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, AUGUST 2013, 64, 4, 305—326 doi: 10.2478/geoca-2013-0022
Large-volume gravity flow deposits in the Central
Carpathian Paleogene Basin (Orava region, Slovakia):
evidence for hyperpycnal river discharge in deep-sea fans
DUŠAN STAREK
1
, JÁN SOTÁK
2
, JOZEF JABLONSKÝ
3
and RÓBERT MARSCHALKO
1
1
Geological Institute, Slovak Academy of Sciences, Dúbravská cesta 9, 842 28 Bratislava, Slovak Republic; dusan.starek@savba.sk
2
Geological Institute, Slovak Academy of Sciences; Branch: Ďumbierska 1, 974 01 Banská Bystrica, Slovak Republic; sotak@savbb.sk
3
Department of Geology and Paleontology, Faculty of Natural Sciences, Mlynská dolina 15, 842 15 Bratislava, Slovak Republic;
jozef.jablonsky@fns.uniba.sk
(Manuscript received September 3, 2012; accepted in revised form March 14, 2013)
Abstract: The deep-water clastic systems of the Central Carpathian Paleogene Basin contain megabeds, which are devel-
oped in distinctive stratigraphic horizons and can be traced over long distances. These beds are characterized by great
individual thickness (4—13 m), uniform lithology and internal structures. On the basis of their lithology, sedimentary
structures and sequence development, the megabeds are characterized by 15 individual facies and interpreted from the
viewpoint of flow hydrodynamics. The grain-size distribution and internal structures of the megabeds point to their depo-
sition from uniform turbulent flows. The main controlling factor for generation of such large voluminous flows is inferred
in the sea-level changes, when a relative rising of sea level during the Eocene/Oligocene boundary was responsible for
long-lasting accumulation of the clastic supply at the basin margins. The large volume of detritus from river discharge and
ravinement surfaces of flooded land was accumulated on the shore and in the conduit heads where the sediment was
remobilized by other triggers. The flows generated by catastrophic floods during the early Rupelian sea-level lowstand are
thought to be the most probably triggering mechanism. The large highly erosive hyperpycnal flows from flooding rivers
could erode accumulated deposits in the conduit or on steeper basin-margin slopes and could cause progressive increase of
the sand volume in the flow. Conduit flushing appears to be the most probable source of sediment for the very large
voluminous flows that were responsible for deposition of the Orava megabeds.
Key words: Early Oligocene, Central Carpathians, megabeds, megaturbidites, hyperpycnal-flow deposits.
Introduction
Specific sedimentary units in the western part of the Central
Carpathian Paleogene Basin (CCPB) are represented by
“megabeds”. They form a distinctive horizon of basin-fill
formations with a great thickness (4—13 m), uniform litholo-
gy and finely structured sandstones and siltstones, which are
developed in fining-upward beds. In the Orava region of the
CCPB, the megabeds display a relatively broad lateral extent
and distribution over a distance of 25 km (Fig. 1C). More-
over, similar occurrences of very thick sandstone beds from
the identical stratigraphic level of other parts of the CCPB
are recorded in the Podhale area (Kozinec beds sensu Golab
1959) and Spišská Magura area (Janočko & Jacko 2001;
Soták et al. 2001; Sliva 2005).
Descriptions of megabeds, often referred as “megaturbid-
ites” (e.g. Ricci Lucchi & Valmori 1980; Mutti et al. 1984;
Bouma 1987; Labaume et al. 1987; Reeder et al. 2000;
Remacha & Fernández 2003 and others), have been presented
over the last decades in several studies. The megabeds with a
great thickness and wide lateral extent, are considered to be a
product of large-volume gravity flows ( > 2.5 km
3
sensu
Talling et al. 2007). They have been recognized mainly in
acoustic and high-resolution seismic profiles (e.g. Piper et al.
1988; Rebesco et all. 2000; Reader et al. 2000; Anastasakis
& Pe-Piper 2006; Gee et al. 2006), drill cores (Reeder at al.
2000; Piper & Normark 2001 ) or in well exposed areas (e.g.
Ricci Lucchi & Valmori 1980; Remacha & Fernández 2003;
Talling et al. 2007; Muzzi Magalhaes & Tinterri 2010).
However, the deposition of a very thick bed by a single sed-
imentary event is relatively rare in deep-marine turbidite sys-
tems (Piper & Normark 2001). Furthermore the processes
which can trigger voluminous turbidity currents able to influ-
ence depositional processes are still poorly understood and
discussed (e.g. Piper & Normark 2009). The megabeds might
be important as excellent markers for stratigraphic and seismic
correlations over long distance, and useful in basin analyses
(e.g. Doyle 1987; Pauley 1995; Remacha & Fernández 2003).
This study deals with the Orava megabeds in the Central
Carpathian Paleogene Basin, with the aim of providing
petrofacies, lithological and stratigraphical data and detailed
descriptions of their sedimentary structures. These data en-
able us to recognize 15 individual facies characterizing the
megabeds, with grain-size distribution, sequence organiza-
tion and lateral transformation also reflecting the hydrody-
namic conditions of their deposition. Our data in comparison
with modern analogues of megaturbidites and other mega-
beds have been used for definition of the most probably
model of large-volume flow initiation and deposition of
megabeds in the western part of the CCPB.
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Background
Initiation processes for turbidity currents
Turbidity currents may be initiated by a wide range of pro-
cesses at the edge of the continental shelf, which are con-
nected generally with transformation of slumps, hyperpycnal
flows from rivers, storm generated flows near the shelf edge,
or combinations of these initiating processes (e.g. Normark
& Piper 1991; Piper & Normark 2009). A very simplified
description of these general processes is presented below.
Initiation by slope failure: The turbidity currents might
be initiated by slope failures immediately in the seaward
areas of a distributary mouth, where predominantly sandy
sediments are accumulated. Such failures are particularly
common in fan deltas (e.g. Hein & Syvitski 1992; Lrnne &
Nemec 2004; Carter et al. 2012). The form of sand failure
depends on the efficiency of packing in the original sand
body, which is a critical parameter for division of large bank
failures into two types: liquefaction slope failures and breach
failures (Van den Berg et al. 2002). Slump-generated turbidity
currents show particular down-flow evolution. The sand and
silt of the initial failures is essentially unlithified to form
slumps and debris flows, which overcome sufficient distance
and may be transformed through hydraulic jumps into turbi-
dity currents on the steep slopes (e.g. Piper et al. 1999; Mohrig
& Marr 2003).
Initiation by sediment-laden flows: Turbidites formed by
direct freshwater sediment-laden flows are characteristic for
steep and high-bedload rivers. These flows discharge on the
narrow shelves, or prograde onto the basinal slopes to accu-
mulate the fan deltas. Some sediments were deposited near
Fig. 1. A – Location of the study area within the Alpine-Carpathian orogen. B – The Central Carpathian Paleogene Basin system depicting
structural sub-basins, basement massifs and surrounding units. C – Geological sketch of the Orava region (after Gross et al. 1993; Biely et
al. 1996, modified) with situated megabeds studied. Key: 1 – location of studied sections (A – Zázrivá (N49°15
’11”, E19°10’29”);
B, C, – localities near Dolný Kubín area (Ve ký Bysterec – N49°12
’24”, E19°17’05”, Záskalie locality – N49°12’51”, E19°17’30”);
D – Kňažia (N49°13
’55”, E19°19’45”); E – Pucov (N49°13’14”, E19°22’43”); F – Ve ké Borové (Svorad plató section – N49°11’06”,
E19°29
’48”); G – Jobová Ráztoka (N49°12’10”, E19°32’1”); H – Huty (N49°12’27”, E19°32’56”); 2 – Mesozoic of the Inner
Carpathians (undivided); 3 – Mesozoic of the Klippen Belt (undivided); 4 – Cretaceous and Paleogene of the Outer Carpathians (undivided);
5 – Magmatic rocks (Tatricum basement); the Central Carpathians Paleogene (6 – Borové Formation; 7 – Huty Formation; 8 – Zuberec
Formation; 9 – Biely Potok Formation); 10 – geological boundaries, faults and overthrust lines.
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the mouth and subsequently reworked, but most hypercon-
centrated coarse bedload would continue to flow seaward
inertially (Prior & Bornhold 1989). Turbidity currents in fan
deltas may often result from hyperpycnal flow (e.g. Mulder
& Syvitski 1995; Mutti et al. 2000; Piper & Normark 2001;
Mutti et al. 2003), although submarine landslides and remo-
bilization of sediment are also widespread (e.g. Hein &
Syvitski 1992; Carter et al. 2012).
Initiation by storms: These turbidity currents are initiated
by storm resuspension of sediments near the shelf edge,
mainly when the heads of the submarine canyons occur in
the surf zone, from which they cross the narrow shelves. Sus-
pension of sand and activation of seaward flows is considered
to be the main initiation process of turbidites in canyon heads
during storms (e.g. Fukushima et al. 1985; Prior et al. 1989;
Mulder et al. 2001).
