GEOLOGICA CARPATHICA, OCTOBER 2010, 61, 5, 393—418 doi: 10.2478/v10096-010-0024-1
The Late Eocene was a transitional period between the Middle
Eocene Climatic Optimum and the Oligocene icehouse
(Fig. 1). By the Oligocene, the climate system tended to an
“icehouse” world. Gradual cooling of the Earth’s climate re-
sulted in the expansion of Antarctic glaciation. Climatic dete-
rioration already began from the Middle/Late Eocene
boundary (Oberhänsli 1996), and was followed by the major
cooling event in the Early Oligocene (Biolzi 1985; Kennet &
Barker 1990; Diester-Haass 1991; Miller et al. 1991; Zachos
et al. 1993, 1996; Prothero 1994; Diester-Haass & Zahn 1996,
2001; Salamy & Zachos 1999; Wade & Pälike 2004; Tripati et
al. 2005, etc.). This climatic cooling led to significant pale-
oenvironmental changes in the Carpathian Paleogene basins
(Fig. 2). From the Early Oligocene, the Carpathian basins
provided the first records of the isolation from the open sea
(Early Paratethys – Báldi 1980, 1984, 1986; Protoparat-
ethys – Russu 1988), which enhanced gradually during the
Late Oligocene and Miocene (Eoparatethys—Mesoparatethys—
Neoparatethys, see Seneš & Marinescu 1974; Nagymarosy
1990; Popov et al. 1993; Rögl 1998, 1999; Kováč 2000;
Steininger & Wessely 2000, etc.). Therefore, the Carpathian
Paleogene basins mirrored the paleoenvironmental changes of
climatic cooling and Paratethyan isolation quite sensitively.
Their sedimentary records, like the Globigerina Marls, Meni-
lite Beds, Tylawa Marls, Dynow Marls etc., occurred near the
Eocene-Oligocene boundary in practically all the Carpathian
Paleoenvironmental changes across the Eocene-Oligocene
boundary: insights from the Central-Carpathian
Geological Institute, Slovak Academy of Sciences, Ďumbierska l, 974 11 Banská Bystrica, Slovak Republic; firstname.lastname@example.org
(Manuscript received February 1, 2010; accepted in revised form June 10, 2010)
Abstract: The sedimentary sequence of the Central-Carpathian Paleogene Basin provides proxy records of climatic changes
related to cooling events at the Eocene/Oligocene boundary (TEE). In this basin, climatic deterioration is inferred from the
demise of the carbonate platform and oligotrophic benthic biota in the SBZ19 and from the last species of warm-water
planktonic foraminifers in the E14 Zone. Upper Eocene formations already indicate warm-temperate to cool-temperate
productivity and nutrient-enriched conditions (Bryozoan Marls, Globigerina Marls). Rapid cooling during the earliest
Oligocene (Oi-1 event) led to a temperature drop ( ~ 11 °C), humidity, fresh water influx and continental runoff, water mass
stratification, bottom water anoxia, eutrofication, estuarine circulation and upwelling, carbonate depletion, sapropelitic
and biosiliceous deposition, H
S intoxication and mass faunal mortality, and also other characteristics of Black Sea-type
basins. Tectonoeustatic events with the interference of TA 4.4 sea-level fall and the Pyrenean phase caused basin isolation
at the beginning of the Paratethys. The Early Oligocene stage of Paratethyan isolation is indicated by a stagnant regime,
low tide influence, endemic fauna development, widespread anoxia and precipitation of manganese deposits. The episodic
rise in the sea-level, less humid conditions and renewed circulation is marked by calcareous productivity, nannoplankton
blooms and the appearance of planktic pteropods and re-oxygenation. Paleogeographic differentiation of the Carpatho-
Pannonian Paleogene basins resulted from plate-tectonic reorganization during the Alpine orogenesis.
Key words: Paratethys, Central Western Carpathians, Terminal Eocene Event, platform drowning, climatic cooling,
productivity changes, estuarine circulation, anoxia, eutrophication, semi-isolation.
Paleogene basins (Fig. 2). Previous studies provided an im-
portant insight into paleoenvironmental changes in the Outer
Carpathian basins (Van Couvering et al. 1981; Krhovský
1981a,b, 1995; Hanzlíková 1981; Roth & Hanzlíková 1982;
Fig. 1. Climatostratigraphic development of the Paleogene Period
from the Paleocene/Eocene Thermal Maximum (PETM), following
by the Middle Eocene Climatic Optimum (MECO), climatic deterio-
ration during the Late Eocene and culminating in the climatic cooling
of the Terminal Eocene Event (TEE) and during the Early Oligocene.
Krhovský et al. 1993; Leszczyński 1996,
1997; Oszczypko 1996; Oszczypko-
Clowes 1998; Gedl 1999; Gedl &
Leszczyński 2005; Puglisi et al. 2006;
Švábenická et al. 2007; Miclaus et al.
2009, etc.). The purpose of this study is to
review the Eocene/Oligocene events, pro-
viding their paleoenvironmental and cli-
matic proxies in the Central-Carpathian
Paleogene Basin (CCPB – Fig. 3).
Eocene/Oligocene transition in the
The CCPB accommodates a forearc
basinal system of the West Carpathian
Mountain chain (Soták et al. 2001;
Kázmer et al. 2003). The general litholo-
Fig. 2. Lithostratigraphic scheme of the Paleogene formations of the Central and Outer Western Carpathians. The Eocene/Oligocene
boundary is marked by the Globigerina Marls and Menilite Beds, which form a marker horizon dividing the lower deep-water turbidite
complex (Paleocene—Upper Eocene) from the Supra-Menilite formations (Upper Kiscellian—Lower Miocene). These marker horizons of
high-rate productivity and subsequent sapropelitic deposition occur in almost all the Carpathian basins.
Fig. 3. Paleogene basins of the Carpatho-
Pannonian area. The Central-Carpathian Pa-
leogene Basin is accomodated in the Central
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
gy of the CCPB around the Eocene/Oligocene transition is
compiled in type-section (Fig. 4). The sedimentary sequence
is developed from the Lutetian formations, continued across
the Eocene/Oligocene boundary and terminated to the Late
Oligocene/Miocene? formations. The initial phase of deposi-
tion is represented by conglomerates, boulder breccias and
poorly sorted sandstones, representing the sediments of rock-
fall avalanches, alluvial fans and delta-fed fans. These sedi-
ments are overlain by the Lutetian transgressive formations,
composed mostly of sandstones and nummulitic limestones.
Their stratigraphic age ranges within the SBZ16—19 (Köhler
1998; Bartholdy et al. 1999). Apart from the carbonate plat-
Fig. 4. Composite section of the transitional formations through the Eocene/Oligocene boundary
in the Central-Carpathian Paleogene Basin. Successive changes in lithology and life environments
reflect the changes in climatic conditions during the deposition. Abbreviations: NL – nummulit-
ic limestones, BM – bryozoan marls, GL – Globigerina limestones, GM – Globigerina
marls, FSH – fish shales, SOC – sapropelitic organic-rich claystones, NCH – nannochalks,
MCH – menilitic cherts.
form facies, Middle Eocene pelagic sediments are also pre-
served. These contain late Ypresian fauna with a dominance
of subbotinid species, early Lutetian fauna dominated by acar-
ininid species, mid-Lutetian fauna with a dominance of moro-
zovellid species and late Lutetian to Bartonian fauna with
morozovelloid and truncorotaloid species. Biostratigraphic
determination of the Middle Eocene formations is based on
the foraminiferal index species, such as Turborotalia frontosa
(late Ypresian – Biozone E7/P9), Acarinina cuneicamerata,
A. praetopilensis, Morozovella aragonensis, M. gorondatxensis
(early to middle Lutetian – Biozone E8—10/P10—P12), Acari-
nina (T.) topilensis, Morozovella spinulosa (late Lutetian –
Biozone E10—11/P12) and Moro-
zoveloides crassata, Truncorota-
loides rohri (Bartonian – Biozone
classification sensu Soták 2007).
The sedimentary sequence of the
CCPB graded up to the Priabonian—
Rupelian formations (Fig. 5), be-
longing to the E14—O3 Biozones
(Soták et al. 2007). The sequence
started with the Globigerina Marls,
the basal part of which belongs to
the E14/P15 Biozone, marked by its
nominate taxon Porticulasphaera
semminvoluta and associated spe-
cies, like Globigerinatheka index,
Globigerinatheka aff. subconglo-
bata, Subbotina linaperta and S.
corpulenta. The zonal boundary
P15/P16 is shown by the LAD of
followed by the significant increase
in the abundance of globigerin-
athekids. The mid-Priabonian micro-
fauna is dominated by large-sized
tests of Globigerinatheka index,
and their acme clearly denotes the
E15/P16 Biozone (after Gonzalvo
& Molina 1992). Associations in
this zone are very rich in large sub-
botinids, comprising species of S.
corpulenta, S. cryptomphala, S.
gortani, S. pera and S. praeturrili-
na. Stratigraphically important spe-
cies of Turborotalia cerroazulensis
lineage are represented by T.
pomeroli and T. cerroazulensis.
The distribution of these species
indicates the subdivision between
the upper part of the E15 Biozone
(= the lower part of the P16) and
the lower part of the E16 Biozone
(= the upper part of the P16). The
succesive formation contains more
planoconvex species of turboro-
taliids, resembling T. cocoaensis.
The presence of this species pro-
vides evidence of the uppermost Upper Eocene biozone
(E16/P17). These intervals are considerably richer in benthic
foraminifers similar to those recognized in the Buda Marls
(microfauna with Tritaxia szaboi or Cylindroclavulina rud-
ilosta – e.g. Hantken 1875; Sztrákos 1987). Close to the E/O
boundary, the Turborotalia cerroazulensis lineage disap-
peared and its last species T. cunialensis occurred together
with the first appearance of the new species T. ampliapertura
(FAD 33.8 Ma). Based on the dinocysts, the Late Eocene age
of the lower part of the marly sequence is indicated by the FAD
of Reticulatosphaera actinocoronata and by the lack of the Ear-
ly Oligocene taxa (e.g. Chropteridium spp.). Boundary intervals
are very rich in volcanic biotite, similar to the biotite-bearing
layers at the Massignano section (e.g. Coccioni et al. 2000),
Eocene/Oligocene sections in Hungary (Báldi 1984), etc.
The Eocene/Oligocene boundary is indicated in the upper-
most part of the Globigerina Marls. The Oligocene marls and
superimposed Menilite-type formations are strongly reduced
Fig. 5. Biostratigraphic determination of the Eocene/Oligocene boundary in the Central-Carpathian Paleogene Basin showing the vertical
distribution of the foraminiferal index species (data adapted from Soták et al. 2007).
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
in the quantity and diversity of foraminiferal microfauna.
They are dominated by small-size globigerinids of the G.
praebulloides—G. officinalis plexus. The small globigerinids
contain paragloborotaloids, tenuitellids and chiloguembellin-
ids, which are considered to be the index fossils of the early
Rupelian. Most of these species, like Dentoglobigerina tapu-
riensis, D. tripartita, Turborotalia ampliapertura, Para-
globorotalia opima nana, Tenuitella gemma, T. liverovskae,
Chiloguembelina cubensis, Ch. gracillima, etc., exhibit the
FAD or ACME within the O1/P18 Biozone. Some of the asso-
ciated species already indicated the O2/P19 Zone (e.g. Dento-
globigerina selli, Paragloborotalia opima pseudocontinuosa,
P. semivera, Parasubbotina carpathica, etc.). The mid-Rupe-
lian formation is barren in content of foraminiferal fauna, but
it contains the endemic nannofossil species of Reticulofenes-
tra ornata (NP23 Zone). In this part of the sedimentary se-
quence, the Menilitic shales are associated with tuffaceous
horizons, Tylawa-type limestones and biosilicite horizons.
The Upper Rupelian sequence became flysch-like in charac-
ter, since it contains poor microfauna. Planktonic foramini-
fers suggest their possible attribution to the O2—O3/P19—P20
Biozones and this is based on the presence of younger spe-
cies such as Tenuitellinata angustiumbilicata, Tenuitellinata
postcretacea, Tenuitella brevispira, Tenuitella evoluta, etc.,
and abundance of chiloguembelinids. Superimposed units of
the Sub-Tatras Group are stratigraphically shifted to the late
Rupelian (LAD of chiloguembelinids in the basal part of the
Huty Formation related to the NP23/NP24 boundary – Van
Simaeys et al. 2004), early Chattian (FAD of “Globigerina” ci-
peroensis angulisuturalis in the Zuberec Formation – O4/P21a
Biozone) and late Chattian to early Aquitanian (FAD of Dis-
coaster drugii and Helicosphaera scissura in the topmost
part of the Biely Potok Formation = Ostrysz Beds). The bio-
stratigraphy of the CCPB was updated by Molnár et al.
(1992), Vass et al. (1993), Olszewska & Wieczorek (1998),
Soták (1998a,b), Starek et al. (2000), Gedl (2000a,b), Soták
et al. (2001, 2007), Garecka (2005), etc.
Proxies and methods
Paleoenvironmental changes in the CCPB are indicated by
various proxies. Their identification is based mostly on the
foraminiferal microfauna, using climatic index taxa, species
abundance and diversity, life-mode strategies, habitat-groups,
morphotypes, trophic preferences, dissolved-oxygen index,
coiling directions and paleobathymetry (e.g. Premoli Silva &
Boersma 1988; Spezzaferri & Premoli Silva 1991; Keller et
al. 1992; Spezzaferri 1995; Van Eijden 1995; Spezzaferri et al.
2002; Bicchi et al. 2003; Molina et al. 2006; Wade et al.
2007; Wade & Pearson 2008, etc.). Besides foraminifers, the
paleoenvironmental conditions have also been inferred from
calcareous nannofossils, organic-walled dinoflagellates, dia-
toms, moluscs, pteropods and fish fauna (e.g. Rusu 1995;
Diester-Haas & Zahn 1996; Monechi et. al. 2000; Pross &
Schmiedl 2002; Van Simayes et al. 2004). Geochemical prox-
ies have been used to approximate the isotopic paleotempera-
tures (after Craig 1965 – T = 16.9—4.2 [
] + 0.13
), atmospheric CO
productivity (trace metals), organic-rich deposition, trophic
conditions and low-oxygen environments (e.g. Zachos et al.
1996; Murphy et al. 2000; Van Breugel 2006; Vetö et al.
Eocene/Oligocene events in the CCBP
The multiproxy study allows us to interpret the paleoenvi-
ronmental changes of platform drowning, climate cooling,
biomass productivity, eustatic events, CCD deepening, conti-
nental run-off and riverine input, eutrophication, widespread
anoxia, water-mass stratification, sapropelitic deposition, vol-
canogenic activity and semi-isolation in the CCPB.
Demise of the carbonate platform
The carbonate platform – or ramp-type basin – (cf.
