GEOLOGICA CARPATHICA
, JUNE 2018, 69, 3, 283–300
doi: 10.1515/geoca-2018-0017
www.geologicacarpathica.com
Towards better correlation of the Central Paratethys
regional time scale with the standard geological
time scale of the Miocene Epoch
MICHAL KOVÁČ
1,
, EVA HALÁSOVÁ
1
, NATÁLIA HUDÁČKOVÁ
1
, KATARÍNA HOLCOVÁ
2
,
MATÚŠ HYŽNÝ
1
, MICHAL JAMRICH
1
and ANDREJ RUMAN
1
1
Department of Geology and Palaeontology, Faculty of Natural Sciences, Comenius University in Bratislava, Mlynská dolina, Ilkovičova 6,
842 15 Bratislava, Slovakia;
kovacm@uniba.sk, eva.halasova@uniba.sk, natalia.hudackova@uniba.sk, matus.hyzny@uniba.sk,
michal.jamrich@uniba.sk, winchestersk@yahoo.com
2
Institute of Geology and Palaeontology, Faculty of Sciences, Charles University, Albertov 6, 128 43 Praha 2, Czech Republic;
holcova@natur.cuni.cz
(Manuscript received November 23, 2017; accepted in revised form March 15, 2018)
Abstract: Depositional sequences originating in semi-enclosed basins with endemic biota, partly or completely isolated
from the open ocean, frequently do not allow biostratigraphic correlations with the standard geological time scale (GTS).
The Miocene stages of the Central Paratethys represent regional chronostratigraphic units that were defined in type
sections mostly on the basis of biostratigraphic criteria. The lack of accurate dating makes correlation within and between
basins of this area and at global scales difficult. Although new geochronological estimates increasingly constrain the age
of stage boundaries in the Paratethys, such estimates can be misleading if they do not account for diachronous boundaries
between lithostratigraphic formations and for forward smearing of first appearances of index species (Signor-Lipps
effect), and if they are extrapolated to whole basins. Here, we argue that (1) geochronological estimates of stage
boundaries need to be based on sections with high completeness and high sediment accumulation rates, and (2) that
the boundaries should preferentially correspond to conditions with sufficient marine connectivity between the Paratethys
and the open ocean. The differences between the timing of origination of a given species in the source area and timing of
its immigration to the Paratethys basins should be minimized during such intervals. Here, we draw attention to
the definition of the Central Paratethys regional time scale, its modifications, and its present-day validity. We suggest that
the regional time scale should be adjusted so that stage boundaries reflect local and regional geodynamic processes as
well as the opening and closing of marine gateways. The role of eustatic sea level changes and geodynamic processes in
determining the gateway formation needs to be rigorously evaluated with geochronological data and spatially-explicit
biostratigraphic data so that their effects can be disentangled.
Keywords: Neogene, Central Paratethys, regional and standard time scales, sea-level changes, geodynamics.
Introduction
Since the definition of the Paratethys Sea realm nearly a cen-
tury ago (Laskarev 1924) the correlation between the regional
Paratethys stages and the standard (Mediterranean) stages
remains poorly documented and validated (Magyar et al.
1999; Brzobohatý et al. 2003; Kováč et al. 2007; Hohenegger
et al. 2014; Sant et al. 2017b). The geochronological and bio-
stratigraphic definition of stage boundaries is one of the criti-
cal problems that limits understanding of the climatic, oceano-
graphic, and ecosystem history of the Paratethys (Rögl et al.
2003; Piller et al. 2007; de Leeuw et al. 2013;
Silye & Filipescu
2016). Although the regional time scale needs to be supported
by absolute dating, geochronological point-based data used
without sufficient knowledge of the local lithostratigraphic
nomenclature and not accounting for forward and backward
smearing of first and last appearances (Signor & Lipps 1982),
sequence- and bio-stratigraphic correlations within the Para-
tethys and between the Paratethys and other regions can be
misleading.
The Central Paratethys (CP) time scale was defined on the
basis of lithostratigraphy of sedimentary sequences belonging
to distinct transgressive–regressive cycles. Lithostratigraphic
boundaries coincide with immigration or evolutionary events
to some degree: the stages were defined on the basis of immi-
gration of new taxa from other bio-provinces (Atlantic,
Mediterranean, Indo–Pacific, and Eastern Paratethys) or by
evolution of endemic biota. Initially, the stage boundaries were
often represented by hiatuses or discordances that correspond
to boundaries between lithostratigraphic formations (Cicha et
al. 1967; Steininger & Seneš 1971; Papp et al. 1973, 1974,
1978, 1985; Báldi & Seneš 1975; Stevanović et al. 1989).
Later, definitions of regional stages have begun to prefer bio-
stratigraphic criteria (Piller et al. 2007). However, first appea-
rances of marine organisms in isolated basins tend to be
affected by significant delays relative to their first appearance
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in open ocean environments owing to geographical and envi-
ronmental constraints (e.g., Kennett et al. 1985; Holcová et al.
2015; Jones & Murray 2017; Sant et al., 2017a). Therefore,
the accuracy of correlations of stage boundaries based on
palaeontological criteria in type sections, without any geo-
chronological verification and without understanding of tem-
poral changes in biogeographic distribution of index species,
is unclear.
The most complete contemporary regional time scale pre-
sented by Krijgsman & Piller (2012), primarily based on bio-
stratigraphic criteria, is still in use (Fig. 1). New point-based
geochronological estimates were measured in order to deter-
mine the age of the regional stage boundaries and to allow
broad-scale correlations (Vasiliev et al. 2010; de Leeuw et al.
2013; Roetzel et al. 2014; Zuschin et al. 2014; Palcu et al.
2015; Sant et al. 2017a). The essential condition of such esti-
mates is that they should be derived from complete sedimen-
tary sequences, with a relatively high sediment accumulation
rate and conditions allowing preservation of index species.
However, these two conditions are rarely met in the CP.
Extrapolating the age of a stage boundary gained by geo-
chronological methods from a single site to a whole basin can
be a major problem in intra-basinal correlations based on sedi-
mentological criteria because depositional facies (e.g., deltas)
and benthic biofacies are frequently diachronous, and the first
and last appearances of planktic species can also be dia-
chronous due to migration, environmental and sampling con-
straints (MacLeod 1996). The point-based geochronological
data should thus be supported by biostratigraphic and biogeo-
graphic distribution of planktic organisms, with their first (FO)
or last (LO) occurrences and vice versa. However, the usage of
Atlantic biozonations of planktic foraminifers (Berggren et al.
1995) and calcareous nannoplankton (Martini 1971) is limited
in the Mediterranean and adjacent CP realm because planktic
foraminifers tend to be rare in isolated basins and some geo-
chronological data indicate age offsets between times of their
appearance in different basins (Iaccarino & Salvatorini 1982;
Iaccarino 1985; Mărunțeanu 1999; Andrejeva Grigo rovič et
al. 2001; Turco et al. 2002, 2017; Lirer & Iaccarino 2005;
Iaccarino et al. 2011; Paulissen et al. 2011; Gonera 2013;
Bartol et al. 2014; Di Stefano et al. 2015; Holcová et al. 2015;
Sant et al. 2017b). Therefore, it is necessary to date
the FO of planktic taxa in the CP by geochronological methods
and ensure that the extent of forward smearing will be assessed
with taphonomic, palaeoecological, and palaeobiogeographic
criteria. For example, if sample sizes are small, palaeoenviron-
mental conditions do not match preferences of index species
closely, and/or if circulation barriers exist between provinces,
the geochronological ages defined on the basis of FO in
a given section are likely to be younger than the real timing of
origination of a given species.
