Tec/onophwcs, 88 (1982) 61-102

Elsevier Scientific Publishing Company, Amsterdam--Printed in The Netherlands

61

DEVELOPMENT OF THE PANNONIAN BASIN BASEMENT THROUGH THE CRETACEOUS-CENOZOIC COLLISION: A NEW SYNTHESIS

Z0LTkN BALLA

Hung~nan Geophysrul Institute ‘Lorand EB~otis’. Columbus 17.23. Budapmt 11.1145 (Hungar~~

(Received June 3, 198 1: revised version accepted February 4. 1982)

ABSTRACT

Balla. Z., 1982. Development of the Pannonian basin basement through the Cretaceous-Cenozoic

collision: a new synthesis. Tecronophysics. 82: 6 I- 102.

A geodynamic analysis was carried out going back in time from the present. The collision took place

during the Cretaceous and Cenozoic, Based on the subduction origin of the Neogene talc-alkalic

volcanites three microplates are delineated. The southern and the northern ones are approximately the

same as in previous plate tectonic syntheses but the third one is outlined here for the first trme as the

basement of a Miocene island arc. Restoring the former position of microplates a basin of oceanic crust is

outlined on the site of the present Pannonian basin of continental crust. At the beginning of the Neogenr.

the present northern microplate was located somewhere west of this basin and the present southern one

was located on its southern shelf. The reconstructed Cretaceous-Paleogene histories of these microplates

were different. During the whole Barremian to Oligocene time interval the southern microplate wa\

located on the passive margin of an oceanic basin which was coming into being through contmental rifting

of the European plate margin in the Valanginian to Hauterivian. The northern microplate. located on the

African plate margin, was undergoing collision with an island arc in the Early Cretaceous. The oceamc

basin of marginal-see type behind this island arc was closed along a southerly directed subduction zone in

the Eocene. This resulted in the appearance of a talc-alkalic volcano-chain along this microplate and in a

new basin opening behind. This marginal sea basin divided the present northern microplate from the

African plate and joined the oceanic basin north of the present southern microplate. The present

Pannonian basin was assembled as a lithospheric mosaic through convergent and transcurrent movements

of the defined microplates during the Neogene. Sediments of the former oceanic basin now strike along

the contacts of continental microplates as flysch belts.

INTRODUCTION

The Pannonian basin lies in the curving part of the Alpine system between the

structural belts of the Eastern Alps-Carpathians and the Southern Alps-Dinarides

of contrasting polarity. The so-called ‘Pannonian Median Mass’ was assumed as a

means of explaining this situation. The Mesozoic and Cenozoic of this ‘Mass’.

however, are not different from those of the same age in the surrounding area. This

0040-195 l/82/0000-0000/$02.75 C 1982 Elsevier Scientific Puhlishrng Company

62

is the reason why some believed that this area was also of the ‘geosynclinal--nro-

genie’ type (for historical review see Balogh, 1972). Both concepts are also recog-

nized in the tectonic interpretations of the last decade: the basement of the

Pannonian basin was considered to be an assemblage either of rigid continental

microplates (SzBdeczky-Kardoss, 197 1, 1976a; Channel1 and HorvBth, 1976; Chan-

nell et al., 1979) or of structural belts continued from the folded frame (Dank and

Bodzay, 1971; Bodzay, 1975. 1977: Szepeshby, 1975, 1977, 1979. 1980; Wein.

1978a, b).

In the earliest models incorporating the plate tectonic concept, the Carpathian

region (= Carpathians + Pannonian basin) was a detail in the kinematic schemes for

the Mesozoic and Cenozoic development of the whole Mediterranean (Szhdeczky-

Kardoss, 1971, 1976a; Dewey et al., 1973; Channel1 and HorvBth, 1976). In the

Mediterranean geodynamic reconstructions in general, the main focus of attention is

on studies of the pre-collisional complexes. I think that a true picture of geodynamic

processes cannot be produced without analysis of the collision itself which took

place in Cretaceous-Cenozoic times.

GEODYNAMIC ANALYSIS OF THE CRETACEOUS-CENOZOIC COMPLEXES

A history of collision can be understood only if it is started from the present

situation and an investigation is made going backwards in time. The time-period

under discussion will be divided into two phases: a Neogene and a Cretaceous-

Paleogene one.

Neogene

The intensive talc-alkalic volcanism is the most important peculiarity of the

Neogene development of the Carpathian region, and its explanation is one of the

principal tasks of geotectonic syntheses. Two main areas of the Neogene talc-alkalic

volcanites are delineated (Balla et al., 1977; Balla, 1980): the East Carpathian

volcanic belt (Fig. 1) and the Inner Carpathian volcanic area (Fig. 2). The genetic

relations between the young alkalic basalts and the collisional structure of the

Pannonian basin is poorly understood (Balla, 1980) and, therefore, these rocks will

not be discussed.

The East Carpathian volcanic belt

The East Carpathian volcanic belt strikes along the innermost side of the Outer

Carpathians and is divided into two sections: the Vihorlat-Gutin volcano-chain

lying in the Northeastern and the C&limani-Harghita chain in the Eastern

Carpathians.

In the Vihorfat-Gutin volcano-chain (Fig. 1, I), the volcanites were dismembered

into several stratigraphic units: their number, content and age were all disputed. Not

Fig. I. The East Carpathian volcanic belt in the structure of the Carpathian region (after Balla. 1980.

modified). Numbers on the map (in circles): I = Vihorlat-Gutin volcano-chain: 2 = Calimani-Harghita

volcano-chain.

a single one of the numerous K-Ar dates (~erlich and Spitkovskaya, 1974;

~~khay~ova et al., 1974), however, is younger than 8 m.y., and these data give

evidence of the existence of two main cycles in the Vihorlat-Gutin range: the

64

younger of them is %I0 m.y. old, the older one 1 I- 15 my. old. ‘I‘hc youngc~-

volcanites are regarded only as rocks of the East Carpathian volcanic belt.

In the Crilimuni-Hurghitu volcano-chain (Fig. 1, 2). the volcanites are separated

into two stratigraphic units. The lower of them covers sediments of the uppermost

Miocene, and the upper overlies the lower one with erosional unconformity. and the

upper boundary of its age is debated. The K-Ar age of the andesitic basalt from the

upper part of the lower complex is 7.4 .k- 0.7 m.y., and the age of the youngest

pyroxene andesite from the upper complex is 3.9 * 0.2 m.y. (RHdulescu et al.. 1973).

Thus, the volcanism occurred in the period 4-8 m.y.

