Sequence stratigraphy and paleogene tectonic evolution of the Transylvanian Basin (Romania, eastern Europe)

Sedimentary Geology ELSEVIER

Sedimentary

Geology

105 (1996) 117-140

Sequence stratigraphy and Paleogene tectonic evolution of the Transylvanian Basin (Romania, eastern Europe) Jean-No8

Proust a,*, Alexandru How b

(’Sciences de la Terre, Universiti de Lille I, URA CNRS 719 Stfdimentologie et GPodynamique, 59655 Villeneuve d’Ascq Cedex, France ’Catedra de Giologie-Mineralogie,

Universitatea

Received

Babes Bolyai; St,: Kogalniceanu

13 July 1994; accepted

31 October

I, 3400 Cluj Napoca, Romania

1995

Abstract The Transylvanian Basin of Romania belongs to the 800 x 400 km wide Pannonian domain of the European Alpine megasuture bordered to the east and the north by the Carpathians. It represents a digitation of the epicontinental Tethyan seaways locally connected during the Palaeogene to the peripheral foredeep troughs. During that time, it was filled up by a 500-m-thick sediment pile organized into three shallow marine and non-marine facies alternations. Each evolved from alluvial fans to restricted marine and outer marine environments. They are dated from Lutetian to Chattian times. The study is focused on the lowermost alternation onlapping the basal, post-Maastrichtian unconformity. This alternation consists of the superimposition of a thick retrogradational and a thin progradational depositional system. The retrogradational depositional system grades upwards from stacked, fault-controlled deposits of alluvial fan, ephemeral stream, salina and sabkha, and restricted marine bioclastic shales. The progradational depositional system is composed of outer marine to estuarine sandstones and shales. The two depositional systems are bounded at their tops by two baselevel change unconformities underlain by highly mobile, low relief sandstone bodies that were deposited in shoal belts. These two unconformities mark significant changes in the regime of the subsidence. These are, respectively, a baselevel rise or ‘drowning’ unconformity where the shoal deposits were associated with oolitic ironstones and glauconitic shales that typify basin starvation during a period of maximum basin drowning, and a baselevel fall or ‘uplifting’ unconformity where the bioclastic shoal deposits were buried by alluvial flood plain deposits that characterize periods of relief rejuvenation tentatively attributed to compressive events. This bimodal succession is interpreted in terms of underfilled~verhlled stages related to intraplate tectonic deformation. The underfilled stage corresponds to the fault-controlled drowning of the basin and the overfilled stage to the increasing flexural rigidity of the substrate culminating in differential uplift in the area. Three of these successions comprise the Palaeogene sedimentary record of the Transylvanian Basin. They attest to a large-scale pulsating evolution of this continental microplate during its northward migration towards the European plate.

1. Introduction The general understanding of the origin and the evolution of flexural basins has made a lot of ad* Corresponding

author.

0037-0738/96/$15.00 Copyright ssDI0037-0738(95)00144-1

vances in the last ten years. This progress mostly concerns the following: (1) the semi-quantitative modelling, based on geologic data, of viscoelastic or elastic continental lithosphere deflection under the load of a nearby fold belt (Beaumont, 1981; Jordan, 1982; Quinlan and Beaumont, 1984; Stockmal

0 1996 Elsevier Science B.V. All rights reserved.

II8

J.-N. Proust, A. Ho.w/Sedirnentary

et al., 1986; Beaumont et al., 1988; Flemings and Jordan, 1989, 1990; Sinclair et al., 1991; Stockmal et al., 1992); and (2) qualitative interpretations of geological data tending to improve the modelling hypotheses (Jacobi, 1981; Quinlan and Beaumont, 1984; Allen and Homewood, 1986; Pfiffner, 1986; Tankard, 1986; Heller et al., 1988; Cant and Stockmal, 1989; Stockmal et al., 1992). These basins are classically filled by deep marine ‘Flysch’ sediments shoaling upward in shallow marine and non-marine ‘Molasse’ deposits (Covey, 1986; Heller et al., 1988; Homewood and Lateltin, 1988; Caron et al., 1989; Allen et al., 1991; Cant and Stockmal, 1993). Commonly, the ‘Flysch’ or ‘underfilled’ stage corresponds to increasing flexural subsidence caused by the acceleration of the thrust load in the adjacent fold-thrust belt. The deep marine, turbiditic facies commonly lies unconformably atop passive margin carbonates along a surface that reflects the initial time lag between the onset of flexural subsidence and the subsequent higher rate of sedimentation. The ‘Molasse’ or ‘overfilled’ stage marks significant decreasing flexural subsidence and isostatic uplift related to the deceleration of thrusting. It exhibits marine to non-marine facies alternations possibly controlled by the migrations of the forebulge associated with overthrusting events in the adjacent fold belt. During this last stage of evolution, in order to maintain the equilibrium of the thrust wedge, basement-involved thrusts occur at the back of the wedge. Another basin may appear behind the main thrust front with a similar sedimentary and flexural evolution and may commonly experience extensive subaerial erosion during isostatic uplift. Such a retroarc foredeep basin has been already described in the Palaeogene North Hungarian basins, which evolved at the back of the compressional Palaeogene arc of the Carpathians (Sztano and Tari, 1993; Tari et al., 1993). The sedimentologic investigations are here applied to an extensive, young, simple, Palaeogene basin from Transylvania (Pannonian Alpine Megasuture, Romania, eastern Europe) filled by three, recurrent, well-dated and well-diversified marine and non-marine facies alternations onlapping the Late Cretaceous carbonates. Our objective is to establish the sequence stratigraphic architecture of the first non-marine to marine facies alternation and to com-

Geology 105 (1996) 117-140

pare it to the elementary underfilled-overfilled cycles associated with the presumed flexural evolution of the Transylvanian Basin. 2. Geological

setting

The Transylvanian Basin of Romania belongs to the 800 x 400 km-wide Pannonian Megasuture of the Alpine mountain belts (Fig. 1) that resulted from the collision and the dextral escape migration of numerous microplates located between Europe and the Dinaric (African) area. It can be viewed as the vertical superposition of Neogene back-arc basins located on a crust thinned by extensional processes (sediments >7 km thick) (Debelmas et al., 1980; Royden and Horvath, 1988; Nagymarosy, 1990; Csontos et al., 1991, 1992; Einsele, 1992) over poorly known Palaeogene basins locally floored by thick continental crust (restricted areas of sedimentation >2 km thick) (Tapponnier, 1977; Mitchell, 1986; Royden and Horvath, 1988; Tari, 1992; Tari et al., 1993). Based on the Mesozoic rock record, the Transylvanian Basin is made up of the juxtaposition of the North Pannonian unit with African affinities, and the South Pannonian unit (Tisza and Dacides subunits) with North Tethyan affinities, along the MidHungarian wrench-tectonic line (Gtczy, 1973; Balla, 1984; Kazmer and Kovacs, 1985). The Transylvanian Basin is bordered by the Apuseni Mountains to the west and by the Carpathian Mountains to the east. The Apuseni represent the Cretaceous ophiolitic sutural belts of the northern Tethyan margin microplates during the tectogenesis of the Alpine orogeny in eastern Central Europe (Fig. 1). The East Carpathians correspond to the intermediate-crust, Neogene fold and thrust belt of the easternmost Carpathian domain. From Eocene to Early Miocene times, sedimentation in the Transylvanian Basin was driven by the suturing of the North and South Pannonian units and by the closure of the Magura flysch trough (Sandulescu, 1988; Csontos et al., 1992). The Transylvanian Basin evolved behind the Magura accretion wedge in a flexural basin, which experienced uplifts, thrusting (Ciulavu et al., 1994), and peripheral wrench tectonic movement along the North and South Transylvanian transform zones (Nagymarosy, 1990) (Figs. 2 and 3). By the end of the Early Miocene (Badenian, Sarmatian, and Pliocene