The calcareous nannofossil zonation (sensu Martini 1971;
e.g., FO of Helicosphaera ampliaperta defines the base of
the Eggenburgian (Burdigalian) in the NN2 Zone; LO of
Sphenolithus belemnos defines the Ottnangian in the NN3
Zone, LO of H. ampliaperta defines the boundary between
the NN4/NN5 zones, LO of Sphenolithus heteromorphus
defines the termination of “Early Badenian” in the NN5 Zone,
and FO of Discoaster kugleri occurs in the Sarmatian NN7
Zone), and planktic foraminiferal markers (sensu Piller et al.
2007; Filipescu & Silye 2008; Catapsydrax appears in the
Ottnangian, Trilobatus bisphericus (=Globigerinoides bisphe
ricus) appears in the late Karpatian, while Praeorbulina
glomerosa and Orbulina suturalis appear in the Early
Badenian) are used. The usage of benthic and/or endemic mol-
luscs or foraminifera for the definition of regional stage
boundaries can be unreliable (e.g., the Ottnangian/Karpatian
boundary is marked by FO of Uvigerina graciliformis).
Stratigraphic correlations based on the composition of benthic
assemblages can be biased by diachronous occurrence of
the benthic taxa temporally tracking their preferred environ-
ments, and can simply reflect an existence of environment that
is optimal for a given taxon during a certain time span (e.g.,
delta, shelf, basin slope, basin floor).
Geochronological methods can accurately estimate the onset
and duration of deposition of some specific sedimentary
facies. This can be achieved by dating points in sections
arrayed in vertical and horizontal transects across the basin
— generally, multiple such point estimates are required. For
example, Šujan et al. (2016) documented the diachronity of
sedimentation of the Pannonian formations in the Upper
Miocene infill of the Danube Basin.
The
spatial shift of facies
types within a given depositional system was demonstrated on
the basis of multiple point-based geochronological age esti-
mates of sediments belonging to the same facies type (and
lithostratigraphic formation) across the basin. The point-based
ages showed that the sedimentation along a shelf-slope-basin
transect lasted for more than 3 Ma, namely the time needed
until the basin was filled up. These results thus documented
the diachronity of sedimentation of the Pannonian formations,
with a lower boundary equal to the base of the Pannonian
regional stage. Therefore, the point-based age data can be
reliably used in order to correlate sections within a basin, but
also clearly show many inconsistencies in the correlation
between several CP basins.
In addition to biostratigraphy, the CP time scale published
by Piller et al. (2007) and Krijgsman & Piller (2012) also
comprised correlations with the global sea-level curve, and
the individual stage boundaries were correlated with the boun-
daries of the 3
rd
order cycles of the global sequence strati-
graphy (after Haq et al. 1988; Hardenbol et al. 1998). However,
the research carried out in semi-enclosed basins has shown
that the global sea-level change is captured by the sedimentary
record only to some degree (Kováč et al. 2004; Krézsek &
Filipescu 2005; Strauss et al. 2006). The active tectonics and/
or a huge amount of material input can intensify, reduce, or
completely hide the signatures of the global sea-level changes
(Kováč & Zlinská 1998; Kováč et al. 1998, 1999a,b, 2004;
Hlavatá-Hudáčková et al. 2000; Kováč 2000; Catuneanu
2006). In addition, the 3
rd
order sequence stratigraphic cycles
recorded in the CP respond not only to the effects of the Medi-
terranean, but also to the Eastern Paratethys water masses.
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REGIONAL TIME SCALE CORRELATION OF THE CENTRAL PARATETHYS IN THE MIOCENE
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, 2018, 69, 3, 283–300
Fig.
1.
Standard Neogene
chronostratigraphy
and biostratigraphy;
selected
regional
Miocene
time
scales
— an
overview
(for references
see figure); proposed regional
time
scale
and Central
Paratethys marine gateways. Explanatory notes: hatched area — not studied; BuSC — Burdigalian Salinity Crisis.
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Moreover, the relative sea-level curve of the Eastern Paratethys
differs significantly from the global eustatic sea-level curve
(Popov et al. 2010). Therefore, the impact of the Miocene
global sea-level changes on sequence-stratigraphic architec-
ture of basins in the CP should be re-evaluated.
In an inland sea, such as the CP, the global factors affected
the palaeogeographical evolution only partially, while impacts
of the local geodynamic processes were more critical (Kováč
2000; Kováč et al. 2016, 2017b; Sant et al. 2017b). The geo-
dynamic development of the Alpine–Carpathian–Dinaride
oro genic systems determined the distribution and extent of
terrestrial and marine environments and significantly shaped
the sequence architecture, palaeoclimate, and water masses
circulation regime of the CP (e.g., Kováč et al. 2004, 2017a;
Grunert et al. 2010, 2014; ter Borgh et al. 2013; Palcu et al.
2015). Intensity of marine currents and oceanic circulation
patterns strongly impacts biogeographic distribution of plank-
ton and benthos (e.g., Kennett et al. 1985; Peters et al. 2013;
Holcová et al. 2015; Jones & Murray 2017; Kováč et al. 2017a;
Sant et al., 2017a). The timing of faunal appearances thus
principally depends on the opening of gateways towards
the Mediterranean, or Eastern Paratethys as it is documented
from the (sub)recent Mediterranean or Black Sea (e.g.,
Kouwenhoven & van der Zwaan 2006; Karami et al. 2011;
Palcu et al. 2015; Kováč et al. 2017a; Sant et al. 2017b).
Therefore, the onset of regional stages should correspond to
conditions with a relatively high marine connectivity between
the Mediterranean and the CP, or at least the connection with
a substantially larger sea-covered area (Eastern Paratethys).
The present location of gateways that represent migration
corridors for marine organisms between the Mediterranean,
Central Paratethys and Eastern Paratethys realm, as well as
the distribution of the individual CP basins does not correspond
to their original position. The sedimentary fill of the Miocene
basins forms part of fold and thrust belts, or is dissected by
transform faults and the individual parts of basins were trans-
ported several hundreds of kilometres from their site of origin.
These changes in the location and configuration of sedimen-
tary basins were not taken into account in palaeogeographical
reconstructions for more than decades (e.g., Hámor & Halmai
1988; Popov et al. 2004; Sant et al. 2017b).
The view that the Outer Carpathian thrust belt was shor-
tened more than 150–200 km during the Miocene (e.g., Kováč
et al. 2017b) can be used as an example. The wide marine
realm in front of the moving orogenic wedge gradually shifted
towards the European platform margins; basins on the top of
the accretionary wedge were folded and thrust ahead (gene-
rally north- and east-ward). Basins on the platform margin
(foredeep depocenters) were diachronously filled up (e.g.,
Meulenkamp et al. 1996). This shortening had a massive
impact on the extent and distribution of marine and terrestrial
environments. Similarly, in the orogenic hinterland system,
the Upper Oligocene–Lower Miocene retro-arc basin was situa-
ted at least 200 km towards the southwest with respect to its
recent position (e.g., Tari et al. 1992; Kováč et al. 2016,
2017b). The basin was later divided into two parts due to
extrusion of the northern Pannonian crustal fragment from
the zone between the Alps and Dinarides, and reached its
present position in the late Early Miocene (e.g., Fodor et al.
1998; Kováč et al. 2016, 2017b). Therefore, geographic maps
not accoun ting palaeogeographic shifts are misleading and
cannot represent baselines for broader palaeogeographic
reconstructions. To evaluate changes in the configurations of
basins through time, an accurate palinspastic modelling based
on interdis ciplinary approach reflecting original position and
extent of basins is needed.
The considerable problem of the regional CP time scale is
that the individual stage boundaries are seldom supported by
up-to-date geochronological data and by biostratigraphic data
that would account for temporal changes in biogeographic dis-
tribution of index species, what makes correlation within indi-
vidual basins of the CP area, as well as with the Mediterranean,
troublesome. The use of regional stages without point-based
geochronological age data and sufficient knowledge of local
lithostratigraphic nomenclature, tectonics, and sequence stra-
tigraphy can be misleading in interregional correlations at
European scale. As we show below, the use of a standard time
scale is more appropriate.
