GeolCarp_Vol54_No2_119

GEOLOGICA CARPATHICA, 54, 2, BRATISLAVA, APRIL 2003

119136

ORIGIN AND SEQUENTIAL DEVELOPMENT

OF BADENIAN-SARMATIAN CLINOFORMS

IN THE CARPATHIAN FORELAND BASIN (SE POLAND)

SZCZEPAN J. PORÊBSKI

, KAJA PIETSCH

, RYSZARD HODIAK

and RONALD J. STEEL

Institute of Geological Sciences, Kraków Research Centre, Polish Academy of Sciences, Senacka 1, 31-002 Kraków, Poland;

ndporebs@cyf-kr.edu.pl

Department of Geology, Geophysics and Environmental Protection, Academy of Mining and Metallurgy, Al. Mickiewicza 30,

30-059 Kraków, Poland

University of Wyoming, Department of Geology and Geophysics, Laramie, Wyoming, 82071, USA

(Manuscript received April 30, 2002; accepted in revised form October 3, 2002)

Abstract: Upper Badenian-lower Sarmatian (Miocene) strata along the active (southern) margin in the Carpathian Foreland

Basin reveal seismic-scale deltaic clinoforms that grew from the south and developed a shelf-to-basin floor relief of over

300 m. Two architectural types and six 4

-order sequences were distinguished along the clinoforms on the basis of

correlation of an extensive network of seismic lines and geophysical well logs. A Type A clinoform consists of steeply

dipping (34°), planar-oblique strata that occur in narrow (24 km wide) belts composed chiefly of aggrading thick-

bedded massive sandstones with onlap upper terminations. Type B clinoforms occur in wide (812 km) belts of less steep

(

≤

2°), strongly tangential strata composed of mouth-bar/prodeltaic increments that lack thick turbidites and show down-

ward-stepping to retrogradational stratal arrangements. Both clinoform types document shelf-margin accretion, chiefly

during periods of relative sea-level fall and early rise. None of these types predicts a coeval basin-floor fan development,

either because the slope was too narrow to ignite delta-fed hyperpycnal flows into high-efficiency turbidity currents

(type A), or the bulk of the sand delivered to the shelf edge was trapped within slope-perched mouth bars shifting 6

10 km around the freshly formed shelf margin (Type B) during relative-sea level oscillations. The lower three sequences

(upper Badenian) reveal a strong aggradational component, whereas starting from the angular unconformity at base of

the Anomalinoides dividens Zone (lower Sarmatian) the offlap break assumes an increasingly flat trajectory basinwards.

This change is thought to reflect a decreased rate of addition of accommodation due to cessation of thrust loading, and

faster progradation of the clinoform. It is concluded that large-scale clinoformed shelf margins are not limited to rifted

continental shelf margins, but can also be present in foreland basins. In such an environment, flexural and fault-induced

subsidence promotes long-term relative sea-level rise in the hangingwall and, consequently, the generation of long and

high deltaic clinoforms on the accreting shelf margin, whereas the actively rising footwall precludes the preservation of

paralic facies and provides an abundant sediment supply for delta growth.

Key words: Western Carpathians, Carpathian Foreland Basin, clinoform, shelf-margin delta.

Introduction

Although Miocene deposits in the Carpathian Foreland Basin

(Fig. 1) have been the subject of vigorous research for many

decades, an understanding of their sedimentary environments

and depositional architectures has developed rather slowly.

The conventional view of these deposits as molasse has led to

the erroneous belief in their invariably shallow-water origin.

Whereas this is essentially true for Miocene basin fill along

the cratonward (northern) side of the basin (e.g. Czapowski

1994; Kasprzyk 1993; Roniewicz & Wysocka 1999), there is

growing evidence of repeated development of a marked

bathymetric gradient downdip of the active (southern) margin.

The evidence includes the presence of Badenian submarine

fans (Maksym et al. 1997), shallow neritic to bathyal foramin-

iferal assemblages (Czepiec & Kotarba 1998; Gonera 1994),

and seismic-scale clinoforms (Krzywiec 1997). Evidence pre-

sented earlier (Porêbski 1999) and expanded here has shown

that these clinoforms descended from the south and developed

a shelf-to-basin floor relief 300400 m in height. Similar cli-

noforms are associated with the shelf break of Quaternary

continental margins (e.g. McMaster et al. 1970; Suter & Ber-

ryhill 1985; Tesson et al. 2000), but can also be an important

constituent of any delta-fed subaqueous platforms located at

the margins of rapidly subsiding marine basins (Porêbski &

Steel in press). Deltaic clinoforms that grew onto bathyal

slopes have, surprisingly, rarely been reported from the pre-

Quaternary record; notable exceptions include examples de-

scribed by Mayall et al. (1992), Steel et al. (2000) and Blink-

Bjorklund et al. (2001).

The main goals of the present work are (1) to reconstruct

clinoform architecture via detailed geophysical log and seis-

mic correlations, in order to get insight into the evolution of

such clinoformed active margins and (2) to better asses the

role of shelf-margin deltas as the potential staging areas for

supply of sand to the deep sea.

