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,  ,  , 

doi: 10.1515/geoca-2016-0020

Stratigraphy, plankton communities, and magnetic proxies 






























Earth Science Institute of the Slovak Academy of Sciences, Dúbravská cesta 9, P.O. Box 106, 840 05 Bratislava, Slovakia; 


Comenius University, Faculty of Science, De t. of Geology and Palaeontology, Ilkovi ova 6, 842 15 Bratislava, Slovakia;; 


Polish Geological Institute – National Research Institute, Rakowiecka 4, 00-975 Warsaw, Poland;; 


Comenius University, Faculty of Science, Dept. of Economic Geology, Ilkovi ova 6, 842 15 Bratislava, Slovakia; 


Charles University in Prague, Faculty of Science, Institute of Geology and Palaeontology, Albertov 6, 128 43 Prague, Czech Republic;


Czech Academy of Sciences, Institute of Geology, Rozvojová 269, 165 00 Prague 6, Czech Republic;;

(Manuscript received December 9, 2015; accepted in revised form June 7, 2016)

A well preserved Upper Tithonian–Lower Berriasian Strapkova sequence of hemipelagic limestones 

improves our understanding of environmental changes occurring at the Jurassic/Cretaceous boundary in the Western 
Carpathians. Three dino agellate and four calpionellid zones have been recognized in the section. The onset of the 
Alpina Subzone of the standard Calpionella Zone, used as a marker of the Jurassic/Cretaceous boundary is de ned by 
morphological change of Calpionella alpina tests. Calpionellids and calci ed radiolarians numerically dominate in 
microplankton assemblages. The  rst occurrence of Nannoconus wintereri indicates the beginning of the nannofossil 
zone NJT 17b Subzone. The FO of Nannoconus steinmannii minor was documented in the lowermost part of the Alpina 
Subzone. This co-occurrence of calpionellid and nannoplankton events along the J/K boundary transition  is typical of 
other Tethyan sections. Correlation of calcareous microplankton, of stable isotopes (C, O), and TOC/CaCO


 data distri-

bution was used in the characterization of the J/K boundary interval. 


C values (from +1.09 to 1.44 ‰ VPDB) do not 

show any temporal trends and thus show a relatively balanced carbon-cycle regime in sea water across the Jurassic/
Cretaceous boundary. The presence of radiolarian laminites, interpreted as contourites, and relatively high levels of 
bioturbation in the Berriasian prove oxygenation events of bottom waters. The lower part of the Crassicolaria Zone (up 
to the middle part of the Intermedia Subzone) correlates with the M19r magnetozone. The M19n magnetozone includes 
not only the upper part of the Crassicollaria Zone and lower part of the Alpina Subzone but also the FO of Nannoconus 
 and Nannoconus steinmannii minor. The reverse Brodno magnetosubzone (M19n1r) was identi ed in the 
uppermost part of M19n. The top of M18r and M18n magnetozones are located in the upper part of the Alpina  Subzone 
and in the middle part of the Ferasini Subzone, respectively. The Ferasini/Elliptica subzonal boundary is located in the 
lowermost part of the M17r magnetozone. A little bit higher in the M17r magnetozone the FO of  Nannoconus steinman-
nii steinmannii
 was identi ed.   

 J/K boundary, pelagic limestones, microfauna, nannoplankton, stable C and O isotopes, magnetic 

 susceptibility,  northern  Tethys.


Collection of sedimentological, geochemical and palaeonto-
logical data from complete stratigraphic sections, which can 
be used for correlation among candidate stratotypes of stage 
boundaries, is one of major goals of the Berriasian Interna-
tional Commission on Stratigraphy (ICS) program. A net-
work of regional stratotypes can provide a continuous record 
of both sedimentation and biotic events across the Jurassic/
Cretaceous boundary, and a precise evaluation of all proxies 
necessary for exact discrimination of the boundary position. 
In the Western Carpathians, the Brodno section (Michalík et 
al. 1990, 2009; Houša et al. 1996, 1999) represents the regional 
stratotype of the J/K boundary. However, ammonites are rare 
and sediment thickness is somewhat reduced in the Brodno 

section, and complementary J/K boundary sections were thus 
recently studied in the Western Carpathian (Fig. 1): 


the Strá ovce section (Borza 1984; Michalík et al. 1990);  
the Hlbo a section (Grabowski et al. 2010) and Po rednie 
sections in the Tatra Mts. (Grabowski   Pszcz kowski 
2006; Grabowski et al. 2013).

Remarkable advances in calpionellid and nannoplankton 

biostratigraphy across the J/K boundary interval have been 
published on the basis of Tethyan Jurassic/Cretaceous boun-
dary (JKB) sections (Lukeneder et al. 2010, 2015; Wimble-
don et al. 2013; Svobodová   Koš ák 2016). An opening of 
the Tethyan/Panthalassa passage between Gondwana and 
North America enabled phyletic evolution of small plank-
tonic protozoans and autotrophic algae in a renewed circum-
equatorial oceanic current. This evolution led to a high 

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, 2016, 67, 4, 303–328

number of bioevents useful for global correlation of pelagic 
carbonate sequences. In spite of their broad spatial extent, 
these events differed in details due to regional palaeoenvi-
ronmental changes (Michalík & Reháková 2011). We think 
that the boundary level should be situated within a bundle of 
such events, allowing good correlations in the absence of the 
primary ammonite markers. 

The warming in combination with eustatic oscillations 

could result in diverse changes of the fauna in the Panboreal 
Realm (Wimbledon et al. 2013; Zakharov et al. 2014). During 
prominent sea-level rise, connection of the Boreal sea with 
the Panthalassa Ocean opened as indicated by occurrences of 
Middle Volgian ammonites with Pacific affinity. Disturbance 
of the marine ecosystem indicated by green algae blooms 
correlates with a negative excursion of C


 isotope near the 

Volgian/Ryazanian boundary. The stratigraphic correlation is 
difficult between the Boreal and Tethyan bioprovinces 
because they underwent different evolutionary pathways, the 
only connecting link seems to be the magnetostratigraphy 
(Houša et al. 2007; Grabowski 2011; Schnabl et. al. 2015). 

Our paper discusses the results of an integrated biostrati-

graphic study using three microplankton groups (calpionel-
lids, calcareous dinoflagellates and nannofossils), stable iso-
tope data (




O), microfacies and sequence stratigraphy, 

as well as the study of magnetic record in the Strapkova sec-
tion, which is regarded here as an auxiliary West Carpathian 

regional JKB section. The 
distribution of the strati-
graphi cally-important  plank-
tonic organisms revealed 

veral coeval calpionellid 

and nannofossil bioevents 
recorded in the pelagic car-
bonate sequence of the JKB. 
The bioevents can be inte-
grated with magnetostratigra-
phy. In addition, magnetic 
susceptibility helps us to 
interpret early depositional 
history of the sediment, 
namely the amount of supply 
of fine-grained terrigenous 
material to a basin. 

Location of the studied 


An important section 

exposing the JKB sequence in 
the western sector of the Pie-
niny Klippen Belt (Western 
Carpathians, Slovakia) is 
named the Strapkova section 
(Fig. 1). It can be well cor-
related with the principal 

Brodno section that is located about 40 km NE of the Strap-
kova section. Two additional JKB sections studied in detail 
(Strá ovce and Hlbo a) are located in the Krí na Nappe of 
the Central Western Carpathians (Michalík et al. 1995; 
Grabowski et al. 2010). The Strapkova section 
(49°04’09.34”N; 18°10’00.85”E; 589 m a.s.l.) is exposed on 
a steep SE slope of the Strapkova hill below the Mount 
Vršatec (Biele Karpaty Mountains, Fig. 1). It is located below 
a local road leading from Vršateck  Podhradie to  erven  
Kame , approximately 1250 m NE from the Vršateck  
Podhradie village, westwards of the middle Váh Valley. The 
Brodno section (49°16’02.16”N; 18°45’12.16”E; 353 m a.s.l.) 
has been described by Houša et al. (1996), Michalík et al. 
(1990, 2009) as the parastratotype section of the JKB in the 
Western Carpathians. It is situated in an abandoned quarry 
north of  ilina town on the eastern side of the narrow straits 
of the Kysuca River Valley (known as the “Kysuca Gate”). 

Geological setting

The tectonic contact of two principal superunits of the 

Western Carpathians (the Outer and Central Carpathians) is 
rimmed by the Pieniny Klippen Belt (Fig. 1). This unit, ori-
ginally rimming the European shelf, is typical of tensional 
basins-and-ridges development from the Early Jurassic until 

Fig. 1. A — Situation sketch of the studied area (gray ellipse) in the frame of Slovakia (arrow indicates 
the Hlbo a section); B — Situation sketch of the Middle Váh Valley section of the Pieniny Klippen 
Belt with indication of the Brodno and Strapkova sections (arrows).    — Simpli ed geological sketch 
of the Pieniny Klippen Belt area between  erven  Kame , Prusk  and Krivoklát villages.