For example, in an inspiring paper by Sant et al. (2017b),
the “Ottnangian Sea” at 18 Ma (see fig. 4A in Sant et al.
2017b) extends from the Alpine Foredeep (Molasse Basin)
across the hinterland of the Central Western Carpathians
(Novohrad–Nógrád Basin) to the area of the Eastern Slovakia
Basin. However, this time slice should be referred to as
the “early Burdigalian” because most of the Ottnangian strata
are not formed by marine sediments in the Novohrad–Nógrád
and Eastern Slovakia basins in the Western Carpathians.
The Ottnangian sediments are represented by terrestrial deposits
or by hiatuses in these basins (e.g., Vass et al. 1979; Rudinec
1989, 1990; Kováč et al. 1995; Vass 2002; Vass et al. 2007).
We note that marine sediments of the age around 18 Ma are
present in both basins, but they are assigned to the Eggenburgian
stage (Vass et al. 1979, 2007; Vass 2002; Fordinál et al. 2014;
Kováč et al. 2017a).
The “Karpatian Sea” with the age of ~16.5 Ma, and the
“Badenian Sea” with the age of less than ~14 Ma depicted in
figs. 4B and 4C (Sant et al. 2017b) represent a new period in
the CP development, prior to the Badenian Salinity Crisis
(BSC) and prior to the onset of the “Late Badenian Sea” trans-
gression, respectively. However, the base of the Badenian
regional stage is traditionally dated to ~16.4–16.3 Ma (e.g.,
Piller et al. 2007; Filipescu & Silye 2008; Hohenegger et al.
2014), whereas the “Karpatian Sea” in fig. 4B (Sant et al.
2017b) is dated to 16.5 Ma. Similarly, as in the previous
case, the standard Miocene chronostratigraphic terminology
should be used; this map captures remnants of the upper
Burdigalian sediments deposited prior to the Langhian
transgression.
Finally, the suggestion of Sant et al. (2017b) “the establish
ment of the “Badenian Sea” (<15.2 Ma), triggered by sub
ductionrelated processes in the Pannonian and Carpathian
domain, is significantly younger (by ~1 Myr) than usually
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estimated” cannot be accepted. The onset of the NN5 Zone
with Orbulina suturalis in the Early Badenian, coexisting with
Praeorbulina ssp. around 15 Ma is well documented in most
basins of the CP (e.g., Kováč et al. 2017a and references
therein).
To conclude, the regional stage boundaries in the CP need to
be dated by biostratigraphic approaches validated by geochro-
nological methods, and the role of gateways should be eva-
luated as a function of local tectonics and not only as a function
of broad-scale eustatic sea-level changes.
The definition of the Central Paratethys regional
stages and their validity
The Miocene time scale applied for the CP area in the
19
th
century compiled Mediterranean and regional stages such
as the Burdigalian and Helvetian for the Early Miocene,
Tortonian for the Middle Miocene, and Levantian stage for
the Pliocene (e.g., Mayer-Eimar 1858). In the 1970’s, the cur-
rently used CP regional stages were defined in the series of
books Chronostratigraphie und Neostratotypen (Cicha et al.
1967; Steininger & Seneš 1971; Papp et al. 1973, 1974, 1978,
1985; Báldi & Seneš 1975; Stevanović et al. 1989). However,
the lower boundaries of these stages were not precisely geo-
chronologically constrained, and in some cases the biostrati-
graphical definition of the stage boundaries was also
insufficient. These stages are applied for the sedimentary
record from the Alps across the Pannonian Basin System up to
the Carpathians, Dinarides, and Balkans. The conversion from
old to new stratigraphic nomenclature led to discrepancies in
duration of the sedimentary record assigned to the same stage
among different basins of the Eastern Alps and Western
Carpathians. For example, the sediments formerly assigned to
the Helvetian stage (Fig. 1) were partly correlated with the
Ottnangian and partly with the Karpatian stage (e.g., Rutsch
1958; Cicha & Tejkal 1959; Rögl et al. 1978; Roetzel et al.
2006). The same problem holds true for the “Tortonian” which
was ambiguously subdivided into sub-stages that did not cor-
respond to the Badenian biozones defined previously by Grill
(1943).
In the following text, the actual definition of CP regional
stages is summarized for the time span from 20.4 to 11.6 Ma.
Time scale modifications suggested over the last decades and
the validity of chronostratigraphic estimates of the boundaries
gained by point-based dating are discussed. Attention is also
drawn to deficiency in definition of stages often caused by
an ecostratigraphic approach.
The base of the Eggenburgian was situated by Piller et al.
(2007) coevally with the base of the standard Burdigalian
stage within the calcareous nannoplankton NN2 Zone, at
the sequence boundary Bur1 (~20.4 Ma). The Eggen burgian
transgression can be detected in the Alpine Foredeep and
in the Vienna Basin, but not in the northern realm of
the Pannonian Basin, where the NN1/NN2 boundary was in
the past correlated with the Egerian/Eggenburgian boundary at
22.8 Ma (sensu Vass & Elečko 1989). However, the FO of
Helicosphaera ampliaperta (correlated with the Aquitanian/
Burdigalian boundary at 20.43 Ma; sensu Gradstein et al.
2012) can be recognized in sediments in most CP basins
(Holcová 2002; Krijgsman & Piller 2012). Therefore, this bio-
stratigraphical event at the Bur1 boundary can be accepted as
a reliable level enabling correlation between the regional and
standard zonation (Fig. 1).
The Ottnangian regional stage (~18.3–17.3 Ma; sensu Piller
et al. 2007) lower boundary was placed in the NN3 Zone,
while the upper boundary was situated within the NN4 Zone,
bounded approximately by the Bur3 and Bur4 3
rd
order
sequence boundaries (after Haq et al. 1988 and
Hardenbol et
al. 1998).
The Karpatian stage (~17.3–16.4 Ma; sensu Rögl et al.
2003) was situated inside the NN4 Zone as well, and its lower
boundary was defined by the FO of endemic Uvigerina
graciliformis. Nevertheless, the new magnetostratigraphic
constraints provided by Sant et al. (2017a) dated the transition
from the Ottnangian marine to brackish sediments in the south-
German part of the Alpine Foredeep (Molasse Basin) to
~17.7–17.5 Ma. Termination of the brackish depositional
environment in the Austrian part of the foredeep was dated to
~17.2 Ma (Roetzel et al. 2014), while the Karpatian marine
sedimentation in the Korneuburg Basin was dated by astro-
nomical tuning of the gamma ray record to the time interval
from 17.0 to 16.3 Ma by Zuschin et al. (2014). These results,
supported by
87
Sr/
86
Sr isotope dating from the Vienna Basin
(Hudáčková et al. 2003) point to an insufficiently defined
boundary between the Ottnangian and the Karpatian. The fora-
minifera tests from deposits assigned to the Ottnangian in
the Cunín-21 borehole provided the Sr-age of 17.01–16.9 Ma.
Foraminifera from the Karpatian strata of the Gbely-100 bore-
hole provided Sr-age of 16.3–15.9 Ma (Hudáčková et al.
2003). Moreover, Sr-age gained from the Cerová-Lieskové
site assigned to the Karpatian is 16.26–15.47 Ma (Less et al.
2015; Kováč et al. 2017a).