Regional setting

The studied succession is located within the outermost, little

to non-deformed unit of the Carpathian orogenic belt (Fig. 1),

which formed the peripheral part of the Carpathian foreland

basin system. This system was generated on the interior side

120 PORÊBSKI et al.

of a collision zone during lithospheric flexure of the East Eu-

ropean Platform in response to a northward-stepping thrust

load (Kotlarczyk 1985; Oszczypko & ¯ytko 1987; Oszczypko

1997). The Miocene fill of the Carpathian Foreland Basin is

up to 3.5 km thick, and consists of deep-water, shallow-ma-

rine and deltaic siliciclastic sediments and evaporites. In addi-

tion to its northerly progradation, the basin fill shows also a

large-scale offlap and thickness increase towards the east

along the convergent margin (Ney 1968), reflecting an in-

creased accommodation in that direction. Such a pattern may

possibly indicate either an easterly increase in loading along

the strike of fold-thrust belt or, together with an echelon ori-

entation of Outer Carpathian nappes, the creation of tectonic

accommodation due mainly to a left-lateral transpression.

Cratonwards, individual clinothem bodies onlap a south-dip-

ping unconformity with an associated hiatus that embodies

the Paleoceneearly Oligocene (e.g. ¯ytko 1999). The hiatus

spans the main subsidence phases in the foreland basin, but

the thrust loading continued through Miocene time, giving

rise to flexural subsidence in the basin (Oszczypko 1997). The

unconformity surface follows a regional, SE-draining valley

system extending across the entire cratonic margin, with val-

ley relief locally up to 1000 m (Karnkowski & Ozimkowski

2001). This craton-sourced system was active mainly during

the underfilled phase of foreland development (e.g. Picha

1996) and possibly during the early Badenian (Po³towicz

1998), but remained largely dormant during late Badenian

Sarmatian time (Krzywiec 1999; Porêbski 1999).

Study area

Structural setting

The study area, extending some 50 km along strike and

30 km downdip, is located immediately east of the Krakow

salient (Ney 1968) and north of the main Carpathian over-

thrust (Figs. 1 and 2). Along this thrust, flysch rocks together

with a narrow strip of Miocene sediments, are superposed

onto the weakly to non-deformed Miocene basin fill. Within

the study area, the fill is up to 1300 m thick and consists mainly

of thin-bedded sandstones, siltstones and mudstones intercalat-

ed with thicker sandstone tongues. Seismic mapping has shown

that the pre-Miocene cratonic basement dips ca. 3° southeast-

wards (Fig. 3), and is dissected by the Szczurowa paleovalley

(Po³towicz 1998). The valley is up to 4 km wide and 400 m

deep in the southeast; towards the NW, it shallows and splits

into three branches, each less than 0.5 km in width (Fig. 3).

The basement is affected by extensional and possibly strike-

slip faults. Major faults trend NW-SE, downthrowing mainly

to the SW. Minor faults strike NNW-SSE and show different

throw directions. Along the main fault zone in the area, re-

ferred to as the £ysokanie-£apczyca fault (Fig. 3), the base-

ment was displaced 250370 m southwards. During Sarma-

tian compression, this fault formed a frontal ramp on which

older Miocene sediments, together with slivers of the base-

ment, were detached and formed a steeply dipping, NNE-

verging anticlinal thrust stack. Another stack forms the core of

the Liplas uplift farther to the SW (Fig. 3). Both stacks contin-

ue eastwards into a narrow band of folded and faulted Mi-

ocene sediments, and all together are distinguished as the

Zg³obice Unit (Kotlarczyk 1985). The unit is bordered to the

south by the main Carpathian overthrust. South of the

£ysokanie-£apczyca fault, the Zg³obice Unit expands in

width within the Gdów re-entrant which is defined by a south-

erly deflection in the overall W-E striking Carpathian over-

thrust (Fig. 2). The clinoform body that escaped deformation

extends north of the Zg³obice Unit, and shows a structural tilt

of 12° NE.

Stratigraphy

The succession is split by an evaporite unit (Fig. 4) that is

distinguished as the Bochnia Formation (Kuciñski 1982) and

forms an excellent correlative horizon traceable across almost

the entire basin. Although the Bochnia Formation was later

redefined to include gypsum and anhydrite within the

Krzy¿anowice Formation (Alexandrowicz 1997) and halite

within the Wieliczka Formation (Garlicki 1994), such a two-

fold subdivision cannot be used consistently in the subsurface

work and, hence is not followed here. Evaporite rocks in the

southern part of the basin provide numerous signs of gravity

collapse (l¹czka & Kolasa 1997) and are considered to repre-

sent deep-water facies (Peryt 2000), whereas those exposed

along the cratonic margin are arranged in a number of up-

ward-shallowing cycles with evidence of subaerial exposure

(Kasprzyk 1993). The evaporite unit is overlain by 20300 m

of shales, mudstones, sandstones and tuffs of the Chodenice

Beds (Fig. 4) that are unconformably followed up by the more

sandy Grabowiec Beds, 300500 m thick. Both units thin bas-

inwards within shales of the Machów Formation (Jasionowski

1997), and are replaced along the northern basin margin by a

thin stack of transgressiveregressive cycles comprised of

mixed, siliciclastic-carbonate nearshore to basinal deposits

(Chmielnik Formation and equivalents; Czapowski 1994; Ro-

niewicz & Wysocka 1999).

Fig. 1. Map showing the location of study area (rectangle) within

the geological framework of the Carpathians.

122 PORÊBSKI et al.

The supra-evaporite succession, discussed here, comprises

three local forminiferal zones corresponding to the NN6

Nannoplankton Zone, and is probably not much more than

1 Ma in duration (Fig. 4). Absolute Ar/Ar dating of a horn-

blende tuff located near the base of the Bochnia Formation

has yielded equivocal results of 1128 Ma (Bukowski 1999).

