infrabucovinian–Ceahlau units (Figs. 16 and

17) rift-related volcanism occurs in the Middle

Jurassic. The rift appears to have propagated north-

westwards. The Bihor–Getic–Serbo–Macedonian

ribbon microcontinent was finally separated from

the European margin by the late Middle Jurassic

(Fig. 23). From the latest Jurassic onwards, a major

left-lateral transpressive contact of the Bihor–

Getic–Serbo–Macedonian and the Dinaric High

Karst margins is needed, therefore the former had

to be located more to the SE. We thus speculate

that after break-up this ribbon-continent quickly

propagated towards the SE, leaving a wide Magura

ocean behind. Southeastward motion of this ribbon

also implies that the Ceahlau–Severin ocean opened

more like the present Gulf of California, leaving a

narrow ocean behind (Sandulescu, 1980a). All in-

cipient oceanic troughs in the Dinaric–Hellenic

sector also expanded at this time.

6.2.3. Latest Jurassic–Aptian

Due to the major left lateral shear between Africa

and Europe, the Meliata–Vardar ocean commenced

within-ocean subduction in the Middle Jurassic (Fig.

24) (Robertson and Karamata, 1994; Csontos, 2000).

In the Late Jurassic, possibly due to the oblique

scissor-like margins, this subduction resulted in obduc-

tion of the accretionary prism and large masses of

ultramafics (Pamic, 1982). As Robertson and Kara-

mata (1994) clearly put out, obduction must have

preceded emplacement of the accretionary prism and

ophiolites on either continental margins. We accept

eastward and westward obduction (Stampfli and Borel,

2002), with the remark that in any case, the final

emplacement of the Vardar ophiolites should have

been directed towards and over the western High

Karst–Austroalpine margin. Many observations sug-

gest that this emplacement happened in the latest

Jurassic (Argyelan and Csaszar, 1998; Csontos,

1988, 2000; Dimitrijevic, 1982; Maluski et al., 1993;

Mandl, 1999; Schweigl and Neubauer, 1997a,b).

Palaeomagnetic data suggest that the Alpine–

Western Carpathian margin was almost straight at that

time (Marton, 1993b). Tectonic transport directions

suggest an oblique or margin-parallel obduction onto

the Bukk–High Karst margin (Figs. 7 and 18). Clastic

material input and gravity gliding of Triassic succes-

sions also occurred in the accretionary and foreland

basins during mid-Late Jurassic (Dimitrijevic and

Dimitrijevic, 1973; Gawlick et al., 1999; Robertson

and Karamata, 1994; Schweigl and Neubauer, 1997a,

b), therefore it is proposed that the Bihor–Getic–

Serbo–Macedonian margin, the host of these succes-

sions, was already close to the Dinaric High Karst

margin at that time. The leading margin of the Bihor–

Getic–Serbo–Macedonian upper plate was probably

denudated, since crystalline basement is transgressed

by latest Jurassic at some places.

During the Early Cretaceous, the eastward move-

ment of Africa relative to Europe and the resulting

left-lateral oblique collision between the Bihor–

Getic–Serbo–Macedonian upper plate and the Dina-

ric High Karst lower plate continued. This soft colli-

sion may account for the 120-Ma metamorphic event

observed in the High Karst margin. The more distal

High Karst–Western Carpathian foreland was marked

by the formation of a turbiditic basin from the Hronic

through Tirolic nappes and Transdanubian Range,

possibly joining the Bosnian flysch (Faupl and

Wagreich, 1992). This basin persisted from earliest

Cretaceous until at least the Albian.

As a result of the lateral shear along the Vardar

suture, the leading edge of the Bihor–Getic ribbon

microcontinent is thought to have blocked and an

oroclinal bend started to form. The internal part of the

oroclinal bend might not only have preserved a small

remnant of the Vardar ocean, but it also experienced

major shortening. The bending is thought to be at the


origin of the centripetal nappes and outward propa-

gating foredeeps of the Tisza and the East Carpathian

Dacia terranes. Early Cretaceous thrusting is weakly

indicated in the southern part of the Bihor micro-

continent (Dallmeyer et al., 1999; Pana, 1998; Pana

and Erdmer, 1994) and suggested in the Getic micro-

continent (Sandulescu et al., 1981a; Pana, 1998). The

innermost nappes received ophiolites and fragments

of the colliding margins from the suture (Persani,

Transylvanides).

6.2.4. Albian–Santonian

During later Early Cretaceous, the Valais oceanic

trough was getting more and more open (Fig. 25). Its

rifting was caused by the northwards propagating

rifting in the Atlantic ocean, then it widened by the

rotation of the Hispanic block relative to Europe

(Stampfli et al., 1998b). The Valais spreading may

have driven the Czorsztyn microcontinent more to the

east, to reach the northern part of the Austroalpine

(Western Carpathians).