Differences in age estimates of the Ottnangian/Karpatian
boundary probably led to incorrect correlations of sedimentary
successions in an interregional context. We assume that
the sediments of the same age were in the Alpine Foredeep
(Molasse zone) assigned to the Ottnangian and in the northern
part of the Vienna Basin to the Karpatian, both assigned to
these stages on the basis of NN4 Zone. This assumption can be
supported by a distinct angular unconformity between the two
“Karpatian” sedimentary formations in the northern Vienna
Basin (fig. 9 in Kováč et al. 2004). The lower “Karpatian”
complex was possibly deposited during the “Ottnangian”
closing of the marine connection towards the Mediterranean in
front of the Alps, and the overlying complex was deposited
during the “Karpatian” opening of the new marine connection
via the Trans-Tethyan-Trench Corridor (sensu Rögl 1998;
Mandic et al. 2002; Kováč et al. 2007; Rasser et al. 2008).
To test this assumption, geochronological data obtained from
basins situated close to the gateways between the CP and
the Mediterranean can be used.
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The Karpatian/Badenian boundary, set within the NN4
Zone, was initially correlated with the boundary between
the Burdigalian and the Langhian, thus, with the boundary
between the Early and Middle Miocene (sensu Blow 1969;
Rögl et al. 1978; Piller et al. 2007). Currently, there is no con-
sensus on the placement of the Early/Middle Miocene boun-
dary in the CP. Piller et al. (2007) correlated it with the Bur5/
Lan1 sequence boundary, while Hohenegger et al. (2014)
shifted the base of the Badenian into the Burdigalian stage,
corresponding to the FO of Praeorbulina at ~16.4 Ma in
the Styrian Basin (Hohenegger et al. 2009). De Leeuw et al.
(2013) placed the FO of Praeorbulina glomerosa on 16 Ma
in the Transylvanian Basin. Krijgsman & Piller (2012)
placed the Karpatian/Badenian boundary at 15.97 Ma (Fig. 1).
The Globigerinoides–Praeorbulina lineage is conti nuously
recorded in the Styrian, Sava, and Transylvanian basins
(Krézsek & Filipescu 2005; Hohenegger et al. 2009, 2014;
Premec Fućek et al. 2017). The FO of Praeorbulina cannot
always be estimated in the Western Carpathian basins because
they are extremely rare or absent (e.g., Andrejeva Grigorovič
et al. 2001; Kováč et al. 2007; Rögl et al. 2008). The definition
of the Badenian stage lower boundary was designated by
the onset of Praeorbulina sicana (currently accepted as
Trilobatus sicanus; erroneous synonym “Globigeri noides
sicanus” is used by some authors either for G. bisphe ricus or
for Pr. sicana) within the NN4 Zone (16.303 Ma at
the top of chron C5Cn.2n) at the Wagna site in the Styrian
Basin by Rögl et al. (2003) and confirmed by Hohenegger et
al. (2009, 2014).
The original sub-division of the Badenian regional stage
into the Early (Moravian), Middle (Wielician) and Late
(Kosovian) sub-stages (Papp et al. 1978; Piller et al. 2007)
remains problematic as well. According to Hohenegger et al.
(2014), the Wielician sub-stage, namely the evaporite sequence
at/below the base of the NN6 Zone, cannot be simply coeval
with the “Middle Badenian” zone with agglutinated forami-
nifera in the western part of the CP because this foraminifera
zone covers a much longer time span (upper part of the NN5
and the lowermost part of the NN6 zones; Andrejeva Grigo-
rovič et al. 2001). Therefore, instead of referring to the
“Wielician sub-stage”, it is more appropriate to use the term
Badenian Salinity Crisis (BSC).
The BSC is a reasonable correlation interval, with the dura-
tion of approximately 500 kyr between ~13.8–13.3 Ma, which
is well documented in the eastern part of the CP (e.g., Filipescu
& Gîrbacea 1997; Krézsek & Filipescu 2005; Peryt 2006;
de Leeuw et al. 2010, 2013). This interval has also been
detected in the sediments of the Pannonian realm (Báldi et al.
2017) and in the wider Mediterranean area (Ied et al. 2011).
The base of the BSC (when dated by the geochronological
methods) can thus be a reliable correlation level for the CP
because it seems to be synchronous with the Langhian/
Serravallian boundary, corresponding to a major glacioeustatic
sea-level drop (sensu Gradstein et al. 2012).
Hoheneggers’ et al. (2014) attempt to solve the “Badenian
conundrum” brought even more confusion into the CP
stratigraphy (Fig. 1). Although the “Middle Badenian” sub-
stage Wielician was not accepted by Hohenegger et al. (2014)
and the BSC range was assigned to the base of the Late
Badenian (Hohenegger et al. 2014), the term “Wielician sub-
stage” is still used in studies from the Eastern Carpathian
region (e.g., de Leeuw et al. 2013; Gonera et al. 2014; Palcu et
al. 2015). It is, however, improper to consider the Moravian
sub-stage within the NN5 Zone, introduced for the “Early
Badenian” by Papp et al. (1978), as the (re)established “Mid
Badenian” (Hohenegger et al. 2014).
Another attempt to correlate the Badenian regional stage
with the standard Mediterranean time scale resulted in the divi-
sion of Badenian into lower and upper parts, roughly corre-
sponding to the Langhian and early Serravallian (Kováč et al.
2007). This definition, placing the BSC at the top of the Early
Badenian, led to a shift of the Late Badenian lower boundary
to 13.63 Ma (instead of 13.82 Ma).
The Sarmatian stage defined as Sarmatian s.s. and Sarmatian
s.l. is difficult to correlate even between the western and
eastern part of the CP (e.g., Suess 1866; Papp et al. 1974).
In the eastern part, the Sarmatian s.l. is divided into sub-stages
Volhynian, Bessarabian, and Khersonian, thus a subset of
the Sarmatian s.l. corresponds to the regional Pannonian stage
in the west (e.g., Piller et al. 2007; Popov et al. 2010; Gozhyk
et al. 2015). For the subdivision of the Sarmatian
s.s.
(12.7–11.6 Ma; sensu
Piller et al. 2007), four successive zones
(Anomalinoides dividens, large elphidia, E. hauerinum,
Porosononion granosum) are used (sensu Grill 1941, 1943;
Papp 1951; Harzhauser & Piller 2004). However, preliminary
analyses of foraminiferal assemblages from boreholes cores
positioned in a 3D seismic model in the northern Vienna Basin
indicate that these assemblages track tempo rally shifting envi-
ronments and their temporal distribution depends strongly
on the former basin topography (Hudáčková et al. 2013).
In the Transylvanian Basin the Badenian/Sarmatian boundary
was dated by the
40
Ar/
39
Ar method to 12.80±0.05 Ma
(de Leeuw et al. 2013). This datum is similar to the one brought
by Harzhauser & Piller (2004) on the basis of sequence strati-
graphy from the western margin of the Central Paratethys, and
correlates with the magnetostratigraphic results of Paulissen et
al. (2011) from the Vienna Basin.
The lower boundary of the Sarmatian s.s. was set by
Harzhauser & Piller (2007) to the extinction event at the Badenian/
Sarmatian boundary (BSEE). However, it seems that the BSEE
timing is diachronous due to complex tectonic evolution of
the Carpathian–Pannonian region, reflecting the final isolation
of the CP from neighbouring basins — the Mediterranean and
Eastern Paratethys (e.g., Magyar et al. 1999; de Leeuw et al.
2013; Palcu et al. 2015; Kováč et al. 2017a,b and references
therein). Therefore, point-based data are necessary to support
this hypothesis, as already suggested by Silye & Filipescu
(2016).