Material and methods

In total, about 110 seismic lines, and geophysical logs from

ca. 160 wells provided by Polish Oil and Gas Co., were used

in the present study (Fig. 2). The available seismic grid ex-

tends some 20 km downdip; dip and strike line spacings over

most of the mapped area are 0.52 km. The vertical resolution

does not generally extend below 20 m. Because of the scarcity

and poor preservation of core material, information from

small and scattered outcrops was also used, giving a better in-

sight into facies development. Unconformities defined by dif-

ferent stratal lapouts and erosional truncations on seismic sec-

tions were traced as close as possible to geophysical wells

(gamma-ray, resistivity, neutron-porosity, spontaneous poten-

tial). Such correlations aided by stacking patterns recognized

on well logs, were used to define key surfaces (maximum

transgression and regression surfaces, erosional unconformi-

ties and their correlative conformities). These formed a basis

for the definition and mapping of depositional systems tracts

(geometry, thickness and net sand contents) and the unravel-

ling the paleogeographic evolution of the studied margin for

selected time slices.

Seismic data interpretation, time/depth conversions and

seismostratigraphic mapping were carried out using Schlum-

bergers GeoFrame 3.8.1 package for 2D seismic interpreta-

tion (Charisma). Well log correlation and mapping of strati-

graphic attributes were performed using Landmarks

StratWorks module. Seismic data are of different vintages

(19781993) and, hence, of variable quality. In order to stan-

dardise and improve the quality of seismic data, an advanced

reprocessing scheme using the Omega system was applied to

selected sections. The crucial element in the reprocessing,

which was done by Geofizyka-Kraków Ltd., involved (1) the

application of COHERENCY STACK and the RNA proce-

dures that improved signal/noise ratio (S/N); (2) the frequen-

cy-distance migration F-X that enhanced the horizontal reso-

lution, particularly in tectonically disturbed segments; and (3)

the application of wavelet processing that permitted the acqui-

sition a zero-phase record that is especially useful for extract-

ing stratigraphic information from seismic data. The repro-

cessed sections, with real amplitude relationships preserved,

were hung on the common reference level set up at 170 m a.s.l.

Clinoforms

Geometry

Seismic sections reveal the presence of a large-scale clino-

formed body that descends towards the north and appears to

wedge out by onlap onto the top of the Bochnia evaporites

dipping southwards (Figs. 5 and 6). However, a closer inspec-

tion of these relationships shows that the top of the Bochnia

Formation does not represent a single onlap unconformity.

The body consists of a series of clinothems that thin north-

and northeastwards towards the basin centre, and are bounded

by lap-out unconformities. The lower ends of successive bas-

inward clinothems do not abut onto the evaporite top. Instead,

they tend to wedge out basinwards at progressively higher

levels above this surface forming a thin zone of ascending

contacts corresponding to condensed, shale-prone sediments.

Clinothem reflectors in the landward (southern) side of the

clinoform body are planar with short angular to tangential

lower ends and can have truncated tops. Basinwards, succes-

sive clinothem bundles assume more tangential shapes with

long, locally hummocky toes and grade occasionally into sig-

moidal forms. The latter display well-developed offlap breaks

that mark the change in declivity at the top of the prograding

sediment prism. Clinoform slopes dip generally at 23°, but

locally the steepest, upper segments in planar reflectors can

reach 46° in inclination. The maximum height that the slope

attains is 330 m, but smaller-scale clinoform bodies (mouth

bars or deltas) perched on the main slope rarely exceed sever-

al tens of metres in amplitude (Fig. 6). Upper clinoforms seg-

ments are almost invariably associated with the most sand-

prone sediment (Fig. 5). These shelf-edge sandbodies

generally tend to pinch out landwards within a more heteroli-

thic facies.

The well-developed slope break is backed by a flat to gently

N-dipping platform (shelf) that can be locally incised by N to

NE-trending channels, 0.51 km wide (Fig. 7F). The pre-

served width of the shelf between the successive offlap breaks

and the frontal thrust of the Zg³obice Unit does not exceed

8 km, but it may be far greater in the uppermost 100150 m of

the succession that has not been seismically imaged. Although

clearly progradational in character, the clinoform set reveal a

strong aggradational component reflected in the rising to land-

Fig. 4. Stratigraphic scheme of Miocene succession in the study

area, showing the sequence stratigraphic subdivision against the

biostratigraphic framework. (*) after £uczkowska (1964, 1995),

Czepiec & Kotarba (1998), Olszewska (1999); (**) after Garec-

ka & Jugowiec (1999). MFS maximum flooding surface; SB

sequence boundary; au angular unconformity; Mau base-of-

Miocene unconformity.

ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 123

ward-stepping trajectory of the offlap breaks (Fig. 5). Inter-

nally, the clinoform set reveal a number of downlap and onlap

unconformities that are particularly discernible within the

slope reaches of the shelf-margin succession (Fig. 6) and have

been used for subdivision of the studied succession into depo-

sitional sequences.