J.-N.

m

Alpine

Carpathian

Proust,

A. Hosu/Sedimentary

Geology

10.5 (1996)

117-140

119

Arc

Great Hungarian

Apuseni

Transylvanian

East

Fig. 1. Geological location map of the Transylvanian Basin in the Alpine Pannonian basins (a) and E-W cross-section (Einsele, 1992)(b) from the East Carpathian mountain belt across the Transylvanian Basin and the Great Hungarian Plain. Note the differences in thickness of the crust between the Transylvanian Basin and the Great Hungarian Plain.

stages), the Transylvanian Basin was involved in the collisional processes of the East Carpathian margin (Sandulescu, 1984) (Fig. 3). Palaeogene deposits (Fig. 4) are 500 m thick in the northwestern part of the Transylvanian Basin and up to 1700 m thick in the deepest areas, close to the outer Carpathian foredeep and to the North Transylvanian fault zone. They are made up of a complex

interfingering of evaporites, carbonates, shales, and fine- to coarse-grained siliciclastics deposited in the proximal parts of the Transylvanian Basin. They schematically correspond to a crude alternation of three, tens of metres thick, continental to marine sequences, which progressively evolved from alluvial deposits and sabkhas to carbonate shoals and outer marine sediments. These alternations are dated by

J.-N. Proust, A. Hosu/Sedirnentar~

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Carpathlan

Pliocene

105 (1996) 117-140

Tectonic activity

Sedimentary record

N e 0

Geology

foreland

Calcalkaline volcanism

evolution

9 e n e

Miocene Suture of the Magura

North and South Plysch trough.

Pannonian

units and closure

of the

Carpathian internides and gradual shift towards the

Oligocene

P

Uplifls and thwstings

in Transylvania (northern pari of the basin).

Escape tectonic movement, wrench faulting. and rigid block rotations (clockwis

; e

Overthrusts in the accretion wedge of the Magura Flysch Trough (south tc north propagation of the closure).

Eocene

0

g e n e

Paleocene

___ Cretaceous

South

_-

unil formation

Pannonian

f

(Laramlde

phase)

Tisca and Dacides subunits suturation. Mulbple compresslonal events. OveRhrusts. foldings, and strong uplifts in Transylvania (Preluca, Apusenl).

Major unconformity Minor unconformity n

Sediment

Fig. 2. Reconstructed sequence of the main tectonic events in the peri-Transylvanian sector of the Intracarpathian area. Data from G~czy (1973), Balla (1984), Sandulescu (1984, 1988). Csontos et al. (1992), Fodor et al. (1992), Tari et al. (1993), Ciulavu et al. (1994), Sztano (I 994). among others.

nannozones (Meszaros and Moisescu, 1991) (Fig. 4), and successively correspond to the Lutetian, Priabonian, and Rupelian substages. The top of the section includes the non-marine rocks of Chattian age. This study focuses on the 120-m-thick PaleoceneUpper Eocene deposits cropping out in the Cluj and Huedin areas in the northwestern Transylvanian Basin (Figs. 4 and 5). These deposits correspond to the first continental-marine alternation onlapping the Laramide unconformity at the top of the South Pannonian overthrust sheets. 3. Facies and depositional environments The facies are depicted from more than 25 crosssections studied along 30- to 50-km-long transects in the Cluj and Huedin areas (Fig. 5). The facies are grouped in order to describe the most typical and complete successions which characterize the different depositional environments. They are described from the landward to the seaward parts of a single theoretical depositional profile.

3.1. Facies association plain deposits

I: Alluvial fan andjood

This facies association is 15 m thick on average and occurs in four repetitive sequences (base of sections 4 and 5 and top of sections 6 and 7, Fig. 5). Each comprises four facies end-members. 3.1.1. Description Facies 1. Poorly sorted matrix-supported, pebbleboulder conglomerate, 2 or 3 m thick, and several metres in lateral extent (Gm lithofacies; Miall, 1977). Most of the clasts are imbricated and range from less than 1 cm to more than 30 cm. The matrix is made up of a red silty mudstone with a lo-cm-scaled, reverse-graded sole along its erosional basal contact. The internal stratification of the sediment body increases upward and exhibits towards the top, tens of cm thick, fining-upward, pebble-sheets interbedded with laminated sandstones. These matrix-supported conglomerates are interbedded laterally with minor clast-supported conglomerates (Gm lithofacies), pla-

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121

East Europe

,

’

/

Moesia

Marine seties II

100

la~oon.bicclasbc shoals,opeo cartnnate platfwm deposits

km

Red beds II fluviakoastal

East Europe

plain deposits +_____

s T

I

Marine saws lagoon.confin& carbonate platform deposits

Li \ k

Redbeds I alluvial fan.saMha

Substrate South Pannoman

Moesia 100

El

MOLASSE

--b--

II FLYSCH

I

TERRANES

--w

THRUST towards (arrows

A N

*______

plate)

SUBDUCTION (arrows towards

deposits

Fig. 4. Simplified section of the Palaeogene deposits in the northwestern part of the Transylvanian foreland basin. It exhibits three non-marine to marine facies alternations. This paper is focused on the basal one.

km

upper

: C 1

upper plate)

Fig. 3. Simplified sketch exhibiting the northeastern propagation of the East Carpathian terranes during the Late Eocene and the Early Miocene. Note the complex foreland evolution of the Transylvanian Basin from Eocene to Miocene (at the front of post-abduction overthrusts and at the back of an active margin with intermediate crust subduction). Modified from Csontos et al. (1992), Tari et al. (1993)and Balintoni (pers. commun., 1994).

nar cross-bedded conglomerates (Gp lithofacies), massive sandstones (Sm lithofacies), and massive siltstones (Fm lithofacies). Facies 2. Stratified, clast-supported, poorly to moderately sorted, cobble-boulder conglomerates (Gm, Gt, Gp). Horizontal stratification, trough and planar cross-bedded and imbrications are common.

