GeolCarp_Vol65_No4_307

www.geologicacarpathica.com

GEOLOGICA CARPATHICA

, AUGUST 2014, 65, 4, 307—326 doi: 10.2478/geoca-2014-0021

Introduction

The Western Carpathians are the north-easternmost
extension  of  the  European  Alps.  They  form  a
northward-convex E—W trending mountain arc ap-
proximately 500 km long and 300 km wide (Fig. 1)
belonging  to  the  Alpine  orogenic  belt  in  Central
Europe.  In  the  west,  the  Western  Carpathians  are
linked  to  the  Eastern  Alps  and  in  the  east  to  the
Eastern Carpathians. The Western Carpathians are
divided  into  three  principal  zones  –  the  External
Western Carpathians (EWC), Central Western Car-
pathians (CWC), and Internal Western Carpathians
(IWC)  (e.g.  Andrusov  et  al.  1973;  Mahe   1986;
Kozur & Mock 1996, 1997; Plašienka et al. 1997;
Plašienka 1999; Froitzheim et al. 2008).

The Tatra Mts cover a relatively small area of

785 km

in the northern portion of the CWC

Cretaceous—Quaternary tectonic evolution of the Tatra Mts

(Western Carpathians): constraints from structural,

sedimentary, geomorphological, and fission track data

SILVIA KRÁLIKOVÁ

, RASTISLAV VOJTKO

, UBOMÍR SLIVA

, JOZEF MINÁR

3,4

BERNHARD FÜGENSCHUH

, MICHAL KOVÁČ

and JOZEF HÓK

Department of Geology and Paleontology, Faculty of Natural Sciences, Comenius University, Mlynská dolina G, SK-842 15 Bratislava,

Slovak Republic; kralikova@fns.uniba.sk; vojtko@fns.uniba.sk; kovacm@fns.uniba.sk; hok@fns.uniba.sk

Nafta Inc., Votrubova 1, SK-815 05 Bratislava, Slovak Republic; lubomir.sliva@nafta.sk

Department of Physical Geography and Geoecology, Faculty of Natural Sciences, Comenius University, Mlynská dolina B1,

SK-842 15 Bratislava, Slovak Republic; minar@fns.uniba.sk

Department of Physical Geography and Geoecology, Faculty of Sciences, University of Ostrava, Chittussiho 10, CZ-71000 Ostrava,

Czech Republic; josef.minar@osu.cz

Institute of Geology, University of Innsbruck, Innrain 52, A-6020 Innsbruck, Austria; bernhard.fuegenschuh@uibk.ac.at

(Manuscript received July 17, 2013; accepted in revised form June 5, 2014)

Abstract: The Tatra Mts area, located in the northernmost part of Central Western Carpathians on the border between
Slovakia and Poland, underwent a complex Alpine tectonic evolution. This study integrates structural, sedimentary, and
geomorphological data combined with fission track data from the Variscan granite rocks to discuss the Cretaceous to Quater-
nary tectonic and landscape evolution of the Tatra Mts. The presented data can be correlated with five principal tectonic
stages (TS), including neotectonics. TS-1 ( ~ 95—80 Ma) is related to mid-Cretaceous nappe stacking when the Tatric Unit
was overlain by Mesozoic sequences of the Fatric and Hronic Nappes. After nappe stacking the Tatric crystalline basement
was exhumed (and cooled) in response to the Late Cretaceous/Paleogene orogenic collapse followed by orogen-parallel
extension. This is supported by 70 to 60 Ma old zircon fission track ages. Extensional tectonics were replaced by transpression
to transtension during the Late Paleocene to Eocene (TS-2;  ~ 80—45 Ma). TS-3 ( ~ 45—20 Ma) is documented by thick
Oligocene—lowermost Miocene sediments of the Central Carpathian Paleogene Basin which kept the underlying Tatric
crystalline basement at elevated temperatures (ca.  > 120 °C and  < 200 °C). The TS-4 ( ~ 20—7 Ma) is linked to slow Mio-
cene exhumation rate of the Tatric crystalline basement, as it is indicated by apatite fission track data of 9—12 Ma. The final
shaping of the Tatra Mts has been linked to accelerated tectonic activity since the Pliocene (TS-5;  ~ 7—0 Ma).

Key words: Western Carpathians, Tatra Mts, tectonics, sedimentology, geomorphology, fission track analysis.

Fig. 1. a – Location of the Tatra Mts, b – Simplified
tectonic map of the Tatra Mts broader area (modified
according to Bezák et al. 2004).