Facies spectrum

The main features of the clinoform prism include:

Major sandbodies are mainly strike-oriented and they tend

to wedge both landwards (to the S) and basinwards (to the N)

within marine mudstones and shales (Fig. 5);

The maximum sand development tends to be associated

with the shelf edge or offlap break (Fig. 5);

Fine-grained sediment provides microfossil assemblages of

bathyal, neritic, nearshore to freshwater aspects (e.g. Gonera

1994; Czepiec & Kotarba 1998; Gedl 1998, unpublished;

Olszewska 1999);

Sands include both classical and high-density turbidites,

deltaic mouth bars, and shell-bearing, cross-bedded shoreface

sandstones (Porêbski & Oszczypko 1999; Porêbski 1999)

(Fig. 8);

There is little, if any, evidence of coastal plain (paralic) fa-

cies.

Facies landwards of the offlap break: Valley and other

channel fills are typified by a blocky to bell-shaped gamma-

ray response and vary in thickness from 5 to 55 m. They con-

sist of medium-grained, locally conglomeratic sand and poor-

ly consolidated sandstones that are massive or display

large-scale trough cross-stratification associated with mollusc

detritus. They are interbedded with intraclasts breccias and

dm-thick heterolithic beds (Fig. 8A). A valley recognized on a

strike-oriented seismic section shown in Fig. 9, is 5 km wide

and displays a compound fill, up to 60 m thick, with clearly

recognizable lateral accretion strata. Either a fluvial or estua-

rine origin can be inferred for this valley fill. Some cross-bed-

ded shell-bearing sandstones may represent sharp-based

Fig. 6. Seismic dip-section showing the growth anticlines and associated mini-basins within the clinoform that is backed southwards by

deformed Miocene sediments.

Fig. 5. Interpreted dip-oriented seismic line showing the concentration of sand bodies along the accreting shelf edge.

124 PORÊBSKI et al.

shoreface facies, as suggested by the bimodal dip directions of

cross-sets, with the modes directed towards the NE (basin-

wards) and SSE (Porêbski & Oszczypko 1999).

Mouth-bar facies form upward-coarsening units that are

usually less than 20 m thick. They consist of non- to very

weakly bioturbated laminated mudstones and linsen to wavy-

bedded siltstone/very fine sandstone/mudstone heteroliths

(Fig. 8C). The latter become interbedded upwards with dm

thick, fine-grained Tbc turbidite sands. Upper levels of the

mouth bar units reveal sharp-based fine- to medium-grained

sandstones rich in plant detritus, which are predominantly rip-

ple and flat laminated. The sandstones, 17 m thick, are com-

monly extensively deformed into balls and pillows, but swa-

ley and hummocky cross-stratifications is discernible in

places. Mixtures of open-marine and fresh-water dinocysts are

common in this facies (Gedl 1998, unpublished).

Lower-shoreface/prodeltaic shelf facies consists mainly of

laminated to occasionally bioturbated mudstones interbedded

with cm-thick, fine-grained graded sand beds. These com-

monly show gutter casts, groves and prod marks that are

aligned generally W-E, i.e. parallel to the inferred shoreline

trend, with the prods pointing to the easterly-directed geo-

strophic storm currents.

Facies basinwards of the offlap break: Coarsening-upward

units typified by gamma-ray response of a strongly serrated

pattern passing upwards into a weakly serrated to blocky seg-

ment, that are located at or just below the former shelf edge,

are interpreted as deltaic mouths bars perched on the upper

slope. These units vary in thickness from 30 to 75 m. The

dominant lithotype comprises amalgamated beds of massive,

fine- to medium-grained and granule-bearing sandstones that

are occasionally rich in plant detritus, shell hash and redepos-

ited foraminifers (Fig. 8B). Dm-thick intervals of inversely-

graded bands (traction carpet), flat and ripple-lamination are

common. Soft-sediment deformation includes large-convolu-

tions near bed tops and metre-scale, basinward-verging imbri-

cated shear planes.

Slope sediments with which these mouth bars interfinger

consist of two main facies. (1) A fine-grained host is repre-

sented by bioturbated mudstones interbedded with heterolith-

Fig. 7. Time maps showing the evolution of the prograding clinoform geometry along selected sequence boundaries and their correlative

conformities. Note a valley breaching the shelf along unconformities SB3 (B) and SB7 (F), and the growth anticlines with intervening slope

basins in (AC).

ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 125

ic packets that abound in very thin Tbc and Tcd sand turbid-

ites (Fig. 8D). Thin-bedded sandstone/mudstone alternations

display structures indicative of deposition from both waning

and waxing turbulent suspensions, such as climbing ripples in

normally graded and coarsening upwards sets. These fine-

grained rocks are interbedded with (2) blocky sandstone units,

535 m thick, which are arranged either in progradational pat-

terns, or occur in aggradational stacks, up to 125 m thick, with

updip onlap terminations. Cores available from outside of the

study area suggest that these sandstones are fine- to medium-

grained and unstructured (massive) in character. Basinwards,

these blocky sandstones, interpreted as the deposits of hyper-

pycnal flows, either terminate rather abruptly, or thin into het-

erolithic packets (525 m) which show a highly serrated gam-

ma-ray response. The packets can possibly represent

depositional lobes made of classical turbidites. Basin plain fa-

cies consists of shales and mudstones that locally contain con-

centrations of pelagic organisms, such as pteropods, foramini-

fers and radiolarians.