Beds of this facies range from 0.3 to 2.5 m in thickness and typically extend laterally for several tens of metres. They present an erosional base and some crude, fining-upward trends (Fig. 6). Facies 3. Massive, 5 to 10 cm thick, very fine- to fine-grained sandstone (Fl) interbedded with red silty claystones. Individual beds are tabular with a sharp basal contact and extend laterally over hundreds of meters. They are locally sharply truncated by lensshape massive sandstone bodies, and are associated with horizontal, planar-laminated, red silty argilites with small ferruginous and calcareous concretions, current ripples, rootlets and millimetric vertical burrows (Trichichnus) (Fig. 6). Facies 4. Structureless, poorly sorted, blocky or mottled silty red mudstones. Specific horizons are evidenced by rootlets outlined by the grey-greenish discoloured areas. The mudstone matrix may give

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Geology 105 (1996) 117-140

I

5m

Fig. 5. Sections 1-8 in the first marine to non-marine facies alternation in the Palaeogene (Lutetian, Bartonian, early Priabonian) deposits of the Cluj-Huedin area (locations shown on inset map). The 60.m-thick alluvial fan deposits at the lower part of the sequences are not figured. X(J) to left of the stratigraphic columns: X facies association and (J) related facies, described in Section 3 of the text; NP = nannoplankton biochronozones; SB = sequence boundary; RS = ravinement surface; MI3 = maximum flooding surface: Roman numerals only: marine flooding surfaces and corresponding baselevel rise surfaces in the non-marine realm.

rise laterally to nodular dolomite including roottraces. Small polygonal fractures filled with sandy material interpreted as desiccation cracks are present but sparse. The 2- or 4-m-thick, topmost part of the superposed facies succession is composed of a well sorted, clean, grain-supported conglomerate. Pebbles are generally a few centimeters large, well rounded, imbricated, and exhibit trough- or planarcross-bedding tens of centimetres thick. These conglomerates are interbedded with lenticular, horizontally stratified and laminated coarse-grained sandy horizons in a basinward direction. They are interbedded with green shales enriched landward with drift wood (see below).

3.1.2. Interpretation The facies 1 conglomerates are attributed to the deposition by viscous gravity flows in the sense of Lowe (1979, 1982). Some laminated sandstones in their upper parts, and laterally interbedded facies, provide evidence for ephemeral stream reworking and close relationships to ponded waters. This faties is locally reworked to form erosional, lenticular and upward-fining intervening channel-fills associated with imbricated, roughly horizontally stratified and cross-stratified conglomerates that may represent longitudinal gravel bars. However, the coarse grain size, the poor sorting, and the irregularly interbedded laminated facies indicate that deposition occurred primarily during intermittent, ephemeral, and

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Geology 105 (1996) 1I7-140

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b

a

Fig. 6. Facies association 1. Alluvial fan and flood plain deposits. Section 4. Pencil for scale is 14 cm. (a) Facies 2, stratified, clast-supported, poorly to moderately sorted, pebbly conglomerate exhibiting an erosional base and some crude, fining-upward trends. (b) Facies 3, cm-thick, massive fine-grained sandstone (Fl) with horizontal, planar laminations, and small-scale current ripple cross-bedding.

turbulent flooding events. These sandstone bodies which may be interpreted as laminated sand sheets in fine-grained, massive, tabular muddy sandstones representing waning flood overbank fines, leads one to interpret the red mudstones as flood plain deposits with intervening pedogenic horizons such as caliches. This facies association corresponds to one of the four elementary events that occurred during a time of gradual base level rise as shown by the decreasing-upward erosional power of the streams and the fining-upward trend (Proust and Deynoux, 1994). The topmost part of the four superposed faties associations are interpreted as the marine edge of the alluvial fan (fan delta) in which the water dynamics or debris flow displacement in a standing body of water has removed the muddy and sandy matrix of the initial massive debris flows. 3.2. Facies association 2: Supratidal hypersaline lagoon deposits

sabkha and

This facies association is made up of non-cyclic, a few-decimeters-scaled successions of two or three of

the five facies. Almost all combinations between the five facies can be observed in the field. They interfmger with facies association 1, up-section and also towards the southeast in sections 5 and 8 (Fig. 5). 3.2.1. Description Facies 1. Thinly laminated green shales and silts with leguminaceous leaves. Facies 2. Dark-grey massive creamy dolomite with intercalated nodular gypsum. In thin section, the nodular gypsum exhibits prismatic-hemi/bipyramidal shaped monocrystals. Facies 3. Laminated anhydrite made up of couplets of anhydrite layers and dark-grey semi-opaque dolomite. The lamination is sometimes irregular (flaser-like) owing to the occurrence of fine anhydrite lenses commonly replaced by alabaster. Slumps are common and form benches of several metres thick, separated by thinly laminated green shales or black dolomicrites. The latter contain small, monospecific coccoliths (NP15 biochronozone) but entirely lack benthic micro- and macro-organisms. Facies 4. Nodular anhydrite with chicken-wire

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J.-N. Proust, A. Hosu/Sedimentary Geology 105 (1996) 117-140

Fig. 7. Facies association 2. Supratidal sabkhra and lagoon deposits. Unquoted section located between sections 5 and 8. Lens c:ap for matrix. scale (6 cm in diameter). Facies 4, nodular anhydrite with chicken-wire or chicken-mesh structures in a dolo-mudstone

structure in a dolo-mudstone matrix (Fig. 7). Relics of anhydrite are usually replaced without pseudomorphism by microgranular gypsum. Facies 5. Alternations of normally graded beds, 20 cm thick, planar-laminated, pebbly calcareous sandstones with subangular intraclasts and lithoclasts, and red silty marls. The sandstones are usually sheet-like and extend laterally for 10-100 m. They lie unconformably on nodular anhydrite deposits.

3.2.2. Interpretation The green shales are generally associated with overbank deposits and are entirely devoid of fossils. They are considered as abiotic, hypersaline, lagoonal deposits (salina). They are overlain by creamy dolomite, which locally becomes darker because of organic components. Prismatic crystals in the gypsum could have formed subaqueously in brine ponds at the sediment-water interface (see mechanism of crystal formation in Rosen, 1989; Rosen and Warren, 1990). These observations argue for the existence of some lagoonal stages in the gypsum precipitation processes, which is in opposition to an earlier interpretation of these sabkha

deposits (Popescu, 1984). The marine origin of the evaporites is corroborated in the laminated anhydrite by the presence of calcareous nannoplankton (NP15 biochronozone) but the lack of benthic micro-organisms, the small size of the coccoliths and the coccolith mono-specificity confirm the restricted conditions. Nodular anhydrite, especially that displaying a typical chicken-wire structure, is indicative of subaerial exposure of soft sediments in which the nodules were formed by displacement of the host sediment (Kinsman, 1966; Hardie and Eugster, 1971). The anhydrite horizons of facies association 2 are interbedded with coarse-grained sandstones that may correspond to ephemeral stream, flash flood deposits (Williams, 1971; Karcz, 1972; Hogg, 1982; Abdullatif, 1989; Olsen, 1989; Cooper, 1990). This facies succession from green shales to the nodular anhydrite corresponds to a complete drying of the lagoonal-lacustrine system. The successions arc nevertheless generally incomplete and commonly truncated by the laminar ephemeral stream sandstone facies, which may be interpreted as the wetting-up counterpart of the drying-up lagoonal stage.

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Fig. 8. Facies association 3. Estuarine and tidal flat deposits. Section 7. Pencil cap for scale (3 cm in diameter). accreted alternations of thick and thin sandy layers interbedded with a dark shale couplet. The average number thinning-upward cycles deposited during a spring/neap tide cycle is about 25.