The Inner Carpathian volcanic area

The arrangement of the Inner Carpathian volcanic units (Fig. 2) does not show

any links with the Carpathian arc so that this volcanism cannot be connected with

the tectonic processes in the Carpathians themselves. The volcanoes of this area are

grouped into several units.

Most of the talc-alkalic volcanites are concentrated in the Middle Hungarians-

Transcarpathian volcunic belt (Fig. 2, a-11). The volcanism started along the whole

belt roughly at the same time 16-18 m.y. B.P. It finished in the Soviet Trans-

carpathians (Fig. 2, I/ ) and northeastern Hungary (Fig. 2, JO) perhaps lo--l2 m.y.

B.P. (Balogh and Rakovits. 1976; Hamor et al., 1978), but the end of the volcanism

has remained uncertain elsewhere.

The BtikkaIja-Tokqj-PreSoc arc (Fig. 2, S- 7) is characterized by a large number

of chiefly small volcanoes with a great deal of dacite and rhyolite besides andesite,

forming a definite continuous arc. Its convexity contrasts with that of the

Carpathians. The volcanism took place in the 14- 10 m.y. time span.

The Central Slovukian-North Hungarian volcano-group (Fig. 2, 1 -4 ) consists

chiefly of large andesitic stratovolcanoes distributed irregularly. The CserhBt-MBtra

volcano (Fig. 2. 4) developed about 17- 14 m.y. B.P. (Himor et al., 1978. 1979); the

Dunazug-BGrzsGny volcano (Fig. 2, 3) was formed about 15-14 m.y. B.P. (Balla et

al., 1981); the Modrji Kameri range (Fig. 2, 2) came into being about 14-13 m.y.

B.P. and the Stiavnica unit (Fig. 2, 1) arose about 14-12 m.y. B.P.

The small volcanoes of the Trunsylvanian group (Fig. 2, 13) are mostly of the same

age as those in the Stiavnica unit.

In the Mecsekian volcano-group (Fig. 2, J2) there are only some bodies of

uncertain type. Their stratigraphic position is unknown, and the K-Ar age is 20-21

m.y. (.&v&&Sobs and Ravasz, 1978).

Geodynamic interpretation

The following key questions to the geodynamic interpretation need to be answered:

Are the Carpathian talc-alkalic volcanites of subduction origin or not? (For an

analysis see Balla, 1980). One supposes that the East Carpathian volcanic be/t (Fig. 1)

was related to a Benioff-zone dipping southwestwards with a near-surface suture

Fig. 2. The Inner Carpathian volcanic area (after Baila, 19X0. modified). Numbers on the map (in c‘~rcle\):

I = the Stiavmca volcano-group: 2 = tht: Modr$ Kameh range: 3 = the Bijrzxmy~Dunazug volcano:

4 = the Cserh5t-Mgtra volcano; 5 = BlikkalJa: h =Tokaj Hillz; 7 = the PreSov range: K = buried volc.tnic

range in the Middle Transdanubian; Q=buried volcanic range in the northwn part of thz Great

Hungarian Plain; 10 = buried volcanoes of the Nyir-Hajdu region: /I = Bcrepovo Hills: /_? : the Mecsek

volcano-group; 13 = the Transylvanian volcano-group.

66

somewhere within the Carpathian foredeep, beneath the margin of the overthrustctf

Carpathian flysch belt (Naumenko and Goncharuk, 1969; Bleahu et al.. 1973

Gofshtein, 1975). A substantial time gap exists. however, between the folding and

the volcanism.

Within the frame of the subduction concept, the fZy.rch belt is analogous to the

accretionary wedges (‘non-magmatic elevations’) in front of the volcanic arcs. The

deformation takes place in the lowermost part of the accretionary wedges (Seeley et

al., 1974; Kulm and Fowler, 1974; Sorokhtin, 1979), while in their upper and back

parts, sediments, for example, of Miocene age can overlie, with sharp discordance,

complexes folded in the earlier phases of the same subduction (Fig. 3A). Thus. the

discordance itself does not indicate the end of the subduction but only shows that

the subduction has continued for a long time. A very important fact is that the

youngest deformation appears in the topographically fowest position. The upper age

boundary of the Carpathian subduction cannot be determined by the oldest sedi-

ment overlying the folded flysch. because the deformation was continuing in the

lowest part of the flysch complex during its overthrusting on the foredeep.

At oceanic margins, the deformation takes place inside the youngest sediments

beneath the inner trench slope. Along the Carpathians, the continental lithosphere of

the European plate joins that of the ~annonian basin. Consequently, the deforma-

tion was taking place in a well-developed collisional state when the oceanic trench

was replaced by a foredeep. A molasse complex of great thickness fills this foredeep

and separates the deformation level from the sedimentation one (Fig. 3B). The

intensive folding in the deep horizons of the foredeep affects the older sediments and

appears near the surface as overthrusts only. The difference between the age of the

Carpathian folding determined by classical methods, and that of the volcanism does

not support the temporal separation of the volcanism from the subduction.

The geodynamic interpretation of the Inner Carpathian volcanic urea is more

complicated and, in the present state of our knowledge, can only be partly achieved.

The Middle Hungarian-Transcarpathian volcanic belt (Fig. 2, X-II ) is discussed first.

For the establishment of the potassium zoning, petrochemical data are sufficient

only in the central section (Juhasz, 1971). The reconstructed ancient Benioff-zone

dips northwestwards (Fig. 4). In front of this volcanic belt, the Szolnok-Maramures

flysch belt (see Fig. 2) lies as an analogue of accretionary wedges. As mentioned

above, the sediments can overlie the folded flysch synchronously with the continuous

subduction. Tuffaceous-tuffitic layers were penetrated by several boreholes above

the folded flysch and they mark a continuing subduction. In some localities, first o!

all in northeast Hungary and in the Soviet Transcarpathians, the folded flysch may

occur beneath the volcanoes, probably in consequence of the trench migration

toward the subduced slab during the development of a mature accretionary wedge.

The upper age boundaqp ofthe f2y.d is a very important matter. Ten years ago it

was already obvious that the Upper Cretaceous and Eocene sediments of the

discussed belt belonged to the flysch facies, and the Oligocene sediments have also

67

recently been identified as flysch (Baiazs et al., 1980). Thus, the temporal gap

between the oldest talc-alkalic volcanites and the youngest ffysch sediments is

insignificant.

Concerning the structural polarity, only indirect data exist. According to

Szepeshazy’s (1973) results, the boundary between the epicontinental and flysch

sediments undulates (Fig. 5): it swells out towards the northwest on the elevations

and towards the southeast in the depressions. As the denudation is surely deeper on

the elevations, this undulation gives evidence of the flysch overthrusting towards the

southeast, i.e. of the southeastern polarity.