Paleoecological constraints

Although the paleobathymetric significance of Miocene mi-

crofossils from the Carpathian Foreland Basin has long been

studied (e.g. £uczkowska 1964, 1995; Gonera 1994; Czepiec

& Kotarba 1998), discontinuous coring makes any precise

bathymetric fluctuations difficult to detect in single vertical

sections. However, the available results point to an overall up-

ward shallowing throughout the studied clinoforms. Czepiec

& Kotarba (1998) found that the Chodenice Beds abound in

Bulimina, Uvigerina, and Pulenia ranging from bathyal

depths (several hundreds of metres) to the outer shelf, whereas

the lowermost Sarmatian deposits reveal high diversity miliol-

id-elphidiid and elphidiid-nonionid assemblages of more

Fig. 9. Close-up view of seismic-strike section showing the com-

pound fill of the valley associated with sequence boundary SB7.

Fig. 8. Examples of outcrop logs showing facies differences between the shelf and slope segments in the delta (based on Porêbski & Os-

zczypko 1999; Porêbski 1999). A Bogucice Sand, Bogucice; B Bogucice Sand, Zabawa; C Chodenice Beds, Zg³obice; D

Chodenice Beds, Gierczyce. Inset map show outcrop locations.

ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 127

tent. This may reflect autocyclic changes in progradation di-

rection or in the process regime affecting the deltas. It has

been demonstrated on high-resolution seismic data from

Pleistocene deltaic shelf margins that such internal downlap

surfaces can have a regional persistence, and reflect minor

drops and stillstands within the overall falling relative sea lev-

el (Kolla et al. 2000; Tesson et al. 2000). In the present in-

stance, no such persistence could be demonstrated for the in-

ternal downlap surfaces. Seven (B) regional downlap

unconformities that pass through the shaly intervals were

mapped throughout the area, and interpreted as main flooding

surfaces (MSF).

shelf margin as an erosion surface below a blocky sand-

stone; this surfaces passes downdip beneath steep (34°)

planar to oblique-tangential strata composed of aggrading

thick-bedded massive sandstones (Fig. 10A, SB3 between

wells Krz-2 and Pu-6). In some instances, the sandstones can

be demonstrated to onlap upslope onto this unconformity.

An erosional unconformity below a sharp-based sandstone

on the shelf can also pass on the slope into either (D) a

downlap unconformity beneath a prograding mouth-bar, or

(E) an onlap unconformity that separates a prograding

mouth-bar complex below from an aggrading to backstep-

ping one above (Fig. 11). In the latter instance, such a turn-

over surface can be marked below by a downward shift in

onlap (Figs. 11 and 12). In some instances, however, it is un-

clear, due to insufficient seismic resolution, whether the

turnover surface merges with the base or with the top of the

sharp-based sandstone.

Among the unconformities listed above, types C to E are

potential candidates for sequence boundaries (SB), because

they are associated with regional erosional truncation and/or

basinwards facies shift. Types C and D conform to the Exxo-

nian type 1 unconformity (Van Wagoner et al. 1990). Type E

(turnover surface) appears to merge upslope into the type D

unconformity. This reflects a more general dilemma as to

whether a master sequence boundary should be placed at the

base of prograding complex (Posamentier & Morris 2000), or

at its top (Plint & Nummedal 2000). The latter approach was

attempted here. Sequence boundaries SB2 and SB3 are simi-

lar in that they are associated with relatively thin, sharp-based

sandstones that tend to pinch-out rapidly landwards within

shelf shales and are underlain by thick-bedded massive sand-

stones onlapping back onto the slope (Fig. 10). Sequence

boundary SB3 appears to be associated with a valley that dis-

sected the shelf edge (Figs. 7B and 10A). SB1 could not be

traced landwards into the shelf realm because it is cut by the

frontal thrust of the Zg³obice Unit. This boundary is overlain

by thick-bedded massive sandstones and hence, is probably of

the same character as SB2 and SB3. Sequence boundary SB4

coincides with the erosive base of the Bogucice Sand tongue

representing the basal member of the Chodenice Beds (Fig.

4). Farther to the southwest, this surface follows the base of a

valley system (Skoczylas-Ciszewska & Kolasa 1959; Porêbs-

ki & Oszczypko 1999) that still further landwards cuts down

into the uplifted substratum (Garlicki 1968). The valley inci-

sion probably did not reach the coeval shelf break. On the

slope, SB4 appears to correlate with a turnover surface that

Origin of the clinoforms

The existing interpretation of the clinoform body under

consideration (e.g. Po³towicz 1997) is that it is a tectonic

monocline developed above thrust faults. This interpretation

is untenable in the light of the above evidence. In particular, it

is inconsistent with (1) the presence of the distinct offlap

breaks defining the platform-to-slope morphology and with,

(2) the offlap geometry of the entire clinoform body, (3) a re-

current major sandbody development along the prograding of-

flap breaks, and with (4) the facies spectrum that documents a

marked bathymetric gradient down the clinoforms. Hence, the

observed topography is interpreted to reflect a shelf-slope-ba-

sin plain setting, i.e., a shelf margin that evolved in response

to deltaic sediment delivery from the south. The nearly 400 m

relief of the clinoform slope represents a reasonable minimum

estimate of the paleowater column depth below the shelf

break. This is compatible with the water depth range postulat-

ed on paleoecological ground (Czepiec & Kotarba 1998). The

concentration of sand at and beyond the former offlap breaks

indicates that the clinoformed shelf margin was constructed

mainly by welding of successive deltas that periodically

spilled over the shelf edge and fed turbidites located on the

slope itself and down to its toe.