Wright & Burchette 1998) was established by the sea-level
rise during the Middle Eocene Climatic Optimum. Climatic
control of nummulid-rich production implies the tropical-type
carbonate ramps in the CCPB (cf. Schlager 2003). Their de-
velopment in the CCPB culminated during the SBZ17 Zone.
Sedimentary units of the carbonate ramp are formed by keep-
up cycles of nummulitic bank, back-bank, lagunar, rubble-flat
and fore-bank facies.
Nummulitic ecosystems were adapted to clear-water and
oligotrophic conditions, requiring temperatures above 20 °C
(Sarangi et al. 2001). Such temperatures dominated during the
late Lutetian—early Priabonian time, which is regarded as
a thermochron in global climate (Rusu 1995). The oxygen iso-
topic composition of the nummulitic limestones in the CCPB
O = —2 ‰) indicates a seawater temperature of around
22 °C. This isotopic temperature corresponds to the tempera-
ture derived from nummulitic tests in Hungary using the Ca/Mg
method (21.4—25 °C, see Berlin et al. 1976). As well as tem-
perature, the nummulitids were highly sensitive to even a
small increase in nutrient availability and primary productivity
(Hottinger 1983). Their life strategy was strongly dependant
on a trophic regime, which had to be oligotrophic, due to their
algal-bearing symbionts. Therefore, the nummulitids were not
competitive in more nutrient-rich waters (Hallock 1987;
Hallock et al. 1991). Open-marine microfossils from the num-
mulitic formations in the CCPB, such as porticulosphaerids
and discoasterids, belong to the warm-water and low-nutrient
taxa (Figs. 6, 7). Nummulitic facies are barren in dinoflagel-
lates, which were not favoured by oligotrophic waters (Gedl
2000a). The mollusc fauna of the Lutetian formations in the
CCPB consists of thermophile forms, such as spondylids, pec-
tinids, ostreids, etc. (Papšová in Gross & Köhler et al. 1980).
Nummulitic fauna of the CCPB began to disappear in
stressful conditions, inferred in climatic cooling and nutrient
excess. Cooler and mesotrophic conditions are indicated from
the Bartonian (EBi 1 shift to lower temperatures sensu Abreu
& Anderson 1998), when the sequence increases in the abun-
dance of orthophragminids and heterosteginids (Fig. 6, after
Bartholdy et al. 1999). The early Priabonian transgression re-
newed growth production on carbonate ramps. Their re-estab-
lishment occurred during the last warm-temperate conditions
of the SBZ19 Zone. The eustatic rise during the early Priabon-
ian is well documented by aggradation of carbonate ramp over
structural highs and this is recognized by the incipient flood-
ing surfaces on the Mesozoic strata (e.g. in Western High
Tatra sections). This transgressive phase reached maximum
flooding in formations with Nummulites millecaput (Köhler
1998), which as giant nummulitids benefitted from the low
energy environment with lower light intensity and fewer nutri-
ents (Hallock & Glenn 1986). A major turnover to climatic
cooling occurred around the demise of the nummulitids (cf.
Geel 2000). The nummulites-bearing formation of the CCPB
reveals the deepening-upward tendency in basin paleobathym-
etry (e.g. in the Jobová Rázstoka section, Važec-Priepady sec-
tion). Platform drowning is demonstrated by the distribution
of benthic foraminifers in the Borové Formation, showing the
gradual decrease in shallow-water taxa (Elphidium, Parrellina,
Fig. 6. Productivity changes in the Central-Carpathian Paleogene Basin. The Late Eocene—Oligocene sequence exhibits the upward de-
crease in carbonate productivity, increase in organic-rich productivity with episodical blooms of calcareous plankton and after it the pre-
dominance of biosiliceous productivity (results derived from foraminiferal, coccolithophorid, dinoflagellate and diatom-based data).
Abbreviations see Fig. 4.
Calcarina, etc.) and the increase in deep neritic to shallow
bathyal taxa (e.g. Uvigerina, Bolivina, etc.). As a consequence
of cooling, nutrification and drowning, the warm-water car-
bonate deposition in the CCPB ceased with the latest nummu-
lites in the P16 Zone = SBZ20 Zone (Köhler 1998), or in the
SBZ21 Zone (Buček & Filo 2004).
Paleoenvironmental changes in the CCPB are apparent
from the development of the temperate-type carbonate ramp
with foramol or foralgal facies in the Priabonian (sensu Car-
annante et al. 1988). Unlike the tropical-type carbonates, the
foramol facies of the CCPB (e.g. in Hybica section) are poor
in nummulitids and rich in skeletal grains, benthic foramini-
fers, crustose algae and bryozoans. Widespread foramol-type
carbonates in the Priabonian also resulted from progressive
eutrophication of the CCPB, since cooler waters are usually
richer in nutrients. The large foraminifers are represented
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
mainly by orthophragminids (e.g. Discocyclina), which toler-
ated lower temperature and higher dispersion, and preferred a
lower energy environment, greater depth and a soft substrate
(Rasser et al. 1999). In addition to orthophragmids, the Pria-
bonian carbonates are also rich in bryozoans, which are recog-
nized as the most abundant organisms on the cool-water
shelves (Surlyk 1997; Light & Wilson 1998). The CCPB
shelves were bypassed by the Bryozoan Marls, consisting of
numerous species and various growth forms (Zágoršek 1992,
2000). The mass abundance of bryozoans indicates temper-
ate cool-water and mesotrophic growth conditions (Fig. 6),
and these also facilitated the bryozoan reef expansion in gla-
cial periods during the Quaternary (cf. James et al. 2000;
James & Bone 2000). Prosperity of bryozoan fauna usually
indicates elevated resources of land-derived nutrients, or the
availability of nutrients via upwelling (e.g. Light & Wilson
Fig. 7. Climatic index taxa of the planktonic foraminiferal microfauna from the Central-Carpathian Paleogene Basin. Temperature indices
of planktonic foraminifers provide evidence of three-fold cooling, marked by increased colour intensity, which culminated in cold and high-
latitude climatic conditions in the Middle Oligocene.
1998). The Late Eocene was a time of widespread appear-
ance of bryozoan-rich sediments, as in the Buda Basin
(Kázmer et al. 1993), North Apennine Basin (Braga &
Barbin 1988), Adriatic Basin (Marjanac et al. 1988), Alpine
Foreland Basin (Rasser et al. 1999), Australian basins
(James & Bone 2000), etc. Therefore, the Bryozoan Marls
imply similar paleoenvironmental conditions of deposition
(climatic cooling, nutrification, upwelling, etc.).
High productivity rate of planktonic foraminifers
Progressive cooling in the CCPB induced changes in the
productivity rate and temperature preferences of planktonic
foraminifers. During the Middle to Late Eocene, the plankton-
ic foraminiferal productivity decreased in warm-water species
and increased in bloom-forming globigerinids (Fig. 6). Lu-
tetian pelagic sedimentation resulted mostly from the warm-
water productivity, which is indicated by muricate species,
that is foraminifers heavily calcified due to a high content of
dissolved carbonates. Muricate morphotypes, such as moro-
zovellids, acarininids and truncorotaloids, are the most com-
mon foraminifers in the Lutetian microfauna of the CCPB.
Warm-water productivity of calcareous nannoplankton is
documented by Discoaster-dominated associations. Muri-
cate foraminifers and discoasterids are oligotrophic organ-
isms (Monechi et al. 2000; Poletti et al. 2004), which
benefited under the nutrient-poor conditions during the Lu-
Foramimiferal microfauna indicate a gradual impoverish-
ment of paleoenvironmental conditions following the Barto-
nian/Priabonian boundary (Fig. 7). In the CCPB, the last
muricates disappeared between the P14/P15 Zones. Priabon-
ian associations of foraminifers differ markedly, with an in-
creasing content of porticulasphaerids and later prevalence of
globigerinathekids. The early Priabonian species of Porticu-
lasphaera semiinvoluta is still considered to represent warm-
temperate habitats, such as the species of the “mexicana”
group (Bolli 1972). On the contrary, the globigerinathekids
were interpreted as subtropical and cooler water forms, which
inhabited middle-latitude sites (Blow 1969; Premoli Silva &
Boersma 1988; Boersma & Premoli Silva 1991). The most re-
cent studies of globigerinathekids (Spezzaferri et al. 2002; Ga-
leotti et al. 2002; Payros et al. 2006) point to their temperate to
cold-temperate preferences and mesotrophic character
(Fig. 7). The abundant increase in globigerinathekids in the
middle Priabonian marls provided the first indication of pro-
Fig. 8. Impoverishment and test-size reduction of foraminiferal microfauna in the
Central-Carpathian Paleogene Basin (Pucov section, Orava region – Soták et al.
2007). Such abrupt change from the large-sized to small-sized globigerinids is
known as the Lilliput event at the Eocene/Oligocene boundary (MacLeod 2006).
gressive climatic cooling in the
CCPB. Consequently, warm-
water species, such as those of
the Turborotalia cerroazulensis
group, became less common, and
truly tropical forms are missing
here (e.g. tubulospinose hantken-
inids). Globigerinathekids disap-
peared abruptly in the middle of
the P16 Zone, and they were re-
placed by cooler late Priabonian
foraminiferal microfauna in the
During the Late Eocene, the
productivity of the CCPB in-
creased considerably in re-
globigerinids. So-called “Globi-
gerina Marls” are considered as
prima facie (Van Couvering et
al. 1981), which denotes a
prominent increase in biopro-
ductivity around the Eocene/
Oligocene boundary. Their mi-
crofauna consists of “globigerin-
ids” with cool-water preferences
(cf. Olszewska 1983). In the
CCPB, the Globigerina Marls
are dominated by subbotinids,
which attained large growth
size, high population density
and lower diversity (Fig. 8).
Subbotinids represented the
cold-temperate habitats of the
Eocene microfauna (Boersma
& Premoli Silva 1991; Pardo et
al. 1999; Spezzaferri et al.
2002, etc.). Therefore, their
was related to climatic cooling,
since the cold waters were su-
persaturated with nutrients.
This increase in trophic re-
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
sources led to a proliferation of spinose taxa, such as subbot-
inids (Boersma et al. 1995).
Globigerina Marls lack an original isotopic record of sea-
water temperature due to their common recrystallization and
reequilibrium during diagenesis. Therefore, their bulk-sedi-
ment isotopic composition attains the more negative
ues (—5 ‰). The carbon isotopic composition of the
Globigerina Marls ranges within normal
C values for ma-
rine carbonates (0 to 0. 8 ‰), thus exhibiting no diagenetic
overprint or vital effect.
Paleoenvironmental conditions of the Globigerina Marls are
also reflected by changes in bottom-water productivity. Benthic
foraminiferal frequency varied significantly (P/B = 4—35 %)
and tended towards peak abundance and diversity in the Tri-
taxia szaboi horizons. These benthic-rich horizons of the Glo-
bigerina Marls imply an enhanced organic matter flux, and
bottom-water ventilation with a preference for epifaunal and
shallow-infaunal species (e.g. Lenticulina cultratus, Marginuli-
nopsis fragaria, Vulvulina haeringensis, Gemellides eocae-
nus, Eponides umbonatus, Cibicides ungerianus, Hanzawia
ammophila, Gyroidinoides soldani, Lagena gracilicosta and
Uvigerina rippensis). Seafloor oxygenation by cold water is
also expressed by occasional red colouring of the Globigerina
Marls, and by their common bioturbation (mainly Chondrites-
type burrows). In some horizons of the Globigerina Marls,
there are no planktonics, and the benthic foraminifers are
dominated by agglutinated taxa (e.g. Ammodiscus polygyrus,
Rhabdammina discreta and Glomospira charoides). The
changing ratios between calcareous and agglutinated foramin-
ifers within the Globigerina Marls indicate the vertical fluctua-
tions of the CCD in the dependance on the productivity,
temperature and CO
concentration (Fig. 9).
The Calcite Compensation Depth (CCD) declines in the
conditions of enhanced productivity like those in the Globige-
rina Marls, which correspond to a peak of productivity near
33.5 Ma (Diester-Haass & Zahn 1996), and this caused the
CCD drop near the Eocene/Oligocene boundary (Thunell &
Corliss 1986; Tripati et al. 2005). Following the downslope
excursion of the CCD, the Globigerina Marls occur suddenly
in carbonate-free deep-water sediments of the Carpathian ba-
sins (e.g. the Terchová Formation in the CCPB). Carbonate
depletion of the Oligocene sediments in the Carpathian basins
could have resulted from shoaling of the CCD due to a higher
saturation of cold bottom-water by CO
. As a consequence,
the seawater became more acid and corrosive to the calcareous
components. The acidification led to a high solubility of the
calcitic microfossils, which disappeared in some horizons of
the Globigerina Marls. Small-scale intercalations of non-cal-
Fig. 9. Productivity driven fluctuation of the CCD during the deposition of the Globigerina Marls (9a – loc.
Údol, scale bar = 1 m). High-rate productivity led to the CCD drop (9c), which corresponds to CCD
deepening across the Eocene/Oligocene transition (9b – Tripati et al. 2005). Productivity (P) decrease
caused a higher rate of dissolution (D), which tend to the CCD rise (9d).
careous black shales with the
Globigerina Marls provide
evidence of the vertical
fluctuations of the CCD,
concentration and acid-
ification, resulted from the
productivity changes, cli-
matic deteriorations and
precessional cyclicity (cf.
Krhovský 1995; Leszczyńs-
ki 1997). Above the Globi-
gerina Marls, the carbonate
dissolution increased con-
siderably, leading to a non-
calcareous deposition of the
Biotic crises and rapid
Climatic changes culmi-
nated in the “Terminal
Eocene Event”, which corre-
sponds to the global cooling
and glacio-eustatic regres-
sion related to the Antarctic
cryosphere expansion. Since
the Late Eocene, a cold cli-
matic phase followed (cryo-
chron sensu Russu 1995)
and the ocean temperature
fell about 2—5 °C world-
wide (Shackleton & Kennett
1975; Pomerol & Premoli
Silva 1986). In the CCPB,
the paleotemperature drop is indicated by the appearance of
cool-water molluscs, such as Nucula and Glycymeris (e.g. in the
Odorín Limestones, Ví az locality – Marschalko & Volfová
1960; Volfová 1964a). The Late Eocene—Early Oligocene
cooling is also manifested in land flora and palynoassocia-
tions, recording the onset of Arctotertiary taxa (Snopková
1980; Konzalová et al. 1993). The cool-water influence in the
CCPB is also documented by pteropods Spiratella (e.g. ptero-
pod shells from the Blatná dolina locality – Orava, Szaflary
Beds – Olszewska 1998), which as the polar epiplanktonic
fauna expanded to the Paratethyan basins from the Boreal
area, the so-called Spiratella Sea (cf. Báldi 1984).
Foraminiferal microfauna of the CCPB indicates a step-
wise cooling from the Eocene/Oligocene boundary. Never-
theless, the earliest Rupelian microfauna still contains some
temperate-water foraminifers, such as dentoglobigerinids,
globoquadrinids and also the last turborotaliids (Fig. 7).