Similarly, the extinction event at the top of the Sarmatian
s.s. (SPEE) has been placed at different levels. Harzhauser
& Piller (2004) placed it at 11.6 Ma. Similar results for
the Sarmatian/Pannonian boundary were brought by Paulissen
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et al. (2011) and ter Borgh et al. (2013) by magnetostrati-
graphy from the Vienna Basin and the southern Pannonian
Basin. Dating of two volcanoclastic layers located approxi-
mately 40 m below the Sarmatian–Pannonian tran
sition
(Transylvanian Basin) yielded
40
Ar/
39
Ar ages of 11.62 ± 0.12
Ma and 11.65 ± 0.13 Ma (de Leeuw et al. 2013). Based on
the sequence stratigraphic correlations to global sea level
curves in the Vienna Basin Lirer et al. (2009) estimated the
Sarmatian/Pannonian boundary at 11.4 Ma. The Sarmatian/
Pannonian boundary in the Transylvanian Basin was dated to
an age of 11.3 ± 0.1 Ma (Vasiliev et al. 2010). The correlation
of the NN8 Zone in the Paratethys domain (based on the pre-
sence of Catinaster coalitus; Galović & Young 2012) with
magnetostratigraphic data points to the Sarmatian–Pannonian
transition in the Croatian Basin around 11.2 Ma.
The perspectives of regional and standard time
scale correlation
On the one hand, the correlation between sedimentary
sequences is relatively simple if only one time scale is used for
a single basin. On the other hand, it is difficult to compare time
scales of basins characterized by their partial isolation from
the World Ocean (WO) with the standard time scale (GTS).
First and last appearances of species represent one of the major
biostratigraphic tools. However, discrepancies in the timing of
first appearances of particular species exist even between the
Atlantic and Mediterranean, and such discrepancies can be
expected to be more pronounced between the CP and other
basins.
For example, the Early/Middle Miocene boundary is appa-
rently correlated with the Burdigalian/Langhian stage boun-
dary (Hilgen et al. 2012; Turco et al. 2017). Using the latest
Time Scale Creator database (Fig. 1) the base of Langhian
stage is correlated with the base of magnetic polarity chron
C5Br (15.97 Ma) and with the FO of Praeorbulina circularis
which is in accordance with the scale of Ogg et al. (2016).
The boundary is not officially established, so the reliability of
such “praeorbulina datum” can be doubtful (Lirer & Iaccarino
2011). In this context, the FO of Praeorbulina glomerosa
glomerosa at 15.2 Ma is the key event in the Mediterranean
(Iaccarino et al. 2011; Turco et al. 2017) while it occurs at
16.4 Ma in the WO (Wade et al. 2011). Orbulina suturalis
appears at 14.6 Ma in the Mediterranean (Abdul Aziz et al.
2008; Di Stefano et al. 2008) and at 15.1 Ma in the WO (Wade
et al. 2011), as reviewed by Sant et al. (2017b). The Langhian
delay of the FO of the planktic foraminifera in the Medi-
terranean compared to the Atlantic Ocean can be explained by:
(i) circulation patterns that did not allow immediate migration
of planktic species to the Mediterranean and/or (ii) the estab-
lishment of conditions for survival of these species in the
Medi terranean, which was influenced by the inflow of CP
water masses into the Mediterranean realm at that time (Kováč
et al. 2017a; Sant et al. 2017b). Moreover, in the Mediterranean,
the last common occurrence (LCO) of Helicosphaera ampliaperta
is dated to 16.1 Ma (Iaccarino et al., 2011). The LO of
H. ampliaperta (~14.9 Ma in WO) defining the top of the NN4
Zone cannot be properly recognized in the Mediterranean
(Di Stefano et al. 2008, 2015). Therefore, using this event for
the NN4/NN5 boundary accompanied by the FO of Orbulina
suturalis at 14.6 Ma in the Mediterranean (Abdul Aziz et al.
2008) while in the WO it appears at 15.1 Ma (Gradstein et al.
2012) is not satisfactory.
Inconsistencies generated by converting the regional stages
to standard ones are partly caused by the lack of multiple
point-based geochronological data, by inadequate biostrati-
graphic data that do not account for temporal shifts in geo-
graphic ranges of index species, and by problems with local
nomenclature in lithostratigraphy. Therefore, the correlation
of individual basins within the CP as well as with the Medi-
terranean or Eastern Paratethys realms without accurate geo-
chronological data remains problematic.
Reflection of eustatic sea-level changes in
the Central Paratethys time scale
The problems of CP sequence stratigraphy are well docu-
mented in the Vienna, Danube, Transylvanian, and other
basins (e.g., Kováč 2000; Kováč et al. 2004, 2007, 2008;
Krézsek & Filipescu 2005). The Miocene depositional sequen-
ces reveal several 3
rd
and 4
th
order cycles that were generated
by eustatic sea-level changes, tectonic evolution of basins, and
local sediment supply delivered by deltas. However, the global
sequence boundaries sensu Hardenbol et al. (1998) were tied
to the regional stage boundaries (Piller et al. 2007). This
sequence–stratigraphic definition of stage boundaries partly
contrasts with the local sequence stratigraphy following local
geodynamic events as demonstrated by Kováč et al. (2004).
Moreover, due to active tectonics and rapid palaeogeographic
changes in the Alpine–Carpathian–Dinaride domain, it is diffi-
cult to discriminate between the 3
rd
and 4
th
order cycles during
the Miocene (Fig. 2).
Around the Aquitanian–Burdigalian transition, the marine
connections of the CP with Mediterranean probably led through
a strait between the Alps and Dinarides. The connections
between the Central and Eastern Paratethys (and possibly up
to Indo–Pacific) via a strait between the Volhynian High and
Moesia were gradually closing (e.g., Popov et al. 2004; Kováč
et al. 2017a,b).
The following early Burdigalian sea-level changes in the CP
were probably influenced by the sea-level rise or fall trans-
ferred from the Mediterranean through a new connection in
front of the Alps (e.g., Rögl 1998; Harzhauser & Piller 2007;
Kováč et al. 2017a,b). The gateway opened at the Bur1 boun-
dary (~20.4 Ma) and closed ~17.7–17.5 Ma, as constrained by
magnetostratigraphic data (Sant et al. 2017a). According to
GTS, the gateway was closed at the Bur4 boundary (17.5 Ma
after Hardenbol et al. 1998; Piller et al. 2007). During this time
interval two 4
th
order cycles (Eggenburgian and Ottnangian;
Fig. 2) were documented in the Vienna Basin and adjacent
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basins to the north-east (Kováč 2000; Kováč et al. 2004).
The regression of the Ottnangian Sea and development of
the “Rzehakia Lake” (e.g., Rögl 1998) is a significant marker
horizon preceding the Karpatian full marine transgression in
the Alpine–Carpathian junction.
The Early Miocene eustatic changes in the Novohrad–Nógrád
Basin represent two 3
rd
order cycles (Egerian–Eggenburgian
and Eggenburgian–Ottnangian; Kováč 2000) and thus cannot
be directly compared with the Vienna Basin 4
th
order cycles
(Fig. 2). A similar situation can be deduced from the lithostra-
tigraphy and micropalaeontology of the Transylvanian Basin
(Filipescu 2011), where an Egerian–Eggenburgian and
Eggenburgian–Ottnangian cycle can be distinguished as well.
The deep-sea equivalent of the Eggenburgian onshore forma-
tions corresponds to the lower part of mid-fan turbidites of
the Hida Formation. The upper part of this formation displays
a regressive trend towards the boundary with the Middle
Miocene. The nannoplankton assemblages indicate Early
Miocene age (NN2 to NN4 zones; Mészáros 1991; Beldean et
al. 2010) while the agglutinated and planktic foraminifera
point to a wider span (Iva et al. 1971; Beldean et al. 2010;
Beldean & Filipescu 2011). Therefore, the correlation between
Fig. 2. Central Paratethys sequence stratigraphy; 3
rd
and 4
th
order cycles of the entire Central Paratethys (this work), as well as the Novohrad–
Nógrád, Transylvanian and Vienna basins (after Kováč 2000; Krézsek & Filipescu 2005; Kováč et al. 2007; Pezlej et al. 2013).