The multiple initiating processes: In submarine conduit
heads and sandy deltas, multiple initiating processes for tur-
bidity currents may interact in a complex manner. Erosion of
deltafront channels or straight-sided prodeltaic channels by
hyperpycnal flows can initiate failure and flushing of earlier
deposited sands (e.g. Mitchell 2005).
Moreover, the deposition of sand during extreme floods is
important on the prodelta, which then might fail due to depo-
sitional oversteepening (Milliman et al. 2007). Many hyper-
pycnal flows from rivers are predominantly muddy (e.g.
Johnson et al. 2001), but may evolve into mixed flows by
erosion of older sands from the conduit.
The initiating mechanisms and factors that trigger the dep-
osition of large volumes and thicknesses of sediment in a
single event are still a matter of discussion in literature. For
example, the relationship between large-volume turbidity
currents and sea-level changes is still poorly known. Several
authors have suggested that a sea-level lowstand is a decisive
factor for a non-steady stage, when terrigenous debris directly
reaches the shelf margin, thus causing instability and sus-
ceptibility to sliding (e.g. Vail et al. 1977; Shanmugam &
Moiola 1982, 1984; Reeder et al. 2000). Unstable environ-
ments can be influenced by different trigger mechanisms,
which cause catastrophic collapse of the basin margin sedi-
ments and lead to the initiation of large-scale gravity flows.
Similarly, when the sea level was at a maximum lowstand,
Yose & Haller (1989) coupled the major collapse of outer
ramps and upper slopes with the impingement of a storm
wave base. The lowering of sea level resulted in prograda-
tion and cannibalization of the delta complex and a large
amount of sand was transported by turbidity currents to the
distal part of the basin (e.g. Postma 1995; Normark et al.
2006). On the contrary, Marjanac (1996) attributed the depo-
sition of megabeds in the Eocene-Miocene flysch formations
of the Central Dalmatia to the periods of accelerated sea-level
rise, when the ground-water table rises, tides amplify, wave
action increases, and the increasing pore-water pressure may
provide favourable conditions for slope instability and col-
lapse of the slope. Similarly, Ricci Lucchi (1990) reported
great thicknesses of turbidites during highstand sea level.
Rothwell et al. (2000) proposed that the generation of mega-
turbidites is more influenced by the rate of the sea-level
changes than the amplitude of these changes.
Many authors regard seismic activity as a dominant trigger
mechanism that causes upslope, large failures of the sedi-
ments. However, other important mechanisms (or their inte-
gration) that have also been suggested for triggering of
large-volume gravitation flows and deposition of megabeds
are connected with volcanic eruptions and earthquake activity
(Anastasakis & Pe-Piper 2006; Anastakasis 2007), meteoric
impacts (Iturralde-Vinent 1992; Dypvik & Jansa 2003), tsu-
nami wave impacts (Cita et al. 1984; Cita et al. 1996; Heike
1984), the release of buried clathrates (Bugge et al. 1987), the
over-supply and under-consolidation of sedimentary material
(Doyle & Bourrouilh 1986), and also the stresses produced by
tidal or river flow (Lowe 1976; Piper et al. 2007).
Geological setting
The CCPB lies inside the Western Carpathian Mountain
Chain (Fig. 1A) and belongs to the basinal system of the
Peri- and Paratethyan seas. The basin accommodated a
forearc position on the destructive Alpine-Carpathian-Pan-
nonian microplate margins and in the hinterland of the Outer
Western Carpathian accretionary prism (Soták et al. 2001).
The basin is mainly filled by turbidite-like deposits, which
overlap the substrates of the pre-Senonian nappe units and
their thickness reach up to a thousand meters. The age of the
formations ranges from the Bartonian (e.g. Samuel & Fusán
1992; Gross et al. 1993) to the latest Oligocene (cf. Soták
1998; Olszewska & Wieczorek 1998; Gedl 2000; Soták et al.
2001, 2007) (Fig. 2). The sediments of the CCPB are pre-
served in many structural sub-basins, including the Žilina,
Rajec, Turiec, Orava, Liptov, Podhale, Poprad, and Hornád
Depressions (Fig. 1B). In the study area, the CCPB sedi-
ments are bounded by the Central Carpathian units in the
south, while the northern boundary is represented by the Pie-
niny Klippen Belt (Fig. 1C), which represents a transpres-
sional strike-slip shear zone related to a plate boundary
(Csontos et al. 1992; Ratschbacher et al. 1993; Potfaj 1998).
The CCPB deposits are commonly divided into four forma-
tions (Gross et al. 1984; Fig. 2). The lowermost, Borové Forma-
tion consists of breccias, conglomerates, polymictic sandstones
to siltstones, marlstones, organodetrital and organogenic lime-
stones. These represent basal terrestrial and shallow-marine
transgressive deposits (Marschalko 1970; Kulka 1985; Gross
et al. 1993; Baráth & Kováč 1995; Filo & Siráňová 1996, 1998;
Bartholdy et al. 1999). This formation is overlain by the Huty
Formation, which mainly includes various mud-rich deep-ma-
rine deposits (e.g. Janočko & Jacko 1999; Soták et al. 2001;
Starek et al. 2004). The overlying sediments of the Zuberec
and Biely Potok Formations consist of rhythmically bedded
and massive sandstones, which represent the various facies as-
sociations of sand-rich submarine fans (Soták 1998; Janočko
et al. 1998; Starek et al. 2000; Starek 2001; Soták et al. 2001).
The megabeds of the Central Carpathian Paleogene Basin
occur in the basal part of the Huty Formation and in the
south-western and southern parts of the Orava region (there-
fore they are informally termed the “Orava megabeds”).
Eight outctops were studied during field research (Figs. 1C
and 3). The major outcrops occur near Dolný Kubín (Ve ký
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Bysterec and Záskalie locality), and in the
bedrock of the Prosiek creek near Ve ké
Borové village (Svorad plató section).
Additional megabeds occur near Pucov,
Kňažia, Huty, Jobová Ráztoka and
Zázrivá villages.
The megabeds appear in the lowermost
part of the Huty Formation, where they
developed from the menilite-type clay-
stones directly above the pre-transgres-
sive and transgressive sediments (e.g. in
Svorad plató section on the northern slope
of the Chočské vrchy Mts – Fig. 3F).
The menilite interval is represented by
non-calcareous,
or
weakly-calcareous
dark mudstones that contain tuffite-bear-
ing beds, biosiliceous horizons (cherts),
breccia beds, and millimeter-scale interca-
lations of light carbonate-rich mudstones.
This interval grades up to grey-calcareous
claystones with siliciclastic megabeds
(Fig. 4B). In another sections, the Globi-
gerina marls and menilites of the Huty
Formation are reduced, and the laminated
sandstone megabeds directly overlie the
Borové Formation (e.g. in Jóbova Ráztoka
and Huty sections – Fig. 3G,H). Near
Pucov (Fig. 3E) and Zázrivá villages, the
megabeds occur within light calcareous
claystones that overlie conglomerates,
which are in some places over 100 m
thick (Gross et al. 1982; Soták et al. 2007;
Starek et al. 2012). The sections near Dolný
Kubín (Fig. 3B,C) pass through mud-
stone-dominated sequences, which are in-
tercalated by turbidite bed-sets, sandstone
megabeds (Fig. 4A,C) and breccias.
The biostratigraphy and stratigraphic
position of the Orava megabeds indicates
that they belong to a distinct marker inter-
val in the basal part of the Huty Forma-
tion, which spread over a great area of the
basin. The megabeds show uniform thick-
ness, usually at a distance of tens of
meters within the outcrops. In spite of the
possibility that these beds can have
a great lateral continuity it is often impos-
sible to correlate the megabeds to each
Fig. 2. Descriptive lithostratigraphy of the filling in the western part of the Central Car-
pathian Paleogene Basin. Nomenclature of the formations according to Gross et al.
(1984, adapted). Biostratigraphy is based on the data from Olszewska & Wieczorek
(1998), Starek et al. (2000), Starek (2001), and Soták et al. (2007).
Fig. 3. Representative logs of mudstone-domi-
nant sedimentary facies with position of mega-
beds in the lowermost part of the Huty
Formation. B – Ve ký Bysterec; C – Záskalie
(localites near Dolný Kubín); D – Kňažia;
E – Pucov; F – Ve ké Borové (Svorad plató);
G – Jóbova Ráztoka; H – Huty. For localization see Fig. 1C. Key: 1 – conglomerates, sandstones and organodetrital limestones of the
Borové Formation; 2 – conglomerates and breccias; 3 – megabeds; 4 – repetitively thin-bedded sequence; 5 – laminated limestones
(Tylawa-type); 6 – brown, dark grey to black non-calcareous claystones (Menilite horizon); 7 – grey calcareous marls.
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other, because of the scattered nature of the key outcrops.