It is thought that the obliquely colliding Czorsztyn

microcontinent caused the Alcapa oroclinal bend, i.e.

the Western Carpathian sector to bend towards the

Dinaric one. The remnants of the Vardar suture were

trapped in the innermost parts of this oroclinal bend in

an uppermost tectonic position. This bending and

related thrusting must have occurred from the Albian

on, when the structurally lower Veporic unit was

metamorphosed (Plasienka, 1998) and when con-

glomerates in the more distant Fatric foreland received

an assemblage of clasts representative of the whole

Gemer nappe pile (Plasienka, 1995). Albian is also an

important metamorphic episode in the Eastern Alps

(Dallmeyer et al., 1996). Early Cretaceous nappe

stacking possibly initiated outward propagating thrust

systems throughout the Cretaceous in the Alps and in

the Western Carpathians (Plasienka, 1998).

Convergent left lateral shear between Europe and

Africa also continued during this interval. With the

change of the rotation pole, however, the main move-

ment vectors slowly turned from E–W to more N–S.

The result was a change from a Dinaric margin-

parallel transpression to a margin-perpendicular short-

ening in the Albian. Tectonic transport directions

suggest a nappe-perpendicular motion with major

folding. This motion produced a bigger underthrusting

of the Dinaric High Karst margin with an Albian

metamorphic event in the Rhodope (Ricou et al.,

1998). The northwestward shift and differential rota-

tion of the Bihor –Getic ribbon microcontinent

resulted in the soft collision of the Bihor–Getic and

the Western Carpathian–Austroalpine oroclinal bends

by the Turonian, another major and common tectonic

episode.

Since space was confined, the Eastern Carpathian

part of the Bihor–Getic microcontinent experienced

collision with the southern margin of Moesia, i.e. the

Coumanian cordillera (Sandulescu et al., 1981a). This

collision is indicated by fossils as Aptian, when

European shallow benthic faunal elements first invad-

ed the southern microcontinents (Voros, 2001). The

Ceahlau oceanic branch was certainly closed by the

Albian, because post-tectonic conglomerates of this

age overlie the nappes (Sandulescu et al., 1981a,b). It

is unclear when the Severin part was closed, but in the

Southern Carpathians Albian shallow-water sediments

transgress an unconformity.

Most of the terranes in the study region were

amalgamated, so in the following the terms Alcapa

and Tisza–Dacia will be used.

6.2.5. Latest Cretaceous–Eocene

This period is characterised by a northeastward–

northward shift and counterclockwise rotation of

Africa relative to Europe (Fig. 26). This movement

derived from the geodynamic framework is thought to

be slightly modified by the Penninic subduction. Its

subduction retreat might have additionally rotated the

amalgamated Alcapa–Adria–Tisza and related ter-

ranes more to the NW. This movement closed the

Pieniny–Valais (e.g. Dewey et al., 1989) and the

Budva–Pindos oceans. These subductions were oppo-

sitely directed and were coupled by a wider north-

northwesterly oriented shear zone characterised by

right lateral shear. This shear zone was probably

distributed between the main ‘‘onion shell’’ faults of

Alcapa including the Pieniny Klippenbelt, and contin-

ued at the northern margin of Tisza, then in the Sava–

Vardar belt of the Dinarides. Most probably more

external zones, like the Dinaric Ophiolite belt, or

Budva were also members of this wider shear zone.

While there are no direct subduction-related magmatic

traces of the Penninic–Valais–Vahic subduction, the

Late Cretaceous–Early Palaeogene banatite belt is

proposed to originate from the subduction of the

Fig. 25. Proposed position of units in the Albian and Santonian times. Same description as for Fig. 22.


atology,Palaeoeco

logy210(2004)1–56

44

Fig. 26. Proposed position of units in the Maastrichtian and Eocene times. Same description as for Fig. 22.


atology,Palaeoeco

logy210(2004)1–56

45


Budva–Pindos ocean. The Pindos was finally closed

by Oligocene (Richter et al., 1995).

The Late Cretaceous northeasterly component of

the main motions resulted in the two oroclinal bends

getting tighter. Because of the position of the Tisza–

Dacia orocline west of Moesia, there was an important

additional ‘‘laramian’’ shortening in the opposing

Tisza and Southern Carpathian sectors but not in the

Eastern Carpathians.

By the Late Cretaceous, the Vah ocean was prob-

ably closed and the Czorsztyn microcontinent collided

with the Austroalpine part. Czorstyn and Alcapa had

an oblique collision, so this could produce the strike-

slip related phenomena and the small Gosau basins

described by Wagreich and Faupl (1994).