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the Early Miocene western and eastern CP basins sequence
stratigraphy remains unclear (Fig. 2).
The late Burdigalian eustatic sea-level rise was transferred
from the Mediterranean into the CP via a new marine gateway
— the Trans-Tethyan-Trench Corridor (sensu Rögl 1998;
Mandic et al. 2002; Kováč et al. 2007; Rasser et al. 2008).
The early Karpatian full marine transgression (NN4 Zone with
Uvigerina graciliformis) represents the next pronounced sea-
level change in the Vienna Basin. The base of this 4
th
order
cycle can be coeval with Bur4 boundary (sensu Hardenbol
et al. 1998). The next late Karpatian 4
th
order cycle inside
the NN4 Zone is
situated above the lower Karpatian deposi-
tional sequence (Kováč et al. 2004). In the Novohrad–Nógrád
Basin, the “Karpatian” transgression started at the top of
Ottnan gian “Oncophora beds” (Vass & Elečko 1992; Holcová
2001) and was followed by offshore deposition inside the NN4
Zone (Kováč 2000). The local 3
rd
order cycle is capped by
sediments of the Early Badenian transgression (Fig. 2).
The Karpatian stage was terminated by a regression causing
large-scale erosion in the northern Vienna Basin around
the sequence boundary of the local VB4/VB5 3
rd
order cycles
(Kováč et al. 2004). The sedimentation above this boundary
continued in the form of lobes of deltaic and alluvial sedi-
ments, followed by the Early Badenian transgression in
the entire basin and in the junction towards the Alpine Molasse
Zone (sensu Kováč et al. 2004). Strauss et al. (2006) correlated
this lower boundary of the Badenian cycle with the Bur5/Lan1
3
rd
order cycle and placed it at 16.4 Ma (after Hardenbol et al.
1998).
The erosion and deposition around the Karpatian/Badenian
boundary fits well with the regressive phase during the late
Karpatian and the Early Badenian transgression. The geochro-
nological point-based data from the Vienna and Novohrad–
Nógrád basins (Hudáčková et al. 2003; Fordinál et al. 2014;
Kováč et al. 2017a) show that the age of the top part of
the Karpatian eustatic cycle (<16 Ma) does not coincide with
the global Bur5/Lan1 boundary (16.4 Ma; sensu Hardenbol et
al. 1998; Krijgsman & Piller 2012), and also does not coincide
with the base of the “earliest” Badenian (sensu Hohenegger
et al. 2014). The relative sea-level fall (prior to the Early
Badenian transgression) in the CP estimated as up to 200 m
was enhanced by the mountain uplift (compiled from Krézsek
& Filipescu 2005; Dellmour & Harzhauser 2012; Filipescu
2011; Kováč et al. 2017a). Therefore, the sea-level fall proba-
bly began after ~16.5 Ma (Fig. 2) and the sea-level low-stand
probably lasted until ~15.5–15.1 Ma, when the Badenian
transgression was initiated. In this case, the absolute age of
the Early Badenian sequence boundary does not simply
coincide with the Bur/Lan1 boundary.
The termination of the Vienna Basin initial rifting led to
a decrease in subsidence rates and to a very indistinct reflec-
tion of the global TB2.3 cycle (16.5–15.5 Ma; sensu Haq et al.
1988; Haq 1991; Hardenbol et al. 1998). Conglomerates at
the base of the local 4
th
order Early Badenian cycle (defined as
the 3
rd
order VB5; Kováč et al. 2004) are overlain by pelites
dated by Kroh et al. (2003) and placed to the uppermost part of
the “Lower Lagenidae Zone”
(sensu Grill 1943) based on
co-occurrence of Praeorbulina glomerosa circularis, Orbulina
suturalis, and Trilobatus bisphericus. The nannoplankton
assemblage with Helicosphaera waltrans, Sphenolithus
heteromorphus, Calcidiscus premacintyrei, Reticulofenestra
pseudoumbilicus, Coccolithus miopelagicus, rare Discoaster
deflandrei and D. variabilis indicates that the pelites belong to
the NN5a Zone of the “Early Badenian” (sensu Kováč et al.
2007).
Strauss et al. (2006) documented the local “Early”
Badenian cycle as an equivalent of the 3
rd
order TB2.3 cycle
(15.97–14.4 Ma; sensu Hardenbol et al. 1998) in the SE
Vienna Basin. This local cycle is situated below the next
“Middle” Badenian cycle, as an equivalent of the 3
rd
order
TB2.4 cycle (14.4–13.65 Ma; sensu Hardenbol et al. 1998).
A similar, local 3
rd
order cycle was documented in the southern
Pannonian Basin System (Pezelj et al. 2013).
The age of the Early Badenian transgression can be deduced
from borehole cores in the eastern Danube Basin (Kováč et al.
2018). The Badenian basal conglomerates and silts, both with-
out volcano-clastics, are overlain by siliciclastics with tuffites.
These sediments contain nannofossils of the NN5a Zone with
common Orbulina suturalis (FO of O. suturalis at 14.56 Ma;
Abdul Aziz et al. 2008). The volcanic activity related to
the basin opening started at 15 Ma (Pécskay et al. 2006) and
points to the age of the marine flooding with Praeorbulina spp.
prior to the deposition of volcanoclastic sequences (Fig. 2).
Similarly, in the Transylvanian Basin, the onset of Dej Tuff
volcanism dated by
40
Ar/
39
Ar method to 14.38±0.06 Ma is also
preceded by the FO of Praeorbulina spp. and Orbulina suturalis
(de Leeuw et al. 2013).
The boundary between the “Early” Badenian (VB5) and
“Middle” Badenian (VB6) 4
th
order cycles (3
rd
order; sensu
Kováč et al. 2004) corresponds to the sequence boundary
within the “Upper Lagenidae Zone” proposed by Weissenbäck
(1996) in the southern Vienna Basin. The “Middle” Badenian
cycle (uppermost Lagenidae Zone – lower Bulimina–Bolivina
Zone; sensu Grill 1943) covers the time span of the NN5 Zone
upper part, and the maximum flooding surface was identified
by Weissenbäck (1996) within the Spiroplectammina carinata
Zone. The high-stand system deposits are capped with the base
of the Late Badenian (VB7) local 3
rd
order cycle (Harzhauser
et al. 2018).
In the Transylvanian Basin, the sedimentary record at
the Early–Middle Miocene transition offers similar proofs in
the form of changes in sedimentary facies and microfossil
assemblages. The upper part of the Lower Miocene sediments
contains foraminifera assemblages dominated by planktic
Streptochilus pristinum associated with rare benthics
(Bulimina, Bolivina, Cibibicidoides) and calcareous nanno-
plankton, probably indicating a late Burdigalian age (Beldean
et al. 2010, 2013). The sea-level drop (100–200 m) is docu-
mented by several deep incised valleys (Krézsek & Filipescu
2005). The overlying Middle Miocene deposits comprising
tuffites and fall-out tuffs interbedded with low density fine
siliciclastics contain typical Lower Badenian planktic forami-
nifera (species belonging to genera Praeorbulina, Orbulina,
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Globigerinoides, Globorotalia). These deposits can be linked
to the main phase of marine transgression that started at
the beginning of the Middle Miocene. The “Early” and “Middle”
Badenian depositional cycles correlated with the 3
rd
order
TB2.3 and TB2.4 cycles of global sea-level change (sensu
Hardenbol et al. 1998) terminated
prior to the BSC (Krézsek
& Filipescu 2005). The plankton
bloom in the Praeorbulina
glomerosa Biozone (M5a) was followed by bloom in the
Orbulina suturalis Biozone (M5b) Krézsek & Filipescu (2005).
The lower boundary of the Late Badenian 3
rd
order cycle VB7
in the Vienna Basin (sensu Kováč et al. 2004) is represented
by subaerial erosion in its NE part.