The impossibility of correlation of individual beds over long
distance, as well as poor knowledge about the paleotopogra-
phy of the Central Carpathian Paleogene Basin due to post-
Paleogene inversion, uplift and tectonic disintegration of the
basin does not allow us to estimate the total volume of sedi-
ment transported during depositional events of individual
megabeds. We can provide at this stage only very crude esti-
mation of the minimum volume of individual megabeds in
the Orava region of the CCPB which varies between 1.2 to
2.8 km
3
.
Results
Sedimentary facies
The Orava megabeds are predominantly formed by sand-
stones and siltstones with varying grain size. They are devel-
oped in the form of fining-upward beds from a coarse-grained
base to fine-grained sandstones, siltstones and gradually to
claystones. One of the most important features of “mega-
beds” is their great thickness (Figs. 3, 4, 5), which is consid-
erably greater than that of time-equivalent stratigraphic
intervals in other parts of the depositional basin. The deter-
mination of the total thickness of the basal sandstone part of
the megabeds was difficult due to a gradual transition into
the overlying mudstones. We have defined the upper limit of
the megabeds usually to the base of the overlying bed that
markedly differs in grain size or to the base of occurrence of
the light calcalerous hemipelagic mudstones. The bed thick-
ness of this predominantly sandstone—siltstone—claystone se-
quence varies between 4—12 m near Dolný Kubín, and reaches
up to 13 m near Ve ké Borové and Huty.
The description and classification of the Orava megabeds
is mainly based on descriptive parameters such as grain size,
roundness, sorting, grain fabric, and sedimentary structures.
In this analysis the grain-size and textural classification was
done according to Blair & McPherson (1999). We used
Powers (1953) standard index for shape classification of par-
ticles. The conglomerates are referred to as lithofacies C; the
sandstones are referred to as lithofacies S, the siltstones as
lithofacies Si, and mudstones as lithofacies M. Within the
megabeds, 15 individual facies with their possible hydrody-
namic interpretation were distinguished. Sedimentary facies
have been depicted on the basis of the following hierarchical
approach: (i) lithofacies based on lithology and (ii) subfacies
distinguished on the basis of sedimentary structures.
Fig. 4. Field exposures of megabeds. A – Finely structured part of the megabed at the Ve ký Bysterec locality. Very thick laminated sand-
stones are dominant and they pass continuously to the siltstone in the uppermost part of the bed. B – Laminated sandstone part of up to
13 m thick megabed at Svorad plató (near Ve ké Borové). C – Bed about 430 cm thick at Dolný Kubín locality.
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C lithofacies: Conglomerate
C
1
subfacies: Massive, non-graded conglomerate
Description: Pebble to fine cobble, poorly sorted conglom-
erate which exhibits a disorganized texture (Fig. 6). They are
compounded exclusively, or most prevailingly, by very angu-
lar to subangular clasts. The fabric varied widely from matrix-
to clast-supported and generally, this subfacies has poorly
sorted gravelly, sandy and muddy matrix (Fig. 6B). The con-
glomerate contains some “intraclasts” of dark non-calcareous
claystones, which are irregularly scattered in a matrix-mixture
of clay, silt and sand. Coherent, plastically deformed frag-
ments of decimeter- to meter-scale blocks of sediments are
a component of the C
1
subfacies (Figs. 6A, 11). The thickness
of this facies varies between 110 and 180 cm.
Interpretation: The internal composition of coarse-grained
C
1
subfacies and the presence of deformed blocks of sedimen-
tary beds indicates deposition from dense flows with visco-
plastic behaviour (non-Newtonian flows). Very similar facies
are commonly referred to as debris-flow deposits or debrite
(F2 facies after Mutti et al. 2003). This type of flow probably
resulted from submarine landslide or multiple landslide
phases
(Gee et al. 2006), when large
slabs of sediment could be de-
tached, fragmented, and were capable of travelling
enormous
distances (Bugge et al. 1988; Masson et al. 1993). Erosion
and disintegration of coherent blocks of sediment as well as
Fig. 5. Schematic log of megabeds. An explanatory text to the individual subfacies is given in the chapter “Sedimentary facies”. A, C – Mega-
turbidites from Dolný Kubín localities. B – Megaturbidites from Svorad plató, Huty, Jóbova Ráztoka localities. D, E – Complex megabeds
from Dolný Kubín localities. F – Complex megabed from Čremoš (Zázrivá).
erosion of the substrate during downslope movement would
produce
debris-flow matrix, which may enable the blocks to
become buoyant and they can be rotated and plastically de-
formed during failure (Lastras et al. 2002; Gee et al. 2006).
C
2
subfacies: Unstructured ungraded to normally graded
conglomerate
Description: Very poorly to poorly sorted, clast-supported,
ungraded to normally graded, medium to coarse pebble con-
glomerate which can form a scoured unit (Fig. 11). This sub-
facies is characterized by angular to subrounded clasts. C
2
subfacies range in thickness from a few cm to 25 cm.
Interpretation: The basal parts of the megaturbidites,
which are composed of a decimeter thick interval of the C
2
subfacies, can be described as R
3
sequence (after Lowe
1982) or F3 facies (after Mutti et al. 2003), which could rep-
resent a record of frictional freezing at the leading edge of
gravelly dense flow.
C
3
subfacies: Unstructured to stratified inversely graded
conglomerate
Description: Poorly sorted, clast-supported, inversely-
graded, fine to coarse pebble conglomerate. This subfacies is
characterized by angular to subrounded clasts. C
3
subfacies
ranges in thickness from a few cm to about 20 cm.
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Fig. 6. A – Coarse-grained, poorly sorted, disorganized conglomerate with coherent fragments of older sedimentary rocks (C
1
subfacies) (Ve ký
Bysterec). B – Matrix- to clast-supported conglomerate with very angular to subangular clasts (detail view of C
1
subfacies). C – C lithofacies
from outcrops near Dolný Kubín contain a large amount of biofragmental particles (detrital components of the Borové Fm). Enlargement 34 .
Fig. 7. Multiple intervals of a parallel lamination (Ve ký Bysterec). A – The detail of parallel-laminated division (S
3
subfacies) with laminae
ordered into “bands”. B – In the middle of the S
3
subfacies, a division of “sinusoidal lamination” or “ripple laminae in phase” ca. 40 cm thick
is developed (S
4
subfacies).
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Interpretation: Grainflow or highly concentrated turbidity
current. Equivalent to R
2
sequence (after Lowe 1982), which
represents the deposits of gravelly high-density flows. Inverse
grading is a result of deposition followed by a traction carpet
stage. The C
3
as well as C
2
subfacies may be interpreted as
the high-concentration flow observed at the base of many tur-
bidity currents (Lowe 1982; Ghibaudo 1992; Mutti 1992).
S lithofacies: Sandstone
S
1
subfacies: Unstructured ungraded to normally graded
coarse-grained sandstone
Description: Poorly sorted, usually ungraded, sometimes
with normal grading, coarse-grained sandstones sometimes
with dispersed granule to pebble-size clasts. S
1
subfacies
ranges in thickness from a few cm to 40 cm.
Interpretation: A rapid accumulation of coarse-grained
sands from dense sandy to gravelly turbidity current which
bypass the zone of deposition of the preceding gravelly flows
(S
3
flow type after Lowe 1982 or F5 facies after Mutti 1992).
S
2
subfacies: Stratified, inversely graded coarse-grained
sandstone
Description: Poorly- to moderately sorted, crudely hori-
zontally-stratified, pebbly/granule coarse-grained sandstones
which form an interval up to 30 cm thick. Small pebbles and
granules are angular to subrounded.
Interpretation: This subfacies indicates traction-carpet
deposition (S
2
flow type after Lowe 1982). S
2
subfacies may
be interpreted as the deposit of the dense sandy to gravelly
flow (Ghibaudo 1992; Mutti 1992).
S
3
subfacies: Parallel-laminated medium- to fine-grained
sandstone
Description: This subfacies is formed by a fining- and gen-
erally thinning-upward series of thin medium- to fine-grained
Fig. 8. The facies of “ripple” laminated sandstones at Ve ký Bysterec locality (Dolný Kubín). A – “Ripple-drift cross lamination“ (S
5
sub-
facies) characterized by its height angle of climb and by the complete preservation of bedforms. B – Vertical changes in lamination from
parallel lamination (S
3
subfacies) to ripple-drift cross lamination (S
5
subfacies) and ripple laminae in phase (S
4
subfacies).
sandstone laminae divided into discrete sets or “bands”
(Fig. 7A). The thickness of these bands ranges from about
10 cm (at the base of the subfacies S
3
) to 2 cm (at the top).
Each sandstone bands consist of three to nine horizontal lami-
nae with small thickness (0.3—1.3 cm) and relatively constant
development. The laminae are inversely graded in the lower
part and structureless in the upper portions. In some places,
the laminae boundaries are highlighted by a concentration of
pyrite, and locally also by accumulations of heavy minerals
(predominantly zircon). Grain-size analysis of the sandstone
laminites shows intervals of medium- to very fine-grained
sand ( 1— 4). The grains are angular and poorly sorted, and
they have variable shapes. The S
3
subfacies ranges in thick-
ness from a few decimeters up to 2 meters (Fig. 5).