Shortening in the internal parts of the Austroalpine

and in the Western Carpathians was accompanied by

perpendicular extension and basin formation. Parallel

belts within the Tisza unit, along former (reactivated)

nappe boundaries, host Senonian sedimentation. This

is especially true not only for the Mures belt (Lupu,

1976) but also for more internal parts (Szentgyorgyi,

1989). By the Late Eocene, a different type of basin

pattern is observed. This latter is either related to the

right-lateral shear or to the shortening events in the

Dinarides.

7. Conclusions

Our reconstruction differs from previous ones in

several important points. These mainly stem from the

importance we give to Carpathian terranes in the

western Tethys. The Intra-Carpathian terranes are all

formed of different Mesozoic geodynamic units, i.e.

(micro)continents and oceans. The Alcapa terrane is

composed of the northern Czorsztyn microcontinent

bordered to the south by the Pieniny–Vah ocean,

followed by the Austroalpine microcontinent. South-

east of these, the remains of the Vardar–Meliata

ocean can be found. The Tisza terrane is built of

nappes of the Bihor microcontinent, flanked to the

south by the Vardar–Mures� ocean. The Urmat unit is

probably derived from the margins of the latter. The

Dacia terrane is composed of sheared off slices of

the European continental margin (Danubian), fol-

lowed to the west and north by the Ceahlau–Severin

ocean and the Bucovinian–Getic microcontinent.

The westernmost element is the Vardar–Mures�ocean.

A second major difference is in the position of

these microcontinents-oceans. The Bihor –Getic

microcontinent originally lay east of the Western

Carpathians and filled the present Carpathian embay-

ment in the Late Palaeozoic–Early Mesozoic. A

major internal ocean, Vardar occupied the region

between the southern margin of the Bihor–Getic

microcontinent and the margin formed by the internal

Western Carpathian–Austroalpine–Transdanubian–

High Karst margin. Both margins are kept almost

linear, because later they enter into a long-lasting

left-lateral transpressive collision, otherwise very dif-

ficult or impossible to explain.

A third major difference is the location of the

‘‘exotic’’ Juvavic, Szilice and Upper Codru nappes.

These units now form gravity nappes, often related to

the Vardar melange. They are thought to have glided

down the now denudated southern margin of the

Tisza–Getic microcontinent and could have been

several times re-emplaced.

A fourth major difference arises from our accep-

tance of the ‘‘Pangea B’’ situation in Permian–Middle

Triassic times. This position and the change into a

‘‘Pangea A’’ situation inMiddle Triassic can explain the

oblique subduction of Palaeotethys; the widespread

Middle Triassic volcanism in the Dinaric–Hellenic

chain and the simultaneous back-arc opening of paral-

lel oceanic branches in the Dinaric–Austroalpine area.

A fifth difference is that we operate with a Vardar

ocean, which disappears by the Early Cretaceous. The

main collision event is a margin-parallel left-lateral

shear imposed by the relative motion of Africa and

Europe, followed by a margin-perpendicular thrust-

ing. In our opinion, the Late Cretaceous–Palaeogene

calcalkaline magmatic rocks widespread in the Bal-

kans are not due to the much earlier subduction of the

Vardar ocean, but to the synchronous subduction of

the Pindos ocean.

The sixth main difference occurs towards the end

of the Mesozoic. Facies belts, tectonic transport direc-

tions and palaeomagnetic data suggest that two oro-

clinal bends, the Alcapa on the Dinaric margin and the

Tisza on the Southern Carpathian–Getic margin were

formed. Their bending in the Albian–Maastrichtian is

due to the blocking of the general left-lateral shear,

and the oblique collision of Alcapa with the Czorsztyn


microcontinent. The two oroclinal bends are finally

opposed and pushed into the Carpathian embayment

during the Palaeogene and Neogene.

The last main difference is the link of the main

Palaeogene shortening in the Alpine sector to the

similarly important shortening in the Hellenic sector.

The oppositely dipping Penninic–Valais and Budva–

Pindos subductions are linked by a major right-lateral

shear belt through former important structural zones.

Acknowledgements

The authors are indebted to many colleagues for

the discussions of earlier oral and written versions of

the manuscript. We would like to thank especially F.

Horvath, A. Galacz, S. Kovacs, M. Kazmer, E.

Marton (Budapest), D. Plasienka, M. Kovac (Bra-

tislava), F. Neubauer and C. Tomek (Salzburg), K.

Birkenmajer (Krakow), M. Sandulescu (Bucharest)

S. Schmid (Basel) and P. Ziegler (Basel). We express

our thanks to W. Frisch, F. Horvath, G. Stampfli,

Alonso-Gutierrez, J. Von Raumer, B. Murphy, A.

Collins and D. Nance who kindly revised and

improved earlier versions of the manuscript. This

version was helped by critical remarks of K.

Birkenmajer, F. Neubauer, B. Sperner, G. Stampfli,

and F. Surlyk. IGCP project 453 is gratefully

thanked for moral and material support. Hungarian

Science Foundation OTKA projects T 043760, T

037595 are also thanked for support.

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