The Late Badenian depositional systems in the SW Vienna
Basin were considered to be of regressive origin (Kreutzer &
Hlavatý 1990; Weissenbäck 1996), while Kováč et al. (2004)
defined a complete 3
rd
order cycle (VB7) in the NE Vienna
Basin. Later on, this cycle was defined for the CP (13.65–12.7
Ma; Kováč et al. 2007) and correlated with TB2.5 global cycle
(13.8–12.6 Ma; Haq et al. 1988). The latest research in the
northern Vienna Basin has confirmed the sequence boundary
between the “Middle” and “Late” Badenian (Harzhauser et al.
2018). Moreover, a huge sea-level drop is correlated with
the base of the BSC and thus with the base of the Serravallian
at 13.8 Ma (Harzhauser et al. 2018). This actually indicates
that the base of the TB2.5 is captured by the Vienna Basin
sequences (Fig. 2). The TB2.5 cycle in the Transylvanian
Basin is correlated with two local 4
th
order depositional cycles:
the Badenian and the early Sarmatian lasting from BSC base
to top of the Anomalinoides dividens Biozone (Krézsek &
Filipescu 2005).
The Badenian/Sarmatian sequence boundary in the Vienna
Basin is placed at the biostratigraphic boundary defined by
molluscs and foraminifera turnover (sensu Harzhauser &
Piller 2007) affected by salinity decrease (Kováč & Hudáčková
1997). On the other hand seismic lines and well-logs show
overlap of the VB7 into the earliest Sarmatian sediments
(Harzhauser & Piller 2004). The base of the local Sarmatian
VB8 3
rd
order cycle (sensu Kováč et al. 2004) is well recorded
by a transgressional overlap on the Upper Badenian sediments.
The falling sea-level in the terminal Sarmatian (uppermost
Porosononion granosum Zone = “pauperization” Zone; sensu
Papp 1956) caused a shift of the littoral zone far into the
basin, indicated by littoral potamidid-bearing sand with
scattered coal in the basin drillings (Harzhauser & Piller
2004). The reg ression at the end of the Sarmatian is also indi-
cated by local erosions and incision of deltaic feeding
channels.
In the Transylvanian Basin, the Sarmatian deposits repre-
sent a single 3
rd
order depositional cycle. In contrast, two
Sarmatian 4
th
order cycles consisting of parasequences were
documented the Vienna Basin (e.g., Harzhauser & Piller
2004). These parasequence sets are present in the entire basin,
as well as in other basins of the Carpathian–Pannonian region
(Styrian and Transylvanian basins), suggesting that they were
governed by orbital impulses — a common feature of different
basins in CP realm (Kováč et al. 2008).
To summarize, extensive erosion characterized the Burdi-
galian–Langhian transition due to sea-level drop in the CP at
~16–15.5 Ma. The Early Badenian 3
rd
order eustatic cycle
ended prior to the BSC (Figs. 1 and 2). The younger 3
rd
order
cycles are marked by the Late Badenian and Sarmatian
transgressions. The three Middle Miocene Central Paratethys
3
rd
order cycles of sea-level changes can be only partly
correlated with the Langhian and Serravallian global sea-level
changes (sensu Kováč 2000; Krézsek & Filipescu 2005;
Strauss et al. 2006; Kováč et al. 2007). Additional geochrono-
logical data are needed to improve correlation of depositional
sequences between the CP basins and to untangle the effects of
regional tectonics from the effects of global eustatic changes.
The Central Paratethys time scale adjusted
to geodynamic development
(i) The geochronological definition of regional stage boun-
daries, (ii) the appropriate application of the point-based data
supported by well-defined biostratigraphic correlation levels,
(iii) the refined CP sea-level changes, (iv) the interpretation of
the plankton and benthos migration driven by opening and
closing of gateways between the Mediterranean, the Central
Paratethys, and the Eastern Paratethys, associated with tapho-
nomic and palaeoecological inferences on the role of rewor-
king, preservation and habitat suitability in determining FO
and LO in individual sections, (v) and the palinspastic
approach should result in the reappraisal tuning of the CP time
scale in respect to geodynamic processes, enabling better cor-
relation with the standard chronostratigraphy of the Miocene
period (GTS). Below, we propose three intervals of the CP
evolution with respect to geodynamic development of the area
and different positions of sea gateways; more likely as a reflec-
tion of geodynamically induced changes and only partially
corresponding to changes in the global sea-level.
The Burdigalian transgression represents the onset of pro-
nounced 3
rd
order sequence stratigraphy cycle in the northern
CP, correlating with the Bur1 boundary (Hardenbol et al.
1998; Piller et al. 2007; Krijgsman & Piller 2012). The base of
the Eggenburgian is dated by the FO of Helicosphaera
ampliaperta, like the base of the Burdigalian stage in the
Mediterranean (20.4 Ma; Piller et al. 2007; Ogg et al. 2016).
During the Eggenburgian and Ottnangian, after the closing
of the earliest Miocene connections towards the Eastern
Paratethys and Mediterranean (e.g., Popov et al. 2004; Kováč
et al. 2017a,b and references therein) a new marine flooding
from the Mediterranean went through the foredeep basin in
front of the Alps. The time span of this connection is docu-
mented by the presence of calcareous nannofossil NN2, NN3,
and a part of NN4 zones (sensu Martini 1971). The gateway
opened around the Aquitanian/Burdigalian boundary and
the sea (Fig. 3A) flooded the foreland and hinterland of the
deve loping Carpathian mountain chain (e.g., Kováč et al.
2017b and references therein). In the distal part of the CP,
the isolation led to development of hypersaline facies, later
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also to hyposaline facies, probably due to spatio–temporal
shifts in rainfall distribution (e.g., Kováč et al. 2017a). In the
eastern segment of the Carpathian Foredeep, evaporites of
the Vorotyshche Formation were deposited during the Eggen-
burgian (e.g., Gozhyk et al. 2015), while a system of brackish
and freshwater lakes with endemic Rzehakia fauna developed
in the late Ottnangian (Harzhauser & Piller 2007; Harzhauser
& Mandic 2008). In the foreland and hinterland of the deve-
loping Carpathian mountain chain, as well as in some parts of
the Pannonian domain, the terrestrial (lake) sedimentation
prevailed between 18 and 17 Ma. In this area, situated along
junction of the Central Western Carpathians and Northern
Pannonian domain a continuous Early Miocene terrestrial
sedimentation is documented by radiometric dating at 17.4–
17.02 Ma (Pálfy et al. 2007). However, along the northern
margin of the Pannonian domain (Novohrad–Nógrád Basin),
the “Ottnangian” sediments containing Rzehakia fauna are
occasionally accompanied by the Karpatian index species
Uvigerina graciliformis (Holcová 2001).
The gateway in front of the Alps disappeared before the end
of the Early Miocene (~17.5 Ma). Roughly at the same time
a new marine strait between the Mediterranean and CP opened
in the hinterland of the Eastern Alps, following the northern
edge of Dinarides (Fig. 3B). The so-called Trans-Tethyan-
Trench Corridor (e.g., Rögl 1998; Piller et al. 2007; Sant et al.
2017b) was active during the upper part of NN4 and NN5
zones (sensu Martini 1971). The base of the local early
Karpatian 4
th
order cycle in the CP can be approximately
correlated with the Bur4 boundary, or slightly above it.