These are already associated with cold-temperate and cold-
preferring taxa (e.g. paragloborotallids, pseudohastigerinids
and chiloguembelinids). Acceleration of cooling in the middle
Rupelian is marked by cold-water species (e.g. tenuitellids and
globigerinids). The cooling response to the planktonic fora-
minifers is most evident from their size reduction, reaching al-
most 80 % in the latest Eocene—Early Oligocene (Fig. 8). This
size reduction is know as a dwarfing or Lilliput effect (MacLe-
od 2006; Wade et al. 2007). Moreover, the Early Oligocene
associations of planktonic foraminifers were dominated by
sinistral forms (Salaj 1998), and this is a coiling mode of fora-
minifers in the cold-water environment of high-latitude oceans
(Norris & Nishi 2001). Considering this, the Early Oligocene
foraminiferal microfauna of the CCPB reveals a dramatic size
reduction, oligotaxic character, S coiling mode and an oppor-
tunistic life-mode strategy (bloom of chiloguembelinids),
which is most likely related to cold-water environments.
Early Rupelian cooling in the CCPB reduced the carbonate
productivity by elimination of calcareous nannofossils, which
occur only as cool-water taxa (e.g. Isthmolithus recurvus,
Zigrhablithus bijugatus and Reticulofenestra lockeri). Later,
the carbonate productivity changed to phytoplankton and bio-
silicite productivity, and this represented the productivity
maximum in the earliest Oligocene (Diester-Haass & Zahn
2001). Phytoplankton of the Lower Oligocene formations of
the CCPB dominated by Wetzelielloideae, whose large-sized
dinocysts imply cool-water conditions (Gedl 2000a).
Late Rupelian microfauna became almost monospecific, and
limited to globigerinid species of G. praebulloides—G. officina-
lis plexus. The small species of Globigerina praebulloides is
considered to represent the subpolar planktonic foraminifers
(Kennett 1982; Pak & Kenett 2002; Bicchi et al. 2003, etc.), or
the upwelling-type foraminifers proliferated from the tempera-
ture drop and nutrient availability (Boersma et al. 1995). In
times of flysch-type deposition in the CCPB (late Rupelian—
Chattian), the planktonic foraminiferal productivity almost
ceased due to water-mass turbidity, low oxygen concentration
and a high sedimentation rate.
The paleotemperature decline in the Early Oligocene is also
recorded in the oxygen isotopic composition of the lime-
stones, interposed between the Globigerina Marls and Meni-
lite Shale Formations. Oxygen isotopes exhibit a positive shift
O values from —2 ‰ in the Middle Eocene lime-
stones to 0.4 ‰ in the laminated limestones, implying a sharp
temperature drop of seawater from 22 °C to about 11 °C
Fig. 10. Paleotemperatures calculated from isotopic composition of the
Nummulitic Limestones and the Globigerina Marlstones by use of Craig’s
equation (Craig 1965). Isotope values
O indicate a sea-surface tempera-
ture drop between the Late Middle Eocene (Nummulitic Lmst) and upper-
most Eocene (Globigerina Mrls) from about 22 °C to 11 °C.
(Fig. 10). An identical
O signal was recorded by
high-resolution isotope stratigraphy at the Eocene/
Oligocene boundary in the ODP Sites (Diester-Haass
& Zahn 1996) and type localities (Oberhänsli et al.
1984). The positive
O excursion links with the Oi-1
shift to cooler temperatures (see Abreu & Anderson
1998), which corresponds to maximum productivity
increase of surface waters (Salamy & Zachos 1999;
Diester-Haas & Zahn 2001), eustatic sea-level fall
and tefrostratigraphic events at the Eocene/Oligocene
boundary (Fig. 11). The carbon isotopic values of
laminated limestones (—4 ‰
due to respiration and planktonic foramin-
iferal metabolism (mainly Globigerina praebulloides).
The higher respiration rate of these foraminifers result-
ed in the incorporation of more
during calcification (Bemis et al. 2000).
From the Eocene/Oligocene transition, the rapid
cooling and high nutrient supply inhibited calcium car-
bonate production in the Carpathian Paleogene basins.
Low activity of CO
resulted in lower seawater pH
and a high organic carbon concentration (Lyle et al.
2005). Therefore, the organic-rich sediments of the
overlying beds indicate the conditions of carbonate un-
dersaturation, high solubility of the silicid acid (e.g.
bloom of silicoflagellates, presence of silicified woods,
etc.), sulphate reduction of organic matter, mobility of
the redox-sensitive elements (e.g. Mn, P, Ba) and
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
Fig. 11. Correlation of isotopic signals from the CCPB with global isotopic curve, eustatic curve and tafrostratigraphic events. This inte-
grated approach allows us to approximate the Eocene/Oligocene boundary. Oxygen isotope curves after Zachos et al. (1993), Diester-Haass
& Zahn (1996), etc.; eustatic curves after Vail et al. (1977) and Haq et al. (1988). Abbreviations see Fig. 4.
methanogenesis. These processes could produce HCO
creasing the seawater alkalinity and resulting in the forma-
tion of the methan- and sulphate-derived carbonates (so-called
Anoxia and eutrophication
The Early Oligocene period was a time of widespread an-
oxia and eutrophication in the Carpathian Paleogene basins,
and this led to sapropelitic and biosiliceous deposition of the
Menilite facies (cf. Roth & Hanzlíková 1982). Similarly, wa-
ter-column anoxia, eutrophication, dinoflagellate- and diatom-
based productivity, high export flux, elevated chemocline,
monsoonal precipitation, interface cycles of coccolithophorid-
rich productivity, reduced precipitation and continental influx
(Fig. 13) also occurred in the CCPB.
The basin-wide anoxia in the CCPB is indicated both in bot-
tom-water and surface-water environments. Poorly oxygenat-
ed bottom sediments were inhabited only by deeper infauna of
benthic foraminifers, such as Virgulinella (Loxostomum)
chalkophilla, Chilostomella tenuis, Ch. cylindroides, Fursen-
koina acuta and Bulimina pyrula, etc. The benthic productivi-
ty of flysch sediments increased temporarily during the
recolonization of the sea-bottom by tubular morphogroups of
agglutinated foraminifers, such as Dendrophrya-Rhabdammina
type associations in the sedimentary formations of the CCPB
(Benešova 1962; Blaicher 1973; Samuel in Gross et al. 1999,
etc.). Oxygenation of the CCPB improved up to the Huty For-
mation (NP24), which represents a highstand phase of mud-
rich deposition with higher calcareous precipitation and
appearance of trace fossils (e.g. Zoophycos ichnoceonoses).
The shelfal sediments above the oxygen minimum zone OMZ
(e.g. in Horehronská kotlina Depression) are much richer in
benthic meiofauna (Samuel 1975; Zlinská et al. 2001).
The expansion of the OMZ to surface-water is inferred from
the impoverishment of planktonic foraminifers, and in prolif-
eration of euryoxibiont forms. Lower oxygen conditions are
expressed by blooms of chiloguembelinids, which occurred in
the CCPB during the early Rupelian. Chiloguembelina-domi-
nated assemblages are indicative of hypoxic to anoxic condi-
tions (Pardo et al. 1999), and this reveals life stress in the
productive zone (similar to the Guembelitria blooms near the
K/T boundary – Keller & Pardo 2004, etc.). Evidence of sim-
ilar conditions is provided by wetzeliellacean dinoflagellates,
diatoms and bacterioplankton.
Eutrophication of the Carpathian Paleogene basins led to
replacement of calcareous productivity by organic-rich and
biosiliceous productivity (Fig. 6), because the dinoflagel-
Fig. 12. Chemostratigraphic data from the Central-Carpathian Paleogene sequence ploted in the TOC, TIC and isotopic curves. Abbrevia-
tions see Fig. 4.
Fig. 13. A model of sapropelitic deposition in the Carpathian Paleogene basins. Menilitic sapropelites were deposited under a high conti-
nental runoff, increased nutrients (eutrophication and elevated nutricline), dinoflagellate and diatom-based productivity, oxygen depletion
(anoxia and elevated chemocline), etc. Calcareous layers is considered to represent the nannofossil limestones, deposited under a high rate
of coccolithophorid productivity, mesotrophic conditions (deeper nutricline), weakening anoxia (deepened chemocline), etc.
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
lates and diatoms utilized nutrients more efficiently than for-
aminiferids and coccolithophorids. Therefore, the eutrophi-
cation in the CCPB was responsible for dinoflagellate
blooms which were dominated by peridionids (Gedl
2000a,b). This trend culminated in catastrophic eutrophica-
tion and the mass expansion of diatoms (Menilite cherts in
upper NP23 – Nagymarosy 2000). The diatomite origin of
the Menilite cherts was documented by discoveries of frus-
tules several times (e.g. Kaczmarska & Kilarski 1979;
Krhovský 1981b, etc.), and by molecular record of the C
highly branched isoprenoids (Rospondek et al. 1997). Pyri-
tized diatoms are also frequently detected in the supra-meni-
lite formations of the CCPB (e.g. Šambron Beds, Huty
Formation, Brzegi Beds). As well as the evidence provided
by diatoms, the increase in seawater Si(OH)
is marked by
the appearance of spicules, benthic sponges and silicified
woods in sediments of the Menilite facies (Starek 2001;
Soták et al. 2007). Silica availability in the CCPB resulted
from the riverine input of nutrient-rich water from adjacent
continents. Consequently, the diatomites from the Menilite
Formation consist predominantly of fresh-water species,
such as Melosira distant, M. islandica, etc. (Řeháková in
Krhovský 1981a). The marine-continental interference in the
CCPB is documented by tidal-marsh foraminifers (mainly
Trochammina species) in the coastal plain sediments of the
Spiš area, abundant presence of land flora and riparian-type
vegetation (leaves, woody fragments) and even by rare avi-
fauna (finding of a bird leg at the Bystré locality, Eastern
Slovakia – Fig. 14). Salinity decrease in the surface-water
layer is also indicated by Deflandrea-dominated associations
of dinocysts in the Early Rupelian sediments of the CCPB
(Gedl 2000a). Nevertheless, the silica production from a vol-
canogenic source could also not be excluded, because the
Menilite Shales in the CCPB are occasionally intercalated by
tefra horizons (Gross 1981) or biotite-rich layers (Soták et al.
2007). The increase in primary productivity from volcanic-de-
rived nutrients is considered to be important at the Eocene/
Oligocene boundary (e.g. Massignano section – Monechi et
al. 2000), Early/Middle Miocene boundary (e.g. Styrian Basin
– Spezzaferri et al. 2002, etc.). Therefore, the excess of nutri-
ents may have also been responsible for the impoverishment
of the Oligocene foraminiferal microfauna in the CCPB.
Primary productivity in the CCPB culminated in hyper-
trophication and oxygen depletion. The euxinia of the photic
zone is considered to be an important factor in the deposition
of organic-rich sediments, such as sapropels (Koopmans et
al. 1996; Bosch et. al. 1998). Sapropelitic character of the
Menilite-type facies in the CCPB is indicated by lowered
Fig. 14. Rich fossil fauna of fishes (Pl. 12a,c), bird’s limb (Pl. 12b – first discovery) and leaves (Pl. 12d) from the Menilite shales of the
Central-Carpathian Paleogene Basin (Bystré locality). Excellent preservation of fossils is caused by bottom-water anoxia.
TIC and elevated TOC contents (Fig. 12). They approach up
to 8 wt. % TOC, and the organic matter yields a high content
of bituminite and alginites (Kotulová 2004). Anaerobic con-
ditions in the photic zone prevented the expansion of bacteri-
oplankton, which required light and hydrogen sulphide. The
photosynthetic activity of green sulphur bacteria in the anoxic
sediments is revealed by isorenierante derivates (see Koopmans
et al. 1996; van Breugel et al. 2005), and their presence was
identified in the sediments of the Menilite facies (e.g. Schulz
et al. 2003). In the CCPB, the anoxygenic water column and
microbial photosynthesis is indicated by carbon isotope com-
position of the Menilite Shales, which
C values are nega-
tive (—0.8 to —7.9 ‰) or extremely low close to isorenierante
(up to —21 ‰). Moreover, the abundance of pyrite framboids
in the Menilite Shales implies a sulphidic water column (see
Wilkin et al. 1997). Sulphate-reducing bacteria react with fer-
rous iron to form monosulphides and pyrite. Anoxic condi-
tions promoted bacterial decomposition of organic matter and
provided a source of N. Nitrification in the CCPB elevated the
ammonium concentration which supported phytoplankton
biomass and eutrophication.
The nannofossil crisis in the Menilite facies was interrupt-
ed by episodes of coccolithophorid-rich productivity, which
is indicated by the presence of laminated pelagic limestones
(Tylawa Limestone) and related carbonates with an almost
complete dissolution and overgrowth of coccoliths due to re-
crystallization and dolomitization. The increase in coccoli-
thophorid productivity was connected with a decrease in
diversity (Bubík 1978). More abundant nannoplankton spe-
cies in these limestones represent high-nutrient taxa (e.g.
Cyclicargolithus floridanus) and Paratethys-endemic or even
salinity-reduced taxa (e.g. Reticulofenestra ornata, Trans-
versopolis fibula and T. lata). Their monospecific associa-
tions are recognized as Solenovian (Polbinian) nannoflora of
the Central and Eastern Paratethys (Nagymarosy & Voronina
1992; Nagymarosy 2000; Melinte 2005). Nannoplankton
bloom in the NP23 Zone was probably prevented by short-
term increase in eustasy, salinity, trophic resources and re-
newed circulation (cf. Krhovský & Djurasinovič 1992;
Nagymarosy & Voronina 1992; Schulz 2003; Schulz et al.
2004). Calcareous vs. organic-walled plankton productivity
depends most likely on the salinity and precipitation. During
wet climates, the increased water runoff led to silicate con-
centration, reduction of salinity and eutrophication of the
surface water in the CCPB. These conditions proliferated a
high productivity of diatoms and dinoflagellates in the Meni-
lite Shales. However, drier phases with reduced runoff and
enhanced evaporation were responsible for nutrient deple-
tion, improvement in water clarity, an increase in salinity
and the dense blooms of calcareous nannofossils (Fig. 13).