3.3. Facies association

3: Estuarine and tidal facies

These facies are generally poorly exposed and are only present at the topmost part of the sections in a southward direction, just below unconformity II (section 7, Fig. 5). They are made up of mediumto coarse-grained sandstones 8 to 10 m thick, with clay drapes, and can be subdivided into three superimposed facies. 3.3. I. Description Facies 1. Trough-cross-bedded, 2- or 3-m-thick, IO- to 20-m-long, medium-grained carbonate sand bodies that grade laterally to bioturbated, shelly and muddy sandstones. These carbonate sands lie on channelized erosional surfaces outlined by well rounded pebbles and 1-5 cm clay chips. Sometimes the clay chips are concentrated in the upper part of the forestepping laminae. Along strike, the lateral accretion bedding is composed of l- to 3-cm-thick, cyclic alternations of thick-thin mud and sand couplets, which thicken and thin progressively. Facies 2. Vertically accreted, 5- to IO-cm-thick

125

Facies 2, vertically of thickening- and

alternations of thick and thin sandy layers interbedded with a dark muddy couplet. The number of cycles may reach an average of 25 per thickening and thinning-upward cycle (Fig. 8). Facies 3. Cyclic alternations of planar-laminated dark silty shales and medium-grained sandstones l-5 cm thick, with current ripples. The thickness of these alternations increases and then decreases upward and is ordered in a cyclic alternation of about 10 to 12 couplets. 3.3.2. Interpretation Facies 1 is interpreted as laterally accreted bedding. Lateral accreted bedding is the most characteristic of tidal environments (Visser, 1980; Terwindt, 1981). Facies 1 is then considered to have been deposited in a tidal environment. Because of the presence of double clay drapes, interpreted elsewhere by Visser (1980) as typical record of a complete tidal cycle, facies 1 corresponds to subtidal sandwaves that prograded on a flat (facies 2) close to subaqueous distributaries in an estuarine environment. Subtidal flat deposits record most of the tides

in a single vertically accreted neap-spring cycle, as opposed to the deposits of intertidal environments (facies 3) where only the spring tides are recorded in the uppermost part of this shallowing-upward facies succession (see examples in Tessier and Gigot, 1989; Tessier, 1993). 3.4. Fucies association shore-connectedfacies

4: Lagoonal and

This facies association is exposed in section 8 (Fig. 5). It is made up of a few-metres-thick detrital carbonate and marl alternations. 3.4.1. Description Facies 1. Packstone-grainstone carbonates up to 3 m thick with Anomia pelecypods and marl alternations. The carbonates are made up of coarseningupward packages of thinly laminated shales, massive marls, and crudely planar-laminated or cross-bedded packstones-grainstones containing the flat left valves breccias with Anemia of Anomia and microbioclastic and rare Ostreidae. The Anemia valves are well sorted and always flat lying in the stratification. The concave right valve is very rarely present. Facies 2. Gryphaea marls, up to 3 m thick, with alternations of thick, l-m-scaled, normally graded bioclastic marls and a thin, 10 cm thick, coquina layer. The marls exhibit a strong upward increase in the carbonate content. The Gryphaea coquinas are in a concave up in-life position, and are disarticulated but not sorted or reworked except at the bounding disconformity. Facies 3. Green, thinly laminated shales up to 10 m thick with scarce fauna (a few pelecypods). The shales are bounded above and below by very continuous, planar-laminated, 1O-cm-thick glauconitic sandstone beds and oolitic ironstones that wedge out laterally (see facies 5 below). 3.4.2. Interpretation Facies 1 with Anemia is interpreted as the deposit of low relief, prograding, current-swept, sediment bodies lying on a shaly, soft bottom substrate of a protected marine lagoon, as inferred from the very low diversity in fauna1 assemblages and the euryhaline affinities of Anomia. These sediment bodies are covered with microbioclastic breccias that ex-

press the high-energy, coarse, protective, deflation lag responsible for their preservation during minor flooding events. The flooding is confirmed by the replacement of euryhaline Anomia thanatocoenoses by stenohaline Gryphea biocoenoses and also by the decreasing-upward energy on the sea bottom (faties 2). The deepest facies is figured by the green shales, which could be compared to the Verdine faties (Odin et al., 1988) owing to the greenish colour of the shales, the scarcity of the faunas, and their close association with iron-rich deposits. In recent sediments, the ‘Verdine facies’ occurs in subtropical, shallow marine environments (20 to 50 m depth) within areas of relatively rapid sedimentation close to the delta front of major rivers (Odin et al., 1988). 3.5. Facies association

5: Barriers and shoal belts

The facies association 5 represents a generic grouping of shoal deposits, which are not immediately superimposed in the field. Four main facies are recorded from the base to the top of the sections (Fig. 5). 3.5.1. Description Facies 1. Oolitic and pelletal packstones-grainstones up to 5 m thick. They are made up of thickeningand coarsening-upward alternations of well sorted, tens of centimetres thick, trough-crossbedded oolitic grainstones with wave ripples and a few centimetres thick, massive, burrowed, pelletal packstones. The pelletal packstones disappear progressively upward. This facies is generally affected by an oomoldic porosity and selective dolomitization of the pellets and the ooids. The earliest cement is a stoichiometric dolomite (Ca/Mg = 1) formed in the mixing zone and then a non-stoichiometric dolomite. The non-stoichiometric dolomite is more developed towards the southeast when the carbonates sediments wedge out in the laminar anhydrite and prismatic gypsum deposits. At the top of the facies, vadose silts filled the oomoldic porosity along horizons bounded by channelized erosional surfaces. Facies 2. Ironstone ooids sand sheets (sections 3 and 8, Fig. 5) generally 1 m thick and less than 20 km in lateral extension. They lie unconformably on ironrich glauconitic greenish shales and are bounded upward by a glauconitic and phosphatized surface.

J.-N. Proust, A. Ho.w/Sedirnentmy

They are generally roughly bedded and made up of poorly sorted, subangular, sand-sized, brown iron oxyhydrate-oolites, -oncholites, -lithoclasts, -fecal pellets, large and thick shell debris, and drift wood in a muddy matrix. Small broken shell debris or quartzose sand grains constitute the nuclei of the oolites. Facies 3. Nummulites peryhoratus grainstones and packstones (sections 2, 3 and 8, Fig. 5) are respectively made up of macrospheric and mixed microand macrospheric assemblages. The base of the facies is mostly composed of Nummulites striatus (Bruguiere a,b) but towards the top Nummulites perforatus (de Monfort) is dominant. The typical elementary facies succession is thickening up, coarsening up, 1.5 m thick on average, and composed of the vertical superposition of microlaminated shales, packstones with mixed nummulite assemblages, and trough-cross-bedded grainstones with macrospheric nummulites. Three or more of these successions are generally superposed. They become more mature texturally up-section where they exhibit a gradual thickening- and coarsening-upward trend. Facies 4. Bioclastic wackestones-packstones with Mates gastropods (7 m thick, sections 2 and 3, Fig. 5) lying unconformably either on green shales or Nummulites perforatus grainstones and packstones. They are thoroughly bioturbated, well bedded, and exhibit a thinning-upward trend Facies 5. Foraminiferal and pelecypod bioclastic packstones-grainstones less than 8-m-thick and more than 10 km in lateral extent (sections 1,6 and 7, Figs. 5 and 9). They lie on a SO-cm-thick base, made up of a few-centimetres-thick, mudstone/wackestone and laminated clay alternations, which give way vertically to lenticularand then flaser-bedded grainstones. The bulk of the facies is composed of 30- to 40-cm-thick, trough-cross-bedded packstone-grainstones capped by wave-ripples, interbedded with 5-cm-thick marly layers. The facies passes progressively upward to hummocky-crossstratified (HCS) grainstones sharply truncated by 3D-undulatory erosional surfaces that become amalgamated upward. Erosional channels, 1 or 2 m deep and 10 to 30 m wide, occur toward the top of the section and are locally outlined by l-3 cm lithoclasts. This facies grades laterally to sandy bioclastic carbonates (5 m thick) lying on a regionally correl-