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, 2016, 67, 4, 303–328

the Palaeogene, not interrupted by Palaeoalpine tectonic 
movements, when the superficial nappe structure of the Cen-
tral Carpathians originated. Its typical klippen-style tectonic 
contrasts with a slight diagenetic transformation of its rock 
sequences. In contrast to the Pieniny Klippen Belt, sedimen-
tary rocks of Central Carpathian sequences are more strongly 
affected both by diagenesis and tectonic stress. This weak 
degree of diagenetic transformation thus favours a complex 
study of the Jurassic and Cretaceous sedimentary sections in 
the Pieniny Klippen Belt. They formed in two parallel shal-
low, but considerably subsiding marine basins separated by 
the Czorsztyn Ridge. The Strapkova and Brodno sections 
formed in more distal areas of the Kysuca Basin, in the 
neighbourhood of the Penninic rift system, which was gra-
dually invaded by the Mid Atlantic-Penninic Ocean arm 
during the Jurassic and Cretaceous (Michalík 1994; Plašienka 

Hemipelagic succession of the Strapkova section 

(attributed to the Orava Unit by Haško 1978 and Schlögl et 
al. 2000; Fig. 2) starts with the Lower Jurassic spotted lime-
stones. An ammonite fauna indicates the Late Sinemurian 
Raricostatum Subzone. The Kozinec Formation composed of 
red pseudonodular limestones alternating with greenish-grey 
marly limestones contains ammonites spanning the latest 
Sinemurian Macdonelli Subzone of the Raricostatum Zone 
up to the Early Pliensbachian Davoei Zone. Grey-greenish 
finely bedded limestones and yellow-grey marly shales con-
tain abundant ammonites of the Pliensbachian and Toarcian 
Margaritatus and Spinatum zones. Red nodular limestones 
following upwards are considered Toarcian in age. Well- 
bedded cherty spiculitic limestones of the Podzamcze Lime-
stone Formation contain isolated beds of crinoidal pack-
stones, capped by a 1 m thick interval of red nodular lime-
stone (Fig. 2). The Czajakowa Radiolarite Formation is built 
of red radiolarites (1 m thick) with Middle Oxfordian radio-
larians, a thick (1.5 m) layer of pink limestone rich in belem-
nite rostra and “upper” red and green radiolarites with Kim-
meridgian Saccocoma packstones in its upper part. Radiola-
rites pass gradually into thin bedded red cherty and nodular 
limestones intercalated by red marlstones. The marlstones 
are followed by the Czorsztyn and the Pieniny limestone for-
mations, which formed the subject of our study. 


Microfossil and microfacies study

88 limestone beds have been sampled for thin sections. 

Beds have been numbered (mostly at 1 metre intervals) by 
numbers from 279 to 382 in accordance with former sam-
pling (Schlögl 2001) of the entire section. The interval 301 to 
324 has been omitted due to uncertainties connected with 
slump deformation. According to new analyses, it seems that 
distortion of original thickness is not significant. More 
densely sampled intervals (around the JKB) have been 

designated according to a decimal system. Pure limestones 
without cherts and silicified parts were selected for thin- 
sections. The majority of samples have been analysed also 
for stable isotopes (C, O), carbonate and TOC content. 

A set of 110 samples was used for microfacies analyses in 

order to document the succession of calpionellids and calca-
reous dinoflagellates. The thin-section samples were studied 
under the LEICA DM 2500 transmitting light microscope 
and the percentages of selected allochems and bioclasts 
(quartz and lithoclasts, calpionellids, radiolarians, globo-
chaetes, saccocomids, filaments, clasts of benthic organisms) 
were calculated. The quantitative evaluation with the optical 
charts sensu Bacelle & Bosellini (1965) was used. Microfa-
cies and biostratigraphically-important microfossils were 
documented using a LEICA DFC 290 HD camera. Thin- 
sections are stored in the collections of the Earth Science 
Institute of the Slovak Academy of Sciences and in the col-
lections of the Department of Geology and Palaeontology 
(Faculty of Natural Sciences), both in Bratislava. 

Calcareous nannofossils were analysed in 99 smear slides 

prepared using technique reported by Švábenická (2012) — 
decantation method and 7 % solution of H




. To obtain the 

relative sample abundances and semi-quantitative informa-
tion about nannofossil species, all the specimens in at least 
300 fields of view were counted in each slide. The smear 
slides were examined under the Olympus BX51 transmitting 
light microscope using an immersion objective of ×100 
magnifications. Calcareous nannofossils were documented 
with the Olympus DP70 digital camera. The set of smear 
slides is stored at the Department of Geology and Palaeonto-
logy, Faculty of Science (Charles University, Prague) and at 
the Institute of Geology of the CAS in Prague.


Stable isotopes (C, O) and total carbon (TOC) analyses 

were carried out on bulk rock carbonate samples. 64 samples 
were selected to C and O isotope analyses: 31 samples were 
selected in the uppermost Jurassic to the JKB interval (281–
300 M) and the next 33 samples were taken in the Berriasian 
(325–360 M)  interval. 


C and 


O were analysed in CO



after standard decay of bulk rock samples in 100 % phos-
phoric acid. Analyses of carbonate samples were done in 
labo ratories of the Czech Geological Institute in Prague on 
the Finigan MAT-2 Mass Spectrometer and in the Earth 
 Science Institute of the Slovak Academy of Sciences in 
Banská Bystrica on the MAT253 Mass Spectrometer 
equipped with the Gasbench device (Thermo Scientific Sam-
ples). Results are introduced in standard del-notation ( ) in 
promile (‰) being related to the Vienna Pee Bee Belemnite 
(VPDB) standard with 0.01 ‰ accuracy. 

The palaeotemperature calculation from calcite oxygen 

isotopes (Anderson & Arthur 1983) is as follows: 
T(°C) 16.0–4.14(








, where 




O of 

calcite of samples  in ‰ (V-PDB) and 




O of sea water. 

According to Gröcke et al. (2003), the value –1.0 ‰ 

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, 2016, 67, 4, 303–328

Fig. 2

. Correlation of lithostratigraphy

, microfacies , magnetic susceptibility of the Strapkova section sequence. 


 — general lithostratigraphy; 


 — sedimentary rate; 

 — lithological column 

of the sequence studied; 


 — quantitative representation of allochems in microfacies; 


 — magnetic susceptibility


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, 2016, 67, 4, 303–328

(V-SMOW) is characteristic of the post-Jurassic ice-free 
world. The contents of total organic carbon (TOC) and total 
inorganic carbon (TIC) were detected in 50 bulk rock sam-
ples on the C-MAT 5500 device of the Ströhlein Firm in the 
Earth Science Institute of the Slovak Academy of Sciences in 
Banská Bystrica. The TIC content was re-calculated on 



Palaeomagnetism and magnetic properties 

65 stratigraphic layers were sampled for a magnetostrati-

graphic study between beds 292 and 364. Samples were 
taken either with gasoline or electrically powered drills. 
Sampling resolution was usually 0.5–1 m. Standard cylindri-
cal specimens 2.2 cm high and 2.5 cm in diameter were pre-
pared from drill cores. Usually, at least two twin specimens 
were obtained from each drill core. The first specimen was 
subjected to thermal demagnetization and palaeomagnetic 
analysis, and the second one was used for rock magnetic 
analysis. Palaeomagnetic experiments were performed in the 
Palaeomagnetic Laboratories of the Polish Geological Insti-
tute-NRI in Warsaw and Institute of Geology, Czech Aca-
demy of Sciences in Prague (Pr honice). In the Palaeomag-
netic Laboratory of the PGI-NRI natural remanent magneti-
zation (NRM) was measured with the JR6a spinner magne-
tometer. Specimens were demagnetized exclusively by the 
thermal method using the MMTD1 non-magnetic oven 
(Magnetic Measurements, UK, rest field <10 nT). Magnetic 
susceptibility was monitored with a KLY-2 bridge (AGICO, 
Brno; sensitivity 10


 SI) after each thermal demagnetization 

step. NRM measurements and demagnetization experiments 
were carried out in the magnetically shielded space (a low-
field cage, Magnetic Measurements, UK, which reduces the 
ambient geomagnetic field by about 95 %). The thermal 
demagnetization was also performed in the Institute of Geo-
logy CAS using the MAVACS apparatus, the NRM was mea-
sured on the SQUID magnetometer 2G enterprises 755 4K 
SRM with shielded entrance. The magnetic susceptibility 
was measured on KLF4 magnetic susceptibility meter from 
AGICO. The measured data from both laboratories are fully 

Results of measurements were further processed using the 

Remasoft software (Chadima & Hrouda 2006). Rock mag-
netic investigations were performed in the palaeomagnetic 
laboratory in Prague. They comprised mass-normalized mea-
surements of the MS and isothermal remanent magnetization 
(IRM). The IRM was applied along the Z axis in the field of 
1T, and then antiparallel in the field of 100mT (using 
MMPM10 pulse magnetizer). The S-ratio (IRM




calculated as ratio of IRM intensities applied in both fields 
was indicative for proportions of low and high coercivity 
minerals. In samples from selected beds, a stepwise acquisi-
tion of the IRM (in the maximum field of 1.4T) was per-
formed, followed by thermal demagnetization of three axes 
IRM acquired in the fields of 1.4T, 0.4T and 0.1T (Lowrie 
1990) in order to identify magnetic minerals.


Sedimentology and microfacies


 (ca. 11 m thick, 

samples 280 to 291, Fig. 2) is represented by red nodular 
limestones of the Ammonitico Rosso facies. Ammonites are 
affected by corrosion and dissolution and thus are very poorly 
preserved. The formation includes Saccocoma-filamentous 
wackestones to packstones, Saccocoma-Globochaete-fila-
mentous packstones, Saccocoma-radiolaria-Globochaete 
packstones,  Saccocoma-Globochaete-radiolaria packstones, 
Saccocoma-Globochaete wackestones to packstones (Fig. 3), 
and radiolarian wackestones. In addition to dominant bio-
clasts, they contain rare aptychi, crinoids (formed by twinned 
lamellar calcite), echinoids, juvenile ammonites, calcitic and 
agglutinated foraminifera, thick-walled bivalves, sponge 
spicules, the problematicum Gemeridella minuta, and calca-
reous dinoflagellate cysts. Dinocysts are represented by 
Cadosina parvula, Colomisphaera nagyi,  Stomiosphaera 
  Carpistomiosphaera borzai,  Colomisphaera 
,  Carpistomiosphaera tithonica (Fig. 4 A),  Colomi-
sphaera radiata 
(Fig. 4 A),  Colomisphaera  carpathica 
(Fig. 4 C),  Colomisphaera lapidosa,  Parastomiosphaera 
Cadosina semiradiata fusca (Fig. 4 D), and Cado-
sina semiradiata semiradiata
  (Fig. 4 E).  Radiolarians and 
spicules are partially or totally calcified. Nodules are rimmed 
by dense systems of stylolites. Silt-sized muscovite flakes 
and quartz grains are common locally, and concentrated in 
the dissolution zones of stylolites. The matrix contains scat-
tered pyrite aggregates. Saccocomas dominated in the Kim-
meridgian and Lower Tithonian associations (Figs. 2, 3). 
Since the Late Tithonian, they were gradually replaced by 
Globochaete alpina spores.


is formed by pale 

grey to white biomicritic limestones, with variable bed thick-
ness. Radiolaria-Globochaete-calpionellids, radiolarian- 
calpionellid-Globochaete,  Globochaete-Calpionella and 
calpionellid-Globochaete-nannofossil microfacies were 
identified. The abundance of bioclasts in micrite matrix 
 varies between wackestone to packstone (Fig. 5). The lime-
stones contain numerous calpionellids, foraminifera, Invo-
 sp. Lenticulina sp., benthic and planktonic crinoid seg-
ments (Saccocoma  sp.), echinoids, ophiuroids, bivalves, 
juvenile ammonites, aptychi, ostracods, sponge spicules, 
problematicum Didemnoides moreti, and Didemnum carpa-
Local silty quartz grains and scattered (also framboi-
dal) pyrite occur in the matrix (Fig. 3B–F). Layers con taining 
sedimentary breccia and synsedimentary slumps were 
observed locally.