Sapropelitic sediments in the CCPB (Menilite Shales) ex-
hibit the carbon and oxygen isotope lightening (
C —7.9 ‰,
O —7.5 ‰) probably due to bacterial oxidation of organic
matter. Such light carbon-isotope values in sediments of the
Menilite Formation can also suggest a strong influence of
fresh water (e.g. Schmiedl et al. 2002). Contrary to this, the
coccolithic limestones exceeded the background values of
the Menilitic formation, attaining a more positive isotopic
values. These limestones correspond to short-lived positive
C excursions (—0.5 to +3.4 ‰), coupled with less pro-
nounced positive shifts of
O (—3.4 ‰). Calcareous nanno-
plankton flourished in surface water with a high content of
dissolved inorganic carbon ( CO
). Consequently, the car-
bon isotopes became heavier in coccolith-rich productivity,
and this led to a positive shift in surface-water
et al. 1996). Coccolithophorid productivity depends on at-
, availability of which increased during the
Widespread anoxia in the CCPB is also inferred from the
manganese sedimentary ores considering their redox-con-
trolled deposition (Roy 2006). Manganese became a mobile
component in anoxic deep-water environments, whereas
Mn-enriched waters precipitated in oxic environments on the
shallow continental shelves (Nijenhuis 1999). The manga-
nese ore basins of the Paratethys spread over South Ukraine,
Georgia, Azerbayan, northwest Turkey, and northeast Bulgar-
ia, etc. (Stolyarov & Kochenov 1995; Öztürk & Frakes 1995;
Varentsov 2002; Efendiyeva 2004, etc.). The Early Oli-
gocene deposition in the CCPB, like that in the Paratethyan
basins (e.g. Maikop Basin, Tard Sea, Ileada Sea, Melleta
Sea, etc.), took place under basin-wide anoxia, which pro-
duced metalliferous deposits (e.g. manganese ore beds near
Kišovce and Švábovce in the CCPB).
The Menilite Shales of the CCPB contain rich ichthyofauna
(Clupeidae, Serranidae – Chalupová 2000; Gregorová &
Fulín 2001) – Fig. 14. Fish mortality resulted most likely
from the toxicity of H
S-rich water, which surfaced from the
Oxygen Minimum Zone due to water-column stratification
or via upwelling (such as “en masse” extinction of Spiratella
fauna – Báldi et al. 1984). H
S is known to be an extremely
potent neurotoxin causing a cessation of pulmonary function
(Oschmann 1995). The toxic crush of fish fauna in the Meni-
lite Shales is indicated by their curved backbones, unat-
tached mouths and mostly adolescent forms (Pokorný 1992).
The alternation of fish shales with barren sediments in the
lamina scale points to seasonal changes in water circulation
or upwelling activity (cf. Vetö 1987).
Widespread anoxic and eutrophic conditions in the CCPB
most likely resulted from a high runoff and a positive water
balance similar to that in the Black-sea type basin (cf. Schulz
et al. 2005). Dissolved silicate, nitrate, ammonium, manga-
nese may have been supplied from continental sources.
Trophic resources could also have increased via upwelling,
which preferentially regenerated nutrients from organic mat-
ter under anoxic conditions (cf. Ingall & Jahnke 1997; Mur-
phy et al. 2000, etc.). Both, this runoff and upwelling
resulted in eutrophication of surface-waters. Biomass pro-
ductivity of surface waters in the Paratethyan basins resem-
bles a drifting flora and conditions similar to those in the
Sargasso Sea (Jerzmanska & Kotlarcyzk 1976; Kvaček &
Bubík 1990, etc.). This algal biocoenosis is dominated by
Sargassum maikopicum (Efendiyeva 2004).
Eustatic and cyclostratigraphic events
Paleoenvironmental changes in the CCPB could corre-
spond to eustatic cycles of sea-level history (Haq et al. 1988)
– Fig. 15. The basin began to develop from the late Lutetian
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
transgression (SBZ16), and this was followed by two 3
der cycles of shallow-marine deposition (TA 3.5—TA 3.6).
Compositional changes of large foraminiferal associations in
the upper SBZ17 Zone indicate a relative cooling event near
the TA 3/TA 4 boundary (Bartholdy et al. 1999). At that
time, the sea-level fall resulted in channel incision and for-
mation of fan delta systems (e.g. the Pucov Fan). New trans-
gression of the TA 4.1 cycle started during the early
Priabonian climatic optimum (last nummulites in the SBZ19
Zone, warm-temperate porticulosphaerids in P15 Zone),
which preceeded very unstable conditions during the late
Priabonian (TA 4.2—TA 4.3 cycles). The highstand of the
TA 4.3 cycle is marked by the productivity maximum of the
Globigerina Marls, which corresponds to the Terminal
Eocene Event. The eustatic cycle TA 4.4 began with the
largest sea-level fall near the Eocene/Oligocene boundary
(50—65 m after Pekar 2003; Wade & Pälike 2004). In the
CCPB, this eustatic event fits well with the positive
cursion, which marked the Eocene/Oligocene boundary on a
world-wide scale (Cavelier et al. 1981; Poore & Matthews
1984; Kennet & Barker 1990; Miller 1992; Zachos et al.
1996). Boundary sections in the CCPB imply the erosional
unconformities with a mass redeposition of morozovellids
and acarininids (e.g. in Pucov section). Such types of deep-
sea hiatuses, with reworked Eocene species, have been ob-
served near the Eocene/Oligocene boundary in the DSDP
sites (Keller 1985). These hiatuses were formed during the
global cooling and sea-level lowstand, resulting from chang-
es in oceanic deep circulation and bottom current erosion
(Keller 1983). Considering this, the erosive events of muri-
cate foraminifers in the CCPB indicate a lowstand phase be-
tween TA 4.3—TA 4.4 cycles at the Eocene/Oligocene
boundary. This eustatic event was associated with carbonate
depletion of the boundary clays, which are almost non-cal-
careous and are enriched by siliceous microfauna (e.g.
sponge microsklerites in Pucov section).
Early Oligocene transgression in the CCPB is documented
by a stepwise onlap of the ravinement surfaces of the Borové
Formation (Bartonian ramp-type formation) by mudstone
formations (NP22—NP23 Zones). The TA 4.4 cycle tended to
a higher sea level in the horizons of coccolithic limestones,
with the appearance of open-marine biota (e.g. pteropods).
Forced regression of the TA 4.4 cycle culminated in incision
of submarine channel systems of the CCPB during the Early
Oligocene (e.g. Tokáreň Fan). The sea-level lowstand of the
TA 4.4 cycle coincides with the beginning of the Paratethyan
isolation which is marked by endemic nannofossil species
and biosiliceous horizons (Menilitic facies). The transgres-
sion of the TA 4.5 cycle is expressed by the more calcareous
and mud-rich deposition of the Huty Formation, which re-
corded the FAD and ACME of Cyclicargolithus abisectus in
the NP24 Biozone. The falling stage of the TA 4.5 cycle
links with the most pronounced Mid-Oligocene regression,
which introduced the sand-rich turbidite systems in the
CCPB. Late Oligocene deposition of the CCPB persisted in a
lowstand eustasy of the TB 1 cycle, which is inferred from
progradational sequences, an increased abundance of braaru-
dosphaerids and reduced salinity. The sand-rich deposition
should have terminated by the Early Miocene, recorded by
the FAD of Discoaster drugii (23.8 Ma) and the appearance of
rich uvigerinid species in the topmost part of the Biely Potok
Formation which occurred earlier than the Eggenburgian
transgression of the Prešov Formation.
Transitional beds around the Eocene/Oligocene boundary
in the Carpathian Paleogene basins provide evidence of syn-
Fig. 15. Paleoenvironmental changes
in the Central-Carpathian Paleogene
Basin related to the global eustatic and
climatic changes. Sedimentary se-
quence records the global eustatic
events (Terminal Eocene Event and
Mid-Oligocene Event sensu Aubreu &
Anderson 1998), and provides the evi-
dence of the 3
(TA 3.5—TA 4.5) and alteration of ther-
mochrons and cryochrons.
genetic volcanism (Fig. 11). The first evidence of volcanic ac-
tivity was recorded by the vitritic-crystal tuffs in the basal
conglomerates beneath the Globigerina Marls (Glazek et al.
1998). Based on this, the tuffs correspond to the taphrostrati-
graphic event near 40 ± 2 Ma (Sochaczewski 2000). The most
frequent tuffs of the CCPB occur above the Globigerina Marls
(e.g. Prosiek, Huty, Pucov, Bystré), and this may correspond
to the Gasiory Tuff (34.6 ± 1.4 Ma – Wieser 1985) and tuffs
intercalated in the upper part of the Buda Marls and lower part
of the Tard Clay (Báldi 1984, 1986). Tuffs from the Buda Ba-
sin have been approximated to the Eocene/Oligocene bound-
ary (33.7 ± 1.0 Ma – Balogh & Pécskay 2001). As well as
tuffs, there are also tuffaceous sandstones containing a great
portion of biotite and volcanic quartz in the Orava region (e.g.
Veličná, Pucov), and this appears to be equivalent of biotitic
sandstones from the uppermost part of the Buda Marls in Hun-
gary (33.7 ± 1.0 Ma – Balogh & Pécskay 2001). Similar bi-
otite-rich layers are widely distributed (e.g. in Messignano
GSSP section – Montanari et al. 1985; Monechi et al. 2000;
Possagno section – Oberhänsli et al. 1984), and they record
volcanic activity at the Eocene-Oligocene boundary. Biotite-
rich layers in the CCPB are significantly enriched in content of
biosiliceous microfossils (e.g. sponge spicules), and therefore
the syngenetic volcanism may have enhanced the nutrient flux
from products of coeval pyroclastic activity (cf. Spezzaferri et
al. 2002; Vetö et al. 2007). Younger taphrostratigraphic events
in the CCPB correspond to the NP24 Biozone (Uhlík et al.
2002), or even to the 14 marker horizons in the Chocholow
Beds of the Podhale region (Westwalewicz-Mogilska 1986).
Volcanogenic activity is commonly recorded around the
Eocene/Oligocene boundary. At the same time-horizon, there
exists a lot of impactoclastic evidence, such as spherule-rich
layers, iridium anomalies, shocked quartz, Ni-rich spinels, etc.
(e.g. Montanari et al. 1993; Wei 1995; Clymer et al. 1996;
Pierrard et al. 1998). Therefore, the Late Eocene climatic dete-
rioration may have resulted from impact-related cooling (Von-
hof et al. 2000).
Alpine orogenesis and Paratethyan isolation
Synorogenic basins, such as the CCPB, should reflect an
important role of the Alpine tectonics. However, there is a
problem due to the synchroneity of the Alpine tectonic phases
and global eustatic events. The initial transgression in the
Central Western Carpathians corresponds well with the Mid-
dle Eocene transgression in the Mediterranean Tethys, but
also with the Illyrian tectogenesis. The Pyrenean tectonic ac-
tivity changed the distribution of lands and seas in Central Eu-
rope, which was still a part of the North Peri-Tethys (Popov et
al. 2004) and archipelago during the Late Eocene (Steininger
& Vessely 2000). Unlike the high rate of Alpine subduction,
the mountain uplift was less intensive and the Alps were most-
ly submerged (Frisch et al. 1998). Therefore, the isolation of
the Paratethyan basins was also facilitated by the largest
eustatic fall at the Eocene/Oligocene boundary (54 m – Pe-
kar et al. 2003). The subsequent isolation of the Paratethyan
basins led to the widespread distribution of marker horizons
(e.g. planktonic pteropods, fish shales, coccolithic chalks, dia-
toms and manganese deposits – Fig. 18), which most likely
indicates the predominance of global eustasy over regional
tectonics. The late Pyrenean phase at the beginning of the
NP24 Zone provided the strongest compressional event in the
Carpatho—Pannonian basins (Tari et al. 1993). Since the Late
Rupelian, the Carpathian flysch basins were supplied with
sand-rich turbidite systems, implying reorganization from the
passive-margin to active-margin fans (Soták et al. 2001).
Therefore, a higher source activity should increase in response
to tectonics. However, the Rupelian/Chattian boundary links
with the most pronouced change in global eustasy, which was
significantly lowered at the time of the first major glaciation in
Antartica. Considering this, the eustatic force of sand-rich
deposition in the Central Western Carpathians (incl. Krosno
Facies of the Outer Western Carpathians) should also be im-
portant (cf. Krhovský & Djurasinovič 1992).
The isolation of the Paratethys resulted from progressive
rise of the collisional wedge in the Alpine-Dinaric-Balkan in-
ternids (Báldi et al. 1984). Late Eocene—Early Oligocene up-
lift in the Dinarides, which became a landmass barrier, is
recorded by terrestrial and fluviatile-lacustrine deposits (Tari
& Pamić 1998). At that time, the Alpine collision achieved the
Pyrenean phase, which is considered to be the main orogenic
phase in the Alps (Trümpy 1980). The Alpine Foreland Basin
lost connection with the Tethys, and this is recorded by the ap-
pearance of the endemic fauna in Solenovian time (Steininger
& Wessely 2000). The Alpine collision induced a large-scale
continental escape of the Carpatho-Pannonian terranes. The
Pannonian Basin was detached from the Slovenian Basin and
emplaced to the Intra-Carpathian area along the Insubric—
Balaton and Mid-Hungarian Lineaments (Csontos et al. 1992;
Fodor et al. 1999; Schmiedl et al. 2002). Accordingly, the
escape tectonics seem to be responsible for the first isolation
of the Carpathian Paleogene basins. The first isolation in the
CCPB was already indicated by the zoogeographical separa-
tion of the Spiš Basin, where nummulitids are absent, and
mollusc fauna indicate an affinity to the Transylvanian Basin
and Caucasus – Aral area (Volfová 1962). Endemic fauna
of Solenovian-type molluscs (Ergenica? sp., Janschinella?
sp.) and cerites (Tympanotonus sp.) are also recognizable
from the Paleogene sediments of the Hron Valley Depression
The Carpathian Paleogene basins were differentiated in
sea-floor topography, bathymetry and water exchange. Dur-
ing the Eocene/Oligocene transition, the differentiation was
recorded by the Globigerina Marls, deposition of which was
controlled by anti-estuarine circulation, high rate calcareous
productivity and lowered depth of calcium carbonate dissolu-
tion (CCD) – Fig. 16. The CCD declines in the conditions of
enhanced productivity, which correspond to a peak of produc-
tivity near 33.5 Ma (Diester-Haass & Zahn 1996), and this
caused the CCD drop near the Eocene/Oligocene boundary
(Thunell & Corliss 1986; Tripati et al. 2005). Following the
downslope excursion of the CCD, the Globigerina Marls and
red-coloured marls occur suddenly in carbonate-free deep-wa-
ter sediments of the Carpathian Paleogene basins. Successive
cooling resulted in shoaling of the CCD due to a higher satura-
tion of cold bottom-water by CO
. As a consequence, the sea-
water became more acid and corrosive to the calcareous
components. Above the Globigerina Marls, the carbonate dis-
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
Fig. 16. Upper Eocene basins with the Tethyan—Boreal communication, which resulted in cold-temperate productivity of the Globigerina
Marls (GM), thermal water mass stratification (TCL – thermocline), anti-estuarine circulation, deep-water oxygenization (DWV), prolif-
eration of benthic life (B), red-coloured marl deposition (RCM), enhanced evaporation (E), lowered precipitation (P), reduced runoff, etc.
