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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, 2018, 69, 3, 283–300

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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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 

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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).

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, 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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