Biomicrite wackestone of radiolarian-calpionellid micro-

facies (in 29 beds, samples from 298 to 339), contains almost 
the same bioclasts as those observed in the Crassicollaria 
Zone. Saccocomids disappeared (Fig. 5); crassicollarian lori-
cas are currently deformed. Small bioclasts are sometimes 
affected by silicification; phosphatization of fragments was 

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, 2016, 67, 4, 303–328

Fig. 3. Microfacies in the Strapkova section: A — Saccocoma packstone. Sample 287. B — Aptychi 
and crinoid fragments in Calpionella-Globochaete wackestone to packstone. Sample 296.3. 


 — Slightly bioturbated Calpionella-Globochaete packstone. Sample 296.3. D — Biomicrite radio-

laria-calpionellid and Calpionella-Globochaete wackestone. Samples 298; 298.5. E — Biomicrite 
radiolarian wackestone. Sample 300.  F. Biomicrite Calpionella wackestone. Sample 300.2.

Depositional environment 
and sequence architecture

The wackestones and pack-

stones of the Pieniny Lime-
stone Formation are formed 
mostly by tests of planktonic 
microorganisms, while mud-
stone micrites and biomicrites 
consist of nannoplankton 
remains and unidentified cal-
cite test fragments. Although 
stratigraphic differences in 
the rock composition and in 
the granulometry of the “bian-
cone” facies are not exten-
sive, eight 7–16 


cycles can be distinguished in 
the sequence (Fig. 2). Each of 
the cycle starts with pack-
stone beds containing infre-
quent remnants of benthic 
organisms, abundant (some-
times redeposited) tests of 
calpionellids, occasional small 
(eolian) grains of quartz and 
mica leaflets. These beds are 
comparable with the lowstand 
part of the cycle. Upwards, 
limestone beds are characteri-
zed by a higher content of cal-
careous dinocysts and calpio-
nellid tests. The highest part is 
richer in chert and frequently 
includes laminar concentra-
tions of (mostly calcified) 
radiolarian tests. These cycles 
correspond to the eustatic 
cycles (Ti3-Ti6 and Be1-Be4) 
figured in Haq (2014). 

The distribution of calpio-

nellids shows several abundance peaks (Figs. 2, 5). The first 
peak is located in the Late Tithonian, the second peak is 
located in the upper part of the Alpina Subzone, and the third 
peak is located below the onset of the Ferasini Subzone. 
Stratigraphic changes in abundance of calpionellids and radio-
larians (Figs. 2, 5) show that they alternate in discrete peaks: 
a decrease in abundance of  calcareous plankton is associa ted 
with an increase of abundance of silica-secreting organisms 
(Reháková & Michalík 1994; Michalík et al. 2009).



bioclasts are accumulated in thin laminae and small nests, 
some of bioclasts that are slightly phosphatized.

Radiolarians occur in very thin silicified laminae at multiple 

levels (I samples 315.15; 333; 339, 351.1; 381.9; 385.8; 
388.25; 388.6; 393.2; 394.3; 395.15) that can be interpreted 
as contourites (Schlögl et al. 2000). These laminites 

observed more rarely. Clastic admixture is represented by 
rare eolian silty quartz grains only (Fig. 2). 

Silty clastic admixture content (quartz and muscovite in 

samples 340 to 359) is low (Fig. 2). Some layers contain 
lami nae rich in bioclasts. Pyrite is scattered in matrix, it 
occurs as framboids or in nest accumulations. Several bio-
clasts, mainly radiolarians are impregnated by pyrite, some 
other bioclasts were phosphatized. 

The sampling of the sequence has been complicated in the 

two sections in which the stratal geometry is distorted by 
folds (between 300–325, and 347–346). Study of sedimen-
tology, and detailed biostratigraphy indicates that these 
 phenomena originated synsedimentary due to rock sliding. 
Their presence confirms the slope environment of 

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, 2016, 67, 4, 303–328

Fig. 4. Calcareous dinocysts in the Strapkova section: A — Carpistomiosphaera tithonica Nowak. 
Sample 287. B — Colomisphaera radiata (Vogler). Sample 287.   — Colomisphaera carpathica 
(Borza). Sample 287. D — Cadosina semiradiata fusca (Wanner). Sample 290. E — Cadosina semira-
diata semiradiata
 (Wanner). Sample 290. F, G — Colomisphaera fortis  ehánek. Samples 290; 298.3. 
H — Colomisphaera cieszynica Nowak. Sample 300.3. I — Colomisphaera lapidosa (Colom). Sam-
ple 300.3. J — Stomiosphaerina proxima  ehánek. Sample 300.3. K — Gemeridella minuta Borza et 
Mišík. Sample 300. L — Didemnum carpaticum Borza et Mišík. Sample 300.1.

represent a special feature of 
the Strapkova sequence (Fig. 6). 
The abundance of radiolarian 
tests is the highest in each 

lamina (1.9 to 2.2 mm 

thick) with a slightly erosive 
base. Similar limestone layers 
with radiolarian laminae 
occur in the Brodno (Michalík 
et al. 2009; bed C42) and 
Rochovica sections (in the lat-
ter case, they occur in much 
younger, Valanginian to Aptian 
strata; Michalík et al. 2008). 

The layer below the radio-

larian laminite is bioturbated 
(Fig. 6A–C). Traces of Chon-
,  Palaeophycus,  Plano-
,  Thalassinoides, and 
Trichichnus were identified in 
cross-sections perpendicular 
to the bedding plane. Primary 
sedimentary features (cross- 
bedding stratification, lamina-
tion) were mostly destroyed 
by bioturbation. The largest 
burrows (Thalassinoides) are 
on average 5–9 mm in diame-
ter, Planolites and Palaeophy-
 burrows attain diameters 
of 2 to 3 mm. Planolites and 
Thalassinoides burrows are 
pe netrated  by Chondrites (with 
diameters of 0.4 to 0.6 mm). 
Simple vertical pyritic bur-
rows of Trichichnus are 
0.2 mm in diameter. Framboi-
dal pyrite clusters co-occur in 
places with bioturbation 
structures. The size of bur-
rows, different ethological 
character (domichnia, fodi-
nichnia, chemichnia) and tro-
phic levels of these traces 
indicate that the bottom was 
well supported with nutrients 
and oxygen, and inhabited by 
burrowers at different sediment depths. 


Calpionellid and calcareous dinocyst zonations

Red nodular limestone of the Rosso Ammonitico facies 

(Czorsztyn Limestone Formation) is dated as late 

Kimmeridgian (Borzai and Pulla zones) to latest Early Titho-
nian Chitinoidella boneti Subzone (Jach et al. 2012). The 
succession of dinoflagellate bioevents allowed us to deter-
mine the following dinocyst zones: the Late Kimmeridgian 
— Borzai (sample 280) and Pulla zones (samples 281–283), 
and the Early Tithonian — Malmica Zone (samples 284 to 
287). After a thin transitional interval (earliest Late Tithonian 
Praetintinopsella Zone), the sequence continues with the 
Maiolica facies of the Pieniny Limestone Formation (Late 

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, 2016, 67, 4, 303–328

Fig. 5

. Quantitative distribution of microplankton: 


 — total plankton abundance; 


 — composition of plankton groups in total plankton contents.

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, 2016, 67, 4, 303–328

Tithonian to late Berriasian). Our study has been focused on 
the lower part of the Pieniny Formation, ca. 40 m thick (from 
the Late Tithonian Crassicollaria remanei Subzone to middle 
Berriasian Calpionella elliptica Subzone). Successive occur-
rence of biostratigraphically important calpionellids and cal-
careous dinoflagellates is shown in the Fig. 7. 


(sensu Grandesso 1977 and Borza 1984) 

Saccocoma-Globochaete-radiolaria packstones (samples 

288 to 289) contain rare Longicollaria  dobeni (Fig. 8A) 
 Carpathella rumanica,  Borziella slovenica, Dobeniella 
  Colomisphaera  carpathica,  Colom. lapidosa and 
Colom. tenuis.


(sensu Grandesso 1977 and Borza 1984) 

Saccocoma-Globochaete-radiolaria packstones and Sacco-

coma-Globochaete  wackestones to packstones (samples 
289.4–290) with Chitinoidella boneti (Fig. 8B), Chitin. elon-
(Fig. 8C), Longicollaria  dobeni,  Dobeniella cubensis
Popiella oblongata,  Colomisphaera  carpathica,  Colom. 
Colom. tenuis and Colom. fortis (Fig. 4F, G)

 (sensu Grandesso 1977) 

Radiolaria wackestones (sample 291) contain rare chiti-

noidellids, cysts of Colomisphaera  carpathica but also the 
first hyaline calpionellid form represented by Praetintinnop-
sella andrusovi 
(Fig. 8E).