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127

Fig. 9. Facies association 5. Barrier and shoal belt deposits. Section 1. Hammer for scale (40 cm). Facies 5, trough crossbedded foraminiferal and pelecypods packstone-grainstones with wave-rippled surfaces, interbedded with thin marly layers. The facies passes progressively upward to hummocky-cross-stratified (HCS) grainstones sharply truncated by 3D-undulatory erosional surfaces that become amalgamated upward.

ative erosional surface (unconformity II, Fig. 5) just above the terrigenous tidal deposits (facies association 3, see above and sections 6 and 7, Fig. 5). These sandy bioclastic carbonates are composed of crude 2-3 m thick successions of sandy beds with wave ripples, poorly sorted, coarse-grained or pebbly, trough-cross-bedded horizons and wellsorted, medium-grained, planar cross-bedded sandstones with low-angle truncations. 3.5.2. Interpretation Facies 1 corresponds to a prograding oolitic sand shoal deposit that was reworked by wave action in a shoreface environment. Its base probably lies on

128

J.-N. Proust, A. Hosu /Sedimentar)

deeper, more restricted marine deposits represented by the pelletal packstones; but its top was certainly fringing the rim of a hypersaline lagoon responsible for the precipitation of non-stoichiometric secondary dolomite. The oolitic shoal was sharply truncated by renewed alluvial fan sedimentation towards the north and is therefore not genetically related to the overlying deepening-upward lagoonal deposits (facies 1, 2, and 3 of facies association 4). Generally, the main origins assumed for ooid ironstones similar to facies 2 are as follows: (1) coating of particles by microbial films responsible for the reduced conditions required to move ferric elements (Dahanayake and Krumbein, 1985), (2) diagenetic ferruginization of carbonate elements in a pyritic substratum (Kimberley, 1979), (3) reworking of ooids from lateritic soils (Siehl and Thein, 1978), (4) formation in turbulent water close to river mouths providing the ferruginous material (Young and Taylor, 1989), and (5) winnowing of glauconitic or ‘Verdine’ shales during successive sea-level changes (Einsele, 1992). In our case, the source of material of the ooid ironstones of facies 2 could be either marine sediments reworked from the ‘Verdine’ facies, previously deposited iron-oolites, iron-oncholites, and iron-pellets, or it could be fluvial deposits (Muresan and Stoicovici, 1987) and reworked from deeply weathered lateritic soils (angular iron clasts and ooids, drift wood). This mixing of elements of different sources is typical of shoreproximal areas close to river estuaries. However, because the marine iron allochems require a high energy but a low sedimentation rate, the scattered elements must have been secondarily winnowed, reworked, and concentrated by currents in relict shelf or foreshore sand bars or ribbons far from the source areas. Similar associations are deposited during some sea-level rises (Hallam and Bradshaw, 1979; Bayer et al., 1985; Bayer, 1989). The well-sorted, grain-supported fabrics, the sometimes amalgamated wave ripples (HCS), and the occurrence of strong current energy with coarsegrained megaripples, planar lamination, and lowangle truncations in facies 3 and 5 are typical of high-energy, middle to upper shoreface environments. These facies may represent shoals 1040 km long, bordering in a basinward direction a backstepping muddy lagoon like the one inferred from facies 1.

Geology IO5 (19961 I 17-140

The massive and structureless aspect of facies 4 on the outcrop implies slight sorting without any significant winnowing. These processes may occur during sediment gravity flows either at the base of the shoreface, below fairweather wave base, or in a protected back barrier setting. The latter hypothesis fits well the relative abundance of mud and the numerous well preserved gastropod shells. 3.6. Facies association

6: Outel; open marine facies

The outer marine facies (sections 1, 2, 3 and 8, Fig. 5) are generally monotonous throughout the northwestern Transylvanian margin. They are represented by grey marls with pelecypods (45 m) passing below the regional unconformity (II, Fig. 5) to more sandy units that were deposited in more diversified environments. 3.61. Description This facies is usually shaly, green and massive in the basal part, then microlaminated and again marly, silty, grey and massive in the upper part. Intercalations, less than IO cm thick, of coarse-grained, laminated bioclastic grainstones-mudstones may occur in the basal part (ca. 10 m, section I, Fig. 5). The laminated grainstones-mudstones are replaced upward by 5- to IO-cm-thick shell beds characterized by broken, disarticulated, and well-sorted shell debris lying on swaley erosional bases. This faties is capped by burrowed sandy marls containing Skolithos and Diplocraterion. It is interbedded basinward with bioclastic wackestone-packstone horizons, and interbedded landward with current and wave-rippled sandy carbonates (lenticular bedding). This facies is sharply truncated by erosional unconformity II of Fig. 5. 3.6.2. Interpretation The undulatory erosional base of the facies, the lenticularity, the sorting, the occurrence of some high-energy laminations are very similar to stormgraded layers that have been described by Aigner (1982, 1985). Such layers are thought to have been deposited in low-energy, muddy, offshore marine shelves. In the uppermost part, facies association 6 passes progressively into a wave-reworked sandier shoreface environment (section 1, Fig. 5).

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129

Basinward NW

--38

t4y

- 5OMy

IWes. louease acmmodatii

amOBcTeaae1”

spacerespectively

2

1

pyq Suptat&

lkcd plain Fig.