Regarding the structural zonality along the southern border of the flysch belt.

there are few available data (Fig. 6). Five main zones can be outlined from the

northwest toward the southeast: the Middle Hungarian-Transcarpathian volcanic

belt (an island arc), the Szolnok-Maramures flysch belt (an accretionary wedge). a

crystalline range with fragments of its shallow-marine cover (a ‘marginal anti-

clinorium’), overthrust mafic rocks and flysch sediments (abducted nappes) and a

crystalline area with sh~low-marine sediments (a ‘craton’).

Both the structural polarity and the structural zonality give evidence of a

continent-island arc collision and of the subduction towards the northwest in

accordance with the dip of the Benioff-zone reconstructed from the petrochemical

data. The crustal thickness, however, does not increase in the flysch belt (Szenas,

1965) because a young depression has developed over the whole belt assemblage.

The immediate reason for the sinking is thought to be the crust thinning as a result

of a deep process. The crust thinning may have destroyed the ‘root’ of the

Szolnok-MaramureS flysch belt (in the Northeastern and Eastern Carpathians. the

young depressions are located behind the volcanic belt only, so here the ‘root’ of the

Carpathian flysch belt has not disappeared).

The structural picture of the Szolnok-Maramures flysch belt is disturbed: the

order of the facial zones (Fig. 5) is changed in some places. and the rocks in general

are affected by dynamometamo~hism. The concept of the second-rate role of the

Szolnok~Maramure~ flysch belt compared with the Carpathian one (SzCnas. 1965:

Juhasz. 1970: Szepeshazy, 1973; Kiiriissy, 1977) can no longer be supported. A

significant problem, however, arises: whereas the Middle Hungarian-Trans-

carpathian volcanic belt strikes as far as south of Lake Balaton, the western ending

of the flysch belt seemed to be somewhere at the river Tisza (Fig.6). I think the

following data mark the western continuation of the flysch belt (Fig. 6): (a) the

presence of the analogues of the Senonian pelagic red G~obotrMn~ana marls in the

Kerekegyhhza-5 borehole (Turonian stage: Side. 1969) and near Vckeny

(Cenomanian stage: Side, 1961); (b) the continuation of the Lower Cretaceous mafic

volcanites associated with the flysch belt which are unbroken, according to the

geomagnetic anomalies, and partly located in the nappe-zone along its southern

border. Consequently, the continent-island-arc collision suture is elongated as far as

the Mecsek Hills.

6X

69

LL

0 IO q 3lkm

Fig. 4. Ancient Benioff-zone isopleths in the central section of the Middle Hungarian-Transcarpathian

volcanic belt. Depths in km are calculated from JuhLz’s (1971) data using Ninkovich and Hays’s (1972)

plots. Analyses with volatile-content over 4% or CO, over 0.5% are neglected. Fractions near borehole

groups: first number in the numerator=average depth in km, second one=standard deviation in km; in

the denominator = number of analyses.

As is consistent with the Biikkalja-Tokaj-PreSou volcanic arc (Fig. 2, 5-7)

convexity, a westward-directed subduction is assumed there. If we ignore the data

from the extended tuff and lava horizons because of the uncertainty in their source

locality, petrochemical data (Panto, 1966; Gyarmati, 1977) from the independent

central bodies (extrusive domes, subvolcanoes, etc.), show a westward dipping

ancient Benioff-zone (Fig. 7A). The isopleths, however, form a considerable angle

with the arc strike. The deviation is eliminated on the basis of the same data,

assuming left-lateral faults (Fig. 7B). These assumed faults are well recognized in

geological maps. The collisional suture marked by the volcanic arc must be located

71

Geological data

Tectonic scheme

Fig. 6. Structural zonality of the Szolnok-Maramureq flysch belt in Hungary. Geological data after K.

Szepeshitzy (1971. 1973, 1979); geomagnetic data after I. KomPromy and I. Hati (1966).

Geomagnetic anomalies

100 * ( 200 3Cokm

0 Sketch of the overthrusts

in front of the arc, perhaps near it, in the basinal territories of East Slovakia and the

Soviet Transcarpathians.

The plate boundaries are delineated as follows. The Cilimani-Harghita, the

Vihorlat-Gutin, the B~kkalja--Tokaj-Pre~ov and the Middle Hungarian-Trans-

carpathian volcano-chains indicate four collisional sutures. For the plate movements,

synchronous transcurrent boundaries must be assumed. On the basis of the data

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75

from the East Carpathian volcanic belt, two microplates (Fig. 8A, B) are outlined;

according to the data from the Inner Carpathian volcanic area, the northern one

(Fig. 8B) is also divided into two microplates (Fig. SC, D). The Middle Pannonian

microplate (Fig. SC) forms the basement of an island arc, i.e. of the Middle

Hungarian-Transcarpathian volcanic belt. Only the North Pannonian (Fig. 8D) and

the South Pannonian (Fig. 8A) microplates can be regarded as true (indivisible)

microcontinents.

The kinematic reconstruction is only outlined here. The main task is to restore the

situation before the movements. Before the subduction, basins at least 200 km wide

must be assumed along the present collisional sutures. with their basements capable

of subduction; that is, these basins must have a (quasi) oceanic crust. Going back in

time, first of all the southern (South Pannonian), then the northern (Fig. SB)

microplate must be moved at least 200 km towards the west. At the same time as

these movements and partly going further back in time, the North Pannonian

microplate must be moved at least 200 km toward the west and the South Pannonian

one towards the south. In this way, the present Pannonian continental lithospheric

units will have been scattered towards the west, southwest and south by the

beginning of the Neogene (Fig. 9). The fundamental condition of this process is the

existence of space to allow the elements of the Pannonian continental lithospheric

mosaic to be scattered in these directions.

The Carpathian region is now surrounded by the continental lithosphere and is

practically closed. The present seismicity marks the only active section of the

continental frame along the Adriatic Sea, just where an ‘exit’ has to be found for the

Pannonian microplates on the basis of the outlined kinematic scheme. The following

hypothesis is made to open this ‘exit’: the Dinaride unit reached its present position

only after the Pannonian lithospheric mosaic had assembled.

At the beginning of the Neogene, a basin of (quasi) oceanic crust occupied the

present Pannonian basin territory or its largest part; the North Pannonian micro-

plate was located on the southwestern margin of this basin and the South Pannonian

one on its southern margin (Fig. 9).