Interpretation: The plane-parallel stratified sandstone could
represent the deposit of a near-bed suspension generated by
progressive turbulent mixing at the head of a sandy dense flow
with relatively low rates of deceleration (Mutti et al. 2003).
Each lamina can be considered to represent a traction carpet
that is driven by basal shearing of an overlying turbulent flow.
The cycle of “moving laminae” (i.e. bedwaves with small am-
plitude and long wavelength – sensu Best & Bridge 1992)
was controlled by the growth and collapse of the traction car-
pets. The millimeter-thick ( ~ 1—2 mm) inversely graded divi-
sions of the individual laminae, observed in parallel-laminated
sands were apparently produced as energetic sweeps with fairly
flat erosional surfaces, by the large turbulent eddies. In agree-
ment with Hiscott (1994) we assume, that between local im-
pingements caused by these isolated, more energetic sweeps,
high rates of sediment fallout from suspension may have de-
posited the structureless, non-graded parts of sandy laminae.
S
4
subfacies: Sandstone with sinusoidal laminae in phase
Description: This is a relatively frequent subfacies within the
studied megabeds. The ripple lamination in phase, as it was de-
scribed by McKee (1939, 1965), is formed by two-dimensional
bedforms which climb vertically. Bedforms have wavelengths
between 0.3—0.8 m and their height varies from about 1 cm to
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Fig. 9. Hydroplastically-deformed sandstones (S
7
subfacies).
5 cm (Figs. 7B, 8B). The S
4
subfacies forms intervals a few
decimeters up to more than 2 meters thick (Fig. 5B,D).
Interpretation: This type of sinusoidal ripple lamination
consists of a series of ripples with symmetrical, sine-wave
profiles and continuous laminae across the ripple system
which is associated with a high rate of fallout of very fine co-
hesive sediments as shown by Jopling & Walker (1968).
S
5
subfacies: Sandstone with “ripple-drift cross lamination”
Description: “Ripple-drift cross lamination” (Jopling &
Walker 1968) is type of climbing-ripples, which is character-
ized by a high angle of climb and by complete preservation
of bedforms (Fig. 8A,B). The thickness of subfacies S
5
ranges
from several decimeters to 1.2 m (Fig. 5C,F).
Interpretation: This is a typical traction plus fallout struc-
ture in which the interaction between rate of fallout and bed-
forms migration allows the formation of climbing sets of lee
side laminae and the preservation of sandy stoss side lami-
nae. The results are subvertical climbing ripples (Jopling &
Walker 1968; Allen 1970).
S
6
subfacies: Laminated fine-grained sandstone with water
escape structures
Description: This subfacies is almost identical to the S
3
subfacies of laminated sandstones but S
6
subfacies also in-
cludes common subvertical to vertical pillars or “pipes” that
intersect lamination. Some laminae are moderately deformed
around pipes.
Interpretation: The pillar structures represent fine water-
escape conduits from which water and fluidized particles
move upward, cutting and deforming overlying sediments
(Owen 1987, 1996).
S
7
subfacies: Hydroplastically-deformed medium- to fine-
grained sandstone
Description: Sandstone deformations vary from gentle to
moderately strong upwardly-concave dish structures to con-
volute lamination (see Fig. 9).
Interpretation: Dish structure formation is connected
with compaction and dewatering of unconsolidated sedi-
ments (Lowe & LoPiccolo 1974). The generating mecha-
nism of convolute structures is linked to fluidization
processes, which create gravitational instabilities (Owen
1996; Rossetti 1999; Neuwerth et al. 2006). The triggering
mechanism of these structures is often related to processes of
sediment gravity flows, overloading of sandstone beds, dew-
atering of unconsolidated sediments (e.g. Lowe & LoPiccolo
1974; Lowe 1975; Lowe & Guy 2000), or they should be in-
duced by seismicity (Moretti et al. 1999; Neuwerth et al.
2006).
Si lithofacies: Siltstone
Si
1
subfacies: Laminated siltstone
Description: This subfacies is composed of coarse-grained
to fine-grained laminated siltstone. The laminae of this subfa-
cies are thinner, and finer than those of subfacies S
3
. Si
1
subfa-
cies forms an interval a few decimeters up to 2 meters thick
(Fig. 5B,C,D,E).
Interpretation: The depositional mechanism of siltstone
laminae is most likely the same as that in the laminated sand-
stones (S
3
subfacies) and reflects traction plus fallout pro-
cesses associated with turbulent flows. The Si
1
subfacies can
be considered equivalent to the Bouma T
d
interval (Bouma
1962).
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Si
2
subfacies: Laminated muddy siltstone
Description: Coarse- to medium-grained siltstones with an
increased clay content, which are characterized by fine, dis-
continuous wispy lamination. Si
2
subfacies forms intervals
10—20 cm thick (Fig. 5C).
Interpretation: Suspension fall-out during final deposition
from a sediment gravity flow. The repetitive occurrence of
muddy siltstones within clean laminated siltstones of Si
1
subfacies would be interpreted as a result of fluctuations in
supply of sediment and speed of the flow.
Si
3
subfacies: Hydroplastically-deformed siltstone
Description: This subfacies is composed mainly of coarse-
grained sandy siltstones with moderately to heavily de-
formed wispy lamination. Some deformations are similar to
flat upwardly-concave dish-type structures which associate
with vertical pillars. Although hydroplastic deformations are
similar to those of subfacies S
6
and partly S
7
, because of their
occurrence in siltstone lithofacies we selected them as
a single division.
Interpretation: The hydroplastic deformations within the
Si
3
are probably associated with dewatering of unconsolidated
sediments, when the suspended load settles too rapidly (e.g.
Owen 1996; Lowe & Guy 2000).
M lithofacies: Massive dark mudstone
Description: This facies is composed of massive mud-
stones, having a > 8.0. Although some parts reveal an in-
creased content of silt (graded mudstones), the mudstones
are mostly devoid of structure or grading. The M lithofacies
shows the sedimentary characteristics of the Bouma turbidite
division T
e
or Stow division T
6
and T
7
(graded and ungraded
turbidite muds respectively).
Interpretation: Suspension fall-out from static or slow-
moving mud cloud. Final deposition from a sediment gravity
flow event (e.g. Piper 1978).
Mixed lithofacies: Uniform mixture of sandy/silty claystones
Description: A unit up to 4 m thick with unimodal compo-
sition of sandy/silty claystones ( 3.0— 9.0) is bounded by
flat surfaces that represent primary bedding. Integral com-
ponents of this bedding are formed by isolated, irregularly
distributed eliptical pseudonodules of very fine-grained
sandstones and siltstones (Fig. 10A,B). These pseudonodules
are arranged with their long axes parallel to the bedding and
range in size from a few cm up to a maximum of 40 cm.
They are rounded in ellipsoidal and spiral shapes, and con-
tain disrupted and intensely convoluted laminae (Fig. 10C).
The central parts of the pseudonodules are highlighted by
concentration of pyrite.
Interpretation: The unimodal grain-size composition
(sandy/silty claystones matrix), thick and relatively large-
sized body of the unit, its occurrence between undeformed
strata, the presence of mostly rounded eliptical pseudonodules
irregularly distributed within the unit, horizontal arrangement
of pseudonodules and their deformed internal structures, as
well as unrepeated occurrence of such a unit within the com-
plex beds suggest that the Mixed lithofacies was not formed
by normal sedimentary processes. We speculate that primarily
deposited sediment with unconsolidated and semi-consolidated,
siltstones and claystones as well as sporadic fine-grained
sandstones, was affected by syn- to immediately post-deposi-
tional soft-sediment deformation. Soft-sediment deformation
is often connected with liquefaction of sediment due to seis-
mic shaking that lowered drastically its bulk density and shear
strength allowing sinking and consecutive deformation of the
overlying silty/sandy layers (Rodríguez-Pascua 2000; Davies
et al. 2004). The formation of isolated or detached pseudono-
Fig. 10. A – A uniform mixture of sandy/silty claystones (Mixed lithofacies) with isolated elliptical pseudonodule of very fine-grained
sandstone (Ve ký Bysterec). B – A detailed view of the pseudonodule within sandy/silty claystones. C – An internal structure of many
pseudonodules indicate plastic extruding and strong deformation of laminae.
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dules resulted from the sinking of load casts in water-saturated
fine-grained sediments (Kuenen 1958). The contorted laminae
within the nodules (Fig. 10C) represent the original laminations
of a silt and fine sand accumulation that was deformed as the
beds underwent liquefaction and separation into a number of
pseudonodular packets (Maurya et al. 1998; Davies et al. 2004).
However, similar division comprising a graded muddy
sandstone/siltstone with a distinctive content of small pseudo-
nodules and fragments of microfolded laminae is interpreted
as syn-depositional deformation of primarily deposited thiner
sandy, silty and muddy beds which resulted from multiple
reflections of an upper turbulent suspension deposited from a
large-volume flow (Remacha & Fernández 2003).