10. Schematic

Cluj-Huedin sequence

SaMa

and hypersaline I~~OOO

area. boundary;

Cditic shoals

reconstruction Note

facies

Nummulitic

A&S

= maximum

Velales shoals

YDalS

architecture

unconformity-bounded surface;

4

5

6

7

Outer.0~

marine

8

p?q

Lagw!al green Shales within Ammia and Guyphaea

of the bidimensional

the large-scale,

RS = ravinement

3

m

in the first Palaeogene

retrogradational flooding

4. Facies architecture and event stratigraphy The vertical organization of the successive facies and depositional environments leads to the recognition of two orders of sequences (Figs. 5 and 10). The smaller one is from 1 to 10 m thick, the larger one from 80 to 120 m. The thinner sequences (Fig. 5) are composed of one or two of the following: (1) a deepening-upward (positive) trend or a proximal to distal evolution from source areas (backstepping); (2) a shallowing-upward (negative) trend or a distal to proximal displacement from the sediment sources (forestepping). Both are bounded by erosional surfaces along which some new space is added for the sediment to accumulate. These erosional surfaces represent marine flooding surfaces in the marine domain and in the continental realm they correspond to base level rise surfaces which bound parasequences (Proust and Deynoux, 1990, 1994). In the continen-

stwm-cbminaled

and progradational

shales

continental linkage

tidal wdtars

to marine

alternation

of parasequence

in the

sets. SB =

surfaces.

tal realm, the backstepping trends are numerous at the base of the section. They are less represented at the top of the section, in the open marine system, where they are replaced by shallowing-upward or forestepping trends. In the transitional areas between marine and non-marine realms, a thick shallowingupward trend is overlain by a thin deepening-upward trend ending in a non-depositional, hiatal surface at the maximum depositional palaeodepth or position of the maximum flooding. Marine flooding surfaces or their non-marine equivalents are considered as relative time-lines even if some of them may represent a lot of time (Posamentier et al., 1988). The correlation of these surfaces allows reconstruction of a bidimensional facies architecture (Fig. 10). In the thicker sequence (Figs. 10 and ll), the stacking pattern of the parasequences exhibits a gradual deepening- and thickening-upward trend (retrogradational) followed by a shallowing-

J.-N. Proust. A. Hos~r/Sedirnentu~

130

Geology

105 (1996)

117-140

Alluwal (floud ~lainr

Drowning surface TIME 42 Ma

RELIEF RESUMPTION

(maximum flooding in the basin)

w

ACTIVE FAULTING Drowning S”hSld,

rime lag unconformity (initial)

Ma at maximum) SPACE (BASIN WIDTH)

Fig.

II

Sediment partitioning in the lowermost continental to marine facies alternation of the Palaeogene foreland basin of Transylvania.

Note the occurrences 01‘shoal belts deposits coeval with major changes in sequential to block

organization

and palaeogeographic

turnover related

tilting.

and thinning-upward tendency (progradational). The deepening-upward trend is bounded below by an angular unconformity of regional extent between the karstified Late Cretaceous carbonates and Early Tertiary ‘Gosau facies’ and the coarse-grained alluvial fan deposits (X, Fig. 10). The time gap can be estimated as a few million years (6 Ma at maximum). The deepening-upward succession is bounded at the top by a heavily bioturbated horizon at the lower part of the outer marine shales, which represents the deepest marine facies in the area and the maximum flooding period (IV, Fig. 10). These shales are truncated by an erosional surface (VI, Fig. IO) with oolitic ironstones overlain by the iirst nummulitic and bioclastic shoal belt deposits. This erosional surface is interpreted as a ravinement surface caused by the upward shift of the shoreface wave base on lagoonal muds. The shallowing-upward trend is bounded at the top along tens of kilometers by a regionally correlative unconformity (I, Fig. IO) that separates beach, estuary, shoals and fluvial deposits (Figs. 12. 13). The shallowing-upward trend is however truncated

in its upper part by a sharp erosional surface (II, Fig. 10) underlain by pebbles or granules and overlain by progradational bioclastic shoal belt and beach deposits. This surface has been induced by wave action at the base of the shoreface during its basinwards shift. The regionally correlative unconformity (I, Fig. 10) and the wave reworked surface (II, Fig. 10) are respectively interpreted as a sequence boundary (see definition in Van Wagoner et al., 1988) and a marine regressive surface of erosion (see definition in Nummedal et al.. 1993). In the thicker sequence, the nodal point between the shallowing and deepening trends is roughly expressed by the first deepening up bioclastic shoal belt (Fig. 11). This transition expresses a major palaeogeographic reorganization marked by the first appearance of open, outer marine conditions in this area. A similar scheme is observed at the top of the section. A 10 m-thick, shallowing-upward, bioclastic shoal belt parasequence set marks the end of the overall shallowing up tendency of the large-scale sequence just before the onset of the uppermost alluvial flood plain deposition. This renewed occurrence of

J.-N. Proust,

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105 (1996)

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131

Alluvial flood plain

Marine shoal belt 2 Fig. 12. Detail of the sharp contact

between the marine series I (a) and the red beds II (b). Unquoted

section located between

sections

5

and 6.

shoals predates a second inversion in the sequential organization or ‘facies dislocation’ that occurs during the main palaeogeographic changes expressed by the modification of the surface drainage in the basin (Fig. 11). These highly mobile, quickly shifting, bioelastic sandstone bodies are located exclusively at the boundaries between the main prograding and retrograding parasequence sets. They indicate either basinward or landward reversals in the sediment fluxes, associated with palaeogeographic changes along disconformities. Some disconformities overlain by renewed marine sediments are associated with fauna1 rejuvenations and are thus well dated. The relationships identified between nannoplankton biochronozones and lithologic formations (Meszaros et al., 1987; Meszaros and Moisescu, 1991) imply the following correspondences between disconformities and nannozones boundaries (Figs. 5 and 10): (1) NP15NP16 boundary and surface VI or surface VII; (2) NP16-NP17 and the maximum flooding surface atop surface IV; and (3) NP17-NP18 boundary and surface II. The duration of the retrogradational and

progradational open marine parasequence sets are tentatively estimated at maximum values of 8 and 4 Ma, respectively (Fig. 1 l), by reference to the chronostratigraphic chart of Harland et al. (1989). 5. Discussion 5.1. The bimodal nature of the sediment architecture The change in the stacking pattern from a generalized landward shift of the successive facies belts (retrogradational, deepening up trend) onlapping the basal unconformity on the South Pannonian unit overthrusts, to an overall basinward shift of the successive depocentres (progradational, shallowing up trend) in the basin seems to correspond to the following specific modifications in the basin architecture (Fig. 11): (1) The transition from a period of active creation of accumulation space to a period of passive infill of this space. The maximum drowning period may be assigned to the maximum occurrence of authigenie material (glauconite) and to sea bottom diage-

132

J.-N. Proust, A. Hosu/Sedimenta~

Geology 10s (1996) 117-140

b

a

Fig. 13. Detailed view of the contact between the marine series 1 (light grey) between sections 5 and 6. (a) Coarse-grained sandstones with low-angle planar (b) Medium-grained, massive and trough cross-bedded sandstones with mud poorly sorted, blocky and mottled silty red mudstones. The mudstones may desiccation cracks. Flood plain deposit.