Cretaceous-Paleogene

As the magmatism appeared in the Carpathian region intermittently both in space

and time during the Cretaceous and Paleogene, it cannot be employed as a basis for

comprehensive geodynamic analysis. According to the distribution of two character-

istic groups of sediments, bodies of two sorts need to be delineated: flysch belts and

the continental units between them. The largest flysch belt in the Carpathian region

strikes along the Carpathians, wholly beyond the Pannonian basin. For this reason,

attention will be concentrated not on this belt but on the Szolnok-Maramures belt

crossing the Pannonian basin from the southwest toward the northeast (Fig. 6).

I

c t

I

c P R M N S Fig. 9.

contrnenlol unit Bchemioo massif

QJaSl/ oceanic area Polish plate

* Russm” platform

/micro/ plate margm, c&am

assumed contour ot the Middte

MoesaIl plate .rr- Pannonian mlcroplate

North Pannonian contour d Ihe Carpathion

microplafe fold system

b,,th PWltWniCI” microplate

/-- ;pdl:;;~pJ;; raps on

Position of Pannonian continental microplates at the beginning of the Neogene.

The Szoinok-Maramure~ flysch belt According to Szepesh.Bzy’s (1973) data, several facies types (Fig. 5) occur in the

Szolnok-Maramure~ flysch heft. Among them, the shallow-marine origin of the coarse detritai and sporadicafiy calcareous sediments (type 1) is defnite. Ind~e~dent qualification of the dominantly arenaceous sediments (sporadicaliy with graded bedding or pebbles, type 2) cannot be made, although their transitional position seems to be definite. The arenaceous-argillaceous sediments (type 3) belong to the

turbidites; intermediate and basic rock derivatives are very subordinate in the

detrital material, and granitic and metamorphic rocks and minerals dominate.

Characteristics of the calcareous-argillaceous sediments (type 4) are: the red colour,

the large amount of planktonic foraminifera and the normal bedding. They are

typical pelagic sediments accumulated in the facies succession further from the

shoreline than the turbidites. The pelagic facies is known from the Cenomanian until

the Lower Eocene, consequently, it was formed over a long time-span (at least, 50

m.y.). The turbiditic sediments developed from the beginning of the Senonian.

probably from the Albian (boreholes Soltvadkert-5 and -7: Szepeshazy. 1971: see

Fig. 6), until the end of the Oligocene, over a period not of 50 m.y. but rather of

70-80 m.y.

The turbiditic sedimentation over a long time-span shows the persistency of a

great undersea slope (i.e. the continental slope), and the similarly long pelagic

sedimentation marks the comparatively large dimensions of the basin (probably, of

oceanic crust). That is why I assume flysch sedimentation at a passive continental

margin of Atlantic type, in accordance with the lack, or the very subordinate role, of

volcanic material. From the viewpoint of deformations in accretionary wedges, no

difference exists between the sediments of the trench itself (Seeley et al., 1974) and

sediments that have arrived at the trench from the outer slope, i.e. together with the

subducing slab.

On the basis of its position, this basin is named Inner Curpathian basin. The

Szolnok-Maramureg flysch belt together with its western continuation is to be

regarded, as a closing suture. I believe that the Carpathian flysch belt (see Fig. 1) is

another collisional suture of the same basin: within its sediments both the pelagic

and the turbiditic facies are present; again practically without volcanogenic material.

and the sedimentation took place in the same Early Cretaceous to Oligocene

time-interval. The continental lithospheric mosaic assembled in the Neogene has

covered the original territory of the Inner Carpathian basin, and the sediments of

this basin are now observable only as strips along the former microcontinent

contacts. These strips have been considered for a long time as ‘flysch-trenches’.

In the Szolnok-Maramureg flysch belt, two petrochemical types of basic volcanites

are wide-spread, namely alkalic and tholeiitic basalts (Onuoha, 1979). Age de-

terminations exist from the alkalic basalts only: in the Mecsek Hills (see Fig. 6), they

alternate with shallow-marine sediments of Valanginian to Hauterivian age (Bilik,

1974). Because of the petrographical similarity of the two basalt types, all basic

volcanites are considered to be Lower Cretaceous (Juhasz and Vass, 1974: Szepeshazy,

1977).

The South Pannonian microplate

Only the Hungarian part of the South Pannonian Microplate (see Fig. 8A) will be

discussed. The Oligocene is absent here, older sediments occur in several transgres-

siue cycles. The Middle-Upper Eocene shallow-marine deposits (type 1 in Fig. 5)

.

. . . . . con(ew d the Emre . . . .: rcdimcntr “VestigQw l

decp‘wolel zcmnc*ts D shu1Iuu-molkw ScllllnPnts

Fig. 10. Cretaceous and Eocene sediments of the Szolnok-Maramureg ffysch belt and of the South

Pannonian microplate (Hungarian parts; after Kaszap. 1963; FiilBp, 1966; Sid6, 1961. 1969: Szepeshirzy.

1971; 1973).

Kpen

occur only in a strip lo-20 km wide along the southern border of the Szolnok-

Maramureg flysch belt (Fig. IOA). Shallow-marine Upper Cretaceous appears be-

tween the rivers Danube and Tisza (Fig. 10B). The nature of these Upper Cretaceous

and Paleogene sediments ranges from gravels to psammites. The detrital material is

82

dominantly of crystalline origin but in some places Mesozoic sediments also occm

The Barremian to Albian sediments (Fig. 1OC) are represented by limestone and

marl which are probably of shallow-marine origin. The Valanginian to Hauterivian

gravel-psammitic deposits are similarly wide-spread (Fig. 10D).

The present structure of the South Pannonian microplate is complicated by

folding, imbrication and perhaps overthrusting. The Hauterivian sediments in the

Mecsek Hills and the Aptian to Albian sediments in the Villany Hills have been

deformed together with the Upper Permian, Triassic and Jurassic layers. The

relationship between the sporadic Cenomanian sediments and these deformations is

unknown. The Senonian shallow-marine sediments form complicated structures

together with older Mesozoic ones. Several boreholes have found shallow-marine

Eocene east of the river Tisza but only some of them have reached the underlying

sediments (of Tithonian or older age: Szepeshazy, 1973); its structural relationships

with the Cretaceous are unknown. Thus, no direct structural evidence of compressive

deformation of Cretaceous or Eocene exists.