Facies association
Two types of megabeds were distinguished in the Orava re-
gion (Fig. 5). The first type is represented by thick separated
beds, which have been classified here as “megaturbidites”
(e.g. Svorad plató, Jóbova Ráztoka). Megaturbidites do not
present internal erosive surfaces or amalgamation surfaces as
well as grain size breaks (Fig. 5, logs A,B,C). An interesting
aspect of the studied megaturbidites is the pure absence of
thick, massive, structureless sandstone/siltstone parts within
the beds. Nearly all lithofacies are finely structured. Each
megaturbidite is a discrete bed, which recorded a single sedi-
mentary event. The second type is represented by complex
beds (CB) (e.g. Dolný Kubín, Čremoš; Fig. 5, logs D,E,F) that
consist of the following successive units (from base to top):
(1) coarse-grained, poorly sorted, disorganized breccia (CB
A
);
(2) the heterolithic unit equivalent to the megaturbidite (CB
B
);
(3) the deformed unit of the clayey siltstones (CB
C
).
In general, the megaturbidite-type beds show a typically fin-
ing-upward trend in grain size from conglomerates and
coarse-grained sandstones to fine-grained sandstones, silt-
stones and homogeneous mudstones. This trend reflects a
gradual decrease in flow velocity and competence.
The lowermost parts of the megaturbidites are composed
of relatively thin gradational intervals of C
2
, C
3
and S
1
sub-
facies. The conglomerates of C
2
subfacies usually fill small-
scale, shallow and laterally pinched-out scours. They can
vertically and laterally pass into unstructured pebbly/granule
sandstones of the S
1
subfacies. The S
1
subfacies is commonly
developed directly at the base of megaturbidites (e.g.
Záskalie, Kňažia and Ve ký Bysterec sections), which have
the flat base-bed surfaces with the sporadic occurrence of
flute marks and tool marks. Inverselly graded conglomerates
of the C
3
subfacies are relatively rare. This subfacies is devel-
oped above the C
2
subfacies and usually vertically passes into
the stratified, inversely graded S
2
subfacies. The S
2
subfacies
forms an interval up to 35 cm thick in the uppermost parts of
the gradational interval and upwards usually rapidly fine-
down to medium-grained laminated sandstones of the S
3
sub-
facies. Multiple intervals of parallel lamination form the most
evident internal structure of the sandstone part of these mega-
beds (Fig. 4). The S
3
subfacies can occur as an interval more
than 2 m thick occasionally at the base of the megabeds (e.g.
Kňažia, Záskalie), but more frequently it follows the S
1
and S
2
subfacies (e.g. Svorad, Pucov, Jóbova Ráztoka etc.). These
parallel-laminated sandstones grade upward into laminated
siltstones (Si
1
subfacies) and mudstones (M lithofacies). The
sandstones with ripple lamination in phase (S
4
subfacies) are
relatively common (e.g. Svorad locality) and the upper bound-
ary of this subfacies is also gradual and usually again passes
upward into fine-grained, parallel-laminated sandstones of the
S
3
subfacies (Fig. 7B). A more uncommon type of structure in
the megaturbidites is represented by the ripple-drift cross lam-
ination (S
5
subfacies). This structure has been observed only
in the Ve ký Bysterec and Čremoš sections (Fig. 5, logs C
and F), where the parallel lamination of the S
3
subfacies is lo-
cally associated with intervals of ripple-drift cross lamination.
The transition of parallel lamination to ripple bedforms (S
4—5
subfacies) might be caused by a temporal decrease in the fall-
out velocity of the flows through intermediate increase in their
shear velocity. The bed characterized by an alternation of dif-
ferent bedform types and subfacies may be interpreted as be-
ing deposited from unsteady, pulsating flows. The even and
parallel-laminated siltstones (Si
1
subfacies) of the megaturbid-
ites are usually developed as an interval a few decimeters up
to 2 meters thick which records the transition between sand-
stones and claystones. This trend is typical of several megatur-
bidites at outcrops near Svorad plató, Huty, Kňažia and
Jóbova Ráztoka (Fig. 5, log B; Fig. 4A,B). The upper parts of
the megabeds are occasionally formed by 10—20 cm thick in-
tervals of the laminated muddy siltstones (Si
2
subfacies).
The uppermost parts of the megaturbidites are composed of
mudstone lithofacies. The transition between the Si lithofa-
cies and M lithofacies is gradual. The thickness of the dark
turbidite mudstones varies from about 3 m (at Pucov and
Svorad sections; Figs. 3E,F and 5B) to about 5 m (at Huty,
Jóbova Ráztoka sections; Figs. 3H,G and 5B). However, be-
cause of the position of the megabeds within muddy deposits
of the Huty Formation, the total thickness of mudstones be-
longing to the megaturbidites, cannot be correctly defined.
The hemipelagic interval of lighter massive calcareous mud-
stones occurs locally above the M lithofacies (e.g. Huty,
Jóbova Ráztoka, Svorad localites; Fig. 5, log B), but it was
not possible to distinguish these two types of mudstones at
localities near Dolný Kubín. The rapid deposition of the Orava
megabeds is manifested by the presence of water escape
structures (e.g. Lowe 1982; Lowe & Guy 2000), and other
types of dewatering structures. Some megabeds near Dolný
Kubín contain 20 up to 150 cm thick intervals of the S
6
sub-
facies, which are almost identical with the laminated S
3
sub-
facies (and often associate with this), but they also contain
common subvertical to vertical streaks similar to the water-
escape pillars. These pipe-like linear structures are either ran-
domly distributed, or they exhibit a tendency to occur in the
upper sandstone intervals of the megabeds. Hydroplastically-
deformed sandstones of the S
7
subfacies form centimeter- to
decimeter-thick intervals which show a limited lateral contin-
uation from tenths of meters to a few meters, and they are sur-
rounded by non-deformed laminated sets (S
3
subfacies). The
S
7
subfacies is situated especially in the lower part of the beds
and often occurs in megaturbidites near the Dolný Kubín—
Ve ký Bysterec locality (Fig. 5A,C,D). The coarse-grained
siltstones with deformed wispy lamination, flat dishlike
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structures and fine vertical water-escape conduits (Si
3
subfa-
cies) form intervals that range in thickness from 30 cm to 3 m.
The Si
3
subfacies predominantly develop on the upper parts of
megabeds near the Dolný Kubín—Ve ký Bysterec locality
(Fig. 5A,C,D). Many intervals of siltstones lithofacies display
transition from laminated siltstones (subfacies Si
1
) to hydro-
plastically-deformed siltstones (subfacies Si
3
).
The CB
A
unit of the complex beds, developed on the base of
some composite megabeds (e.g. Ve ký Bysterec and Čremoš
sections), represents deposits from debris flows. This unit is
characterized by relatively coarse-grained conglomerates with a
disorganized texture and presence of deformed fragments and
blocks of sediments (C
1
subfacies). The thickness of the CB
A
unit ranges from 110 cm to 180 cm (Figs. 5E,F and 11). The
base of the conglomerates exhibits a slightly erosive character.
The upper part of C
1
subfacies was furrowed by small-scale ero-
sional scours which are developed at the base of clast supported
conglomerates of the C
2
subfacies (Ve ký Bysterec; Fig. 11).
The furrows point to the erosion of the soft
upper surfaces of de-
bris-flow deposits by subsequent gravelly dense flows.
The CB
B
unit is composed of the facies, which in general
are very close to those originally described within megatur-
bidites. The vertical arrangement of the subfacies of the CB
B
unit corresponds to the fining-upward grain-size trend of
megabeds. The uniform mixture of sandy/silty claystones
with pseudonodules (Mixed lithofacies) is well documented
as the CB
C
unit in the Ve ký Bysterec section (Figs. 5D,E
and 10A), where they are developed above thick, laminated
siltstones (Si
1
subfacies) or hydroplastically-deformed silt-
with prevalence of the carbonate rocks has also been found
in the C
2
and C
3
subfacies, in which a significant portion of
detrital particles (up to 10 %) is composed of foraminiferal
tests (mainly nummulites; Fig. 6C). The bioclasts are com-
mon mainly in the megabeds near Dolný Kubín and Pucov,
but they are more rare in the gradational intervals of the
megabeds at the others localites.
The content of quartz and siliciclastic components increases
markedly over the carbonates in the upper intervals of the S
1
and S
2
subfacies, and these clasts became dominant in all
sandstone and siltstone subfacies. The results of petrographic
analysis show that the amount of quartz and stable rock
grains reaches up to 47—52 % in the parallel-laminated sand-
stones (S
3
subfacies), and their proportion increases moder-
ately from the base to the top of this subfacies. The content
of unstable rock grains, which predominantly involve car-
bonates, is approximately 17—23 %. The remaining propor-
tion of sediment is composed of a matrix and carbonate-fill
cement from the intergranular pores. The medium- to fine-
grained sandstone lithofacies and siltstone lithofacies often
contain large amounts of coal particles to millimeter-thick
coal seams and up to millimeter-sized flakes of mica.