(sensu Remane et al. 1986)

Radiolarian wackestone (sample 293) with very rare sec-

tions of microgranular chtitinoidellids contains Tintinopsella 
remanei (Fig. 8D), Calpionella alpina, Crassicollaria inter-
 and cysts of Colomisphaera carpathica.


(sensu Remane et al. 1986)

Calpionella-Globochaete locally slightly laminated wacke-

stones to packstones (8 beds from 294 to 294.6) with Crassi-
collaria intermedia 
(Fig. 8F), Crass. parvulaCrass. massu-
(Fig. 8G), Calpionella alpina,  Calp. grandalpina 
(Fig. 8K), Calp. elliptalpinaTintinnopsella carpathica, and 
cysts of Colomisphaera lapidosa, Colom. carpathica, 
Stomio sphaerina  proxima
, Cadosina semiradiata semiradiata, 
and Cadosina sp.


(sensu Reháková & Michalík 1997)

Radiolarian-Calpionella-Globochaete, Calpionella-Globo-

chaeteGlobochaete-Calpionella, locally slightly laminated 
and/or bioturbated wackestones (18 beds from 294.7 to 
296.2) contain Crassicollaria brevis (Fig. 8H), Crassicol-
laria parvula
,  Calpionella alpina;  Crassicollaria massuti-
,  Calpionella grandalpina, Tintinnopsella carpathica
cysts of Colomisphaera lapidosa, Colomisphaera carpathica, 
Stomiosphaerina proxima
, Cadosina semiradiata semiradiata 
and Cadosina semiradiata fusca are less abundant if com-
pared with the Intermedia Subzone. 


(sensu Reháková & Michalík 1997)

The FO of Crassicollaria colomi was identified in slightly 

bioturbated biomicrite (16 beds from 296.3 to 297) with 
Calpionella-Globochaete  (Fig. 3B, C), radiolarian-Calpio-
nella, radiolarian-Calpionella-Globochaete and radiolarian 
microfacies.  Crassicollaria parvula dominates over Crass. 
(Fig. 8I), Crass. brevis,  Crass. massutiniana, and 
Calpionella alpina (Fig. 8J), which prevails over Calpionella 
 and Tintinnopsella carpathicaColomisphaera 
lapidosa, Colom. carpathica
, Stomiosphaerina proxima, 
Cadosina semiradiata semiradiata
, Cados. semiradiata 
cysts were also identified. 

(sensu Pop 1974; Remane et al. 1986; 

Reháková & Michalík 1997; Lakova et al.1999; 

Boughdiri et al. 2006; Andreini et al. 2007; 

Lakova & Petrova 2013). 

Fig. 6. Bed 393.1 of the maiolica limestone  with accumulations of 
radiolarian tests arranged in laminae (A, the uppermost part). Under-
lying rock is penetrated by burrows of infaunal organisms (B) indi-
cating more intensive oxidation (B) of dysoxic sediment by bottom 
current.   — Another bed (351.1) of the maiolica limestone with 
radiolarian tests in distinct laminae deposited from a contour 

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, 2016, 67, 4, 303–328

Biomicrite wackestone is composed of radiolarian, radio-

larian-calpionellid and calpionellid microfacies (29 beds 
between 298–339; Fig. 3E, F). The sample 298 shows transi-
tion from the microfacies rich in Crassicollaria parvula to 
microfacies in which spherical forms of Calpionella alpina 
dominate. Thus, the J/K boundary interval according to 
Remane’s et al. (1986) definition is situated in bed 298. Four 
crassicollarian abundance events influenced by synsedimen-
tary erosion were documented (samples 298.1–298.4). Since 
the sample 298.6, Calpionella-Globochaete wackestones 
prevail (Fig. 3D), radiolarians appear in high portion in sam-
ples 299, 332 and 338. The dominant Calpionella alpina 
with rare Crassicollaria parvula and Tintinnopsella car-
with Tint. doliphormis create a calpionellid associa-
tion typical of the Alpina Subzone. Calpionellids are accom-
panied by rare to seldom cysts of Colomisphaera carpathica, 
Col. cieszynica 
(Fig. 4H), Col. lapidosa, (Fig. 4I), Col.  cf. 

fortis, Col. sp., Stomiosphaerina proxima (Fig. 4J) and 
 Cadosina semiradiata semiradiata, 
microproblematica of 
Gemeridella minuta (Fig. 4K) and Didemnum carpaticum 
(Fig. 4L).


(sensu Remane et al. 1986)

Biomicrites, locally slightly bioturbated with calpio nellid-

Globochaete, calpionellid-Globochaete-radiolarian and 
radio larian wackestones (studied in samples 340–343). In the 
calpionellid association Calpionella alpina dominated over 
the infrequent Remaniella ferasini, R. catalanoi,  R. duran-
delgai, R. borzai, Tintinopsella carpathica, Crassicollaria 
. The dinoflagellate cyst association consists of Colo-
misphaera lapidosa
,  Col. carpathica, Stomiosphaerina 
 Cadosina semiradiata fusca and Cad. semiradiata 

Fig. 7. Vertical range of calpionellid species and cysts of calcareous dino agellates.

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, 2016, 67, 4, 303–328


(sensu Pop 1974)

Biomicrites, locally bioturbated wackestones (samples 344–

359) with radiolarian-Calpionella-Globochaete and Calpionel-
la-Globochaete microfacies contain calpionellid associations 
of  Calpionella alpina, Calp. elliptica (Fig. 9B, C),  Calp. 
minuta, Tintinopsella carpathica 
(Fig. 9L), Tint. longa 
(Fig. 9A), Lorenziella hungarica (Fig. 9I), Remaniella catala-
noi, Rem. ferasini 
(Fig. 9D,  E), Rem. durandelgai (Fig. 9F), 

Rem. borzai  


(Fig. 9 H), Rem. 

colomi, Rem. filipescuiRem. 
(Fig. 9G).  Cysts 
are represented by Colomi-
sphaera lapidosa
, Colom. 
carpathica, Cadosina semira-
diata semiradiata 
and  Cad. 
semiradiata fusca

Calcareous nannofossils and 
nannofossil zonation

In the samples studied, cal-

careous nannofossils are 
rather rare and their preserva-
tion ranges from moderate 
(only in a few samples) to 
extremely poor, heavily etched 
by dissolution. In total, 29 
calcareous nannofossils taxa 
were indentified. A compara-
ble diversity has been 
observed in the Barlya section 
(Lakova et al. 1999) and in 
the Nutzhof (Reháková et al. 
2009). A slightly lower diver-
sity has been reported from 
the Brodno (Michalík et al. 
2009) and Hrušové sections 
(Ondrejí ková et al. 1993), 
conversely higher diversity 
and abundance also have been 
observed, for example, in the 
Puerto Escaño (Svobodová & 
Koš ák 2016) or Torre de 
Busi and Foza sections 
(Casellato 2010). Successive 
distribution of nannofossils 
along the lithological column 
is shown in the Fig. 



Watznaueria (more than 55 %), 
Cyclagelosphaera (nearly 


20 %), Conusphaera (14 %), 
and  Nannoconus (7 %) are 
the most abundant compo-
nents of the assemblage 


(Fig. 10). The occurrence of 

these most abundant genera is in accordance with previous 
studies of calcareous nannofossils of the JKB interval (e.g., 
Michalík et al. 2009; Reháková et al. 2009; Lukeneder et al. 
2010; Wimbledon et al. 2013). Nannoliths represented by 
Polycostella beckmannii,  Hexalithus noeliae and Assipetra 
 are less present. The species indicative of 
eutrophic environments such as Zeugrhabdotus erectus and 
Diazomatholithus lehmannii occur only sporadically. Despite 
the poor preservation, several biostratigraphically important 

Fig. 8. Calpionellids in the Strapkova section: A — Longicollaria dobeni  (Borza). Sample 288.  
B — Chitinoidella boneti Doben. Sample 290.   — Chitinoidella elongata  ehánek. Sample 289.  
D — Tintinnopsella remanei Borza. Sample 293. E — Praetintinnopsella andrusovi Borza. Sample 
291. F — Crassicollaria intermedia (Durand Delga). Sample 294. G — Crassicollaria massutiniana 
(Colom). Sample 293. H — Crassicollaria brevis Remane. Sample 295. I — Crassicollaria colomi 
Doben. Sample 298.1. J — Calpionella alpina Lorenz. Sample 295. K — Calpionella grandalpina 
Nagy. Sample 294. L — Tintinnopsella carpathica (Murgeanu and Filipescu). Sample 347.

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, 2016, 67, 4, 303–328

species have been recorded: Nannoconus wintereriN. stein-
mannii minor
,  N. kamptneri minor,  N. steinmannii  stein-
. The full list of calcareous nannofossil taxa found is 
given in alphabetical order in the “Appendix” and strati-
graphically important taxa in the Fig. 11. 

The abundance of calcareous nannofossils in the sequence 

is generally low. On average throughout the section, about  
50 specimens per sample were observed. It means one speci-
men per six fields of view of the microscope. Due to the low 
abundance and prevailing bad preservation of calcareous 
nannofossils, only several biostratigraphic events have been 
defined. The first occurrence (FO) of N. wintereri was 
recorded in the bed 298.1, close to the expected JKB interval 
based on calpionellids (this study, see above). This bioevent 
represents the base of the NJT 17b Subzone, which Casellato 
(2010) considered to cover the JKB interval. The FO of  
N. steinmannii minor was recorded in bed 300.0 in middle 
part of the M19n magnetozone. Casellato (2010) indicates it 
as the base of the NKT Zone. N. kamptneri minor occurs spo-
radically from bed 343 upwards. The FO of N. steinmannii 
, namely the base of the NK-1 Zone sensu 

Bralower et al. (1989) was 
recorded in bed 352, in the 
lower part of the Elliptica 
Subzone (Fig. 10). N. kampt-
neri kamptneri
 was not found 
in the samples studied. 