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along the state border of Slovakia and Poland (Fig. 1). Their
impressive  landform  is  highlighted  by  Gerlachovský  štít
(peak)  (2655 m a.s.l.),  the  highest  peak  in  the  whole  Car-
pathian  arc.  Geomorphologically,  the  mountains  are  divided
into the Eastern (High Tatra Mts and Belianske Tatry Mts) and
Western  Tatras.  The  Tatra  Mts  form  an  asymmetrical  horst-
like  W-E  trending  megaanticline.  The  southern  boundary
against the Liptov and Poprad subbasins belonging to the Cen-
tral  Carpathian  Paleogene  Basin  (CCPB)  is  marked  by  the
sub-Tatra fault (Figs. 1 and 2 – Mahe  et al. 1967, 1986). As
a result of their unique morphometric character, they were re-
cently  defined  as  markedly  the  smallest  basic  morphostruc-
tural region of the Western Carpathians (Minár et al. 2011).

In spite of numerous publications on the structural geology

of the Tatra Mts area (e.g. Nemčok et al. 1993; Janák et al.
2001;  Sperner  et  al.  2002;  Jurewicz  2005;  Vojtko  et  al.
2010), their Alpine geodynamic evolution is still not precisely
revealed.  Moreover,  the  great  majority  of  published  studies
focus  only  on  partial  time  and  thematic  aspects.  Since  the
Alpine post-collisional tectonic evolution of the Tatra Mts is
still  not  fully  understood,  a  comprehensive  synthetic  model
using  a  multidisciplinary  approach  is  needed.  For  this  pur-
pose the structural, sedimentary, and geomorphic data com-
bined with zircon and apatite fission track analysis (ZFT and
AFT) were applied.

To achieve a comprehensive model of the tectonic evolu-

tion  of  the  Tatra  Mts  during  the  Cretaceous  to  Quaternary,
we have focused on the following problems: (i) even though
common  features  of  the  evolution  of  the  Tatra  Mts  and  the
other CWC mountains exist, the recent unique character of the
mountains  is  conditioned  by  specific  features  of Cretaceous
to Quaternary  development,  revealed  from  structural,  sedi-
mentary,  geomorphological,  and  geochronological  data;
(ii) on  the  other  side,  relationships  between  the  extreme  re-
lief and kinematics of the sub-Tatra fault system and paleo-
stress  changes  probably  exist.  They  can  be  revealed  by
comparison  of   paleostress  variations  and  development  of
asymmetry  on  the  basis  of  lithological,  structural,  and  geo-
morphological data; (iii) the time of the neotectonic morpho-
logical  individualization  of  the  Tatra  Mts  that  should  be  in
line  with  geomorphologic  evidence  can  be  estimated  from
the sedimentary, structural, and geochronological records.

Regional geological and geomorphological settings

The Tatra Mts are bounded by the denudation remnants of

the CCPB (Figs. 1 and 2). The region, located near the
boundary zone between the CWC and IWC (boundary: Pie-
niny Klippen Belt) was affected by strong Miocene deforma-
tion (e.g. Ratschbacher et al. 1993; Nemčok & Nemčok
1994; Kováč & Hók 1996; Plašienka et al. 1997; Pešková et
al. 2009; Vojtko et al. 2010).

The core of the Tatra Mts is formed by the Tatric crystal-

line basement (Fig. 2) which is composed of two thick-
skinned Variscan tectonic units (Kahan 1969; Janák 1994).
Both tectonic units differ in metamorphic grade and lithology.
The Jalovec Unit (lower one) occurs only in a tectonic inlayer
in the south-western portion of the mountains (Western Tatra

Mts).  This  unit  is  built  up  by  a monotonous  complex  of
schists  and  gneisses  with  intercalations  of  quartzite.  The
Baranec Unit (upper one) has a more variegated composition
and consists of two complexes. Metamorphic rocks are pre-
dominantly composed of migmatite, orthogneiss, paragneiss,
and  occasionally  amphibolite  with  relics  of high  pressure
metamorphic  rocks.  The  main  mass  of the  upper  unit  com-
prises different types of Variscan granites (Kahan 1969; Nem-
čok et al. 1993, 1994; Janák 1994; Janák et al. 1996, 1999).