Fig. 17. Lower Oligocene basins with the Black-sea type hydrography, estuarine circulation, inflow of Boreal waters, surface-water over-
flow (FWI – fresh water influx), closure of the Mediterranean seaway, semi-isolation, water column stratification (CHM – chemocline),
deep-water stagnation, bottom-water anoxia, upwelling (UPW), eutrophication (EF), dinoflagellate- and diatom-based productivity, export
flux, fish fauna intoxication, precipitation (P), humidity, monsoonal activity (M), etc.
solution increased considerably, leading to non-calcareous
deposition of the Menilite Formation.
Carpathian and Hungarian Paleogene basins record a differ-
ent rate of the Paratethyan isolation or Tethyan communica-
tion. While the Buda Basin reveals a full-marine deposition,
richness of foraminiferal microfauna (Tritaxia szaboi beds,
Buda Marls, Kiscell Clays), later termination of low-latitude
habitats (e.g. latest nummulites, lepidocyclinids and globo-
quadrinids), the hiatuses (“infra-Oligocene” regression) and
proximity of volcanogenic activity, the Central-Carpathian
Paleogene Basin implies restricted environments, impoverish-
ment of foraminiferal microfauna (Huty Formation, Zuberec
Formation, Brzegi Member), predominance of oxygen-defi-
cient taxa, sapropelitic deposition and biosilicite productivity
(Menilite black shales and cherts), calcareous nannoplankton
blooms, manganese beds and tuffaceous admixture, etc. Con-
sidering this, the Oxygen Minimum Zone (OMZ) implied the
southward weakening of anoxia towards the Tard Sea in the
Hungarian Paleogene Basin and its northward widening to-
wards the Carpathian Flysch Sea (Fig. 19). The northern part
of the CCPB seems to be more isolated, humid, eutrophicated,
oxygen depleted and cooled by Boreal waters. On contratry,
the southern part of this basin represents the neritic and near-
shore zone of the North Buda Paleogene Basin (Gross 1978;
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
Fig. 19. Tethyan vs. Paratethyan marker facies in the Carpathian Paleogene basins. The figure exhibits northward increase in the intensity
of the Paratethyan isolation and southward amplification of the Tethyan influence.
Central Paratethys, the Transylvanian Paleogene Basin was
connected rather with the Eastern Paratethys, considering im-
migrations of mollusc fauna (e.g. Lenticorbula sokolovi, L.
helmerseni – Báldi 1980; Moisescu 1995), evaporitic and la-
goonal-lacustrine deposition (Proust & Hosu 1996), oolitic
ironstones such as those in the Maikop Basin (Popov & Stol-
yarov 1996) and even migrant Asian mammals (Baciu &
Hartenberger 2001). The affinities of the Transylvanian Paleo-
gene Basin to the Solenovian Sea and Boreal Sea (Nucula
compta Level) proved its more northern position within the
The Intra-Carpathian system of the Paleogene basins was
disturbed during plate-tectonic reorganization of the ALCA-
PA terranes (Figs. 20, 21). The Slovenian and Hungarian Pa-
leogene basins were accommodated more southerly, and later
on they have been shifted to their present position (Csontos et
al. 1992). The Hungarian and Central-Carpathian Paleogene
Basins are quite different (epicontinental-type basin vs. mar-
ginal basin of the Carpathian Flysch Sea), exhibiting no direct
paleogeographical connection between them. The northern
limit of the Hungarian Paleogene Basin is inferred in the Šahy
Antiform, which represents a nearshore zone of the Kiscellian
Sea containing sebhka-type facies (Vass 2003). The Veporic
borderland was submerged in some places by transversal de-
pressions (Soták et al. 2004). The sediments of the Hungarian
Paleogene Basin in the Štúrovo, Lučenec and Rimavská kotli-
na depressions show a closer proximity to the Mediterranean
Sea, revealed by their full-marine facies, tide-influenced depo-
sition, abundance of foraminiferal microfauna, presence of
lepidocyclinids and miogypsinids, etc. (Samuel & Vaňová
1967; Vaňová 1975; Sztanó 1995; Holcová-Šutovská 1996;
Nagymarosy 1990), influenced by the Tethyan waters. The
circulation model of the Paratethyan sea (Dohman 1991;
Schulz 2003) presumed a mixture of the Boreal deep water
and Tethyan surface water. The restricted seaway connection
with the Mediterranean Tethys during the Early Oligocene re-
sulted in fresh water overflow in the Paratethyan basins. The
high runoff and separation of the CCPB by the intrabasinal
high (Low Tatra Highland), indicates an estuarine-type circu-
lation with an inflow of marine bottom waters and an outflow
of fresh surface waters (Fig. 17). Here, the CCPB revealed
features of the Black Sea-type basins similar to other basins in
the Paratethys Sea (Schulz et al. 2005).
During the Early Oligocene, the oxygen crisis affected the
whole Central Paratethys, including the Hungarian Paleogene
Basin (Tard Clays), Slovenian Basin (Fish Shale – Tegel
Unit), Carpathian Paleogene basins (Menilite Beds) and Aus-
trian foreland basin (Schöneck—Ottenthall Formations), Tran-
sylvanian Basin (Ileanda Shale), etc. Nevertheless, some of
the Paratethyan basins were reconnected with the Mediterra-
nean Tethys, as is indicated by the Lower Oligocene bioher-
mal limestones containing nummulitids (e.g. N. vascus) in the
Slovenian and Hungarian Paleogene Basins (Gornji Grad
Beds – Nebelsick et al. 2000; Szépevölgy Limestones – Ko-
rpás et al. 1999), but not/or very rarely in Central-Carpathian
Paleogene Basin (Buček & Filo 2004). On the other side, the
Spiratella-rich fauna of the Hungarian Paleogene Basin (Tard
Clay) provide evidence of the cold-water influence of the Bo-
real Sea. This implies that Tethyan-Boreal communication,
most likely via the Mid-Hungarian corridor and Slovenian
Strait, which is indicated by the mixed mollusc fauna in the
Kiscellian Clay (Báldi 1984). In contrast to other basins in the
Holcová 2001, etc.). It is most likely that the Hungarian and
Central-Carpathian Paleogene Basins came into tectonic jux-
taposition due to the NE-directed displacement of the Pelso
Unit (Nagymarosy 1990). This unit attained its present posi-
tion by tectonic and rotational movement
from the Late Oligocene to Early Miocene
(Vass et al. 1996). This is also the case in the
Hungarian and Transylvanian Paleogene ba-
sins, which were jointed by large-scale tec-
tonic movement of the ALCAPA and
Sedimentary sequences of the Central-
Carpathian Paleogene Basin provide evi-
dence of following paleoenvironmental and
climatic events related to the Eocene/Oli-
EBi-1: growth potential of carbonate
platform progressively weakened due to
drowning, temperature drop and overfeed-
ing stress of oligotrophic organisms (e.g.
EPi-1: demise of the carbonate platform
in response to climatic deterioration and
mesotrophication in the Carpathian basins
during the Late Eocene;
EPi-2: temperate- to cool-water produc-
tivity of Bryozoan Marls and Globigerina
Marls, indicating a rise of nutrients from
runoff or via upwelling. Plankton produc-
tivity of the Globigerina Marls implies the
changes in life-mode strategy, diversity,
temperature preferences and water-depth
TEE: peak abundance of bloom-forming
globigerinids corresponds to climate-pro-
ductivity maximum at 33.7 Ma, which re-
veals the increase of
O (cooling) and the
respired due to me-
tabolism); – the positive shift of oxygen
isotopes indicates a decrease in seawater
boundary (12—15 °C); – contemporaneous
TA 4.4 sea-level fall and increase in car-
bonate dissolutional rate; – taphrostrati-
graphic events near 33.7 ± 1.0 Ma;
Oi-1: the major climatic turnover took
place in the Early Oligocene; – cooling
and humidity led to temperate/wet condi-
tions; – fresh water discharge, water-col-
umn stratification, brackish surface water,
dinoflagellate-rich productivity, estuarine
circulation, oxygen depletion, sulphidic
bottom water and seasonal upwelling activ-
S intoxication and a high mortality);
Oi-2: eutrophication, expansion of OMZ,
photic zone anoxia and stagnant regime of
Fig. 20. Palinspastic scheme of the Peri-Tethyan domain in the Late Eocene. Basinal
systems show a plate-tectonic reorganization due to the escape tectonics of the ALCA-
PA Unit. They were full-marine basins with southward connection to the Mediterranean
Sea and northward connection to the Boreal Sea.
Fig. 21. Palinspastic scheme of the basinal systems of the Proto-Paratethys in the be-
ginning of the Oligocene. Paleogeography was changed in response to the Mesoalpine
tectogenesis, which led to the uplift of the Dinaroalpidic borderland and first isolation
of the Alpine foreland, Carpatho-Pannonian and Transylvanian basins.
the CCPB during the mid-Early Oligocene; – sapropelitic
deposition of menilite shales, fish shales and manganese
ores; – episodical blooms of calcareous nannoplankton in
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
the period of runoff stillstand, anti-estuarine circulation and
higher carbonate resources (Tylawa-type limestones); – coc-
colithophorid-rich productivity links with a positive
cursions, indicating short-lived greenhouse events and/or
decrease photosynthetic rate of
Earliest Paratethyan stage: the first isolation event in the
NP23 Biozone, the appearance of Paratethys-endemic nanno-
fossils, diatom-based productivity and biosiliceous deposition;
differentiation of basin-floor topography and paleogeographic
communication; – plate-tectonic reorganisation of the intra-
Carpathian basins during the Alpine orogenesis.
Acknowledgments: The author wishes to thank to Przemys-
law Gedl, András Nagymarosy and Diego Puglisi for their valu-
able comments and improvements in earlier drafts of the
manuscript. This research was supportedby the Scientific Grant
Agency of the Ministry of Education of the Slovak Republic
and the Slovak Academy of Sciences (VEGA Grant 2/0140/09).
Abreu V.S. & Anderson J.B. 1998: Glacial eustasy during the Ceno-
zoic sequence stratigraphic implications. AAPG Bull. 82, 7,
Baciu C. & Hartenberger J.-L. 2001: Un exemple de corrélation
marin-continental dans le Priabonien de Roumanie. Remarques
sur la Grande Coupure. C.R. Acad. Sci., Earth Planet. Sci., Paris
Balogh K. & Pécskay Z. 2001: K/Ar and Ar/Ar geochronological
studies in the Pannonian-Carpathians-Dinarides (PANCARDI)
region. Acta Geol. Hung. 44, 2—3, 281—299.
Bartholdy J., Bellas S.M., Cosovic V., Fucek V.P. & Keupp H. 1999:
Processes controlling Eocene mid-latitude larger Foraminifera
accumulations: Modelling of the stratigraphic architecture of a
fore-arc basin (Podhale Basin, Poland). Geol. Carpathica 50, 6,
Báldi T. 1980: The early history of the Paratethys. Földt. Közl. 110,
Báldi T. 1984: The terminal Eocene and Early Oligocene events in
the Hungary and the separation of an anoxic, cold Paratethys.
Eclogae Geol. Helv. 77, 1, 1—27.
Báldi T. 1986: Mid-Tertiary stratigraphy and palaeogeographic evo-
lution of Hungary. Akad. Kiado, Budapest, 1—201.
Báldi T., Horváth M., Nagymarosy A. & Varga P. 1984: The Eocene-
Oligocene boundary in Hungary. Acta Geol. Hung. 27, 1—2, 41—65.
Bemis B.E., Spero H.J., Lea W.D. & Bijma J. 2000: Temperature in-
fluence on the carbon isotopic composition of Globigerina bul-
loides and Orbulina universa (planktonic foraminifera). Mar.
Micropaleont. 38, 213—228.
Benešová E. 1963: Microbiostratigraphical study of the Vlachy-1
borehole. In: Chmelík et al. (Eds.): Reference borehole Vlachy-1.
Práce Výsk. Úst. Čs. Naft. Dolu, XX, Publ. 92, 49—57.
Berlin I.S., Barhatova N.N. & Habakov A.V. 1976: Ca/Mg methods
for estimation of the Eocene paleotemperature conditions in
the Hungarian and Slovak seas. Magy. Áll. Földt. Intéy. Evi. Je-
lent. 1974, 477—485 (in Hungarian).
Bicci E., Ferrero E. & Gonera M. 2003: Palaeoclimatic interpretation
based on Middle Miocene planktonic Foraminifera: the Silesia
Basin (Paratethys) and Monferrato (Tethys) record. Palaeo-
geogr. Palaeoclimatol. Palaeoecol. 196, 3—4, 265—303.
Biolzi M. 1985: Late Eocene—Early Miocene boundary in the selected
Atlantic, Mediterranean and Paratethyan sections based on bios-
tratigraphic and stable isotope evidences. Mem. Soc. Geol. Ital.
Blaicher J. 1973: Microfauna of the Podhale Flysch in the Zakopane
IG 1 Borehole. Biul. Inst. Geol. 265, 105—133.
Blow W.H. 1969: Late Middle Eocene to Recent planktonic foramin-
iferal biostratigraphy. Proc. 1
Int. Conf. Plankt. Microfossils,
Geneva, 1967 (Vol. I.), 199—422.
Boersma A. & Premoli Silva I. 1991: Distribution of Paleogene
planktonic foraminifera: analogies with the recent? Palaeo-
geogr. Palaeoclimatol. Palaeoecol. 83, 29—48.
Boersma A., Premoli Silva I. & Hallock P. 1995: Trophic models for
the well-mixed and poorly mixed warm oceans across the Pale-
ocene/Eocene epoch boundary. In: Aubry M.-P., Lucas S. &
Berggren W.A. (Eds.): Late Paleocene—Early Eocene climatic
and biotic events in the marine and terrestrial records. Columbia
University Press, New York, 204—213.
Bolli H.M. 1972: The genus Globigerinatheka Brönnimann. J. Fo-
ram. Res. 2, 3, 109—136.
Bosch H.-J., Sinninghe Damsté J.S. & Leeuw J.W. 1998: Molecular
palaeontology of Eastern Mediterranean sapropels: evidence for
photic zone euxinia: In: Robertson A.H.F., Emeis K.C., Richter
C. & Camerlenghi A. (Eds.): Proceedings of the Ocean Drilling
Program. Sci. Results 160, 285—295.
Braga G. & Barbin V. 1988: Les Bryozoaries du Priabonien strato-
typique Province Vincenza, Italie. Rev. Paleobiologie 7, 2,
Bubík M. 1992: Low diversity calcareous nannoplankton assemblages
from the Oligocene Štibořice member of the Menilitic formation
(West Carpathians, Czechoslovakia) from Bystřice nad Olší.
Knihovnička ZPN 14b, 2, 223—245.
Buček S. & Filo I. 2004: Oligocene larger foraminifers in Paleogene
sediments westward of Banská Bystrica (Middle Slovakia). Slo-
vak Geol. Mag. 10, 4, 277—283.
Carannante G., Esteban M., Milliman J.D. & Simone L. 1988: Car-
bonate lithofacies as paleolatitude indicators: problems and limi-
tations. Sed. Geol. 60, 1—4, 333—346.