Carbonate and C



The CaCO


 content in the 

uppermost part of the Czorsz-
tyn Fm. is relatively high 
(Fig.12). In the basal part of 
the Pieniny Fm. (from the 
beds 291 to 300) it decreases 
below 80 %. The decrease is 
in accordance with microfos-
sil analysis which pointed to 
raised silica bioproduction. 
The CaCO


 content reaches 

up to 90 % again in the Pieniny 
Fm. (325–360), where nanno-
conid and calcareous micro-
plankton remnants become 
abundant (Tremolada et al. 
2006; Michalík et al. 2009; 
Grabowski et al. 2013).

However, locally — as in 

bed 334 (Fig. 12), the CaCO



content decreases below 
50 %. Microfacies study have 
suggested that the main 

source of silica in the whole sequence came from radio larians 
and only a very low amount came from detrital minerals 
(quartz, clays, accessories; Figs. 2 and 5). Silica (opal– 
chalcedony) came from radiolarian tests, replaced by calcite 
and concentrated in cherts. 

TOC content is low (0.08–0.31 %) in all samples (Fig. 12). 

The C


contents slightly increases (more than 0.1 %) in the 

top of the Czorsztyn Fm and at the base of the Pieniny Fm., 
where CaCO


 content decreases. Similarly, in beds 350 and 

353, slight TOC accumulation could result from selective 
sorption of (dissolved) C


 by fine grains with more active 

surface, but probable also from raised fossil production. 

Stable carbon and oxygen isotopes

Both C and O isotopes of bull rock samples show a rela-

tively small variation and shift within a relatively narrow 
range (


C range from +1.09 to +1.96 ‰ VPDB, 


O from 

–2.93 to –1.20 ‰ VPDB). Late Tithonian 


C values (+1.96 

to +1.46 ‰ VPDB) show a slightly decreasing trend 
(Fig. 12). Next, higher up section (beds 188–300) values 

Fig. 9. Calpionellids in the Strapkova section: A — Tintinnopsella longa (Colom). Sample 345.  

 — Calpionella elliptica Cadisch. Samples 344; 359. D, E — Remaniella ferasini (Catalano). 

Sample 346; 354. F — Remaniella durandelgai Pop. Samples 344; 346. G — Remaniella cadischiana 
(Colom). Sample 359. H — Remaniella borzai Pop. Sample 353. I — Lorenziella hungarica Knauer 
and Nagy. Sample 343.

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, 2016, 67, 4, 303–328

Fig. 10

. V

ertical ranges of nannofossils and quantitative distribution of nannoplankton genera (open circles indicate uncertain specie

s identi


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, 2016, 67, 4, 303–328

Fig. 11. Important nannofossils species. Scalebar represents 5 μm. A — Watznaueria barnesiae (Black in Black & Barnes) Perch-Nielsen. 
Sample 291. B — Watznaueria manivitiae (Bukry) Moshkovitz & Ehrlich. Sample 330.   — Watznaueria britannica (Stradner) Reinhardt. 
Sample 299.5. D — Cyclagelosphaera margerelii Noël. Sample 293. E — Cyclagelosphaera 

 (Manivit) Roth. Sample 300.1.  

F — Diazomatolithus lehmanii Noël. Sample 297.3. G — Hexalithus noeliae (Noël) Loeblich & Tappan. Sample 299.4. H — Conusphaera 
 (Trejo) subsp. minor (Bown and Cooper) Bralower in Bralower et al. Sample 295. I — Conusphaera mexicana (Trejo) subsp. 
mexicana Bralower in Bralower et al. Sample 299.3. J — Zeugrhabdotus embergeri (Noël) Perch-Nielsen. Sample 298.4. K — Nannoconus 
globulus (Brönnimann) subsp. minor Bralower in Bralower et al. Sample 330. L — Nannoconus globulus (Brönnimann) subsp. globulus 
Bralower in Bralower et al. Sample 350. M — Nannoconus wintereri Bralower and Thierstein in Bralower et al. Sample 352. N — Nanno-
 kamptneri (Brönnimann) subsp. minor Bralower in Bralower et al. Sample 356. O — Nannoconus steinmannii (Kamptner) subsp. 
minor Deres and Achéritéguy. Sample 300.   — Nannoconus steinmannii (Kamptner) subsp. minor Deres and Achéritéguy. Sample 300.  
Q — Nannoconus steinmannii (Kamptner) subsp. minor Deres and Achéritéguy. Sample 381. R — Nannoconus steinmannii (Kamptner) 
subsp. steinmannii Deres and Achéritéguy. Sample 358.

achieving a range of +1.14 to +1.38 ‰ (in average +1.24 ‰ 
VPDB) show a new (balanced) isotope C composition of 
marine water during sedimentation of the Pieniny Fm. The 
high resolution carbon isotope record resembles the typical 
(stable or smooth) trend worldwide documented in the J/K 
boundary sequence (Weissert & Mohr 1986; Weissert & 
Channel 1989; Weissert & Lini 1991; Gröcke et al. 2003; 
Tremolada et al. 2006;  ák et al. 2011; Price et al. 2016). The 


C values between +1.3 to +1.5 ‰ (VPDB) occur in 

the Brodno (+1.3 to +1.6 ‰), Hlbo a (+1.0 to +1.5 ‰), and 
Strá ovce (+1.0 to +1.3 ‰) sections (Michalík et al. 1995; 
2009). At the Nutzhof section that was also located on the 
north Tethyan margin (Lukeneder et al. 2010) which is in an 
equivalent position on the north Tethyan margin, bulk carbon 
isotope values range between +0.49 and +2.10 ‰. 

The range of 


O data is not larger than 2 ‰, 


O values 

are relatively high in nodular limestones of the Czorsztyn 
Formation, and sharply decline at the onset of the Middle 
Tithonian. However, on average they remain close to 2 ‰ 
VPDB in the Pieniny Formation  (Fig. 12). Although 


values strongly vary over small stratigraphic scales, namely 
between individual beds, they do not show clear large-scale 


O values can be diagenetically modified more 

strongly than 


C. In the basal part of the Pieniny Fm (beds 



O data shift from –1.48 to –2.48 ‰. 


 values in higher part of the section (325–360) reach –1.84 to 
–2.93 ‰. 

Rock magnetism and demagnetization 

Samples were moderately to weakly magnetic with NRM 

intensities in the lower part of the section (up to sample 335 
including Tithonian and lower part of the Berriasian) mostly 
between 1 and 5×10


 A/m. Sample 296.5 revealed the high-

est NRM intensity around 9.5×10


 A/m. Higher up, in the 

upper part of the lower Berriasian, the NRM values fluctu-
ated around 1×10


 A/m (Supplementary Fig. S1).

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, 2016, 67, 4, 303–328

Low coercivity minerals dominate within the section 

which is manifested by negative values of S-ratio, mostly 
between –0.9 and –0.7 (Supplementary Figs. S1 and S2). 
Three samples (ST 333, 337.5 and 356.5) reveal slightly 
higher values of S-ratio: between –0.5 and –0.3. A single 
sample ST 296.5 reveals an extremely high value of S-ratio: 
0.54. The sample 296.5 also contains an unusually large 
amount of ferromagnetic minerals which is manifested by 
a very high intensity of the IRM


 (Supplementary Fig. S1). 

The sample 296.5 might also be distinguished by relatively 
high unblocking temperatures (up to 620 


C) and slightly dif-

ferent direction of C component (see below). Results of Low-
rie’s (1990) analyses (Supplementary Fig. S3A) confirm that 
medium and high coer 

civity minerals dominate in this 

sample. The maximum 
unblocking temperature of 


C unambiguously indi-

cates the presence of hema-
tite. Samples with moderately 
negative values of S-ratio 
(between –0.3 and –0.7) 
reveal presence of magnetite 
which is a dominant magnetic 
carrier. Its presence is docu-
mented by the maximum 
unblocking temperature of 


C in the 0.1T curve. 

However, the contribution of 
hematite is still significant as 
can be seen on the 1T curve 
(Supplementary Fig. S3B–D). 
Samples with low negative 
values of S-ratio contain 
almost exclusively magnetite 
(Supplementary Fig. S3E).



might be observed from the 
vertical log of S-ratio (Sup-
plementary Fig. S1) that the 
contribution of hematite is 
slightly more distinct in the 
lower half of the section.

During thermal demagneti-

zation, three characteristic 
NRM components were 
revealed. The least stable 

component is unblocked 

between 20 and 150–200 


(Fig. 13).  An  intermediate 

component is demagne-

tized in the temperature range 


C. Finally, a double 

polarity C component might 
be identified between 420 and 


C. Unfortunately, 

abrupt MS rise is observed 
during thermal treatment 

between 400 and 450 


C (Fig. 13) and sometimes the C com-

ponent cannot be demagnetized to the origin. 

Age of magnetization components and palaeotectonic 

The A component clusters match better in the present day 

coordinates (Table 1 and Fig. 14A). Their direction in geo-
graphic coordinates is close to the present day normal pola-
rity geomagnetic field direction in the area of investigation. 
Therefore, they are interpreted as recent viscous remanent 
magnetization of no geological importance. The mixed pola-
rity   component  is interpreted as the primary one. After 
bedding correction, the normal (Cn) and reversed (Cr) 

Fig. 12. Late Jurassic and Early Cretaceous C and O isotope stratigraphy calibrated against quantity of 
selected groups of microfossils and siliclastics.

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, 2016, 67, 4, 303–328

Fig. 13. 

Thermal demagnetization of typical specimens. Upper left: stereographic projection of demagnetization path; upper right: orthog

onal projection; lower left: NRM decay during thermal 

treatment; lower right: MS changes during thermal treatment. Components 

A, B and C are indicated in stereographic and orthogona

l projections. 


 — sample 293, Late 


ithonian Crassicollaria 

remanei Subzone, M19r magnetozone; 


 — sample 297A, Late 


ithonian Crassicollaria colomi Subzone, M19n2n magnetozone; 

 — sample 329_5, Early Berriasian Calpionella alpina Sub-

zone, M19n1r (“Brodno”) magnetozone; 


 — sample 334, Early Berriasian Calpionella alpina Subzone, M18r magnetozone; 


 — sample 338_5B, Early Berriasian Calpionella alpina Subzone, 

M18n magnetozone; 


 — sample 346_6C, Early Berriasian Calpionella elliptica Subzone, M17r magnetozone.