Paleocurrent orientation
The studied megabeds basically lack paleocurrent indica-
tors, or they are relatively poor in typical current indicators
such as flute marks and current cross-bedding structures. In
Fig. 11. The complex beds (CB) (Ve ký Bysterec). CB
A
unit characterized by coarse-grained, poorly
sorted, disorganized conglomerate (C
1
subfacies) with plastically deformed, meter-scale blocks of
sediments (white arrows). Take note of an indication of primary bedding preserved within the
blocks of sediments. The upper part of C
1
subfacies was furrowed by small-scale erosional scours
(black arrows), which points to the erosion of subsequent gravelly dense flows represented by clast
supported conglomerates of C
2
subfacies at the base of the CB
B
unit.
stones (Si
3
subfacies) of the
CB
B
unit. The CB
C
unit is over-
lapped by thick massive mud-
stones (M lithofacies).
Lithology
A lithological analysis of the
Orava megabeds highlights the
differences in their petrographic
composition in dependance on
their grain-size composition.
The coarse-grained conglomer-
ates with a disorganized struc-
ture and angular clasts (C
1
subfacies) are composed almost
exclusively of carbonate rocks
from older Mesozoic and Paleo-
gene units. The clast composi-
tion reflects the source areas
formed by the Mesozoic base-
ment units (Choč and Krížna
Nappes) and Late Eocene to
Early Oligocenne sedimentary
units (detrital components of
the Borové Formation, and intr-
aclasts of dark non-calcareous
claystones of the Huty Forma-
tion). The similar composition
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most of the outcrops, the megabeds are surrounded by mud-
stones, siltstones and fine-grained thin sandstones with trace
fossils on the flat bed surfaces. The exception to this occurs
in area near Dolný Kubín, where the megabeds are surrounded
by alternating thin-bedded to medium-bedded turbidites,
which show more intensive current bottom erosion. Paleo-
current data derived from these flute casts, longitudinal fur-
rows, ridges and tool marks provide flow orientation from
SW and W to NE and E. The determination of typical paleo-
current indicators from individual megabeds was difficult,
since in the majority of localities their basal bottom surfaces
are not visible or high-concentration flows at the base of tur-
bidity currents did not allow the development of any current
marks. Additional paleocurrent data have been identified
from climbing-ripple structures in cross lamination of the
megabeds (SE transport direction), grain imbrication in lami-
nae (to E and NE), parting lineation (SW—NE orientation)
and the orientation of the long axis of the pseudonodules in
the CB
C
division (WSW—ENE orientation).
Stratigraphic position of megabeds
An important aspect of this study is the stratigraphic posi-
tion of the megabeds. The biostratigraphic classification of the
megabeds has been carried out by foraminiferal and nanno-
plankton study of claystone intervals. The foraminiferal mi-
crofauna of these claystones consists of small-sized forms of
tenuitellids, pseudohastigerinids, paragloborotalids, cassiger-
inelinids, chiloguembelinids and catapsydracids (see Fig. 12).
They contain several index species, like Tenuitella gemma
(Jenkins), T. munda (Jenkins), T. clemenciae (Bermudez),
T. brevispira (Subbotina), Pseudohastigerina micra (Cole),
P. naguewichiensis (Myatliuk), Paragloborotalia nana (Bolli),
Chiloguembelina cubensis (Palmer), Cassigerinella chipo-
lensis (Cuschman & Ponton), Dentoglobigerina rohri (Bolli),
Globorotaloides suteri Bolli, Catapsydrax martini Blow &
Banner, Protentella (Bolliella) navazuelensis Molina, etc.,
which are common in the Early Oligocene associations (see
Berggren et al. 1967; Li 1987; Leckie et al. 1993; Spezzaferri
1994; Pearson & Wade 2009). The presence of Pseudohasti-
gerina shows that the age of the megabeds should not be
younger than the early middle Rupelian, which in the P-series
zonation corresponds to the P18 Biozone (Chiloguembelina
cubensis—Pseudohastigerina spp. sensu Berggren et al. 1995),
and in the O-series zonation corresponds to the O1 Biozone
(Pseudohastigerina naguewichensis HOZ sensu Berggren &
Pearson 2005). Tenuitella-rich associations with other micro-
perforate foraminifers are common in the Oligocene forma-
tions of the Carpathian Flysch units (Olszewska 1985; Bąk
Fig. 12. Representative species of the foraminiferal microfauna from the megabeds in the Ve ký Bysterec locality. 1 – Tenuitella gemma
(Jenkins); sample DK-1, scale bar 20 µm. 2 – Tenuitella munda (Jenkins); sample DK-3, scale bar 50 µm. 3 – Tenuitella clemenciae
(Bermudez); sample DK-3, scale bar 50 µm. 4 – Paragloborotalia nana (Bolli); sample DK-3, scale bar 40 µm. 5 – Pseudohastigerina
micra (Cole); sample DK-1, scale bar 40 µm. 6 – Pseudohastigerina naguewichiensis (Myatliuk); sample DK-3, scale bar 40 µm.
7 – Cassigerinella chipolensis (Cuschman & Ponton), DK-4, scale bar 40 µm. 8 – Protentella (Bolliella) navazuelensis Molina; sample
DK-2, scale bar 40 µm. 9 – Globorotaloides suteri Bolli; sample DK-3, scale bar 40 µm. 10 – Dentoglobigerina rohri (Bolli); sample
DK-3, scale bar 40 µm. 11 – Catapsydrax martini (Blow & Branner); sample DK-3, scale bar 50 µm. 12 – Chiloguembelina cubensis
(Palmer); sample DK-3, scale bar 40 µm.
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2005; Bubík & Poul 2010, etc.). Olszewska (1997) defined
their acme zones for the Rupelian—lower Chattian sediments.
Dominance of tenuitellids provides evidence of cold-water
condition of the Early Oligocene (e.g. Pippèrr & Reichen-
bacher 2010).
The foraminiferal stratigraphic data correspond to the re-
sults of the nannoplankton stratigraphy. The nannofossil asso-
ciations (Fig. 13) imply the coexistence of Transversopontis
fibula Gheta and Reticulofenestra ornata Müller, which is in-
dicative for the NP23 Biozone (Nagymarosy & Voronina
1992; Krhovský et al. 1992; Švábenická et al. 2007, etc.).
However, the low frequency of these species points rather to
the upper part of the NP22 and lower part of the NP23 Zones.
Other nannoplankton species include Helicosphaera compacta
Bramlette & Wilcoxon, Helicosphaera bramlettei Müller,
Chiasmolithus oamaruensis (Deflandre), Isthmolithus recur-
vus Deflandre, Lanternithus minutus Stradner, Coccolithus
formosus
(Kamptner),
Transversopontis
pulcheroides
(Sullivan) and abundant reworked species of Eocene and
Cretaceous nannofossils (e.g. Discoaster barbadiensis Tan,
Discoaster saipanensis Bramlette, Tribrachiathus orthostylus
(Bramlette & Riedel), Eifellithus sp.
The stratigraphic data obtained from the foraminiferal and
nannofossil study allow us to properly determine the age of
the Orava megabeds as early middle Rupelian (P18/O1—
NP22/NP23 Biozones) (Fig. 14).
Discussion
Interpretation of the Orava megabeds
Sedimentological literature provides a lot of ideas concern-
ing the processes that initiate the turbidity currents. However,
displacement of enormous amounts of material in single flows
is rare and megabeds, like those in the Orava region, which are
characterized by great thickness, considerably exceeding the
thickness of related beds in other parts of the depositional sys-
tem, suggests catastrophic events on the basinal margins.
The presence of the Orava megabeds makes it possible to
discuss possible paleogeographical scenarios explaining how
to deliver the required volume of sediments to the Central Car-
pathian Paleogene Basin and confront our data with published
modern analogues of such a type of megabeds. The sedimen-
Fig. 13. Representative species of calcareous nannofossils from the megabeds in the Ve ký Bysterec locality. 1 – Reticulofenestra ornata
Müller; sample DK-1. 2 – Reticulofenestra umbilica (Levin); sample DK-4. 3—4 – Transversopontis fibula Gheta; sample DK-4. 5 – Lan-
ternithus minutus Stradner; sample DK-3. 6 – Helicosphaera bramlettei Müller; sample DK-1. 7 – Helicosphaera compacta Bramlette &
Wilcoxon; sample DK-3. 8 – Isthmolithus recurvus Deflandre; sample DK-4. 9 – Discoaster barbadiensis TAN; sample DK-4. 10 – Dic-
tyococcites bisectus (Hay, Mohler & Wade) Bukry & Percival; sample DK-1. 11 – Coccolithus formosus (Kamptner) Wise; sample 4.
12 – Braarudosphaera bigelowii (Gran & Braarud) Deflandre; sample DK-3. Scale bar for all figures: 1 µm.