Cavelier C., Chateauneuf J.-J., Pomerol Ch., Rabussier D., Renard M.
& Vergnaud-Grazzini C. 1981: The geological events at the
Eocene/Oligocene boundary. Palaeogeogr. Palaeoclimatol.
Palaeoecol. 36, 223—248.
Chalupová B. 2000: Eocene fish fauna from the Menilite Beds (Huty
Formation) of the Central Carpathian Paleogene Basin (Orava
region, NW Slovakia). Slovak Geol. Mag. 2, 3, 168—171.
Clymer A.K., Bice D.M. & Montanari A. 1996: Shocked quartz from
the late Eocene: impact evidence from Massignano, Italy. Geol-
ogy 24, 483—486.
Coccioni R., Basso D., Brinkhuis H., Galeotti S., Gardin S., Monechi
S. & Spezzaferri S. 2000: Marine biotic signals across a late
Eocene impact layer at Massignano, Italy: evidence for long-
term environmental perturbations? Terra Nova 12, 258—263.
Craig H. 1965: The measurement of oxygen isotope paleotempera-
tures. In: Tongiorgi E. (Ed):
Stable isotopes in oceanographic
studies and paleotemperatures. Consiglio Nationale Ricerche
Lab. Geol. Nucl., Pisa, 161—182.
Csontos L., Nagymarosy A., Horváth F. & Kováč M. 1992: Tertiary
evolution of the Intra-Carpathian area: a model. Tectonophysics
Diester-Haass L. 1991: Eocene-Oligocene paleooceanography in the
Antartic Ocean, Atlantic sector (Maud Rise, ODP Leg 113, Sites
689 and 690). Mar. Geol. 100, 249—276.
Diester-Haass L. & Zahn R. 1996: Eocene-Oligocene transition in the
Southern Ocean: history of water mass circulation and biologi-
cal productivity. Geology 24, 2, 163—166.
Diester-Haass L. & Zahn R. 2001: Paleoproductivity increase at the
Eocene—Oligocene climatic transition: ODP/DSDP sites 763 and
592. Palaeogeogr. Palaeoclimatol. Palaeoecol. 172, 153—170.
Dohmann L. 1991: Die unteroligozänen Fischschiefer im Molasse-
becken. PhD Thesis, Ludwig-Maximilian-Universität, Munich,
Efendijeva M.A. 2004: Anoxia in waters of the Maikop paleobasin
(Tethys Ocean, Azeri sector), with implications for the modern
Caspian Sea. Geo-Marine Lett. 24, 177—181.
Fodor L., Csontos Bada G., Györfi I. & Benkovics L. 1999: Tertiary
tectonic evolution of the Pannonian Basin system and neigh-
bouring orogens: a new synthesis of palaeostress data. In: Du-
rand B., Jolivet L., Horváth F. & Séranne M. (Eds.): The
Mediterranean Basins: Tertiary extension within the Alpine
Orogen. Geol. Soc. London, Spec. Publ. 156, 295—334.
Frich W., Kuhlemann J., Dunkl I. & Brügel A. 1998: Palinspastic re-
construction and topographic evolution of the Eastern Alps dur-
ing late Tertiary tectonic extrusion. Tectonophysics 297, 1—15.
Galeotti S., Coccioni R. & Gersonde R. 2002: Middle Eocene—Early
Pliocene subantarctic planktic foraminiferal biostratigraphy of
Site 1090, Agulhas Ridge. Mar. Micropaleont. 45, 357—381.
Garecka M. 2005: Calcareous nannoplankton from the Podhale Fly-
sch (Oligocene—Miocene, Inner Carpathians, Poland). Stud.
Geol. Pol. 124, 353—369.
Gedl P. 1999: Palynology of the Eocene-Oligocene boundary in the
Polish Flysch Carpathian). Przegl. Geol. 47, 394—399 (in Polish).
Gedl P. 2000a: Biostratigraphy and palaeoenvironment of the
Podhale Palaeogene (Inner Carpathians, Poland) in the light of
palynological studies. Part II. Summary and systematic descrip-
tions. Stud. Geol. Pol. 117, 155—303.
Gedl P. 2000b: Palaeogeography of the Podhale Flysch (Oligocene;
Central Carpathians, Poland) – its relation to the neighbour-
hood areas as based on palynological studies. Slovak Geol. Mag.
6, 2—3, 150—154.
Gedl P. & Leszczyński S. 2005: Palynology of the Eocene-Oligocene
transition in the marginal zone of the Magura Nappe at Folusz
(Western Carpathians, Poland). Geol. Carpathica 56, 2, 155—167.
Geel T. 2000: Recognition of stratigraphic sequences in carbonate
platform and slope deposits: empirical models based on microfa-
cies analysis of Palaeogene deposits in southeastern Spain.
Palaeogeogr. Palaeoclimatol. Palaeoecol. 155, 211—238.
Glazek J., Przybycin A. & Sochaczewski A. 1998: Tuffite between
Upper Eocene conglomerates of the Tatra Mountains (Car-
pathians, Poland) and its stratigraphic importance. Przegl. Geol.
46, 7, 622—630.
Gonzalvo C. & Molina E. 1992: Biostratigraphy and chronostratog-
raphy of the Eocene-Oligocene transition in Torre Cardela and
Massignano (Italia). Rev. Espan
~ola Paleont. 7, 2, 109—126 (in
Gregorová R. & Fulín M. 2001: Fossil fish fauna at Inner Car-
pathian Paleogene from the locality Bystré nad Top ou. Natura
Carpathica XLII, 43—54.
Gross P. 1978: Paleogene beneath Central-Slovakian neovolvanic
rocks. In: Vozár J. (Ed.): Paleogeographical evolution of the
West Carpathians. Dionýz Štúr Institute of Geology, Bratislava,
Gross P. 1981: Tuffite sandstone in Inner Carpathian Paleogene of
Orava. Geol. Práce, Spr. 84, 157—164.
Gross P., Köhler E., Biely A., Franko O., Hanzel V., Hricko J.,
Kupčo G., Papšová J., Priechodská Z., Szalaiová V., Snopková
P., Stránska M., Vaškovský I. & Zbořil L. 1980: Geology of
Liptovská Kotlina (Depression). Dionýz Štúr Institute of Geo-
logy, Bratislava, 7—236.
Gross P., Buček S., Durkovič T., Filo I., Maglay J., Halouzka R.,
Karoli S., Nagy A., Spišak Z., Žec B., Vozár J., Borza V.,
Lukáčik E., Janočko J., Jetel J., Kubeš P., Kováčik M., Žaková
E., Mello J., Polák M., Siráňová Z., Samuel O., Snopková P.,
Raková J., Zlinská A., Vozárová A. & Žecová K. 1999: Expla-
nations to geological map of the Popradská kotlina and Hornad-
ská kotlina Depressions, Levočské vrchy Mts., Spiš-Šariš High-
land, Bachureň Mts. and Šarišská vrchovina Highland. Geol.
Surv. Slovak Republic, Bratislava, 1—239.
Hallock P. 1987: Fluctuations in the trophic resource continuum:
a factor in global diversity cycles? Paleoceanography 2, 457—471.
Hallock P. & Glenn E.C. 1986: Larger Foraminifera: a tool for pale-
oenvironmental analysis of Cenozoic carbonate depositional fa-
cies. Palaios 1, 55—64.
Hallock P. & Schlager W. 1986: Nutrient excess and the demise of
coral reefs and carbonate platforms. Palaios 1, 389—398.
Hallock P., Premoli Silva I. & Boersma A. 1991: Similarities be-
tween planktonic and larger foraminiferal evolutionary trends
through Paleogene paleooceanographic changes. Palaeogeogr.
Palaeoclimatol. Palaeoecol. 83, 49—64.
Hantken M. 1875: Die Fauna der Clavulina szaboi-Schichten. I. Fora-
miniferen. Mitt. Jb. K. Ungar. Geol. Anst., Budapest 4 (1), 1—93.
Hanzlíková E. 1981: Biostratigraphy and ecology of the Menilite
beds in the Moravia. Zemní Plyn Nafta 26, 1, 29—62.
Haq B.U., Hardenbol J. & Vail P.R. 1988: Mesozoic and Cenozoic
chronostratigraphy and cycles of sea-level change. In: Wilgus et
al. (Eds.): Sea-level changes – an integral approach. SEPM
Spec. Publ. 42, 71—108.
Holcová-Šutovská K. 1996: Foraminiferal assemblages: indicators of
the paleoenvironmental evolution of marine basins and eustatic
changes (Kiscellian—Karpathian of the South Slovak and Danube
basins). Geol. Carpathica 47, 2, 119—130.
Holcová K. 2001: New methods in foraminiferal and calcareous nan-
noplankton analysis and evolution of Oligocene and Miocene
basins of the Southern Slovakia. Slovak Geol. Mag. 7, 1, 19—41.
Hottinger L. 1983: Processes determining the distribution of larger
foraminifera in space and time. Utrecht Micropaleont. Bull. 30,
Ingall E. & Jahnke R. 1997: Influence of water-column anoxia on the
elemental fractionation of carbon and phosphorus during sedi-
ment diagenesis. Mar. Geol. 139, 219—229.
James N.P. & Bone Y. 2000: Eocene cool-water carbonate and biosil-
iceous sedimentation dynamics, St. Vincent Basin, South Aus-
tralia. Sedimentology 47, 761—786.
James N.P., Feary D.A., Surlyk F., Toni Simo J.A., Betzler Ch., Hol-
bourn A.E., Li Q., Matsuda H., Machiyama H., Brooks G.R.,
Andres M.S., Hine A.C. & Malone M.J. 2000: Quaternary bryo-
zoan reef mounds in cool-water, upper slope environments. Ge-
ology 28, 7, 647—650.
Jerzmanska A. & Kotlarczyk J. 1976: The beginnings of the Sargasso
assemblage in the Tethys. Palaeogeogr. Palaeoclimatol. Palae-
oecol. 20, 297—306.
Kaczmarska I. & Kilarski W. 1979: The structure of Melosira sulcata
(EHR.) KÜTZ. Var. sulcata frustules from Lower Oligocene di-
atomites from futoma (Carpathians, Poland). Ann. Soc. Geol.
Pologne XLIX, 1—2, 185—194.
Kázmér M., Monostori M. & Zágoršek K. 1993: Benthic communities
on the Upper Eocene slope at Budapest, Hungary: A progress re-
port. Discussiones Palaeontologicae, Budapest 39, 79—89.
Kázmér M., Dunkl I., Frisch W., Kuhlemann J. & Ozsvárt P. 2003:
The Palaeogene forearc basin of the Eastern Alps and Western
Carpathians: subduction erosion and basin evolution. J. Geol.
Soc., London 160, 413—428.
Keller G. 1983: Paleoceanographic implications of Miocene deep-sea
hiatuses. Geol. Soc. Amer. Bull. 94, 590—613.
Keller G. 1985: Eocene and Oligocene stratigraphy and erosional un-
conformities in the Gulf of Mexico and Gulf Coast. J. Paleontol-
ogy 59, 4, 882—903.
Keller G. & Pardo A. 2004: Disaster opportunists Guembelitrinidae:
index for environmental catastrophes. Mar. Micropaleont. 53,
Keller G., MacLeod N. & Barrera E. 1992: Eocene—Oligocene faunal
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
turnover in planktic foraminifera, and Antarctic glaciation. In:
Prothero D.R. & Berggren W.A. (Eds.): Eocene-Oligocene cli-
matic and biotic evolution. Princeton University Press, 218—243.
Kennett J.P. 1982: Marine geology. Prentice-Hall, Englewood Cliffs,
Kennett J.P. & Barker P.F. 1990: Latest Cretaceous to Cenozoic cli-
mate and oceanographic developments in the Ocean Drilling
Program. Scientific Results, Vol. 113, College Station, Ocean
Drilling Program, Texas, 937—960.
Konzalová M., Rákosi L. & Snopková P. 1993: Correlations of Paleo-
gene palynoflora from Bohemia, Hungary, Slovakia. In: Plande-
rová et al. (Eds.): Paleofloristic and paleoclimatic changes
during Cretaceous and Tertiary. GÚDŠ, Bratislava, 63—76.
Koopmans M.P., Köster J., Heidy M.E., van Kaam-Peters H.M.E.,
Kenig F., Schouten S., Hartgers W.A., De Leeuw J.W. & Sin-
ninghe Damsté J.S. 1996: Diagenetic and catagenetic products
of isorenieratene: Molecular indicators for photic zone anoxia.
Geochim. Cosmoch. Acta 60, 22, 4467—4496.
Korpás L., Lantos M. & Nagymarosy A. 1999: Timing and genesis of
early marine caymanites in hydrothermal paleokarst system of
Buda Hills, Hungary. Sed. Geol. 123, 9—29.
Kotulová J. 2004: Oligocene Menilite black shales – geochemical
and maceral analysis. 32
International Geological Congress,
Florence 2004, Abstracts, 1—752.
Kováč M. 2000: Geodynamic, paleogeographic and structural devel-
opment of the Carpatho-Pannonian region during the Miocene:
a new insight on the Neogene basins of the Slovakia. SAS Publ.—
VEDA, Bratislava, 7—202.
Köhler E. 1998: The last occurrences of large nummulites in the
Western Carpathian Eocene. Zemní Plyn Nafta 43, 1, 163—171.
Krhovský J. 1981a: Stratigraphy and paleoecology of the Menilitic
Formation of the Ždánice Unit and the Diatomites of the
Pouzdřany Unit (the Western Carpathians, Czechoslovakia).
Zemní Plyn Nafta 26, 1, 45—62.
Krhovský J. 1981b: Microbiostratigraphic correlations in the Outer
Flysch Units of the southern Moravia and influence of the eusta-
sy on their palaeogeographical development. Zemní Plyn Nafta
26, 4, 665—688.
Krhovský J. 1995: Early Oligocene palaeoenvironmental changes in
the West Carpathian flysch belt of Southern Moravia. Geol. Soc.
Greece, Spec. Publ. 4, 209—213.
Krhovský J. & Djurasinovič M. 1992: The nannofossil chalk layers in
the Early Oligocene Štiborice Member in Velké Němčice (the
Menilitic Formation, Ždánice Unit, South Moravia): orbitally
forced changes in paleoproductivity. Knihovnička ZPN 15, 33—53.
Krhovský J., Adamová M., Hladíková J. & Maslowská H. 1993: Pale-
oenvironmental changes across the Eocene/Oligocene boundary
in the Ždánice and Pouzdřany Units (Western Carpathians, Czech-
oslovakia): the long-term trend and orbitally forced changes in
calcareous nannofossil assemblages. In: Hamršmid B. & Young
J.R. (Eds.): Nannoplankton research. Vol. II, Proc. 4
ference, Prague 1991. Knihovnička ZPN 14b, 105—155.
Kvaček Z. & Bubík M. 1990: Oligocene flora of the Štibořice Mem-
ber and geology at Bystřice nad Olší (NE Moravia). Věst. Ústř.