Sequence stratigraphic framework

A prevailing depositional motif in the stratigraphic architec-

ture of the Chodenice and Grabowiec clinoforms is a coarsen-

ing-upward unit interpreted to reflect progradation of a deltaic

shore (Fig. 10). Because of the predominantly geophysical na-

ture of the database available, a sequence stratigraphic frame-

work for the upper Badenian-lower Sarmatian supra-evaporite

succession was based on the nature of stratal lap outs seen of

the seismic record and on parasequence stacking patterns, tied

to the foraminiferal and nannoplankton zonation (Fig. 4).

Unconformities and sequences boundaries

Five types of unconformity have been distinguished in the

studied clinoforms (Fig. 11). The most common are (A) inter-

nal downlap unconformities that partition the clinoform into

individual stacked and laterally offset clinoform sets. Howev-

er, many such unconformities appear to have only local ex-

Fig. 11. Summary of stratal lap-out patterns derived from the seis-

mic record.

128 PORÊBSKI et al.

separates an upper backstepping mouth-bar set from a lower

progradational set (Fig. 10A). Sequence boundaries SB5 and

SB6 display a similar character (Fig. 10B). Sequence bound-

ary SB7 was mapped only in the northeastern corner of the

study area, where it appears as toplap truncation surface be-

low a major valley (Fig. 9) that breached the former shelf edge

(Fig. 7F).

Depositional sequences

In total, six 4th-order depositional sequences have been dis-

tinguished within the stratigraphic framework of upper Bade-

nian-lower Sarmatian strata (Fig. 4). Each sequence forms a

northward thinning wedge that displays a sigmoidal shape in

dip cross-section, with the thickest part located immediately at

or below the former shelf edge (Fig. 13). The six sequences

are stacked vertically to form an offlapping prism, within

which the distal pinch-outs prograded some 15 km north-

wards during the time span involved (Fig. 13).

Transgressive systems tracts (TST) overlie a regional maxi-

mum regressive surface (MRS) and are characterized by a

modest thickness (<20 m thick in the area landwards of the

shelf edge) and by major shoreline backstep (generally greater

than the preserved width of the shelf). Deposits in these tracts

include plankton-enriched shales in the basinal realm and ma-

rine mudstones and prodeltaic heteroliths on the shelf. Trans-

gressive shoreface or deltaic sands are remarkably rare. High-

stand systems tracts (HST) tend to be thick and shaly on the

shelf, but their thickness decreases upwards through the suc-

cession. Highstand deposits include mainly marine shale and

progradational deltaic deposits, and there is a remarkable lack

of paralic coastal-plain deposits within the study area. Depos-

its assigned to forced regressive systems tracts (FRST) in-

clude prograding mouth bars that show landward pinch-outs

Fig. 13. Thickness map of depositional sequence S4 and the successive locations of the shelf break during progradation of the clinoforms.

Fig. 12. Portion of a seismic dip-line showing the updip transition of onlap (turnover) onto the SB5 unconformity on the slope into the

sharp-based sandstone at the shelf edge.

ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 129

by onlap back towards the shelf and pass basinwards into

thin-bedded prodelta/slope heteroliths locally intercalated

with thick-bedded blocky sandstones. Lowstand systems

tracts are characterized within the slope either by thick-bed-

ded blocky sandstones, or aggrading to backstepping deltaic

mouth bars, or both. Thin (less than 30 m) blocky to fining-

upwards sandstones located on the shelf have been interpreted

as valley fills, although some blocky unit may in fact repre-

sent forced regressive mouth bars. The inferred valley fills

(Fig. 9) may contain either fluvial or estuarine deposits; in the

latter case, these fills should rather be included within trans-

gressive systems tract (comp. Mellere & Steel 1995).

Forced regressive systems tract

The systems tract assignment of forced regressive deposits

and the stratigraphic position of the associated sequence

boundary are still a matter of controversy (Posamentier &

Morris 2000). However, there is a growing consensus that de-

posits formed during the fall of relative sea level should not be

lumped together with those formed during the sea-level rise,

but warrant a distinction within a separate, forced regressive

or falling-stage systems tract (Hunt & Tucker 1992, 1995;

Mellere & Steel 1995; Plint & Nummedal 2000). The base of

the forced regressive systems tract should be placed at the first

downlap surface recording initial sea-level drop from its max-

imum highstand (sequence boundary sensu Posamentier &

Morris 2000; see also Posamentier et al. 1992). In the present

case, however, the proximal part of the systems tracts is not

preserved; hence, the positions of the highstand shorelines are

uncertain. Therefore, it is difficult to evaluate whether the first

observed downlap surface, or the bypass zone between down-

dip detached mouth bars, reflects the initial drop in relative

sea level from its highstand location, or a pulse within the al-

ready falling sea level, as exemplified by Pleistocene eustatic

stepwise fall (e.g. Kolla et al. 2000). Nonetheless, a distinc-

tion of the forced regressive systems tract was attempted dur-

ing well correlations (Fig. 14). For mapping purposes, howev-

er, highstand and forced-regressive deposits were included

within a single tract (HSTFRST). In the approach adopted

here, the sequence boundary is placed at the base of onlap-

ping, thick-bedded slope turbidites (Fig. 11) and where these

are absent, at the turnover from the progradational to retrogra-

dational parasequence sets within the slope realm (Fig. 10). In

both cases, the sequence boundary defined in this way can

commonly be seen to pass into the base of sharp-based sand-

stones inboard of the shelf (e.g. Fig. 14).