The deposition of the “upper Karpatian–lowermost Badenian”
sediments was associated with significant changes in geomor-
phology, especially with the uplift of mountain chains, accom-
panied by local fluctuations in humidity. During the relative
sea-level fall by up to 200 m, a huge erosion of mostly
Karpatian strata and development of a pronounced 3
rd
order
sequence boundary (placed above Bur5/Lan1 boundary of
GTS) is assumed. The development of a new river network
caused the input of voluminous masses of fresh water into
the sea which probably triggered the switch of circulation
regime, at least in the western part of the CP during this time
(Fig. 1). The shift from an anti-estuarine to an estuarine
circulation regime (during the latest Burdigalian and early
Langhian) delayed plankton immigration into the CP (Kováč
et al. 2017a) and probably also influenced the marine environ-
ment in the adjacent Mediterranean area (e.g., problem with
the Burdigalian/Langhian boundary definition in the Medi-
terranean; Iaccarino et al. 2011; Lirer & Iaccarino 2011).
During the Early Badenian transgression, an anti- estuarine
regime between the CP and the Mediterranean was re-estab-
lished again (Kováč et al. 2017a). The base of this Cen tral
Paratethys 3
rd
order cycle can be placed inside the Langhian
3
rd
order sequence of GTS (below the maximal flooding on
the global sea-level curve) bordered by the Bur5/Lan1 and
Ser1 boundaries. This assumption is supported by the Lower
Badenian sediments with the FO of Orbulina suturalis together
with the NN5 Zone at ~14.6 Ma. However, the occurrence
of the Praeorbulina spp. in several CP basins could
suggest the Middle Miocene transgression around 16–15 Ma
(Fig. 2).
From what has been discussed above, a couple of questions
arise: where is the boundary between the Lower and Middle
Miocene in the sedimentary record, and how should we under-
stand the Karpatian regional stage? In other words: Does
the sedimentary sequence assigned to the Karpatian belongs to
the Early Miocene time span? We suggest that the lower part
of the deposits assigned to the Karpatian regional stage
belongs to the Early Miocene, whereas the upper part belongs
to the Middle Miocene.
The temporal span of the Karpatian regional stage remains
unclear. The Karpatian marine transgression is documented
from the southern Vienna Basin at ~17 Ma (Zuschin et al.
2014), while the base of the Badenian is placed at ~16.4 Ma
(e.g., Piller et al. 2007; Hohenegger et al. 2014), thus the Kar-
patian stage is just limited to 600 ky. We note that the lower
boundary of the Badenian stage, as suggested by Papp et al.
(1978), should be placed at the beginning of the NN5 Zone,
which means ~15 Ma, whereas Hohenegger et al. (2014)
considered the time interval between 16.3 and 15.1 Ma as
the “lowermost” Badenian. Following Papp et al. (1978),
the resulting time interval would last ~2 Ma (i.e. “Karpatian–
lowermost Badenian”; sensu Hohenegger et al. 2014).
Significant changes in CP palaeogeography took place
between 17 and 15 Ma, controlled predominantly by geody-
namic development of the Alpine–Carpathian–Dinaride oro-
genic systems (Kováč et al. 2017b), and therefore it would
be appropriate to define a new regional stage between the
Ottnangian and (re)defined Badenian on the basis of geochro-
nological dating and constrained by the Central Paratethys
3
rd
order sequence stratigraphy.
The Late Badenian and Sarmatian s.s. sub-stages repre-
sented a period when the connection to the Mediterranean was
gradually closed (or at least its existence has not been suffi-
ciently proved). The connection to the Eastern Paratethys most
likely became opened (e.g., Popov et al. 2004; Bartol et al.
2014; Palcu et al. 2015; Silye & Filipescu 2016; Kováč et al.
2017a; Harzhauser et al. 2018). The view that the marine
connection from the east was opened is also induced by a sea-
level rise in the Eastern Paratethys during this time (Popov et
al. 2010). The presence of the NN6 Zone is common in all CP
basins due to marine connection with the Mediterranean in
the west until the base of the Sarmatian (sensu Bartol et al.
2014) and with the Eastern Paratethys during the Late
Badenian and Sarmatian (e.g., Popov et al. 2004). The NN7
Zone was identified only in several basins (Palcu et al. 2015;
Kováč et al. 2017a).
Sedimentary sequences of the western part of the CP can
therefore be roughly correlated with the early and late
Serravallian (Fig. 3C), with the base at ~13.82 Ma (e.g.,
Hilgen et al. 2009; Iaccarino et al. 2011; de Leeuw et al. 2013)
and the top at ~11.6 Ma (Hilgen et al. 2005; Vasiliev et al.
2010; de Leeuw et al. 2013). The connection with the Eastern
Paratethys in front of the Carpathians persisted even longer
294
KOVÁČ, HALÁSOVÁ, HUDÁČKOVÁ, HOLCOVÁ, HYŽNÝ, JAMRICH and RUMAN
GEOLOGICA CARPATHICA
, 2018, 69, 3, 283–300
Fig. 3. Palinspastic – topographic maps of the Central Paratethys; modified after Hámor & Halmai 1988; Rögl 1998; Kováč 2000; Popov et al.
2004; Kováč et al. 2017a,b; Sant et al. 2017b): A — the Burdigalian CP with gateway in front of the Alps; B — the late Burdigalian–Langhian
CP with gateway between the Alps and Dinarides; C — the Serravallian CP with gateway towards the Eastern Paratethys. Data were handled
using the PostgreSQL Server v. 9.4; topology and spatial geometry using the GRASS-GIS v. 7.2.1; GRASS-GIS software was used to compute
the location of each cell of the DTM using regularized spline with tension for approximation from vector data (module v.surf.rst; GRASS
Development Team 2017).
295
REGIONAL TIME SCALE CORRELATION OF THE CENTRAL PARATETHYS IN THE MIOCENE
GEOLOGICA CARPATHICA
, 2018, 69, 3, 283–300
and it is thus possible to correlate this time interval with
the upper part of the Eastern Paratethyan Sarmatian s.l.
(e.g., Popov et al. 2004; Gozhyk et al. 2015).
Conclusions
The CP time scale defined by biostratigraphic data remains
poorly constrained by geochronological and spatially-explicit
biostratigraphic methods, making the correlation with the stan-
dard GTS problematic (Fig. 1). The increase in spatial and tem-
poral coverage of point-based geochronological data is there fore
an essential task. In broad-scale palaeogeographic analyses
requiring correlation of the CP with the Medi terra nean,
the stan dard geological time scale should be used as a reference
to avoid problems with the definition of regional stages.
The rise or fall of the sea-level, as well as climate changes
in the semi-enclosed CP realm often have a local character and
were influenced by global sea-level changes only to some
degree. The differences between the global, Mediterranean or
the Eastern Paratethys sea-level curves indicate that the 3
rd
order
sea-level cycles in the CP need to be further validated and
the climate evolution should be better resolved (Fig. 2).
The complex geodynamic evolution of the Alpine–Carpathian–
Pannonian and Dinarides domains causes difficulties in corre-
lation with GTS, even between individual CP basins. It would
be beneficial to revise the regional time scale in respect with
the geodynamics of the orogenic system, as well as the ope-
ning of gateways between the CP, Mediterranean, and Eastern
Paratethys (Fig. 3A–C). The palaeogeographical reconstruc-
tions should reflect the original position and extent of basins
which fill was later deformed by folding and thrusting in front
of the orogenic system or by the movement of crustal frag-
ments along several hundred km long transform boundaries.
These observations were not taken into account for decades
leading to palaeogeographical misconceptions on a European
scale. To understand changes in the layout of sedi mentary
basins during distinct time spans an improved palinspastic
model based on an interdisciplinary approach is needed in
the future.
Acknowledgements: The research was supported by the Slovak
Research and Development Agency under the contracts
APVV-16-0121, APVV-15-0575, APVV-14-0118, APVV
SK-AT-2017-0010 and Progres Q45. We would like to thank
Rastislav Vojtko and Tomáš Klučiar for creating palinspastic
models. Our gratitude goes to Adam Tomašových and to two
anonymous reviewers for insightful comments.
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