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, 2016, 67, 4, 303–328

directions cluster in the NW and SE quadrant, respectively, 
of a stereonet with moderate inclination (Fig. 14C). Its pri-
mary nature is also supported by the fact that polarity changes 
of the C component correlate well with the Global Polarity 
Time Scale (see below). The clustering of the C component 
does not improve after tectonic correction (Table 1) as might 
be expected in the case of a primary component. The McFad-
den & McElhinny’s (1990) reversal test gives negative results 
(critical angle 14.5





). It might be explained either 

by contamination of the intermediate B component or incom-
plete demagnetization of the samples containing hematite. 
The B component must be regarded as secondary magneti-
zation, as it always reveals a normal polarity (Fig. 14B). Sig-
nificant spread of both B and C components might result 
from overlapping of unblocking temperature spectra and 
from incomplete demagnetization of hematite. The position 
of the B component is usually close to Cn primary compo-
nent (Table 1). Therefore, the B component most probably 
represents pre-folding or early synfolding remagnetization of 
normal polarity. It might be acquired during the maximum 
burial or early phase of Late Cretaceous folding and thrus-
ting, alike abundant secondary magnetizations documented 
in the Central Western Carpathians (Grabowski 2005; 
Grabowski et al. 2009).

In the pre-folding coordinates, the declination of C compo-

nent reveals a moderate 46


 counter-clockwise (CCW) rota-

tion from the present-day north (Table 1). However, cluster-
ing of the C component is too weak for significant palaeotec-
tonic application (the value of precision parameter k >10 is 
required; see Van der Voo 1993). Having applied some selec-
tion (rejecting specimens deviating from the main cluster), 
the amount of the CCW rotation slightly decreases to 35


Clustering of the C component (both normal and reversed 
populations) improves after tectonic correction, although the 
precision parameter k is still slightly below 10 (see Table 1 
and Fig. 13d) and the reversal test is still negative. Declina-
tion of the C component is concordant with the general CCW 
of the study area. A counter-clockwise rotation of 47


 ( 18


was reported by Márton et al. (2013) from the Upper Creta-
ceous pelagic marls in the PKB in the neighbouring locality 
of Vršatec. 

Palaeoinclination of the Strapkova sec-

tion (41


), corresponding to palaeolatitude 



N  5

is slightly shallower than Titho-

nian–Berriasian palaeoinclinations from 
the PKB and Central Carpathian reference 
sections (Márton et al. 2015) which results 
from incomplete cleaning of the primary 
C component. 

Magnetostratigraphy and correlation 
with the Global Polarity Time Scale 

According to the polarity of the C com-

ponent, four normal (N1–N4) and four 

reversed polarity intervals (R1–R4) were documented 
(Fig. 15). Samples 292 and 292.5 revealed normal polarity of 
the C component (N1 interval). The subsequent four samples 
between 293 and 294.5 were of reversed polarity (R1 inter-
val). The long normal polarity (N2) interval was indicated 
between 295 and 328.5. It is followed by quick polarity 
changes manifested by the R2 (329 and 329.5) and N3 (330) 
intervals. The sample 330.5 was of undefined polarity. Three 
distinct polarity intervals were distinguished in the upper part 
of the section: reversed R3 interval (samples 331.5–334), 
normal N4 interval (334.5–341.5) and reversed R4 


(343–363.9) interval. 

The N1 interval is interpreted as the topmost part of the 

M20n magnetozone (Fig. 15). It is situated between the 
Tithonian Praetintinopsella Zone and the bottom of the 
Remanei Subzone. R1 interval is correlated with the M19r 
magnetozone. It covers the Remanei and Intermedia sub-
zones. The long normal N2 interval must be interpreted as 
the M19n2n. The boundary between Crassicollaria and 
Calpionella zones is usually situated within this magneto-
zone (for review, see Ogg et al. 1991; Grabowski 2011; 
Satolli et al. 2015). The short R2 and N3 intervals, in the 
lower part of the Alpina Subzone, are respectively correlated 
with the M19n1r (“Brodno”) and M19n1n magnetosubzones. 
The R3 interval is interpreted as the M18r magnetozone. This 
magnetozone is situated entirely within the Alpina Subzone 
(Houša et al. 2004; Grabowski & Pszcz kowski 2006; 
Pruner et al. 2010). The next normal N4 interval is paralleled 
with the M18n. The boundary between the Alpina and Ferasini 
subzones falls in the upper part of this magnetozone. It is 
concordant with the FAD of Remaniella ferasini which is 
observed usually in the M18n magnetozone (Ogg et al. 1991; 
Houša et al. 2004). A long reversed R4 interval in upper part 
of the section is interpreted as the M17r. It starts in the mid-
dle part of the Ferasini Subzone and continues into the Ellip-
tica Subzone. It is concordant with abundant data from  Italian 
sections (Ogg et al. 1991) and from the Po rednie sections 
(Grabowski & Pszcz kowski 2006), where the FO of 
Calpio nella  elliptica is observed also in the M17r magneto-
zone. The FO of Remaniella cadischiana is noted in the bed 
359 (Fig. 9G). As this taxon usually appears in the upper part 





































Cn+ Cr








Cn select








Cr select








Cn+Cr select






: Characteristic magnetizations from the Strapkova section. Palaeopoles:  (Cn + Cr 

population) — Pole latitude: 44.6


N; Pole longitude: 268.2


E; dp 6.0, dm 10.1  

(Cn + Cr selected population) — Pole latitude: 52.4


N; Pole longitude: 257.9


E; dp 5.1, 

dm 8.4 (dp, dm — con dence oval of palaeopole estimation).  Explanations: 

 — decli-

nation/inclination before tectonic correction, 

 — declination / inclination after  

tectonic correction; 


, k — Fisher statistics parameters, 

 — number of beds inves-

tigated/used for calculation of mean direction.

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, 2016, 67, 4, 303–328

of the M17r (Grabowski & Pszcz kowski 2006; Grabowski 
et al. 2010), it seems that the top of the section is quite close 
to the M17r/M17n magnetozones boundary.

Magnetic susceptibility

There is a moderately good correlation of MS with IRM 

(Supplementary Fig. S2A, B). This indicates that there might 

be a significant contribution of 
ferromagnetic minerals to the 
MS. Although a long term 
decrease of the IRM



observed as in the case of the 
MS (Supp. Fig. S1), the two 
curves are not identical which 
indicates that the contribution 
of paramagnetic minerals to 
MS cannot be neglected.

Magnetic susceptibility re 


veals a long term decreasing 
trend (Supplementary Fig. S1 
and Fig. 16). Its values are 

relatively high in the lower 
half of the section, between 




/kg in the Titho-

nian and lowermost Berriasian 
(the Alpina Subzone, up to 
sample 338). Large MS varia-
tions are also observed in that 
part of the section. The MS 
decreases by 50 % throughout 
the Tithonian, up to the JKB. 
Then it fluctuates between 


4 and 10×10




/kg in the 

lower part of the Alpina Sub-
zone, in M19n and M18r mag-
netozones. Significant increase 
up to 12×10




/kg is observed 

in upper part of the Alpina 
Subzone, in the bottom part of 
the M18n magnetozone. Then 
MS again decreases through-
out the M18n magnetozone to 




/kg. Within the Fera-

sini and Elliptica Subzones, 
MS values gently fall from 4 to 




/kg, with only two 

minor positive excursions in 
the M17r magnetozone.


Environmental proxies

In contrast to other Meso-

zoic system boundaries, the JKB time span was related to 
less dramatic environmental changes, generating problems 
with the definition of the JKB position, reflected in contra-
dictions of its determination in the Brodno and Strapková 
sections (Fig. 17) but also in its definition worldwide 
(Lukeneder et al. 2010; Michalík & Reháková 2011; 
 Wimbledon et al. 2013; Schnabl et al. 2015; Price et al. 
2016). Use of complex proxy parameters is inevitable. 

Fig 14. Stereographic projections of the magnetization components A, B and C. Left column: before 
tectonic correction (in situ); right column: after tectonic correction (in bedding coordinates). Full 
symbols – lower hemisphere projection; open symbols – upper hemisphere projection.

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, 2016, 67, 4, 303–328

Fig. 15. Magnetostratigraphy (NRM intensity; declination — D; inclination — I of the component C).

Fig. 16. Final summarization of data from the Strapkova section.

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, 2016, 67, 4, 303–328

Fig. 17. Integrated correlation of the O and C isotopes curves between Strapkova and Brodno sections.

Fluctuations of O isotope composition and distribution of 
nannoliths in “Maiolica-type” limestone, where sedimentary 
record, including C isotope composition, has not been 
influen ced by clastic input, can closely reflect original envi-
ronmental conditions. 

A link between the abundance of calcareous nannofossils 

and the CaCO


 content can be observed. The calcareous 

nanno fossils abundance noticably increases from bed 298.6 
and this trend continues up to the bed 299.6. This is an inter-
val with the highest nannofossil abundance in the succession 
studied. This peak is associated with a higher CaCO



(Figs. 12, 16). Conversely, the lowest calcareous nannofossil 
abundance has been recorded in bed 334, with only six speci-
mens. This event correlates with the most remarkable 
decrease of the CaCO


 content, with the radiolarian event and 

with the negative 


O excursion (Figs. 12, 16). 

The observed shift in the 


O values (2 ‰) throughout the 

section could indicate a relatively strong temperature change 

in the Early Cretaceous ice-free world (Anderson & Arthur 
1983; Gröcke et al. 2003; Shurygin et al. 2015) in the JKB 
sequence upwards. However, the 


O signal can have been 

modified by other characteristics of water in the basin and by 
diagenetic processes in sediment. Bulk-rock analyses can be 
rather informative about relative changes in temperature and 


O composition (Michalík et al. 2009; Lukeneder 

et al. 2010, 2015). 