The North Pannonian microplate

Only the Hungarian part of the North Pannonian microplate (see Fig. 8D) will be

discussed. On the basis of the facies and distribution of the Paleogene deposits, three

sections along the present strike are outlined (Fig. 11). The western section includes

the Bakony facies zone (Balazs et al., 1980) with Oligocene in the continental

terrigenous molasse facies. The Eocene shallow-marine terrigenous-calcareous sedi-

ments are Early to Late Eocene in age. Bauxite and coal mark the unbroken

southern shoreline of the sedimentary basin and show the absence of significant

transport of detrital material from the adjacent mainland. That part of the sedimen-

tary basin which could be investigated was of complicated and varied topography

with islands and small depressions (Kopek et al., 1970), but no lithological evidence

of the northern shoreline exists, and the so-called Raba-line (Fig. IlB) limits the

Eocene sediment distribution area as a later fault. At the same time, a mainland with

strong denudation is thought to have existed in this direction, being the source of the

terrigenous material. In addition, the Eocene fauna of the Bakony facies zone shows

distinct pelagic influences (Szots. 1956; Kecskemtti, 1973, 1978). In the Eocene, this

zone seems to have been located on the southern shelf of a large basin, but between

this basin and the Bakony shelf, which was of archipelago type, an island chain must

have been situated. It was built up of crystalline rocks and was steep and high

topographically. This island chain is assumed to have been able to provide a source

of the terrigenous material in the north and a possibility for the pelagic connections

in the same direction.

The middle section includes the Buda facies zone (Balazs et al., 1980) with

calcareous-terrigenous marine sediments of Oligocene and Upper Eocene age. A

clear paleogeographic regularity exists here: the terrigenous material has originated

from the north, the transgressions have developed to the north and interruptions in

83

the stratigraphic sequences have disappeared southwards. In the south, the sedi-

ments now disappear along a line without any lithological traces of the shoreline.

Consequently, the Buda facies zone is a remnant of the northern shelf of a large

marine basin, and later tectonic movements have torn this zone from the main part

of the basin. The Upper Eocene and Oligocene sediments are known also in a

narrow strip being located south of Lake Balaton. This strip is supposed by BalBzs et

al. (1980) to be the western continuation of the Buda facies zone. The Paleogene

sediments are imbricated (Fig. 1lC) along the continuation of the tectonic line

limiting the Buda facies zone itself from the south. This tectonic line can be

identified with the so-called ‘Balaton-line’.

In the eastern section, Paleogene deposits are absent.

The Bakony and the Buda facies zones originated in the marginal parts of ~M’O

different Paleogene basins: originally the Bakony facies zone lay on the southern shelf

of a large basin and the Buda one on the northern shelf of another basin situated

south of the first one. According to their positions, they are named northern and

southern basins respectively.

The present relationships between the facies zones, as well as between the sections,

are not clear. It is possible that the sections are independent lithospheric units

assembled in the Miocene only. Because of this I shall discuss the Cretaceous

deposits separately in three sections.

In the western section, the Senonian (Haas, 1979) the Albian (Ftilop, 1975:

Csaszar, 1978) and the Aptian (Ftilop, 1964, 1975) paleogeography is similar to that

of the Eocene: bauxite and coal mark the unbroken and flat southern coast:

marshes, lagoons and a shallow sea with reefs, islands and peninsulas are thought to

have existed along this coast. The terrigenous material originates from the north and

the pelagic connections (Gtczy, 1954; Sidb 1963, 1966, 1974. 1975; Czabalay. 1975.

1976) are similar to the Eocene ones. Thus, the paleogeographic situation seems to

have been permanent over the whole Aptian-Eocene period. Transgressions and

regressions were probably caused by crustal movements of a small (max. 100-200 m)

amplitude and by eustatic oscillations.

A very sharp change is observable in the composition of the detrital material.

During the Eocene and Senonian, it consisted almost exclusively of granitoid,

metamorphic and sedimentary rocks and minerals. However, in the Albian and the

Aptian sediments, the derivatives of basic rocks occur: diabase pebbles, plagioclase,

magnetite, ilmenite, augite, enstatite. Additionally, in the Albian sediments, deriva-

tives of ultramafics, i.e. chromite and leukoxene (85% of heavy minerals: Fulop.

1975), and derivatives of intermediate-acidic volcanites, i.e. a great deal of light

glass-detritus (FoldvBri et al., 1973) are present. Under the permanency of the

paleogeographical conditions, this feature marks a radical change in the proportions

of the rocks in the denudation area. The source of chromite can be specified only as

ultramafics which appeared on the surface from the deep levels of the oceanic

lithosphere, and the intermediate to acidic volcanites must have been derived from a

84

talc-alkalic volcanic zone. The basic rocks could be members of both series. The role

of this rock-association sharply decreased before the Senonian when the crystalline

(and sedimentary?) rocks became predominant in the source area.

The Neocomian (Ftilop, 1964) overlies the Tithonian everywhere with transitions.

The sediments are of pelagic facial types. Their spatial distribution gives no clue to

the direction of either the transport of terrigenous material or the nearest shoreline.

Among the heavy minerals (Ftilop, 1975), ilmenite and magnetite occur accompanied

by enstatite and diopside in the Tithonian, but in view of their small amount (0.2%

of the terrigenous material being contained in 5-6% of the whole rock), these

minerals must have been transported by air.

In the Cretaceous-Paleogene history of the western section two sharp paleogeo-

graphical changes are outlined: at the Neocomian-Aptian and at the Eocene-Oligo-

cene boudaries. The Neocomian seems to be the direct continuation of the Jurassic.

in some places having Jurassic facies types (‘Ammonitico Rosso’, ‘Hierlatz’) in the

lower horizons. During the Aptian-Eocene phase, the western section was lying in

the southern part of an asymmetric trench-like sedimentary basin. The southern

coast of this basin was unbroken and flat, being built up of carbonate rocks. The

northern coast must have lain beyond the present Raba-line cutting the northern

part of this basin, and it seems to have been represented by an active island arc

during the Aptian and the Albian and by a crystalline-sedimentary island chain

during the Senonian and Eocene. In the Oligocene, the whole area became mainland.

and coarse detrital material, probably also of northern origin was accumulated here.

In the structural development, a single sharp change took place. The Neocomian

sediments together with the Jurassic ones have been folded, whereas the Aptian and

younger sediments have only suffered disjunctive dislocations. The folding is of

southern vergence.

In the middle section, the Cretaceous sediments occur sporadically. The Senonian

sediments are known on the southern slope of the Uppony Hills (Balogh, 1964; see

Fig. 11). Their fauna1 connections are similar to those of the western section (Side.

1974). The Neocomian sediments occur in the Gerecse Hills (Ftilbp, 1958; see Fig.