Depositional systems tracts and paleogeography

The lowstand systems tract in sequence S1 (LST1) forms

two depocentres (Fig. 15A). The southwestern depocentre is a

slope-toe linear wedge, 4 km wide and over 100 m thick,

which is built mainly of thick-bedded turbidite sandstones

that downlap onto the top of the Bochnia evaporites (Fig. 10).

The other depocentrum consists of similar sand turbidites that

fill and onlap from the south the SE-plunging Szczurowa Val-

ley incised into the basement (Figs. 3 and 15B). The top of the

Bochnia Formation is interpreted as a maximum flooding sur-

face (MFS0) because of the deep-water nature of the evapor-

ites (Peryt 2000) and, because they are immediately overlain

by pteropod-bearing shales that represent a deep-water basinal

facies. Sequence 1 has a very thick (75125 m) highstand and

forced-regressive systems tract (HSTFRST1) that is domi-

nated by mud-prone sediment and has a maximum flooding

surface (MFS1) associated with radiolaria-bearing shales

(Fig. 4). At its very top there occurs a prograding sandbody,

interpreted as a forced regressive mouth bar, that extends

15 km along the WNW-ESE trending shelf edge (Fig. 15B). It

wedges outs downdip into rather gently dipping slope mud-

stones that are overlain (across a sequence boundary) by hy-

perpycnal turbidites of LST2 (Fig. 15C). The turbidites form a

fan-shaped body, up 6070 m thick and 14 km wide along the

strike. The body was derived from the southeast and was in a

large part ponded in a slope mini-basin located between

growth anticlines (Fig. 15C).

The mud-prone HSTFRST2 has near its top a thin series of

prograding to downstepping mouth bars striking WNW-ESE

that wedge out downdip at the former shelf break (Fig. 15D).

Fig. 14. Interpretation of depositional systems tracts in selected dip cross-sections.

ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 133

ing valley system (Fig. 10). Only the northeastern part of this

system was mapped here (Fig. 7F), because its proximal part

lies above the upper limit of the seismic record.

Clinoform types

Two end members in the large-scale clinoform architecture

can be distinguished within the studied shelf margin (Fig. 16).

Type A clinoform is characterized by a steep, (34°) and nar-

row (24 km) slope of a relatively low height (150200 m).

Slope sediments are dominated by thick-bedded massive

sandstones that show downlapping lower ends and onlapping

upper terminations. The coeval shelf margin can locally be

dissected by distributary channels or incised valleys. Such

shelf-margin architecture appears to have developed during

the maximum fall and early rise in relative sea level.

The Type B clinoform is less steep (2°), tangential in cross-

section and wide (>12 km), and shows a considerable height

(up to 400 m). Clinoform strata are constructed mainly of

mouth bars perched on the slope. These show initially a pro-

gradational to downstepping arrangement over a distance of

610 km downdip from the former shelf edge; this pattern is

followed by a rise and landward migration of the offlap breaks

(comp. Fig. 10). Thick-bedded massive sandstones bodies are

noticeably absent, although thicker, shingled turbidites can lo-

cally be present, being associated mainly with the downstep-

ping segment of the shelf-edge trajectory. There is no evi-

dence for shelf edge dissection accompanying the

development of this type of shelf-margin architecture. The

Type B clinoforms are formed during a stepwise fall and early

rise in relative sea level.

Although no basin-floor fan was documented within the

limits of the study area, it is predicted that such fan deposits

may develop basinwards of the Type A clinoforms, provided

the initial slope downdip width was sufficiently large (larger

than that documented here) to ignite low-density turbidity cur-

rents from river-born hyperpycnal flows (see below).

Lower-order sequences

The six 4th-order sequences described are stacked basin-

wards and vertically to form an aggradational to pro-

Fig. 15. Net-sand isolith maps with paleogeographic interpretation added, showing clinoform evolution for lowstand systems tracts (LST)

and combined highstand and forced regressive systems tracts (HSTFRST). (Fig. 15 begins on page 130.)

134 PORÊBSKI et al.

gradational succession that get sandier both upwards and bas-

inwards. Two 3rd-order sequences have been distinguished

within the succession (Fig. 4). The lower sequence (Ba2, up-

per Badenian) consists of sequences 1 to 3, and its upper

boundary is defined at the base of valley system that underlies

the Bogucice Sand tongue. The upper sequence (Sa1; lower

Sarmatian) consists of sequences 4 to 6; its top has tentatively

been placed at the surface SB7 associated with a valley system.

Sequence Ba2 consists of Type A clinoform bundles in

which the offlap break rises steeply basinwards throughout

the development of the component high-order sequences

(Fig. 10). This is associated with thick, mud-prone transgres-

sive and highstand systems tracts landwards of the shelf

break, with the highstand tracts commonly capped by thin

forced regressive mouth-bar/shoreface sands. Lowstand tracts

are typified by thick massive turbidites deposited on a narrow,

steep slope, and their maximum strike extent is 1430 km. In

some instances (LST3 and possibly LST2), the slope turbid-

ites appear to be linked up dip to valleys that breached the

shelf edge (Fig. 15C). The slope is locally affected by anti-

clines that grew above the tips of blind thrusts. The resultant

mini-basins perched on the slope acted as traps for hyperpyc-

nal turbidites and mouth bars spilling over the shelf edge.