The Tatric sedimentary cover (Tomanová cover sequence)

is largely of Late Permian to Cretaceous age (Permian—Early
Turonian).  The  autochthonous  cover  sequence  contains  the
Javorinská Široká, Tomanová, and Osobitá successions. The
allochthonous cover sequences are located predominantly in
the  central  part  of  the  mountains  and  include  the  Červené
vrchy, Giewont, and Široká partial nappes. Overall the Tatric
cover  sequences  display  a monoclinal  structure  with
a moderate  northward  inclination  (Fig. 2).  The  thickness
ranges  from  several  hundred  meters  to  a maximum  of
~

2000 m in the Kominy Tyłkowe area (for further informa-

tion see Nemčok et al. 1993).

The tectonically overlying the Tatric Unit is a nappe derived

from  the  area  between  the  Tatric  and  Veporic  realms  (Biely
& Fusán  1967).  It  can  be  subdivided  into  several  smaller
nappes (predominantly the Bobrovec, Suchy Wierch, Havran,
and  Bujačí  partial  nappes)  or duplexes  that  differ  mainly  by
lithostratigraphic  content,  position,  and  regional  distribution
(Fig. 2). The stratigraphic range of the Fatric Unit is Early Tri-
assic—Early Cretaceous. The age of thrusting is constrained by
the deposition of a synorogenic flysch – the Poruba Forma-
tion (Albian to Early Turonian) in the Fatric and Tatric units
(e.g. Andrusov et al. 1973; Plašienka 2003).

The Hronic Unit represents structurally the highest nappe

system.  It  appears  only  in  the  western  portion  of  the  Tatra
Mountains  (Fig. 2).  This  nappe  contains  sedimentary  strata,
from Anisian to Toarcian in age. It is composed of two par-
tial  nappes,  the  Siwa  Woda  (lower)  and  the  Furkaska-
Koryčiská  (upper)  (Nemčok  et  al.  1993).  The  lithological
composition  of  these  partial  nappes  refers  to  the  Ludrová
succession  which  is  interpreted  as  a slope  between  the  Rei-
fling Basin and the Mojtín-Harmanec carbonate platform (cf.
Havrila 2011).

These nappes form the substratum of the Eocene to earliest

Miocene CCPB sedimentary succession (Podtatranská skupina
Group;  cf.  Gross  et  al.  1984,  1993;  Sliva  2005  –  Fig. 2)
which is predominantly composed of deep-marine siliciclas-
tics  up  to  4 km  in  thickness  (Soták  et  al.  2001).  The  Pod-
tatranská  skupina  Group  is  divided  into  four  formations
(Golab  1959;  Roniewicz  1969;  Gross  et  al.  1984)  –  the
Borové  Formation  a.k.a.  “Numulitic  Eocene”,  Huty  Forma-
tion  a.k.a.  Zakopane  Member,  Zuberec  Formation  a.k.a.
Chocholów  Formation,  and  Biely  potok  Formation  a.k.a.
Ostrysz Formation (Fig. 3 column “Lithostratigraphy”) with
stratigraphic  span  from  Middle  Eocene  to  earliest  Miocene
(Olzsewska  &  Wieczorek  1998;  Soták  1998,  2010;  Gedl
2000; Garecka 2005). Traditionally, the CCPB is interpreted
as  a fore-arc  basin  located  behind  the  Carpathian  accretion-
ary  wedge  (e.g.  Royden  &  Báldi  1988;  Tari  et  al.  1993;
Kázmér et al. 2003) and now surrounding the Tatra Mts.

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The Miocene and Pliocene sediments are not preserved in

the Tatra Mts and they are not confirmed from its immediate
surrounding  as  well.  However,  the  Middle  Miocene  to
Pliocene terrestrial and fresh water sediments, up to 1300 m
thick  form  the  fill  of  the  Orava—Nowy  Targ  Basin  (Fig. 3
column  “Lithostratigraphy”)  located  several  kilometers
north (Watycha 1976; Gross et al. 1993). Their original ex-
tents probably reached further southward and covered a large
part of the adjacent sediments of the CCPB (Wagner 2011).

Significant amounts of Quaternary sediments can be found

in the southern foothills of the High Tatra Mountains where
massive Pleistocene moraine and fluvioglacial sediments are
more than 400 m thick. These sediments are thought to have
been deposited in a graben which was related to normal fault-
ing along the Ružbachy and sub-Tatra faults (e.g. Nemčok
et al. 1993, 1994).