1 1), overlying the Tithonian and older rocks with erosional disconformity and

containing diabase in the detrital material. In the north, the thickness is greater than

300 m, and these sediments are at least partly of turbiditic origin. Toward the south

the thickness decreases rapidly and the detritus becomes finer. Thus, the paleogeo-

graphic situation reminds one of Aptian-Eocene in the western section. No litho-

logical evidence of the southern shoreline, however, exists; this fact shows an inner

position within the trench-like basin in accordance with the occurrence of the

flysch-like sediments.

In the Hungarian part of the eastern section of this microplate, Cretaceous

sediments are, like Paleogene ones, unknown.

6 0

:*+++++++++ + +

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X9

Geodynamic interpretation

At the beginning of the Neogene, the main elements of the Pannonian basins

continental lithosphere were located several hundred kilometres towards the west.

southwest and south from their present position and made room for the Inner

Cur-path&n basin (see p. 80) of Cretaceous-Paleogene age. On the basis of the

lithologic and paleogeographic analysis. a conclusion has been drawn on the (quasi)

oceanic crust of this basin. Accordingly, the following question is to be considered:

are the basic volcanites now associated with the sediments of this basin (see Fig. 6).

components of this crust or not? Because of petrochemical peculiarities. the alkalic

basalts are thought to be products of continental rifting (Bilik. 1980) of

Valanginian-Hauterivian age (Bilik, 1974). The tholeiitic basalts can be related to

this continental rift opening, so they are Barremian or younger in age. With a mean

spreading rate of some cm/yr, the basin of several hundred kilometres width could

be formed during one or two Early Cretaceous stages. so that the sedimentation is

expected to have been extended to the whole basinal territory in the Albian or the

Cenomanian. The age of the flysch belt sediments agrees well with this concept.

The South Pannoniun microplute was located on the southern shelf of the Inner

Carpathian basin. Within this unit, compressive dislocations known certainly to be

of Cretaceous age occur only in Transylvania (see Fig. 1). Here, two independent

magmatic complexes of Upper Jurassic-Neocomian and of Senonian-Paleocene age

mark the cause of dislocations, i.e. the subduction (Herz and Savu. 1974). At the

same time, the South Hungarian Lower Cretaceous magmatic complex (alkalic and

tholeiitic basalts) differs sharply from the Transylvanian ones and is related to the

dilatational tectonics. The dislocations in the South Hungarian Mesozoic can now be

explained only by subduction which resulted in the Inner Carpathian basin closing.

Accordingly, these dislocations seem to be Late Oligocene-Miocene in age.

The differences raise the question as to when the present unity of the microplate

came into being. At the beginning of the Cretaceous. the South Hungarian basic

volcanites and the Transylvanian island arc complex and nappe-system may have

lain on the opposite sides of the same microcontinent and together may have marked

a southerly directed relative motion. Because of the lack of data on the position and

direction of the subduction zone related to the Transylvanian banatites (i.e.

Senonian-Paleocene talc-alkalic magmatites), there is no possibility of deducing the

South Hungarian features which were perhaps connected with this subduction. Thus,

it is more probable that this microplate was assembled only after the Paleocene.

from two or more elements of different origin.

The North Punnoniun microplate was probably located west or southwest of the

Inner Carpathian basin, but their paleogeographic relationships need an intensive

analysis. In the Neocomian, the western section was lying within a pelagic realm. By

the Aptian, this continental lithospheric unit had got in front of an active island arc

in consequence of a northerly-directed subduction. Pyroxenes, magnetite and ilmenite

in the Tithonian and Neocomian sediments must have been derived from this island

50 i km

d km

km

- ~--_--___-.._ ._ ~~_.. ., (

K,sen - EC, r 50

xm

----------TG i

Ec,.~ :

;r? km

0 WOkm

Fig. 13. Development of the western section of the North Pannonian microplate. Note: The upper mantle

also participates in the subduction together with the crust.

arc, so that the subduction started not later than in the Late Jurassic. During the Aptian and the Albian, a (micro) continent-island arc collision took place. The disappearance of the rock derivatives of island arc in the Senonian could have been caused by the emergence of the (micro) continent edge subducted under the island arc, during the postcolhsional thermic and isostatic equalization after the end of the subduction (Fig. 13). This makes the lack of the sediments in the Turonian understandable when, about S-10 m.y. after the end of the collision, the maximum emergence might have been expected. During the Late Cretaceous and the Eocene the inactive island arc was lying on the passive southern border of a marginal sea, having previously been located behind the island arc.

According to the data from the Gerecse Hills (see Fig. 1 l), the western part of the middle section had already got in front of an active island arc by the Neocomian, thus the collision started much earlier here than in the western section. At the same

NW

1 &tit - K,nc

504 km

91

SE

-0

i

K,apt -alb,

i M km

K,alb, -K,cm, [

!-so km

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i 50 km

rate, the collision is also expected to have been finished much earlier, probably by

the beginning of the Aptian. The further development must have been of postcolli-

sional type. The Uppony (see Fig. 11) Senonian deposits dominated by granitic,

metamorphic and sedimentary rocks and minerals in the detrital material were

probably accumulated at that time.

The last marine sediments of the Bakony facies zone (see Fig. 11) were deposited

in the Late Eocene. Synchronously, the sedimentation began in the Buda facies zone.

It can thus be seen that the sediments of the northern basin (see p. 83) disappeared

simultaneously with the appearance of the southern basin sediments. At the same

time, a talc-alkalic volcano-chain (Fig. I LB), probably of subductional origin, was

coming into being around the boundary between these basins, along the whole

western and middle sections. The lithosphere of the basin, with sediment disappear-

ing in this time, is assumed to have been consumed during this subduction.

Therefore, the northern basin was closed and the ancient Benioff-zone dipped from

92

Fig. 14. Structural sketch of northwestern Hungary based on geomagnetic anomalies. Geomagnetic

anomalies after I. Kom&romy and I. Ha& (1966X

this basin toward the volcano-chain, i.e. toward the south. The newly-formed

southern basin lay in the dipping direction of the Benioff-zone, i.e. at the back of the

volcano-chain. Because of its position, this basin qualifies as a typical marginal sea.

The Oligocene continental molasse seems to mark the postcollisional emergence.

In the Cretaceous-Paleogene history of the North Pannonian microplate, one

collision of Early Cretaceous and another one of late Eocene age and a postcolli-

sional emergence can be recognized. Both collisional sutures are to be expected along

the northern rim of the microplate but the Benioff-zones dipped in the opposite

directions. The volcanic belt related to the second Benioff-zone is located on the

microplate itself; the island arc marking the first Benioff-zone has been deduced on

the basis of the lithological data only.