The shifts between samples are frequently smaller than 

0.5 ‰ (Fig. 12). The first shift that occurred between the top 
of the Ammonitico Rosso and base of the Maiolica beds is 
relatively large (–1.20 to –2.18 ‰) and could indicate rela-
tively continual and intensive temperature rise by 3.5 to 4 
degree over a relative short (185–191 m)  interval.  The 


signal is more stable (–1.48 to –2.15 ‰) at the JKB interval 
(192–300.6 m) and suggests stabilization of possible higher 
temperature values. Data obtained by detailed study of 25 
samples fluctuate in a narrower interval (less than 0.7 ‰) 

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, 2016, 67, 4, 303–328

do not show any clear trend and imply a stable temperature 
regime at the J/KB. Negative 


O trend in the Maiolica 

sequence (325–360 m) continued and it indicates graded 
warming (approx.1–2.5 degree). Similar 


O trend and tem-

perature range has been observed in the Brodno (Michalík et 
al. 2009) or in the Nutzhof sections (Lukeneder et al. 2010) 
or other section ( ák et al. 2011, Price et al. 2016). 

Analysis of the data introduced above enabled stratigra-

phical and palaeoecological correlation of the Strapkova and 
Brodno sections (Fig. 17).


Bed to bed fluctuation of 


coincides with nanno- and microplankton events. During the 
Late Tithonian, an increase of Polycostella and Conusphaera 
was recorded. Increasing abundance of Nannoconus occurred 
during the Early Berriasian. Nannoconid peaks (beds 300; 
334; 350–352) correspond to negative 


O excursions and 

higher calcite accumulation (Fig. 16). Nannoconids are 
regarded as warm-water taxa (Street & Bown 2000; Melinte 
& Mutterlose 2001; Tremolada et al. 2006; Michalík et al. 
2009; Svobodová & Koš ák 2016). According to the O iso-
tope data, changes in calcareous nannofossil assemblage 
composition and the appearance of NannoconusPolycostella 
and Conusphaera (Figs. 10, 16) were followed by warming 
of 2–3 °C. On the other hand, the abundance of radiolarian 
tests indicates colder oxygenated and eutrophic upwelling 
intervals. Occasional colonization of hemipelagic bottom by 
infaunal trace maker assemblage indicates that the bottom 
water layer was not stagnant but periodically affected by con-
tour intermediate and low velocity currents. Hüneke & Stow 
(2008) characterized contourite beds by fine lamination and 
remnants of micro-cross-lamination of silty particles, domi-
nance of skeletal fragments of planktonic organisms, paucity 
of benthic shells, thorough bioturbation and burrowing of the 
underlying layer, dominance of microfacies of packed biomi-
crites, including wackestones and (foraminiferal) packstones, 
calcilutites with calcisiltite lenses, and so similar to the 
 situation observed in the Strapkova section. The contourites 
should have been deposited on the foot of continental slope 
at a depth of more than 300 metres. Enhanced water dynamics 
could be responsible for microfossil redeposition and for 
 several apparent blooms of crassicollarians. Temperature, 
salinity changes and raised trace metal contents in sea water 
could result in thinning and deformation of crassicollarian 
loricas (Tappan 1993; Reháková 2000b; Vandenbroucke et 
al. 2015) observed in several beds (296.3–297 m). 

Sedimentation rate

The overall sedimentation rate increases up section, from 

7.7 m/Myr in magnetozone M19r, through 9.5–9.8 m/Myr in 
M19n and M18r to 12.7 m/Myr in M18n and at least 


15 m/Myr in M17r (see Table 2). This trend is in agreement 
with the data of Grabowski & Pszcz kowski (2006), who 
also documented an increasing sedimentation rate between 
the magnetozones M19r and M17r in the Po rednie section 
(Fig. 18). Sedimentation rate in the Strapkova section is 
gene rally higher than in the Po rednie section although the 

shape of curves does not exactly coincide (Fig. 18). It is also 
much higher than in the Brodno section, where it does not 
exceed 3 m/Myr in the M19r and M19n magnetozones 
(Houša et al. 1999). Compared with the South Alpine sec-
tions, the sedimentation rate in the Strapkova section is com-
parable to that in the Torre de Busi section in the Lombardian 
Basin (Channell et al. 2010; Grabowski 2011). It is higher 
than the sedimentation rate calculated for Trento Plateau sec-
tions (mostly 2–6 m/Myr in the M19r–M18n interval). It could 
indicate a more distal depositional setting of the Strapkova 
section in comparison with the Brodno section (Fig. 18). 

Magnetic susceptibility and stratigraphic correlations 

Magnetic susceptibility in pelagic and hemipelagic car-

bonates of Late Tithonian–Berriasian age is usually confined 
to lithogenic influx into a basin (Grabowski et al. 2013; 
2014). This is most probably also the case in the Strapkova 
section, although geochemical data (e.g., correlation between 
lithogenic elements and MS) are not available. The MS curve 
obtained in the Strapkova section might be well correlated 
with coeval intervals in the Brodno section (Houša et al. 
1999) and in the Po rednie III section from the Tatra Mts. 
(Grabowski et al. 2013). A long term decreasing MS trend 
between the Upper Tithonian and upper part of the Lower 
Berriasian (i.e. top of M20n and M17r) is well constrained in 
the Po rednie and Strapkova sections (Fig. 19). A part of this 
trend is also observed in the Brodno section between 


M19r and M18r. Short-term MS fluctuations might be 
 compared as well. 

Fig. 18. Sedimentation rates in the Strapkova section compared with 
the Brodno section (PKB) and Po rednie section (Tatra Mts., Fatric 









rate (m/My)




1.44 (142.57–144.0)

15.0 (at least)




0.63 (144.0–144.63)





0.37 (144.64–145.01)





1.27 (145.01–146.28)





0.26 (146.28–146.54)


 Attempt of sedimentation rate estimation in the Strapkova 

section (timescale after Ogg, 2012).

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, 2016, 67, 4, 303–328

A decreasing trend I occurs between the uppermost part of 

M20n, through M19r to the lower part of M19n2n. It is fol-
lowed by a gently increasing trend II, which terminates in 
the upper part of M19n2n in the Strapkova and Po rednie 
sections and in the middle part of this magnetozone in the 
Brodno section. Trend II in the Po rednie III and Brodno sec-
tions terminates exactly at the J/K boundary, but in the Strap-
kova section — in the lower part of the Alpina Subzone. 

The next decreasing trend III falls in the uppermost part 

of M19n2n, approximately up to the M19n1r (Brodno) mag-
netozone. Trend IV reveals a generally increasing character 
and culminates in a local MS maximum in the magnetozone 
M18n. It is well resolved in the Strapkova and Po rednie sec-
tions, while most probably only the lowermost part of this 
trend is observed in the Brodno section. 

Trend V is related to profound MS decrease in the upper 

part of M18n and in M17r, covering the uppermost part of the 
Alpina Subzone, through the entire Ferasini Subzone and 
a large part of the Elliptica Subzone. Trends I–V are 
 apparently synchronous in relation to magnetostratigraphy. 
They might reflect changes of detrital input to the Pieniny 
and Zliechov (Central West Carpathians) basins controlled 
by regional tectonics and/or climate (Michalík 2007). The 
third, eustatic component (Grabowski et al. 2013), can be 

Fig. 19. Integrated correlation between Strapkova, Brodno and Po rednie III sections based on bio-, magnetostratigraphy and magnetic 
susceptibility. Position of the Jurassic/Cretaceous boundary is indicated according to different de nitions: black arrow - Intermedia/Alpina 
subzonal boundary; gray arrow - Colomi/Alpina subzonal boundary. Trends in MS variations (Roman numerals) are explained in the text.

involved after thorough correlation of well-dated but 
 geographically remote sections. 

MS variations in the Strapkova section negatively correlate 

with CaCO


 content (see Fig. 16). The CaCO


between beds 292 and 333 matches well the MS decrease 
(trends I to III and the lower part of trend IV). The upper part 
of trend IV (with maximum MS values in lower part of 
M18n) might be compared with a slight decrease of CaCO



(beds 334 to 337). The final decrease of MS during trend 
V correlates exactly with increasing CaCO


 in the uppermost 

Alpina, Ferasini and Elliptica subzones. Comparison of MS 
trends with nannofossil data (Fig. 16) indicates climatic 
 control of MS variations. The abundance of Conusphaera 
(relative cooling) coincides with the high MS values of  
trends I to IV. The Nannoconus dominance (relative  warming) 
is related to decreasing MS trend (trend V). Notably, the MS 
also  correlates with sequence stratigraphy. Sequence boun-
daries and total plankton abundance maxima match with 
local MS highs (see Figs. 5a and 16). 

There is an apparent contradiction between magnetic stra-

tigraphy and biostratigraphy especially in the detailed situa-
tion of the JKB. It might be related to the fact, that the JKB 
was defined according to slightly different criteria. In the 
Po rednie III (Grabowski & Pszcz kowski 2006) and 

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, 2016, 67, 4, 303–328

Brodno sections (Houša et al. 1999), the boundary was defined 
as the Intermedia/Alpina subzonal boundary. Alternatively, the 
Colomi/Alpina subzonal boundary was selected as the JKB 
marker in a revised version of the Brodno (Michalík et al. 
2009) and in the Strapkova (this study) biostratigraphy. The 
Colomi/Alpina subzonal boundary is correlated with the top-
most part of M19n2n (Michalík et al. 2009) in the Brodno sec-
tion, but in the Strapkova section with its lower half. Accor-
ding to defined MS trends, the Colomi/Alpina boundary falls 
at the boundary between trends III and IV in the Brodno sec-
tion and between trends I and II in the Strapkova (Fig. 19) 
section. The Intermedia/Alpina subzonal boundary is 

apparently synchronous in the Po rednie and Brodno 
 sections. It is situa ted in M19n2n, almost exactly between 
trends II and III, close to the local MS maximum. However, 
more high-resolution MS curves are desired, in order to test 
if the MS trends observed are also present in other sections 
beyond the Zliechov and the Pieniny basins. 