Úst. Geol. 65, 2, 81—94.
Leszczyński S. 1996: Origin of lithological variation in the sequence
of the sub-Menilite Globigerina Marl at Znamirowice (Eocene-
Oligicene transition, Polish Outer Carpathians). Ann. Soc. Geol.
Pol. 66, 245—267.
Leszczyński S. 1997: Origin of the sub-Menilite Globigerina Marls
(Eocene-Oligocene transition) in the Polish Outer Carpathians.
Ann. Soc. Geol. Pol. 67, 367—427.
Light J.M. & Wilson J.B. 1998: Cool-water carbonate deposition on
the West Shetland Shelf: a modern distally steepened ramp. In:
Wright V.P. & Burchette T.P. (Eds.): Carbonate ramps. Geol.
Soc., Spec. Publ. 149, 73—105.
Lyle M., Lyle A.O., Backman J. & Tripati A. 2005: Biogenic sedi-
mentation in the Eocene equatorial Pacific – the stuttering
greenhouse and Eocene carbonate compensation depth. In: Wil-
son P.A., Lyle M. & Firth J.V. (Eds.): Proceedings of the Ocean
Drilling Program. Sci. Results 199, 1—35.
MacLeod N. 2006: Size, extinction, survivorship and phylogeny in
foraminifera. Geol. Soc. Amer. 38, 7, 1—473.
Marjanac T., Babac D., Benic J., Cosovic Drobne K., Marjanac L.,
Pavlovec R. & Velimirovic Z. 1998: Eocene carbonate sedi-
ments and sea-level changes on the NE part of Adriatic carbon-
ate platform (island of Hvar and Pelješac peninsula, Croatia). In:
Hottinger L. & Drobne K. (Eds.): Paleogene shallow benthos of
the Tethys, 2. Dela – Opera SAZU 4, Ljubljana 34/2, 243—254.
Marschalko R. & Volfová J. 1960: Submarine slide and its macrofau-
na in the Palaeogene of the Central Carpathians. Geol. Práce,
Spr. 19, 95—108.
Melinte M.C. 2005: Oligocene palaeoenvironmental changes in the
Romanian Carpathians, revealed by calcareous nannofossils.
Studia Geol. Pol. 124, 341—352.
Miclaus C., Loiacono F., Puglisi D. & Baciu S.D. 2009: Eocene—Oli-
gocene sedimentation in the external areas of the Moldavide Ba-
sin (Marginal Folds Nappe, Eastern Carpathians, Romania):
sedimentological, paleontological and petrographical approach-
es. Geol. Carpathica 60, 5, 397—417.
Miller K.G. 1992: Middle Eocene to Oligocene stable isotopes, cli-
mate, and deep-water history: the Terminal Eocene Event? In:
Prothero D.R. & Berggren W.A. (Eds.): Eocene-Oligocene cli-
matic and biotic evolution. Princeton University Press, 160—177.
Miller K.G., Wright J.D. & Fairbanks R.G. 1991: Unlocking the ice
house: Oligocene-Miocene oxygen isotopes, eustasy, and mar-
gin erosion. J. Geophys. Res. 96, 6829—6848.
Mitchell S.F., Paul C.R.C. & Gale A.S. 1996: Carbon isotope and se-
quence stratigraphy. In: Howell J.A. & Aitken J.F. (Eds.): High
resolution sequence stratigraphy: Innovations and applications.
Geol. Soc. London, Spec. Publ. 104, 11—24.
Moisescu V. 1995: Considerations on the fauna of Lenticorbula in the
Kiscellian deposits of the Transylvanian Basin (Cluj-Ticu area).
Rom. J. Paleont. 76, 73—76.
Molina E., Gonzalvo C., Ortiz S. & Cruz L.E. 2006: Foraminiferal
turnover across the Eocene-Oligocene transition at Fuelte Caldera,
southern Spain: No cause-effect relationship between meteorite
impacts and extinctions. Mar. Micropaleont. 58, 4, 270—286.
Molnár J., Karoli S. & Zlinská A. 1992: Oligomiocene of Šarišská vr-
chovina Mts. Geol. Práce, Spr. 95, 41—45.
Monechi S., Buccianti A. & Gardin S. 2000: Biotic signals from nan-
noflora across the iridium anomaly in the upper Eocene of the
Massignano section: evidence from statistic analysis. Mar. Mi-
cropaleont. 39, 219—237.
Montanari A., Drake R.E., Bice A.W., Curtis G.H., Turrin D.M. &
DePaolo B.D. 1985: Radiometric time scale for the upper
Eocene-Oligocene boundary based on K-Ar and Rb-Sr dating of
volcanic biotites from the pelagic sequence of Gubbio, Italy. Ge-
ology 13, 596—599.
Montanari A., Asaro F., Michel H.V. & Kennet J.P. 1993: Iridium
anomalies of late Eocene age at Messignano (Italy), and ODP
Site 689 (Maud Rise, Antarctica). Palaios 8, 420—437.
Murphy A., Sageman B.B. & Hollander D.J. 2000: Eutrophication by
decoupling of the marine biogeochemical cycles of C, N, and P:
a mechanism for the Late Devonian mass extinction. Geology
28, 5, 427—430.
Nagymarosy A. 1990: Paleogeographical and paleotectonical outlines
of some intra-Carpathian Paleogene basins. Geol. Zbor. Geol.
Carpath. 41, 3, 259—274.
Nagymarosy A. 1991: The response of the calcareous nannoplankton
to the Early Oligocene separation of the Paratethys. INA News-
letter, London 13, 2, 62—63.
Nagymarosy A. 2000: Lower Oligocene nannoplankton in anoxic de-
posits of the Central Paratethys. J. Nannoplankton Res. 22, 2,
Nagymarosy A. & Voronina A.A. 1992: Calcareous nannoplankton
from the Lower Maykopian Beds (Early Oligocene, Union of In-
dependent States). Knihovnička ZPN 14b, 2, 189—221.
Nebelsick J.H., Bassi D. & Drobne K. 2000: Microfacies analysis and
palaeoenvironmental interpretation of Lower Oligocene, shal-
low-water carbonates (Gornji Grad Beds, Slovenia). Facies 43,
Nijenhuis I.A. 1999: Geochemistry of eastern Mediterranean sedi-
mentary cycles. On the origin of Miocene to Pliocene sapropels,
laminites and diatomites. Geol. Ultraiectina 167, 9—168.
Norris R.D. & Nishi H. 2001: Evolutionary trends in coiling of
tropical Paleogene planktic foraminifera. Paleobiology 27, 2,
Oberhänsli H. 1996: Klimatische und ozeanographische Veränderun-
gen im Eozän. Z. Dtsch. Geol. Gesell. 147, 3, 303—413.
Oberhänsli H., Grünig A. & Herb R. 1984: Oxygen and carbon iso-
tope study in the Late Eocene sediments of Possagno (Northern
Italy). Riv. Ital. Paleont. Stratigr. 89, 3, 377—394.
Olszewska B. 1983: A contribution to the knowledge of planktonic
foraminifers of the Globigerina Submenilite Marls in the Polish
Outer Carpathians. Kwart. Geol. 27, 547—570.
Olzsewska B. 1998: The Oligocene of the Polish Carpathians. In:
Cicha I., Rögl F., Rupp Ch. & Čtyroká J. (Eds.): Oligocene—Mi-
ocene foraminifera of the central Paratethys. Abh. Senckenberg
Naturforsch. Gessell. 549, 23—28.
Olszewska B.W. & Wieczorek J. 1998: The Paleogene of the Podhale
Basin (Polish Inner Carpathians) – micropaleontological per-
spective. Przegl. Geol. 46, 8/2, 721—728.
Oschmann W. 1995: Black shales models: an actualistic approach.
EUROPAL 8, 26—35.
Oszczypko M. 1996: Calcareous nannoplankton of the Globigerina
Marls (Leluchów Marls member), Magura Nappe, West Car-
pathians. Ann. Soc. Geol. Pol. 66, 1—15.
Oszczypko-Clowes M. 1998: Late Eocene—Early Oligocene calcare-
ous nannoplankton and stable isotopes (
O) of the Glo-
bigerina Marls in the Magura Nappe (West Carpathians). Slovak
Geol. Mag. 4, 2, 107—120.
Öztürk H. & Frakes L.A. 1995: Sedimentation and diagenesis of an
Oligocene manganese deposit in a shallow subbasin of the Parat-
ethys: Thrace Basin, Turkey. Ore Geol. Rev. 10, 117—132.
Pak D. & Kennet J.P. 2002: A foraminiferal isotopic proxy upper wa-
ter mass stratification. J. Foram. Res. 32, 3, 319—327.
Pardo A., Keller G. & Oberhänsli H. 1999: Paleoecologic and pale-
oceanographic evolution of the Tethyan realm during the Pale-
ocene-Eocene transition. J. Foram. Res. 29, 1, 37—57.
Payros A., Orue-Etxebarria X. & Pujalte V. 2006: Covarying sedi-
mentary and biotic fluctuations in Lower-Middle Eocene Pyre-
nean deep-sea deposits: palaeoenvironmetal implications.
Palaeogeogr. Palaeoclimatol. Palaeoecol. 234, 258—276.
Pekar S.F., Christie-Blick N., Miller K.G. & Komiz M.A. 2003: Cali-
brating Oligocene eustasy to oxygen isotope data: eustatic ecti-
mates from two-dimensional flexural backstripping from the new
Jersey continental margin (USA). Geophys. Res., Abstr. 5, 13654.
Pierrard O., Robin E., Rocchia R. & Montanari A. 1998: Extraterres-
trial Ni-rich spinel in upper Eocene sediments from Massignano,
Italy. Geology 26, 307—310.
Poletti L., Premoli Silva I., Masetti D., Pipan M. & Claps M. 2004: Or-
bitally driven fertility cycles in the Palaeocene pelagic sequences
of the Sothern Alps (Northern Italy). Sed. Geol. 164, 35—54.
Pokorný V. 1992: General paleontology. University Carolinae, Pra-
gue, 9—296 (in Czech).
Pomerol Ch. & Premoli Silva I. 1986: The Eocene-Oligocene transi-
tion: events and boundary. In: Pomerol Ch. & Premoli Silva I.
(Eds.): Terminal Eocene events. Developments in Paleontol.,
Stratigr., 9, Elsevier, Amsterdam, Oxford, New York, 1—24.
Poore R.Z. & Matthews R.K. 1984: Oxygen isotope ranking of Late
Eocene and Oligocene planktonic foraminifers: implications for
Oligocene sea-surface temperatures and global-ice-volume.
Mar. Micropaleont. 9, 111—134.
Popov S.V. & Stolyarov A.S. 1996: Paleogeography and anoxic envi-
ronments of the Oligocene—Early Miocene Eastern Paratethys.
Isr. J. Earth Sci. 45, 161—167.
Popov S.V., Akhmetev M.A., Zaporozhets N.I., Voronina A.A. &
Stolyarov A.S. 1993: Evolution of the Eastern Paratethys in the
Late Eocene—Early Miocene. Stratigr. Geol. Corr., Moscow 1/6,
Popov S.V., Rögl F., Rozanov A.Y., Steininger F.F., Sherba I.G. &
Kováč M. (Eds.) 2004: Lithological-paleogeographic maps of
Paratethys. Cour. Forsch.-Inst. Senckenberg 250, 1—46.
Premoli Silva I. & Boersma A. 1988: Atlantic planktonic foramin-
iferal historical biogeography and paleohydrographic indices.
Palaeogeogr. Palaeoclimatol. Palaeoecol. 67, 315—356.
Pross J. & Schmiedl G. 2002: Early Oligocene dinoflagellate cysts
from the Upper Rhine Graben (SW Germany): paleoenvironmen-
tal and paleoclimatic implications. Mar. Micropaleont. 45, 1—24.
Prothero D.R. 1994: The Eocene-Oligocene transition: Paradise Lost.
Columbia University Press, New York,1—291.
Proust J.-N. & Hosu A. 1996: Sequence stratigraphy and Paleogene
tectonic evolution of the Transylvanian Basin (Romania, eastern
Europe). Sed. Geol. 105, 117—140.
Puglisi D., Badescu D., Carbone S., Corso S., Franchi R., Gigliuto
L.G., Loiacono F., Miclaus C. & Moretti E. 2006: Stratigraphy,
petrography and palaeogeographic significance of the Early Oli-
gocene “menilite facies” of the Tarcau Nappe (Eastern Car-
pathians, Romania). Acta Geol. Pol. 56, 1, 105—120.
Rasser M.W., Less G. & Báldi-Beke M. 1999: Biostratigraphy and
facies of the Late Eocene in the Upper Austrian Molasse Zone
with special reference to the Larger Foraminifera. Abh. Geol.
B.—A., Wien 56/2, 679—698.
Rospondek M.J., Köster J. & Sinninghe Damsté J.S. 1997: Novel C
highly branched isoprenoid thiophenes and alkane from the Me-
nilite Formation, Outer Carpathians, SE Poland. Org. Geochem.
56, 5/6, 295—304.
Roth Z. & Hanzlíková E. 1982: Palaeotectonic and palaeoecological
position of the Menilitic Formation in the Carpathian Mts. Čas.
Mineral. Geol. 27, 113—126.
Roy S. 2006: Sedimentary manganese metallogenesis in response to
the evolution of the Earth system. Earth Sci. Rev. 77, 273—305.
Rögl F. 1998: Palaeogeographic considerations for Mediterranean
and Paratethys Seaways (Oligocene—Miocene). Ann. Naturhist.
Mus. Wien 99A, 279—310.
Rögl F. 1999: Mediterranean and Paratethys. Facts and hypoteses of
an Oligocene to Miocene paleogeography. Geol. Carpathica 50,
Russu A. 1988: Oligocene events in Transylvania (Romania) and the
first separation of Paratethys. D.S. Inst. Geol. Geofiz. 72—73/5,
Russu A. 1995: Paleoclimatic meaning of Paleogene mollusca in NW
Transylvania. Rom. J. Paleont. 76, 47—52.
Salaj J. 1998: Reflection of paleoclimate in paleogene sediments of
the Middle Váh river valley. Zemní Plyn Nafta 42, 3, 171—187.
Salamy K.A. & Zachos J.C. 1999: Latest Eocene—Early Oligocene
climate change and Southern Ocean fertility: inferences from
sediment accumulation and stable isotope data. Palaeogeogr.
Palaeoclimatol. Palaeoecol. 145, 61—77.
Samuel O. 1975: Foraminifera of Upper Priabonian from ubietová
(Slovakia). Západ. Karpaty, Sér. Paleont., 1, GÚDŠ, Bratislava,
Samuel O. & Vaňová M. 1967: New occurrence about the stratigra-
PALEOENVIRONMENTAL CHANGES ACROSS EOCENE—OLIGOCENE BOUNDARY (CENTRAL CARPATHIANS)
phy of Eocene from Štúrovo. Geol. Práce, Zpr. 41, 41—51.