The Early Cretaceous island arc must have been strongly deformed both by the

granite-gneiss doming during the post-collisional equalization (see Zonenshain et

al., 1976) and the contrary-directed Late Eocene collision. Thus, in the present state,

only its remnants exist, Beyond the Raba-line (Fig. 14) some complexes of an age

and composition that correspond to such remnants are known. These are the

analogues of the Alpine Penninics in the K6szeg Hills (ultramafics, gabbro and

diabases from the basement of the island arc and metamorphosed graywacke and

flysch from the cover) and their northeastern continuation in the Little Hungarian

Plain (see Fig. 1) basement, both of which are marked by geomagnetic anomalies

(Fig. 13). At the back. the gneiss and the crystalline schist of the Sopron Hills (Fig.

13) were related to the Lower East Alpine nappe-system (Wein, 1969; Dank and

Bodzay. 1971). In this case, the Penninics should be expected to be just a few

kilometres beneath them. The total lack of geomagnetic anomalies in the Sopron

area, however, shows that the Penninics of KGszeg type could exist only at a

minimum depth of about 12-15 km. On the basis of this fact and of the position

beyond the remnants of an island arc, the crystalline rocks of the Sopron area are

rather a member of a granite-gneiss doming zone. i.e. the “Zentralgneiss” is

probably their Alpine analogue.

As is seen, beyond the Riba-line real complexes can now be recognized which

had earlier only been assumed on the basis of the paleogeographic and geodynamic

analysis. During the Late Eocene collision, the R&ba-line acted as an overthrust

zone. Along this zone, a significant shortening took place which probably resulted in

the disappearance of the large northern part of the Aptian to Eocene depressions of

the western section of the North Pannonian microplate.

So far as the assemblage time of the microplate is concerned, no definite data

exist. The differences between the Early Cretaceous developments of the western and

the middle sections show that in these places the same collision was taking place at

different times. The Upper Eocene volcanic belt binds the western and the middle

sections (see Fig. 11B). Accordingly. the present arrangement of these sections

perhaps began to form only in the second half of the Eocene. In contrast, the eastern

section could join after the Oligocene,

Historical synthesis

The analysis given above leads to the conclusion that the history of the Panilonian

basin cannot be characterized as an indivisible process because its basement is

represented by a continental lithospheric mosaic which assembled and reached the

present position only in the Neogene. Until the end of the Oligocene, the history of

the main units can only be outlined (Fig. 12, see p. 87-88) and a complete picture

can only be constructed for the later events.

Until the end of the Jurassic, the South Pannoniun microplate had belonged to a

large continental entity. At the beginning of the Early Cretaceou& a continental rift

occurred within this entity. In the second half of the Early Cretaceous, this rift was

being opened and transformed into the Inner Carpathian basin of (quasi) oceanic

crust. Through the opening, the South Pannonian microplate became a part of the

southern shelf of this basin and remained there until the end of the Oligocene.

At the end of the Jurassic, the North Pannoniun microplate belonged to another

large continental entity and was located on the passive southern margin of an

94

oceanic basin. The oceanic lithosphere subducted in the northern direction under an

island arc. Beyond this arc, the northern basin of oceanic crust and of marginal-sea

type was located. The subduction resulted in the collision which occurred partly in

the Neocomian (middle section) and partly in the Aptian and Albian (western

section) followed by postcollisional emergence. In this way, the remnants of the

former island arc became the southern margin of the northern basin, and a changing

linear depression of foredeep type came into being on the remaining free margin of

the subducted (micro) continent. The northern basin was being closed in the second

half of the Eocene through a southerly directed subduction (opposite to the former

one). Along the rim of the previous foredeep-like depression, a volcanic belt came

into being. Behind this belt, the southern basin opened fike a marginal sea. This

basin separated the ~cr~ontinent from the large southern continental entity.

The Late Jurassic and Early Cretaceous development of two microcontinents was

different. During the Late Cretaceous and Eocene, however, their paleogeographic

and geodynamic positions were similar: both were located on the southern passive

margins of (quasi) oceanic basins. Thus, the northern basin is assumed to have been

located in the western continuation of the Inner Carpathian basin (see p. 80), and

both rnicroplates are most likely to have been on two sections, of different origin, of

the same passive continental margin. In the second half of the Eocene, the North

Pannonian microplate must have been north of its position in consequence of

processes taking place in the west. Thus, the southern basin being opened behind

this microplate could have become the western continuation of the Inner Carpathian

basin. In this way, two microplates could be on the opposite shores of the same

(quasi) oceanic basin.

The closing of this basin began along the east-west directed subduction zone. The

connected island arc divided the basin into two parts. The southwestern one was

closed by the collision of the South Pannonian microplate with the island arc. The

northern part was consumed in the following way. The North Pannonian microplate

was torn from the southern basin and was displaced along a transcurrent fault

behind the island arc (i.e. the Middle Pannonian microplate). It then moved with

this unit to collide with the European plate margin in the east. The South Pannonian

microplate lagged behind a little after having turned eastwards along the western

and northern border of the Moesian plate (see Figs. 1 and 9) and also moved up to

collide with the European plate. in this way, the Inner Carpathian (quasi)oceanic

basin disappeared and the continuous continental lithosphere of the present Pannon-

ian basin was formed.

A CHECK: COMPARISON WFTH INDEPENDENT DATA

The results of the geodynamic analysis given above will be considered as

constituting a working hypothesis which should be verified by further investigations.

9s

For our present purposes, further investigations can be substituted by data not

employed in the analysis. Two groups of such data exist, namely results from the

previous geotectonic syntheses and results of paleomagnetic measurements.

Comparison with geotectonic syntheses

In the last decade, Hungarian experts on tectonics established that the well-known

structural zonality of southwest-northeast direction is caused primarily by the

substantially different origin of the northwestern and southeastern parts of the

country. Investigation of both pre-Mesozoic (Szadeczky-Kardoss, 1976b; Jantsky,

1979) and Upper Permian to Mesozoic complexes (Szepeshby, 1975, 1978, 1979.

1980; Channel1 and Horvath, 1976; Wein, 1978a, b; Channel1 et al.. 1979) has

supported this idea. In view of this, the separation of both the South Pannonian and

the North Pannonian microplates can be regarded as being proved.

Geczy (1973a, b) was probably the first to outline the essential difference in origin

of these microplates. According to him, the present southeastern (my South Pannon-

ian) microplate belonged to the European plate in the Early Jurassic paleobiogeo-

graphical zonality, and the present northwestern (my North Pannonian) microplate

was a part of the African plate. During further investigations, only one alternative

has appeared: the present northwestern microplate belonged to an independent

continental unit within the Tethys ocean (V&i%, 1977). My results suggest that the

South Pannonian microplate was a part of a northern continental unit up to the

middle of the Early Cretaceous, and the North Pannonian one belonged to a

southern continental unit until the Late Eocene-which is in agreement with both

possibilities.