nesis (oolitic ironstones) at the ravinement surface (surface VI, Figs. 5 and 10). The period of active creation of accumulation space may be tentatively coupled with early synsedimentary fault activity evidenced by the following: (a) the northward location of the thickest sediment pile with downward pinchout of most of the unconformities in the deepest marine facies; differential sedimentation from north, of massive evaporite-rich deposits in the north, and of interbedded evaporite-alluvial sediments in the south; (b) the apparent superposition of the shoal and shoal belt through time (Fig. 10); (c) the coarsegrained sediments in alluvial fans with poorly sorted pebble-cobble conglomerates. The active faulting ends close to the surface VII atop the Anomia shales which buried the faults and marked the beginning of the first open marine incursion in this area. The role of the fault zone then decreased and only controlled the overall spatial distribution of the depositional environments through time. This crude geometry is parallel to a buried

and the red beds II (dark grey). Unquoted section located laminations and low-angle truncations. Beachface deposit. pebbles. Backshore and fluvial deposit. (c) Structureless, pass laterally into nodular dolomite with root-traces and

perennial composite structure that emerges westward in the pre-Eocene basement of the Apuseni Mountains (inset map, Fig. 5). The activity of this structure may explain the creation of the small-scale bounding surfaces (surfaces VI, V and IV, Figs. 5 and 10) in the immediately overlying shoal belt deposit. (2) The transition from a closed, undrained basin with alluvial fans, lacustrine facies and then restricted marine deposits to an open drainage system represented by the open marine and tide-dominated estuary deposits. The undrained conditions are expressed by (a) the supersaturated phreatic waters responsible for the precipitation of evaporites, (b) the dominant surficial drainage mainly controlled by episodic, rain-controlled ephemeral streams, and (c) the starved shallow sea bottom conditions responsible for the formation of glauconite and diagenetic ironstones in the restricted marine realm. The drained conditions in the progradational part of the section are exhibited by the open seaway required to produce extensive offshore marine storm deposits, and by the major rivers outlets.

J.-N. Prousi, A. Hosu/Sedimentary Geology 105 (1996) 117-140

(3) The global widening of the basin exhibited by the transition from basal overlapping alluvial fans to thinner, sheet-like sediment bodies of larger lateral extent, and also by the transition from converging inflows in the alluvial domain to diverging outflows in the estuaries. The basal retrograding system with alluvial fans, salinas, and restricted marine environments, coeval with fault activity, may then correspond to a gradual decrease of the aggradation rate resulting from a progressive widening of the basin uncompensated by sediment input during block tilting (see similar mechanism in Olsen and Schlische, 1988). If the fluvial input remains constant through time, the sediment load is progressively dispersed in thinningupward, sheet-like, but laterally more extensive, sediment bodies. The water table rises in the basin. The prograding system with open marine deposits ending in fluvial flood plain environments coeval with fault burial expresses, in contrast, a gradual decrease in the subsidence rate (differential uplift?) and then feedback to more shallow marine tidal conditions when the basin is progressively filled up to base level. The bounding surfaces of these two parasequence sets are considered as simply caused by changes in the regime of subsidence between an active (retrograding) and then a passive (prograding) system related to shifts in the regional strain field. These surfaces are underlain by low-relief, highly mobile sandstone bodies associated with changes in the basin-scale palaeogeographic organization. 5.2. The eustatic control on the shape of the regional sea-level curve Comparison of the local relative sea-level curve inferred from the changes in the depositional environments through time in the Transylvania Basin, and the global eustatic sea-level curve for the Eocene depicted in northwestern Europe and all around the North Atlantic margins (Haq et al., 1988) reveal the following: (1) the long-term trends are in opposition with synchronous relative sea-level rise and eustatic sea-level fall, and (2) the short-term oscillations are very similar (Fig. 14). It can be then assumed that, in Transylvania, the long-term trend expresses a local phenomenon and that the short-term oscillations express some global signal superimposed on the long-term variations.

133

The following two different mechanisms are widely accepted to explain global sea-level variations: (1) a change in the ocean water volumes caused by the waxing and waning of the low-altitude continental icesheets (glacio-eustasy) and (2) a change in the volume of the ocean basins caused by variations in length and spreading rate of the ocean ridges, which displace the ocean water onto the continents (tectono-eustasy). The first mechanism has a short-period effect (a few thousands years) that is very apparent in the last few thousands of years with a magnitude over 100 m and a rate of 1 cm per year. The second mechanism has a long-period effect (a few Ma years) with a magnitude that may encompass 100 m and a rate of 1 cm per 1000 years. The mechanisms responsible for regional sealevel changes are numerous and highly diversified. In intraplate basins, the regional sea-level changes are generally related to differential flexures and uplifts caused by peripheral tectonic activity such as orogeny, transform zone creation, changes in the spreading rate, etc. (Cloetingh, 1988a,b). This tectonic activity is invoked to explain regional deviations of the relative sea-level curve from the global sea-level chart when major plate reorganizations occur simultaneously and when rapid shifts in the magnitude of the subsidence-uplift phases occur in the same basin (Cloetingh, 1992). The regional sealevel signature of intraplate stresses is composed of a slow rise followed by a rapid fall correlated to the differential vertical motion of the lithosphere that modulates the long-term basin deflection caused by its thermal cooling (Cloetingh, 1992). 5.3. The tectonic control on the shape of the regional sea-level curve A basin-scale compression best explains the regional tectonic strain field oriented SW-NE in the Eastern Carpathian domain from the Cretaceous to the Neogene (Tapponnier, 1977; Csontos et al., 1992). From the Eocene to the Early Miocene the (micro)plate tectonic activity in the area was driven by the suturing of the North and the South Pannonian units and by the south-to-north diachronism of the closure of the Magura flysch trough (Sandulescu, 1988; Csontos et al., 1992) (Fig. 2). The main consequences were the clockwise rotation of the

J.-N. Proust, A. Hosu/Sedimentary

134

RELATIVE SEA-LEVEL (Transylvanian Basin) Time in Ma

+50m

I

0

-5Om

Geology 105 (1996) 117-140

EUSTATIC SEA-LEVEL (Global) 200m

100

I

OLIGOCENE

15

LUTETIAN

YPRESIAN

PALEOCENE This paper

Haq et al. 1988

Fig. 14. Comparison of the relative sea-level curve depicted in the Transylvanian Basin and the eustatic sea-level curve (Haq et al.. 1988) for the Eocene. Note the discrepancy between the long-wavelength sea-level rise in Transylvania that is contemporaneous with the eustatic sea-level drop, and the similarities between the short-wavelength sea-level changes.

South Pannonian unit (Marton and Marton, 1989) accompanied at its rims by (1) the end of calcalkaline volcanism, (2) wrench tectonics along the mid-Hungarian line, and (3) the formation of an accretionary wedge in front of the subducting slab(?) of the Magura Ocean with west-to-east propagation of the overthrusts (Csontos et al., 1992). These largescale tectonic events are marked in the Transylvanian Basin by growing anticlines in the early Priabonian redbeds (Fig. 1.5) by the presence of overthrusts under the Early Miocene unconformity, and by uplift indicators (fission tracks) in the core of granite in the Apuseni Mountains (Ciulavu et al., 1994). The bimodal (retrogradational and progradational) nature of the sediment infill of the TranSylvania Basin is also very similar to that inferred

for the elongated, asymmetrical sub-basins in the Eocene-Oligocene North Hungarian Basin (Baldi and Baldi-Beke, 1985; Royden and Baldi, 1988; Tari et al., 1993). These sub-basins are composed of initial shallow marine deposits followed by a thick accumulation of bathyal sediments and a gradual shallowing trend ending with evidences of subaerial erosion. In the transpressive context, the sedimentation is considered to have been controlled by growing anticlines on blind thrusts in a retroarc basin (Fodor et al., 1992; Tari et al., 1993). The presence of backthrusts under the anticlines is still to be proven (Sztano, 1994), but the large wavelength deformations seem very similar to the intraplate stress tectonic style. Such a mechanism could be invoked for Transylvania because of the following: (1) the