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Fig. 14. Stratigraphic distribution of foraminiferal and nannofossil species in the megaturbidite se-
quence of the CCPB (Bysterec section). Based on the index species the age of the Orava megabeds cor-
responds to the Biozones O1—O2 and NP22—NP23 (early middle Rupelian). The numbers (1—5) mark
the samples taking position.
tary volume of megaturbidites is frequently used in literature
for their classification as well as for interpretation of the pro-
cesses that triggered these flows (e.g. Mutti et al. 1984; Reeder
et al. 2000; Piper & Normark 2009). In spite of the fact that
the Orava megabeds are considered to represent the deposits
of large voluminous flows, their total sedimentary volume in
the CCPB could not be properly quantified. Therefore, it is
difficult to come to decisive hypothesis about the formation of
such exceptionally thick units as the Orava megabeds, so that
we provide rather a multiproxy approach to interpretation of
their possible depositional mechanisms. The large voluminous
gravity flows are often linked to seismic activity causing ups-
lope slide failure (e.g. Mutti et al. 1984; Sérguret et al. 1984;
Labaume et al. 1987). However, some authors suggest that
only minor turbidity currents
are associated with frequent
failures and significant tur-
bidity currents are only rare-
ly created by liquefaction of
sand or silt beds (e.g. Trip-
sanas et al. 2008; Piper &
Normark 2009). Moreover,
the duration of sediment liq-
uefaction and their simple
slope failures was measured
in minutes or at most hours
(Andresen & Bjerrum 1967;
Karlsrud & Edgers 1982;
De Groot et al. 1987; Emdal
et al. 1996), therefore the
longer lasting events of fail-
ure, capable of producing the
units on the scale of the
megaturbidites, could only
originate by multifold steps
of shoreline retrogression.
Even if distributaries and fan
deltas on the basinal margins
could supply enough sand
and silt, the coarse sediments
are blanketed by the plume-
type mud deposits, occurred
as a thin mudstone interbeds
in sandy deposits. The com-
paction of the mudstones
could establish their resis-
tivity to retrogression. The
beds originated by succes-
sive retrogressive failures
would produce abundant
mudstone intraclasts, but
they are entirely missing in
all of the sandstone/siltstone
facies of the Orava mega-
beds. We suppose, that prob-
ably only coarse-grained
CB
A
unit of the complex
beds, the structures of which
indicate deposition from de-
bris flow, can be explained by slope failure from the basinal
shelf-slope systems. The failures could be generated by exten-
sive breakdowns in the delta front, but the large amount of
bioclasts from the Eocene carbonates as well as the presence
of synsedimentary slumps and coherent blocks suggests that
not only the Oligocene deposits (mudstone intraclasts) but
also unconsolidated or semiconsolidated older deposits of the
Borové Formation were affected by slope failures. These
failures could be connected with oceanographic processes in
canyon head or with straight-sided deeply incised conduits
where erosion by flows might have provided conditions for
breaching (e.g. Piper & Normark 2009).
The megaturbidites, developed as several meters thick,
finely structuralized sandstone and siltstone of the CB
B
unit,
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apparently differ from of the CB
A
unit. The detrital particles
of the Orava megabeds have a polygenic composition, re-
flecting the presence of plutonic, supracrustal, low-grade
metamorphic and sedimentary rocks in the source areas. The
plutonic and metamorphic sources were exposed and denuded
together with the unstable rocks of the carbonate complexes.
Almost 90 % of the sandstone and siltstone sediments of the
megabeds are well sorted and they lack larger boulders and
coarse-grained clastics. Therefore, such fine-grained sedi-
ments with mature siliciclastic composition appear to have
been accumulated by rivers, which fed the basin from tecton-
ically active sources with permanent discharge of finely dis-
integrated particles. The large amount of coal particles, coal
seams and wood fragments in sandstone- to siltstone lithofa-
cies of the Orava megabeds could also indicate a source in
the delta deposits or fluvial discharge to the continental shelf
(cf. Mutti et al. 2003). Riverine input of hyperconcentrated
bedload during catastrophic floods that flowed seaward due
to inertia could have generated voluminous flows which de-
posited very thick sandstone—siltstone beds (e.g. Mutti et al.
2000, 2003, 2009; Normark & Reid 2003; Plink-Björklund
& Steel 2004; Piper et al. 2007). However, the largest ob-
served turbidites are much more voluminous than the sedi-
ment discharge of a large single river flood (Piper &
Normark 2009). Similarly, the storm-initiated deposition of
turbidites are generally supply-limited because of the death
of the many flows inside the canyon (Piper 1970; Paull et al.
2005) and only occasional flows become sufficiently large
and energetic to reach the basin floor (e.g. Piper & Normark
1983). The channel avulsions, which occasionally occurred
due to flow erosion, may also result in deposition of excep-
tionally thick sand beds (Piper & Normark 2001). However,
such beds are usually massive, unstructured and rich in mud-
stone clasts, therefore they differ from structural and textural
features of the sandstones in the Orava megabeds.
Even though each of these processes alone is questionable
as the main generating mechanism of the Orava megabeds,
they would represent effective triggers that in combination
with the suitable condition on the basinal margins could led
to the formation of large-scale gravity flows.
Nevertheless, the main factor which probably played a sig-
nificant role in triggering of the voluminous flows is inferred
in long-lasting accumulation of the clastics on the basinal
margins of the CCPB.
The CCPB as a tectonic-type basin was bordered by active
margins flanked by coarse-grained and sandy sediments of
alluvial fan, fluvial, fluviodeltaic and fan-delta systems,
which after the marine incursion fed the coastal zone
(Marschalko 1970; Baráth & Kováč 1995; Filo & Siráňová
1996, 1998). Interference of eustasy and tectonics led to the
relative sea-level changes in the CCPB, evolution of which
involved a succession of highstand and lowstand phases with
shifting of the coastline and marginal facies (e.g. Soták et al.
2001; Starek et al. 2012). During the Early Oligocene, the
CCPB passed through a highstand stage of relative sea level
(e.g. Soták 1998; Soták et al. 2001), when the majority of the
coarse sediments remained in fan deltas and in actively pro-
graded peripheral beaches (cf. Dabrio 1990). The coastal
erosion of ravinement surfaces on the basinal margins of the
CCPB during the rise in sea-level rising from the Priabonian/
early Rupelian transition, together with high-bedload river-
ine input into the marine realm could have produced a large
quantity of sediments, which were delivered to conduit
heads through longshore drift, or fan delta progradation and
subsequently remobilized by following initiating processes
(e.g. Piper & Normark 2001).
Even though hyperpycnal flows generated by catastrophic
floods are questionable as initiating events of the Orava
megabeds, their deposition would be influenced by these
processes as well. These large-volume and highly turbulent
flows could have accelerated down the continental slopes,
eroding deposits from the conduit and basinal-slope accumu-
lations, progressively increasing the proportion of sand in
the flows, which were transported a significant distance onto
the basin floor (e.g. Hughes Clarke et al. 1990; Piper & Nor-
mark 2001, 2009; Talling et al. 2007). Downslope passing of
the large highly erosive hyperpycnal flows could have been
accompanied by extensive slope failures of straight-sided
deeply incised conduits, which could have led to initiation of
debris flows at the base of the complex beds resulting in
highly disorganized deposit (C
1
subfacies) frozen during the
early stages of their downslope motion. Disorganized pebbly
mudstone facies and massive conglomerates (debris-flow
deposits or debrites) at the base of thick sandy/silty beds
characterized by a fining upward facies sequence (turbulent
portions of the flow) are referred to as highly efficient bipar-
tite flows (e.g. Mutti et al. 2003; Tinterri et al. 2003) in
which sharp grain-size breaks represent a variation in hydro-
dynamic conditions of the gravity flow during its downcur-
rent evolution.
The CB
B
unit, as well as the unit identified as megaturbid-
ites, which probably represent the more distal equivalents of
the composite megabeds, were produced during a single
event and reflect the continual deposition from the turbulent
flows. Their structures suggest basal near bed processes as
follows: – rapid deposition of coarse-grained sediments
from high-concentration flow at the base of the turbidity cur-
rents (C
2—3
, S
1—2
subfacies); – steady basal shear stress
(growth and collapse in the traction carpets) and a steady rain
of suspended sediment (S
3—7
, Si
1—3
subfacies); – intermittent
phases of variable flow velocity (S
4
, S
5
subfacies); – succes-
sive decreases in flow competence (upwardly-fining granu-
lometry) and terminal depositional processes from muddy
suspension (M lithofacies). The rapid deposition of the
megabeds is expressed by the water escape structures and
hydroplastic deformations (S
6—7
, Si
3
subfacies).
The megabeds with a great thickness ( > 1 m) and similar
alteration of plane-parallel lamination with climbing ripple
lamination, and with large amounts of coal seams and wood
fragments could be interpreted as deposits from hyperpycnal
flows (e.g. Mutti et al. 2003; Olariu et al. 2010).