The high-resolution analysis of calpionellid and dinofla-

gellate associations was used in order to characterize the JKB 
interval in the Strapkova section. Three dinoflagellate and 
four calpionellid zones have been recognized. They show 
a Late Tithonian burst and calpionellid diversification and 
a later decrease in diversity of crassicollarians. Such changes 
in plankton composition and diversity across the Jurassic/
Cretaceous boundary were also documented by Reháková 
(2000), Reháková in Michalík et al. (2009), Wimbledon et al. 
(2013), Grabowski et al. (2010). The onset of the Alpina Sub-
zone of the standard Calpionella Zone, used as a marker for 
the JKB, was documented in sample 298. This limit is defined 
by morphological change of Calpionella alpina tests. There, 
medium-sized spherical forms of Calpionella alpina domi-
nate in biomicrite limestone and are accompanied by calci-
fied radiolarians. The successive Ferasini Subzone characte-
rized by the FO of Remaniella ferasini was identified in 
 sample 340. In sample 344, Calpionella elliptica, the bio-
marker of the Elliptica Subzone appeared. 

Nannofossil distribution documents the Tithonian NJT 17b 

Subzone to Early Berriasian NKT and NK-1 nannofossil 
zones (sensu Casellato 2010; and Bralower et al. 1989). The 
first occurrence of Nannoconus wintereri, which indicates 
the beginning of the NJT 17b Subzone and at the same time 
the beginning of the JKB transition has been located in 
 sample  298.1.

Correlation of calcareous microplankton with C and O 

 stable isotopes and TOC/CaCO


 data distribution was used in 

characterization of the JKB interval. 


C values ranging 

from 1.1 to 1.4 ‰ (PDB) indicated a typical balanced regime 
of carbon in the sea water. Negative 


O shift from –1.5 to 

–2.3 ‰ (V-PDB) in the uppermost Tithonian indicates a tem-
perature rise of 2–3° followed by stable temperature regime 
during the JKB with a warming tendency higher up the 

section. Radiolarian laminites interpreted as contourites and 
bioturbation levels prove oxygenation events of bottom 
waters during the Berriasian. 

Primary magnetization of mixed polarity was isolated and 

correlated with the Global Polarity Time Scale. The lower 
part of the Crassicolaria Zone (up to the middle part of the 
Intermedia Subzone) correlates with the M19r magnetozone. 
The M19n magnetozone includes the upper part of the 
 Crassicollaria Zone and lower part of the Alpina Subzone. 
The reversed Brodno magnetosubzone (M19n1r) was 
 identified in the uppermost part of M19n. The tops of the 
M18r and M18n magnetozones are located in the upper part 
of the Alpina Subzone and in the middle part of the Ferasini 
Subzone, respectively. The Ferasini/Elliptica subzonal 
boundary is located in the lowermost part of the M17r 

General MS decrease between the upper Tithonian and 

Berriasian is in agreement with the increasing content of 


 and warming trend documented by nannofossils. It is 

also accompanied by an increasing sedimentation rate resul-
ting from higher carbonate productivity. 

It appears that the MS might be of some importance in pre-

cise correlation of calpionellid bioevents. Minor MS varia-
tions in magnetozones M19r and M19n, related to changes of 
detrital input, correlate between sections from the Pieniny 
Klippen Belt (Strapkova and Brodno sections) and Tatra Mts 
(Po rednie section). They clearly demonstrate a subtle dia-
chronism between Crassicollaria/Calpionella zonal boun-
daries, defined according to different criteria.

 The authors thank T. Sztyrak, 

K.  í ková and K. Fekete for their help during  eld work. 
Dr. V. Šimo is acknowledged for determination and comments 
on ichnofossils. The research was supported by the VEGA 
Project 2/0034/16 and 2/0057/16, as well as by the APVV 
project 14-0118, projects 7AMB14SK201, RV067985831, 
and project No. GACR 16-09979S of Czech Grant Agency. 
Palaeomagnetic and rock magnetic investi gations were  nan-
cially supported by the project  DEC-2011/03B/ST10/05256 of 
the National Science Centre, Poland.


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Longicollaria  dobeni (Borza, 1966)
Carpathella rumanica Pop, 1998
Borziella slovenica (Borza, 1969)
Dobeniella tithonica (Borza, 1969)
Chitinoidella boneti Doben, 1963
Dobeniella cubensis (Furrazola-Bermúdez, 1965)
Popiella oblongata Reháková, 2002
Praetintinnopsella andrusovi Borza, 1969
Crassicollaria intermedia (Durand Delga, 1957)
Crassicollaria massutiniana (Colom, 1948) 
Crassicollaria brevis Remane, 1962
Crassicollaria parvula Remane, 1962
Crassicollaria colomi Doben, 1963
Calpionella alpina Lorenz, 1902
Calpionella grandalpina Nagy, 1986
Calpionella elliptalpina Nagy,1986
Calpionella elliptica Cadisch, 1932
Calpionella minuta Houša,1990
Tintinnopsella carpathica (Murgeanu and Filipescu, 1933)
Tintinopsella doliphormis (Colom, 1939) 
Tintinnopsella longa (Colom, 1939)
Tintinnopsella remanei Borza1969
Remaniella ferasini (Catalano, 1965)
Remaniella catalanoi Pop, 1996
Remaniella duranddelgai Pop, 1996
Remaniella colomi Pop, 1996
Remaniella borzai Pop, 1996

 Pop, 1996

Remaniella cadischiana Pop, 1996
Lorenziella hungarica Knauer and Nagy,1964

Stomiosphaera moluccana Wanner, 1940
Carpistomiosphaera borzai (Nagy, 1966)
Colomisphaera nagyi (Nagy, 1966)
Colomisphaera pulla (Borza, 1964)
Colomisphaera radiata (Vogler, 1941)
Colomisphaera tenuis (Nagy, 1966)
Colomisphaera fortis  ehánek, 1992
Colomisphaera lapidosa (Colom, 1935)
Colomisphaera carpathica (Borza, 1964)
Parastomiosphaera malmica (Borza, 1964)
Stomiosphaerina proxima  ehánek, 1987
Cadosina semiradiata fusca (Wanner,1940)
Cadosina semiradiata semiradiata (Wanner,1940)

Another microfossils
Gemeridella minuta Borza et Mišík 1975
Didemnoides moreti Durand-Delga
Didemnum carpaticum Borza et Mišík 1975
Globochaeta alpina Lombard 1945.

Assipetra infracretacea (Thierstein, 1973) Roth, 1973
Conusphaera mexicana (Trejo, 1969) subsp. mexicana Bralower in 
Bralower et al. 1989
Conusphaera mexicana (Trejo, 1969) subsp. minor  (Bown et 
 Cooper, 1989), Bralower in Bralower et al. 1989

 (Manivit, 1966) Roth, 1973

Cyclagelosphaera margerelii Noël, 1965
Diazomatolithus lehmanii Noël, 1965
Faviconus multicolumnatus Bralower in Bralower et al. 1989
Hexalithus noeliae (Noël, 1956) Loeblich et Tappan, 1966
Lithraphidites carniolensis De andre, 1963
Microstaurus chiastius (Worsley, 1971) Bralower et al., 1989
Nannoconus sp. Kamptner, 1931
Nannoconus erbae Casellato, 2010
Nannoconus  globulus (Brönnimann, 1955) subsp. globulus 
 Bralower in Bralower et al. 1989
Nannoconus globulus (Brönnimann, 1955) subsp. minor Bralower 
in Bralower et al. 1989
Nannoconus infans Bralower in Bralower et al. 1989
Nannoconus kamptneri (Brönnimann, 1955) subsp. minor Bralower 
in Bralower et al. 1989
Nannoconus steinmannii (Kamptner, 1931) subsp. minor Deres et 
Achéritéguy, 1980
Nannoconus  steinmannii (Kamptner, 1931) subsp. steinmannii 
Deres et Achéritéguy, 1980
Nannoconus wintereri Bralower et Thierstein in Bralower et al. 1989
Polycostella beckmannii Thierstein, 1971
Retacapsa sp. Black, 1971
Watznaueria  barnesiae (Black in Black et Barnes, 1959) Perch- 
Nielsen, 1968
Watznaueria biporta Bukry, 1969
Watznaueria britannica (Stradner, 1963) Reinhardt, 1964
Watznaureia fossacincta (Black, 1971a) Bown in Bown et Cooper 
Watznaueria manivitiae (Bukry, 1973) Moshkovitz et Ehrlich, 1987
Watznaueria ovata Bukry, 1969
Zeugrhabdotus embergeri (Noël, 1958) Perch-Nielsen, 1984
Zeugrhabdotus  erectus (Deflandre in Deflandre et Fert, 1954) 
 Reinhardt,  1965

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Supplementary Fig. S1.

 NRM intensity



 and S-ratio in the Strapkova section, against biostratigraphy and magnetostratigraphy

. Details of magnetostratigraphic interpr


tations are presented in the Fig. 15.


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, 2016, 67, 4, 303–328

Supplementary Fig. S2. Rock magnetic correlations in Strapkova sections. A — MS vs. IRM


B — MS vs. IRM


 without anomalous 

sample 296.5;   — MS vs. NRM; D — S-ratio vs. IRM



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, 2016, 67, 4, 303–328

Supplementary Fig. S3. Thermal demagnetization of the IRM acquired in the  elds of 0.1T, 0.4T and 1T in three perpendicular 
A — sample 296.5, Late Tithonian Crassicollaria colomi Subzone, M19n2n magnetozone (S-ratio 0.54)
B — sample 326, Early Berriasian Calpionella alpina Subzone, M19n2n magnetozone (S-ratio –0.63)

 — sample 333, Early Berriasian Calpionella alpina Subzone, M18r magnetozone (S-ratio –0.63)

D — sample 356.5, Early Berriasian Calpionella elliptica Subzone, M17r magnetozone, (S-ratio –0.35)
E — sample 357.6, Early Berriasian Calpionella elliptica Subzone, M17r magnetozone, (S-ratio=–0.85)