Sarangi S., Sakar A., Sarin M.M., Bhattacharya S.K., Ebihara M. &
2001: Growth rate and life span of Eocene—Oligocene
Nummulites tests: inferences from Sr/Ca ratio. Terra Nova 13,
Schackleton N.J. & Kennet J.P. 1975: Paleotemperature history of the
Cenozoic and the initiation of Antarcic glaciation: oxygen and
cabon isotope analyses in D.S.D.P. sites 272, 279 and 281. Init.
Reps. DSDP 29, 743—755.
Schlager W. 2003: Benthic carbonate factories of the Phanerozoic.
Int. J. Earth Sci. 92, 445—464.
Schmiedl G., Scherbacher M., Bruch A.A., Jelen B., Nebelsick J.H.,
Hemleben Ch., Mosbrugger V. & Rifelj H. 2002: Paleoenvi-
ronmental evolution of the Paratethys in the Slovenian Basin
during the Late Paleogene. Int. J. Earth Sci. (Geol. Rundsch.)
Schulz H.-M. 2003: The western Central-Paratethys at the Eocene/
Oligocene boundary – oceanography of a marginal sea and for-
mation of hydrocarbon source rocks. Clausthaler Geowissen-
schaften 3, 1—130.
Schulz H.-M., Sachsenhofer R.F., Bechtel A., Polesný H. & Wagner
L. 2002: The origin of hydrocarbon source rocks in the Austrian
Molasse Basin (Eocene-Oligocene transition). Mar. Petrol.
Geol. 19, 683—709.
Schulz H.-M., Bechtel A., Rainer T., Sachsenhofer R.F. & Struck U.
2004: Paleogeography of the Western Central Paratethys during
Early Oligocene Nannoplankton Zone NP 23 in the Austrian
molasse basin. Geol. Carpathica 55, 4, 311—323.
Schulz H.-M., Bechtel A. & Sachsenhofer R.F. 2005: The birth of the
Paratethys during the Early Oligocene: from Tethys to an an-
cient Black Sea analogue? Global and Planetary Change 49,
Seneš J. & Marinescu F. 1974: Cartes paléogéographique du Néogene
de la Paratethys centrale. Mem. BRGM 78, 785—792.
Snopková P. 1980: Paleogene sporomorphs from the West Car-
pathians. Západ. Karpaty, Sér. Paleont. 5, 7—74.
Sochaczewski A. 2000: Tuffites layers between Upper Eocene con-
glomerates of the Tatra Mountians (Inner Carpathians, Poland)
and their stratigraphic significance. Vijesti 37/3, 116—117.
Soták J. 1998a: Sequence stratigraphy approach of the Central Car-
pathian Paleogene (Eastern Slovakia): eustasy and tectonics as
controls of deep-sea fan deposition. Slovak Geol. Mag. 4, 3,
Soták J. 1998b: Central Carpathian Paleogene and its constrains. Slo-
vak Geol. Mag. 4, 3, 203—211.
Soták J. 2007: Biostratigraphic subdivision of the Eocene and Early
Oligocene formations of the Central Western Carpathians on
the basis of planktonic foraminifera: index forms, biozones
and stage definitions. In: Zlinská A. (Ed.): Proceedings of the
8. Paleontological Conference. Konferencie-Sympóziá-Semi-
náre, ŠGÚDŠ, Bratislava, 86—87.
Soták J., Pereszlényi M., Marschalko R., Milička J. & Starek D.
2001: Sedimentology and hydrocarbon habitat of the submarine-
fan deposits of the Central Carpathian Paleogene Basin (NE Slo-
vakia). Mar. Petrol. Geol. 18, 87—114.
Soták J., Plašienka D. & Vojtko R. 2004: Paleogene sediments of Ve-
poric zone: biostratigraphic data from a new occurrences NNW
of Tisovec. In: Zlinská A. (Ed.): Conferences, Symposia & Sem-
inars, D. Štúr Inst., Bratislava, 99—100.
Soták J., Gedl P., Banská M. & Starek D. 2007: New stratigraphic data
from the Paleogene formations of the Central Western Car-
pathians at the Orava region: results of integrated micropaleonto-
logical study in the Pucov section. Miner. Slovaca 39, 2, 89—106.
Spezzaferri S. 1995: Planktonic foraminiferal paleoclimatic implica-
tions across the Oligocene-Miocene transition in the oceanic
record (Atlantic, Indian and South Pacific). Palaeogeogr. Palae-
oclimatol. Palaeoecol. 114, 43—74.
Spezzaferri S. & Premoli Silva I. 1991: Oligocene planktonic fora-
miniferal biostratigraphy and palaeoclimatic interpretation from
Hole 538, DSDP Leg 77, Gulf of Mexico. Palaeogeogr. Palaeo-
climatol. Palaeoecol. 83, 217—263.
Spezzaferri S., Basso D. & Coccioni R. 2002: Late Eocene planktonic
foraminiferal responce to an extraterrestrial impact at Massig-
nano GSSP (Northeastern Appenines, Italy). J. Foram. Res. 32,
Starek D. 2001: Sedimentology and paleodynamics of the Paleogene
formations of the Central Western Carpathians. Thesis, Slovak
Acad. Sci., Bratislava, 1—152.
Starek D., Andreyeva-Grigorovich A.S. & Soták J. 2000: Suprafan
deposits of the Biely Potok Fm., in the Orava region: sedimenta-
ry facies and nannoplankton distribution. Slovak Geol. Mag. 6,
Steininger F.F. & Wessely G. 2000: From the Tethys Ocean to the
Paratethys Sea: Oligocene to Neogene stratigraphy, paleogeog-
raphy and paleobiogeography of the circum-Mediterranean re-
gion and the Oligocene to Neogene Basin evolution in Austria.
Mitt. Österr. Geol. Gesell. 92, 95—116.
Stolyarov A.S. & Kochenov A.B. 1995: Metalliferous Majkopian de-
posits of the Mangyshlak. Lithol. Mineral. Resour. 2, 161—172.
Surlyk F. 1997: A cool-water carbonate ramp with bryozoan mound:
Late Cretaceous—Danian of the Danish basin. In: James N.P. &
Clarke J.D.A. (Eds.): Cool-water carbonates. SEPM Spec. Publ.
Sztanó O. 1995: Palaeogeographic significance of tidal deposits: an
example from an early Miocene Paratethys embayment, North-
ern Hungary. Palaeogeogr. Palaeoclimatol. Palaeoecol. 113,
Sztrákos K. 1987: Les Foraminiferes bartoniens et priaboniens des
couches a “Tritaxia szaboi” de Hongrie et essai de reconstitution
paléogéographique de la Montagne Centrale de Hongrie au Bar-
tonien et au Priabonien. Cahiers de Micropaleontologie 2, 5—25.
Švábenická L., Bubík M. & Stráník Z. 2007: Biostratigraphy and
paleoenvironmental changes on the transition from the Meni-
lite to Krosno lithofacies (Western Carpathians, Czech Repub-
lic). Geol. Carpathica 58, 3, 237—262.
Tari G., Báldi T. & Báldi-Beke M. 1993: Paleogene retroarc flexural
basin beneath the Neogene Pannonian Basin: a geodynamic
model. Tectonophysics 226, 433—455.
Tari V. & Pamić J. 1998: Geodynamic evolution of the northern Di-
narides and the southern part of the Pannonian Basin. Tectono-
physics 297, 269—281.
Thunell R.C. & Corliss B.H. 1986: Late Eocene—Early Oligocene
carbonate sedimentation in the deep sea. In: Pomerol Ch. &
Premoli-Silva (Eds.): Terminal Eocene events. Developments
in Paleont. Stratigr. 9, 363—380.
Tripati A., Backman J., Elderfield H. & Ferretti P. 2005: Eocene bi-
polar glaciation associated with global carbon cycle changes.
Nature 436, 341—346.
Trümpy R. 1980: Geology of Switzerland. Part A. Wepf Publishers,
Uhlík P., Biroň A., Šucha V., Andrejeva-Grigorovič A., Clauer N. &
Halásová E. 2002: Illite-smectite of bentonite beds from Central
Carpathian Paleogene Basin. Miner. Slovaca 34, 2, 85—92.
Vail P.R., Mitchum R.M. Jr., Todd R.G., Widmier J.J., Thompson S.,
Sangree J.B., Bubb J.N. & Hatelid W.G. 1977: Seismic stratig-
raphy and global changes of sea level. In: Payton C.E. (Ed.):
Seismic stratigraphy – applications to hydrocarbon exploration.
Amer. Assoc. Petrol. Geol., Mem. 26, 49—212.
Van Breugel Y. 2006: Causes for negative carbon isotope anomalies
in Mesozoic marine sediments: constraints from modern and an-
cient anoxic setting. Geol. Ultraiectina 258, 9—119.
Van Breugel Y., Schouten S., Paetzel M., Ossebaar J. & Sinninghe
Damsté J.S. 2005: Reconstruction of
C of chemocline CO
(aq) in past oceans and lakes using the
C of fossil isorenier-
atene. Earth Planet. Sci. Let. 235, 421—434.
Van Couvering J.A., Aubry M.-P., Berggren W.A., Bujak C.W.,
Naeser C.W. & Wieser T. 1981: The Terminal Eocene Event
and the Polish connection. Palaeogeogr. Palaeoclimatol. Palae-
oecol. 36, 321—362.
Van Eijden A.J.M. 1995: Morphology and relative frequency of
planktic foraminiferal species in relation to oxygen isotopically
inferred depth habitats. Palaeogeogr. Palaeoclimatol. Palaeo-
ecol. 113, 267—301.
Van Simaeys S., De Man E., Vandenberghe N., Brinkhuis H. &
Steurbaut E. 2004: Stratigraphic and palaeoenvironmental analy-
sis of the Rupelian—Chattian transition in type region: evidence
from dinoflagellate cysts, foraminifera and calcareous nannofos-
sils. Palaeogeogr. Palaeoclimatol. Palaeoecol. 208, 31—58.
Vaňová M. 1975: Lepidocyclina and Miogypsina from the Faciotype
localities Budikovany and Bretka (South Slovakia). In: Báldi T.
& Seneš J. (Eds.): Chronostratigraphie und Neostratotypen—Mi-
ozän der Zentralen Paratethys. Bd. V. OM Egerien. Veda, Bra-
Varentsov I.M. 2002: Genesis of the Eastern Paratethys manganese
ore giants: impact of events at the Eocene/Oligocene boundary.
Ore Geol. Rev. 20, 65—82.
Vass D. 2003: The Šahy Antiforms and its role in the tectonics and
paleogeography of the Hungarian Paleogene Basin and the No-
vohrad/Nógrad Basin (Souther Slovakia and Northern Hunga-
ry). Acta Geol. Hung. 46, 3, 269—289.
Vass D., Šutovská K., Karoli S. & Janočko J. 1993: Biely Potok For-
mation of Central Carpathian Paleogene in Prešov Basin. Geol.
Práce, Spr. 97, 79—88.
Vass D., Túnyi I. & Márton E. 1996: Young Tertiary rotation of the
megaunit Pelso and neighbour units of the Western Carpathians.
Slovak Geol. Mag. 3—4, 363—367.
Vetö I. 1987: An Oligocene sink for organic carbon: upwelling in
the Paratethys? Palaeogeogr. Palaeoclimatol. Palaeoecol. 60,
Vetö I., Ozsvárt P., Futó I. & Hetényi M. 2007: Extension of carbon
flux estimation to oxic sediments based on sulphur geochemistry
and analysis of benthic foraminiferal assemblages: A case histo-
ry from the Eocene of Hungary. Palaeogeogr. Palaeoclimatol.
Palaeoecol. 248, 119—144.
Volfová J. 1962: Makrofauna aus dem Zentralkarpatischen Paläogen
der Ostslowakei. Geol. Práce, Zoš. 63, 93—97.
Volfová J. 1964a: Annual report about macropaleontological investi-
gations in the Spišská Nová Ves region. Manuscript, Geofond,
Volfová J. 1964b: Macrofauna of the Central Carpathian Paleogene
in the Banská Bystrica and Brezno regions. Zpr. Geol. Výsk.
v r. 1963, Bratislava 2, 121—122.
Vonhof H., Smith J., Brikhuis H., Montanari A. & Nederbragt A.
2000: Global cooling accelerated by early late Eocene impact?
Geology 28, 8, 687—690.
Wade B.S. & Pälike H. 2004: Oligocene climate dynamics. Pale-
oceanography 19, 1—16.
Wade B.S. & Pearson P.N. 2008: Planktonic foraminiferal turnover,
diversity fluctuations and geochemical signals across the
Eocene/Oligocene boundary in Tanzania. Mar. Micropaleont.
Wade B.S., Berggren W.A. & Olsson R.K. 2007: The biostratigraphy
and paleobiology of Oligocene planktonic foraminifera from the
equatorial Pacific Ocean (ODP Site 1218). Mar. Micropaleont.
Wei W. 1994: How many impact generated microsherule layers in the
upper Eocene? Palaeogeogr. Palaeoclimatol. Palaeoecol. 114,
Westwalewicz-Mogilska E. 1986: A new view on the genesis of the
Podhale flysch deposits. Przegl. Geol. 12, 690—698.
Wieser T. 1985: The teschenite formation and other evidence of mag-
matic activity in the Polish Flysch Carpathians and their geotec-
tonic and stratigraphic significance. In: Wieser T. (Ed.):
Fundamental researches in the western part of the Polish Car-
pathians. CBGA XIII, Krakow, 23—36.
Wilkin R.T., Arthur M.A. & Dean W.E. 1997: History of water-col-
umn anoxia in the Black Sea indicated by pyrite framboid size
distributions. Earth Planet. Sci. Let. 148, 517—525.
Wright V.P. & Burchette T.P. 1998: Carbonate ramps: an introduc-
tion. In: Wright V.P. & Burchette T.P. (Eds.): Carbonate ramps.
Geol. Soc., Spec. Publ. 149, 1—5.
Zachos J.C., Lohmann K.C., Walker J.C.G. & Wise S.W. 1993:
Abrupt climate change and transient climates during the Paleo-
gene: a marine perspective. J. Geol. 101, 191—223.
Zachos J.C., Quinn T.M. & Salamy K.A. 1996: High-resolution (104
years) deep-sea foraminiferal stable isotope records of the
Eocene—Oligocene climate transition. Paleoceanography 11, 3,
Zágoršek K. 1992: Priabonian (Late Eocene) Cyclostomata Bryozoa
from the Western Carpathians (Czechoslovakia). Geol. Car-
pathica 43, 4, 235—247.
Zágoršek K. 2000: Eocene bathymetry of Slovakia based on Bryozoa
growth forms. Miner. Slovaca 32, 2, 89—98.
Zlinská A., Andrejeva-Grigorovič A. & Filo I. 2001: Biostratigraphic
analysis of samples from outcrops near ubietová. Geol. Práce,
Spr. 105, 71—76.