From sequence boundary SB4 on, the shelf break recurrent-

ly assumes a noticeably flatter trajectory basinwards (Fig. 10),

and the shelf margin grew mainly through the accretion of

Type B clinoform bundles. Landwards of the shelf break,

transgressive and highstand systems are either thin, or absent,

whereas prograding to downstepping mouth bars form the

bulk of forced regressive systems tracts. Slope-perched mouth

bars that aggraded and stepped back onto the shelf typify low-

stand tracts. Although the feeder system was intermittently

able to arrive close to the shelf edge, there is evidence of a

major incision into the shelf edge only for one sequence

boundary (SB7) (Fig. 7F).

Discussion and conclusions

The best-known examples of shelf-margin or deep-water

deltas (Edwards 1981) are those described from the Pleis-

tocene of the modern continental shelf edges (e.g. McMaster

et al. 1970; Suter & Berryhill 1985; Tesson et al. 1990, 2000;

Kolla et al. 2000). It is commonly believed that deltaic progra-

dation across the shelf to its margin is caused by either forced

regression during relative sea-level fall (e.g. Posamentier et al.

1992), or anomalously large siliciclastic sediment influx

(Poag & Sevon 1989).

It is suggested here that the bathymetric step for such deep-

water deltas does not need to be the topographic break inherit-

ed after extensional faults or rifting, as is commonly the case

for modern divergent margins. Such a topographic or shelf-

edge break can also be the constructional brink where succes-

sive shelf-transiting deltas encounter and descend into a deep

basin. This type of constructional shelf edge is created by re-

peated regressivetransgressive transit of deltas. The shelf has

its maximum width during sea-level highstand, and can have

zero width at lowstand, when delivery of sediment causes re-

newed shelf-margin accretion.

In the present case, the initial shelf break was nucleated

along a thrust or thrust-generated anticlinal front located

southwards from the present basin margin. During the earliest

Badenian, this front was probably situated some 30 km south

of the Carpathian overthrust, as suggested in the palinspastic

Fig. 16. Summary of main types of clinoform architecture.

ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 135

restoration by Oszczypko (1997, his Fig. 15). This may possi-

bly represent a crude estimate of the position of highstand

shorelines with respect to the coeval shelf margin. The latest

Badenian-early Sarmatian thrusting resulted in growth anti-

clines affecting the clinoform body, which led to a steepening

of its slope. A strongly aggradational character to shelf-mar-

gin deltas, as observed in the lower clinoform set (sequence

Ba2), appears incompatible with the striking absence of paralic

deposits. This is because a climbing clinoform trajectory caused

by relative sea-level rise tends to generate space behind, where

a thick tail of paralic deposits can be accommodated. Howev-

er, the absence of such a tail in the present instance reflects a

modest to negative accommodation behind the rising thrust

margin, whereas subsidence in the hangingwall created

enough space to accommodate steep and high clinoforms.

Shelf-margin deltas appear to record periods when the sedi-

ment is stored at the shelf edge rather than is being delivered

into the deep water. The difference between the two clinoform

end-member types distinguished here seems to have been con-

trolled chiefly by rate of relative sea-level change that gov-

erned the position of the seaward end of the feeder system

with respect to the former shelf break, although differences in

depositional regimes on the outer shelf could also have played

a significant role (cf. Steel et al. 2000). The sand-prone Type

A clinoforms appear to record a situation when the feeder sys-

tem incised across the shelf edge in response to a rapid rela-

tive fall in sea level. This caused cannibalisation of the previ-

ous shelf deltas and promoted fluvial input directly onto the

slope. The resultant hyperpycnal flows came to rest on the

slope itself, probably because the small downdip width of the

slope and the presence of perched basins prevented flow igni-

tion into high-efficiency turbidity currents capable for trans-

porting sand into the basin-plain area. The composite, het-

erolithic Type B clinoforms record initially rather slow and

stepwise fall in relative sea level. Incision of the feeder-chan-

nel system progressed gradually behind the prograding deltaic

shelf margin (comp. Sydow & Roberts 1994), but was appar-

ently unable to breach the freshly accreting shelf edge even

during the maximum lowstand. During early sea-level rise,

aggradation followed by backstepping of deltas took place.

Therefore, the development of Type B clinoform sets docu-

ments trapping of sand within deltaic depocentres that shifted

forth and back on the slope and, consequently, a decreased

probability for sand to escape into the basin-floor setting.

The change from the steeply rising to a more horizontal tra-

jectory of the shelf margin, together with the compressional

setting of the studied clinoform, points to a major change in

the tectonic component in the relative sea-level signal during

the clinoform growth. The particularly strong aggradation and

the large thickness of the upper Badenian highstand systems

tracts (sequences 1 to 3) point to strong differential move-

ments across a N-verging thrust-fault margin, that generated a

long-term relative sea-level rise on the hangingwall. The peri-

od of fault-induced subsidence ended with the latest Bade-

nian-early Sarmatian uplift to the southwest. Since then, high-

frequency relative sea-level fluctuations were superposed on a

less vigorous regional flexural subsidence pattern, that result-

ed in faster accretion and increased width of the clinoforms

(sequences 4 to 6).

Acknowledgments: The study was supported by a State

Committee of Scientific Research (KBN) Grant T12B05218

to SJP, and by Landmark Graphics Corporation via the Land-

mark University Grant Program to Institute of Geological Sci-

ences, Polish Academy of Sciences. We thank the Polish Oil

and Gas Co. for permission to use the subsurface data and, in

particular, to Eugeniusz Jawor and Tadeusz Wilczek for their

encouragement to undertake this study. The final manuscript

benefited from critical and insightful review by Guy Plint and

André Strasser.

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