The period of long-lasting proximal deposition and lower
frequency of catastrophic events due to low activity of trig-
gering mechanisms, and following conduit flushing appears
to have been the most common source of sediment for very
large turbidites. Conduit flushing could explain why the vol-
ume of megaturbidites is much greater than the volume of
sediments transported in a single river flood or from usual
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slope failures (Piper & Normark 2009), which is probably
the case of the sediments in the Orava megabeds.
Deposition from turbidity currents played an important
role in basin-forming processes of the CCPB, where the tur-
bidite systems were influenced mainly by tectonics and eus-
tacy (e.g. Soták 1998; Starek et al. 2000; Soták et al. 2001;
Starek 2001). However, only one of the sedimentary forma-
tions of the CCPB reveals the occurrence of overthickened
sandstone—siltstone beds without internal erosive surfaces or
amalgamation surfaces as well as grain-size breaks, which
therefore should be deposit from single depositional events.
The early stages of deposition in the CCPB revealed mainly
continental and shallow-marine conditions, when after the
initial transgression (TA3.5—3.6 third-order Exxon cycles),
the subaerial sediments changed to subaqueous sediments in
the Borové Formation (Baráth & Kováč 1995). This forma-
tion comprises mainly deposits of alluvial fans, fluvial and
fluviodeltaic systems, fan-deltas as well as shoreface and
carbonate ramp environments (Marschalko 1970; Kulka
1985; Baráth & Kováč 1995; Filo & Siráňová 1996, 1998;
Janočko & Jacko 1998; Soták et al. 2001). The distinctive
eustatic fall of sea level in the early Priabonian at the begin-
ning of the TA4 supercycle led to regression, subaerial expo-
sure and erosion of the Middle Eocene formations. The
exposed shelf was prone to developing incised valleys as
a result of fluvial erosion (Starek et al. 2012). During the
Priabonian transgression, the subsequent rise of the relative
sea level accompanied by rapid tectonic subsidence along
the fault-bounded margins of the CCPB (Soták et al. 2001)
led to extensive flooding and deepening of depositional en-
vironments (e.g. Buček et al. 1998). In this stage of the ba-
sin’s evolution, most of the clastics, produced by rivers and
erosion of flooded land, was retained on the coastal plain as
a component of beach complexes and deltaic systems. Direct
riverine sediment-laden discharge to the distal parts of the
basin is rather rare. The large volume of sand could be
trapped within incised-conduit heads and delta distributaries.
During the late Priabonian and Rupelian, the CCPB was
filled up by mud-rich deposits of the Huty Formation, which
represents different sediments of the deep-marine clastic sys-
tems (e.g. Janočko & Jacko 1999; Soták et al. 2001; Starek
et al. 2004). In the lower part of the Huty Formation the
mudstones greatly prevail over sandstones. The sandy tur-
bidites, in dependence on their position in the submarine tur-
bidite systems, mostly do not exceed a thickness of 10 cm,
only rarely reaching up to 50 cm. These deposits of the Huty
Formation are associated with the Orava megabeds, the pos-
sible initiating mechanisms of which are discussed above.
However, numerous sandy turbidites are generated during
falling sea-level stages, which result in progradation and
cannibalization of the fan delta (e.g. Postma 1995) and as the
major riverine input of sediment-laden flows occur seaward
of high bedload deltas (e.g. Milliman & Kao 2005; Normark
et al. 2006; Milliman et al. 2007). Evidence of the large-vol-
ume hyperpycnal-type sediments in submarine turbidite de-
posits of the CCPB may indicate the transitional sea-level
drop during the early Rupelian, which culminated by river-
ine input, brackishing and semiisolation in the NP23 Zone
(Soták 2010).
Incoming regression of the TB1 supercycle in the CCPB,
which resulted from an abrupt sea-level fall at around 30 Ma
at a time of major glaciation in Antartica (Robin 1988;
Kennett & Barker 1990; Zachos et al. 1993), led to sand-rich
submarine fan deposition of the Zuberec and Biely Potok
Formations (Soták 1998; Soták & Starek 1999; Starek et al.
2000). These Upper Oligocene formations show an abun-
dance of thick sandstone beds, especially the sandstone
lithosomes of the Biely Potok Formation up to several
meters in thickness. Their coalescing lobes grade up to form
the mid-fan lobe complex. However, the thick tabular sand-
stones of the Biely Potok Formation are amalgamated from
units, usually up to 1 m thick beds (Gross et al. 1993; Starek
2001), which are usually massive, unstructured and contain
abundant claystone chips. Unlike the Orava megabeds, these
sandstones are laminated only in the topmost part of the
beds. In spite of a good predisposition of basin margins for
initiation of sandy turbidites during the Late Oligocene re-
gression in the CCPB, their triggering frequency probably
did not allow long-lasting accumulation of sandy sediments,
which would be sufficient to generate such large-volume
gravity flows like those, which delivered such a volume of
sand to the megabeds in the lower part of the Huty Formation.
Conclusion
The Orava megabeds provide a record of major deposi-
tional events in the western part of the Central Carpathian
Paleogene Basin (CCPB). They are specific for certain de-
velopment stage of this basin, because of their occurrence
only in the lower part of the Huty Formation, where they
form distinct interval with far-reaching correlation over tens
of kilometers within the basin fill. The megabeds occur in
the middle Rupelian sequence, the biostratigraphic age of
which, derived from foraminifers and nannofossils, corre-
sponds to P18/O1 and NP22/23 Biozones.
Two types of megabeds, slightly dissimilar in their sedi-
mentary features and flow-type deposition were classified
here. They are developed as separated beds (megaturbidites)
and the complex beds (CB). The megaturbidites are repre-
sented by thick finely structured beds with upwardly-fining
trend in grain size from fine-grained conglomerates and
sandstones to siltstones and homogeneous mudstones. The
complex beds include a basal disorganized coarse-grained
unit (CB
A
); heterolithic unit equivalent to the megaturbidite
(CB
B
); and the mixed lithofacies of deformed sandy/silty
claystones (CB
C
).
The complex beds can be interpreted as a result of bipartite
flows (e.g. Mutti et al. 2003; Tinterri et al. 2003) with debris-
flow deposit at the base, and with superincumbent thick finely
structured sandstone/siltstone facies resulting from turbulent
portions of the flow. Individually placed megaturbidites prob-
ably represent the more distal equivalents of the complex beds
and reflect uninterrupted deposition from turbulent flows.
We suppose, that two main factors probably played signif-
icant roles in the generation of the large-volume gravity
flows and deposition of the Orava megabeds in the middle
Rupelian sequence of the CCPB.
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1. Long-lasting accumulation of the clastics on the basinal
margins of the CCPB during a period of rising sea-level and
a highstand stage of relative sea level in the early Rupelian.
2. The catastrophic trigerring mechanisms on the basinal
margins which could have led to the formation of large-scale
gravity flows.
Several triggers, including oceanographic processes in a
canyon head or upslope slide failures linked with seismic ac-
tivity could have initiated these processes, and therefore we
cannot definitely exclude them. However, the basin-scale ex-
tension forming a coeval horizon of sandstone megabeds;
finely structured sandstone/siltstone facies with definite plane-
parallel stratification and climbing ripple lamination; differ-
ences between shelf-derived carbonate material in the basal
part of complex beds and continental-derived sandy/silty si-
liciclastic material of the megaturbidites as well as a great
abundance of continental material (phytodetritus, leaves) sug-
gest that deposition of the Orava megabeds was influenced by
erosive hyperpycnal flows generated by catastrophic floods.
The presence of hyperpycnal deposits in the CCPB could
indicate the connection of the turbidite systems of deep-sea
fans with fluvial systems and shelf-edge deltas, which is a new
insight in the reconstruction of this basin. The occurrence of
megabeds in the lower part of the Huty Formation provide ev-
idence of possible rising discharge from flooding rivers,
which entered the CCPB from the SW and W to NE and E.
Flood discharge and generation of hyperpycnal flows with
deposition in deep-sea should be dominant processes during a
sea-level fall and the extremely large volume of clastics in the
megabeds could be derived by conduit flushing and coastal
erosion of deposits accumulated in shelf areas during previous
rising of the sea level. It is probably also a case of hyperpycnal
megaturbidites in the CCPB, where their accumulation fol-
lowed the sea-level highstand in the Globigerina Marls at the
Eocene/Oligocene boundary and fall to the sea-level lowstand
during the early Rupelian, which culminated in riverine input,
brackishing and semi-isolation in the NP23 Zone.
Acknowledgment: This work was supported by the Slovak
Research and Development Agency (VEGA) by Project
No. 2/0042/12 and funds receiving through the Centre of
Excellence for Integrative Research of the Earth’s Geosphere
(ITMS 26220120064, European Regional Development Fund).
The authors thank to Silvia Ozdinová for providing the data
from nannoplankton stratigraphy. Our thanks also go to David
J.W. Piper and Roberto Tinterri for the detailed review of
this paper and constructive comments.
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