The third (my Middle Pannonian) microplate was previously not considered to

be an independent lithospheric unit, but was assumed to be a tectonic zone of great

significance and of high mobility (Wein, 1969; Dank and Bodzay, 1971; Bodzay.

1975, 1977; Channel1 and Horvath, 1976; Szepeshazy, 1977, 1979, 1980; Channel1 et

al., 1979). Therefore, the analysis given above brings only a little novelty in the

tectonic division. This novelty is connected with the third microplate: it is delineated

as a unit having the same role as the other two, and its contours are changed in the

northeast. In the previous syntheses, the Middle Transdanubian (south of the

Balaton-line, see Fig. 11) structural-facial zone was followed in the Btikk Hills (see

Fig. 1 l), and the Zemplin Hills (see Fig. 11) were thought to belong to the

southeastern microplate. Both concepts were based on the composition of the

Triassic and/or older complexes without taking into account the real structure of the

Btikk and the Zemplin Hills and the tectonics of younger sequences.

The continuation of the Triassic and Paleozoic complexes of the Btikk area in the

Middle Transdanubian was never proved but was always assumed. In this connec-

tion, it may be sufficient to say that Tomor (1957) was the last to review the Middle

Transdanubian drilling data. The section of this zone near Budapest (see Fig. 11) can

really be of ‘Btikkian type’ (Btrczi-Makk, 1978; Szalay et al.. 197g), but this section

belongs in fact to the Paleogene Buda facies zone (Fig. 11). i.e. to the North

Pannonian microplate, just as the Bukk Hills do.

There is undoubtedly a great degree of similarity between the Triassic. Permian

and older complexes of the Zemplin Hills and both the Transylvanian (see Fig. 1 J

and South Transdanubian (see Fig. 1) complexes. This fact, however, grves evidence

only of their common origin but not of their direct connections in the present

structure. The Szolnok-Maramures flysch belt (see Fig. 6) of the Cretaceous-Paleo-

gene age lies between the Zemplin Hills and Transylvania or South Transdanubia

and proves the lack of structural connections. This leads to the question of how the

Zemplin unit of ‘European’ origin reached the eastern end of the North Pannonian

microplate which is of ‘African’ or ‘Intratethian’ origin at least in its western part.

As stated above, in the eastern section of this microplate the Cretaceous and

Paleogene are absent; thus this section may have joined the western ones at the end

of the Oligocene.

Summarizing, good correspondence with the former geotectonic syntheses exists

in the determination of the microplate contours and origin.

Comparison with paleomagnetic data

Paleomagnetic investigations have given the following results concerning the

complexes of the time interval considered (Marton, 1980; Marton and Marton,

1980a, b):

(1) The Bakony Hills (see Fig. 11) on the western section of the North Pannonian

microplate belonged to the African lithospheric slab until the Senonian. Since the

Miocene (after the Paleogene without paleomagnetic data), the middle and the

eastern sections of the same microplate have belonged to the European plate.

(2) The Mecsek Hills (see Fig. 6 or Fig. 10) in the western part of the South

Pannonian microplate belonged to the European plate until the Neocomian and

reached their present position by later displacements.

All conclusions made on the basis of the geological analysis, are in good

agreement with the consequences from the paleomagnetic data and determine more

exactly the geological age and the geodynamic cause of the changes (see Fig. 14). The

Bakony unit was torn from the African plate as a result of the southern basin

opening around the Eocene-Oligocene boundary, and it does not have any connec-

tions with the European plate because the northern basin was between them. It has

joined this plate by the consumption of this basin at the end of the Eocene and got

to its present position by lateral displacement towards the east. The Mecsek unit was

torn from the European plate as a result of the Inner Carpathian basin opening in

the Early Cretaceous and it joined the same plate again, by the consumption of this

basin, only in the Miocene. These correspondences support the kinematic conclu-

sions.

97

CONCLUSIONS APPLICABLE TO SURROUNDING REGIONS

The geodynamic analysis given above not only explains the Cretaceous-Cenozoic

history of the Pannonian basin basement but also contains conclusions which can be

employed in the surrounding regions. The final result of the geodynamic analysis is

that the continental lithosphere of the Pannonian basin assembled and reached its

present position only in the Neogene, and that a basin of (quasi) oceanic crust (the

Inner Carpathian basin) formerly existed in this region. This result and the outlined

history of two main microplates influence the concepts on the development of the

Dinarides, Carpathians and Alps as follows:

(1) The Dinurides could get to their present position relative to the Carpathian

region only in the Neogene.

(2) The folding in the Northeastern and Eastern Carputhiuns is a result of the

subduction being marked by the East Carpathian volcanic belt (Fig. 1). but behind

the Western Carpathians a large left-lateral displacement took place synchronously

with the same subduction.

(3) Some well-known stratigraphic and structural correlations exist between the

A/ps and the western section of the North Pannonian microplate (see Figs. XD and

11). The three main phases of the tectonic movements (at the end of the Early

Cretaceous. at the end of the Eocene and in the Miocene) coincide in both regions.

In the light of the Transdanubian data, a change in the subduction direction from

northerly in the Early Cretaceous to southerly in the Eocene can also be assumed for

the Eastern Alps. Substantial differences between the Eastern Alps and northern

Transdanubia seem to have occurred only in the Miocene: while the first unit moved

toward the north, the second one moved toward the east. This feature caused strong

structural differences: in the Eastern Alps the meridional collision and connected

dislocation developed further, while in the western section of the North Pannonian

microplate. the change in motion of about 90” compared with the former one, and

the fact that subduction took place only at the far eastern border of the microplate,

preserved the pre-Miocene state which in the Eastern Alps can be reconstructed

mainly from indirect data.

ACKNOWLEDGEMENTS

This research was supported by the Central Geological Office of Hungary and is

published with its permission. The theoretical aspects of the geodynamic analysis

were elaborated through very helpful discussions, first of all with L.P. Zonenshain

and then with 1.0. Murdmaa, L.A. Savostin. L.I. Lobkovsky (P.P. Shirshov Institute

of Oceanology of the USSR Academy of Sciences). Many of the scientists of the

Hungarian Geological Institute and the L. Eotvos Hungarian Geophysical Institute

helped in the interpretation of the relevant data.

98

REFERENCES

AfvLne SO& E. and Ravasz. Cs., 1978. A komloi andezit K-Ar kora (K-Ar dating of the andesite ot

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