GEOLOGICA CARPATHICA, 54, 2, BRATISLAVA, APRIL 2003
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;
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.
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.
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).
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.
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.
ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 121
Fig. 3. Depth-map of the top of cratonic basement below the offlaping Miocene succesion. The basement is dissected by the southeasterly
plunging valleys (dotted) and is affected by extensional and strike-slip faults. The faults in the southwestern corner formed ramps for
thrusts during the latest Badenian compression.
Fig. 2. Index maps showing the locations of seismic lines, wells and correlation sections.
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.
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-
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-
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-
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;
Sands include both classical and high-density turbidites,
deltaic mouth bars, and shell-bearing, cross-bedded shoreface
sandstones (Porêbski & Oszczypko 1999; Porêbski 1999)
There is little, if any, evidence of coastal plain (paralic) fa-
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.
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.
126 PORÊBSKI et al.
shoreward aspects. This upward shallowing was associated
with a decreasing salinity from normal marine to hyposaline
levels, indicative of increased flux of freshwater (Czepiec &
Kotarba 1998). The organic matter is dominated by type III
kerogen (Kotarba et al. 1998), consistent with the postulated
overall deltaic setting of the succession.
Evidence of syndepositional deformation
In the southeastern part of the study area clinoform slope is
affected by two, parallel WNW-ESE striking blind thrusts
with growth anticlines located above faults tips (Fig. 6). The
thrust appears to flatten out southwards in a decollement plane
merging into the evaporite unit. These anticlines define two,
strike-elongated mini-basins perched on the slope. The basins
are 11.5 km wide and 80100 m deep, with their axes plung-
ing gently eastwards (Fig. 7AC). Fills include thick-bedded
blocky sandstones and localized mouth bars. The formation of
these growth structures appear to have post-dated the filling
up of the Szczurowa Valley at its southern reach and came to
an end before the development of unconformity SB5 (Figs. 4
Fig. 10. Examples of log-based dip correlations with environmental interpretation added. Note thick-bedded, blocky sandstones in
slope segments below sequence boundary SB4 and the downward-stepping to back stepping mouth bars perched on the slope.
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
(C) Another type of unconformity is expressed on the
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
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-
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
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.
130 PORÊBSKI et al.
LST3 is typified by a thick (over 100 m) accumulation of
thick-bedded hyperpycnal turbidites that form an aggrading
linear body, up to 30 km in strike dimension on the slope
(Fig. 15E). Part of this body was trapped in a mini-basin on
the slope behind a growing anticline, and extends updip for
5.5 km (comp. Fig. 7B). At, and landwards of the WNW-ENE
striking shelf edge, there occur a series of aggrading to retreat-
ing mouth bars that farther to the SW appear to pinch out by
onlap within shelf shales. A valley fill, up to 30 m thick, un-
derlies these mouth bars, and appears to have breached the
shelf break near its western end (Fig. 15C).
HSTFRST3 has a more heterolithic aspect than the previ-
ously described highstand to forced-regressive systems tracts.
Within the shelf realm these heteroliths coarsen upwards into
a sandstone whose top is planar to slightly descending basin-
wards (Figs. 10 and 15F) that is characteristic for forced re-
Fig. 15A,B,C. Continued on pages 131, 132 and 133.
ORIGIN AND SEQUENTIAL DEVELOPMENT OF CLINOFORMS 131
gression (Plint & Nummedal 2000). This sand belt trending
WNW-ESE is formed of prograding mouth bars that descend-
ed onto the slope and were in part ponded in the slope mini-
basin (Fig. 15F). A NNE-SSW striking finger-like sandbody
in the SW part of the area may represent the fill of a distribu-
tary channel or valley. The mouth bars are separated basin-
wards by muddy slope sediments from another sand belt inter-
preted as slope-toe thin-bedded turbidites. Both sand belts,
together with the intervening slope mudstones, are interpreted
as forced regressive deposits.
The overlying LST4 consists of a series of NW-striking
mouth bars that are arranged in a backstepping fashion, and
reveal the highest net sand content just landwards of the
former shelf break, and are partly trapped behind rising anti-
clines piercing the slope in the eastern part of the area (Figs.
15G and 7C). At their landward onlap termination, the bars
Fig. 15D,E,F. Continued on pages 132 and 133.
132 PORÊBSKI et al.
appear to be dissected by channel feeders. West of the study
area, this unit is based by a major incised valley system
(Porêbski & Oszczypko 1999), which apparently did not
reach the former shelf edge. HSTFRST4 is thin and sandy
landwards of the shelf edge and could in fact be entirely clas-
sified as a forced regressive systems tract. This unit is typified
by an E-W-trending perched mouth bar that descended for 5
6 km down the slope; its landward pinch-out is located just
beyond the shelf edge (Fig. 15H). The bar is separated by a
1011 km bypass zone from the preceding (highstand?)
mouth-bar edge (Fig. 15H).
LST5 in turn, displays mouth bars that aggraded on the
slope and retreated 89 km from the newly formed shelf edge
(Fig. 15I). HSTFRST5 displays a narrow belt of mouth bars
located just around the shelf break (Fig. 15J). LST6 is typified
by backstepping mouth bars that are backed updip by finger-
like N-S to NW-ESE trending fills of feeder channels
(Fig. 15K). This tract is bounded at its top by a NE-SW strik-
Fig. 15G,H,I. Continued on page 133.
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.
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).
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
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