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J.-N. Proust, A. Hosu/Seditnenta~

136

regional character of the long-term sea-level change, (2) the slow rise followed by an abrupt sea-level fall, (3) the bimodal nature of the sedimentation, (4) the passage from a narrow undrained to an enlarged and drained sub-basin, (5) the growing anticlines and overthrusts in the Palaeogene deposits and the uplift in the nearby Apuseni fold-thrust belt, and (6) the peripheral wrench- and thrust-tectonic events along the Mid-Hungarian line and in the Magura Basin accretionary wedge. 5.4. Speculative

depositional

model

Based on the intraplate tectonic model, the proposed outline of basic evolution corresponding to the observed elementary stratigraphic events is as follows (Fig. 16): (1) After the initial emersion of the Laramide uplift (1, Fig. 16), the basin was drowned as peripheral tensional stresses led to the progressive burial of the relief, the overall widening of the basin, the onlap of the sediments onto the basal unconformity, and the landward migration of the shoreline (inset, Fig. 16). The basin was filled by coarse-grained, immature, and spatially restricted alluvial fan deposits (2, Fig. 16) near the previously uplifted areas and finer-grained deposition in salinas in more distal areas (3, Fig. 16). The facies succession progressively gave way to deeper and more confined, restricted marine deposits (4, Fig. 16) with increasing drowning through time. These marine deposits finally sealed the growing faults, marking the definite cessation of active fault-controlled subsidence at the minimum of substrate relief in the basin. This period corresponds to the first marine incursion in the basin and to the formation of a marine ravinement surface or drowning surface. The sediment budget became increasingly lower than the subsidence. The accommodation space was progressively unused. The sedimentation was condensed. (2) Relative uplift at the basin flank and subsidence at the basin centre related to peripheral compressional stresses led to the progressive creation of relief and narrowing of the basin (inset, Fig. 16), ending with fluvial deposition (8, Fig. 16), offlap coeval with the formation of the uplifting unconformity and the seaward migration of shoreline belt. The basin was progressively filled by a prograding,

Geology 105 (1996) 117-140

shallowing-upward succession from outer, offshore marine deposits (5, Fig. 16) to tide-influenced estuary (6, Fig. 16) and shoal deposits (7, Fig. 16). Lateral facies changes were rare and the palaeogeography was mainly controlled by progressive flexure of the substratum, partly because of the sediment load redistribution coming from erosion of the adjacent relief. The formation of the uplifting unconformity may be related to the end of the period of offlap at the maximum of differential uplift and subsidence between the basin flank and the basin centre. The sediment budget became increasingly greater than the subsidence. The accommodation space was filled. 6. Conclusion The typical elementary building block of the Palaeogene Transylvanian Basin is made up of a vertically superposed retrogradational and progradational parasequence set, depicted in a single, 10 Ma long, non-marine to marine facies alternation (Fig. 16). The stratigraphic succession of the retrogradational parasequence set is made up of the following: (1) stacked, fault-controlled, poorly extensive alluvial fans; (2) ephemeral stream, and salina deposits; (3) lagoonal and shore-connected sediments. The facies succession of the progradational parasequence set is composed of the following: (1) outer, open marine facies; (2) subtidal to supratidal estuarine deposits. The two parasequence sets are bounded at their tops by low relief sandstone bodies that were deposited in shoal belts, and which mark major palaeogeographic rejuvenations associated with time-transgressive baselevel change unconformities. These unconformities are as follows: (1) A base level rise (or drowning) surface underlain by oolitic ironstones and then glauconitic shales that typify basin starvation during the periods of maximum basin drowning. (2) A base level fall (or uplifting) unconformity overlain by alluvial flood plain deposits, which characterize periods of relief rejuvenation during an uplift event. This typical succession of non-marine to marine strata is one of three such elementary building blocks

J.-N:Pmust,

A. Hosu/Sedimentary

Geology 105 (1996) 117-140

regressive progradational sedimentation (“passive” infilling of the basin)

“uplifting” unconformity \

/

137

“drowning” surface /

TRANSYLVANIAN

nsgressive retrcgradational sedimentation (“active” faulting and drowning)

initial time-lag unconformity

p7;9

platform carbonates

m

coarse-grained

II\^^I/

sebkha

alluvial

and salina

fan

evaporites

rmj

shalylagoon

t-j

detrital

r--Gq

storm dominated outer marine and estuarine deposits

carbonates

in shoal belts

Fig. 16. Idealized cross-section in the continental to marine facies alternation in the Palaeogene of the Transylvanian Basin. The inset cartoon exhibits the relationships that may exist between compressional and extensional stresses at the basin flank (landward of the main sediment load) and the corresponding differential uplifts and drownings in the intraplate domain.

of the Palaeogene sedimentary record of the Transylvanian Basin. Each of them corresponds to single alternations of underfilled-overfilled stages controlled respectively by active faulting and drowning periods followed by uplift events. These three cycles correspond to the preserved sedimentary record of the Palaeogene in the Transylvanian Basin and therefore provide information on the large-scale pulsated evolution of the basin during the period of active compressional activity of the microplates, which collided with the European Platform to the north in Miocene times. Acknowledgements This work has been carried out as part of a PECO mobility action founded by the European Community and the Centre National de la Recherche Scientifique (URA719, Lille, France). I would like to thank B. Tessier, 0. Averbush, F. Chanier, H. Chamley, J. Ferriere, M. Lopez and the reviewers D. Cant, S. Cloetingh, K.A.W. Crook, and L. Csontos, who largely contributed to the scientific improvement of the manuscript. Special thanks also for S. Cloetingh,

K.W.A. Crook and J. Ferriere for the detailed editing work. I am also grateful to E. Hanton, M. Bocquet, and M. Carpentier for their technical support. References Abdullatif, O.M., 1989. Channel-fill and sheet-flood facies sequences in the ephemeral terminal river Gash, Kassala, Sudan. Sediment. Geol., 63: 171-184. Aigner, T., 1982. Calcareous tempestites: storm dominated stratifications in Upper Muschelkalk limestones (Middle Trias, West Germany). In: G. Einsele and A. Seilacher (Editors), Cyclic and Event Stratification. Springer-Verlag, Berlin, pp. 180-198. Aigner, T., 1985. Storm depositional systems. In: G.M. Friedman, H.J. Neugebauer and A. Seilacher (Editors), Lecture Notes Earth Sci., 8, 174 pp. Allen, PA. and Homewood, P., 1986. Foreland basins: an introduction. In: PA. Allen and P. Homewood (Editors), Foreland Basins. IAS Spec. Publ., 8: 3-12. Allen, P.A., Crampton, S.L. and Sinclair, H.D., 1991. The inception and early evolution of the North Alpine Foreland Basin, Switzerland. Basin Res., 3: 143-163. Baldi, T. and Baldi-Beke, M., 1985. The evolution of the Hungarian Paleogene basins. Acta Geol. Hung., 28: 5-28. Balla, Z., 1984. The Carpathian loop and the Pannonian Basin: a kinematic analysis. Geophys. Trans., 30: 313-353.

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