The Tatra Mts form an asymmetrical horst structure associ-

ated with the Miocene exhumation, which occurred along the
sub-Tatra fault system as was confirmed using several geolog-
ical methods (e.g. Burchart 1972; Andrusov et al. 1973; Krá
1977; Mahe  1986; Kováč et al. 1994; Plašienka et al. 1997;
Janák et al. 2001; Baumgart-Kotarba & Krá  2002; Struzik et
al. 2002; Anczkiewicz 2005; Anczkiewicz et al. 2005, 2013;
Śmigielski et al. 2012). The exhumation continued during the
neotectonic  period  (Pliocene—Quaternary)  as  is  indicated  by
the large amount, several hundred meters in thickness, of the
Quaternary glaciofluvial sediments distributed predominantly
at the southern foot of the Tatra Mts (e.g. Nemčok et al. 1993).

The Tatra Mts differ from the rest of the Western Car-

pathians  by  their  exceptional  geomorphological  character.
The  mountains  are  characterized  by  dominant  glacial  relief
formed  during  the  Pleistocene  (Lukniš  1973)  and  the  most
intensive periglacial processes in the recent (Midriak 1983).
Absence  of  significant  remains  of  planation  surfaces  is  an-
other specific geomorphic feature of the Tatra Mts. The vari-
ation of altitude (750—2650 m a.s.l.) is bigger than anywhere
else  in  the  Western  Carpathians.  In  addition,  the  mountains
are characterized by values of mean slopes ( ~ 25°) and avail-
able relief ( ~ 700 m), which are much higher than in the other
mountains of the CWC ( ~ 5° vs.  ~ 150 m).

Distinctive W—E as well as N—S geomorphological differen-

tiation is typical for the Tatra Mts. Altitudinal difference be-
tween the Eastern and Western Tatra Mts reach about 400 m.
However, the Belianske Tatry Mts (north-eastern part of the
Eastern Tatra Mts) are approximately 500 m lower than the
rest of the Eastern Tatra Mts (High Tatra Mts). N—S altitudi-
nal asymmetry is also clearly visible in the High Tatra Mts
alone – the difference between average altitudes of S and N
ridges reaches about 100 m here (Holle 1909 ex Lukniš 1973).

Methods

Structural analysis

Standard procedures for brittle fault-slip analysis and pa-

leostress reconstruction are now well established (Etcheco-
par et al. 1981; Angelier 1990, 1994). Paleostress axes
characterized by the brittle structures were computed by the

WinTensor  software  (Delvaux  &  Sperner  2003)  using  the
method  of  Angelier  (1984).  Fault  data  were  inverted  to
obtain  the  four  parameters  of  the  reduced  stress  tensor:
σ

(maximum principal stress axis),

(intermediate principal

stress axis),

(least principal stress axis), and also the

Φ ra-

tio of principal stress differences. The stress regime was
expressed numerically from the

Φ ratio using an index Φ’

(according to Delvaux et al. 1997).

Sedimentological analysis

The sedimentary and geodynamic development of the Tatra

Mts and their surroundings have been studied also using sedi-
mentological  analysis  applied  to  the  CCPB  strata,  especially
focusing  on  the  Spišská  Magura  Mts  and  eastern  part  of  the
Podhale  region.  Detailed  sedimentary  logs,  available  from
well-preserved outcrops, have been constructed and correlated
to  obtain  information  on  sedimentary  environments.  Paleo-
transport  reconstruction  was  based  on  measurement  of  exist-
ing marks on bedding planes, cross bedding, lamination, and
also  on  measurement  of  channel  axis  and  synsedimentary
folds  as  well.  Provenance  analyses  of  coarse  clastics  were
based on macroscopic and thin-section observations.

Geomorphological analysis

Analysis was performed by use of digital elevation models

derived from the basic topographic maps of the Slovak Re-
public 1 : 50,000, geological maps 1 : 50,000 (Nemčok et al.
1994) and geomorphological map 1 : 50,000 (Lukniš 1968).
ArcGIS 9.3 software was used for processing.

The paragenetic system of the Studený potok stream (in-

cluding not only recent but also all Pleistocene accumulations
of the stream) was the subject of analysis. It includes the area
of  Slavkovský  štít  (peak)  (2452 m a.s.l.)  where  fission  track
analysis was performed. The denudation segment – DS (val-
ley in the Tatra Mts) and accumulation segment – AS (glacial
and  glaciofluvial  sediments  on  the  foothills)  were  distin-
guished. On the basis of outcrops of pre-Quaternary rocks (Pa-
leogene deposits and crystalline rocks) recorded in geological
and geomorphological maps, a few boreholes and geophysical
profiles, a first approximation of the Quaternary base surface
of  the  AS  was  created  (using  Topo  to  Raster  interpolation
from  the  selected  points).  The  final  approximation  was  per-
formed  by  adding  new  points  of  the  Quaternary  base  (—3 m
below recent surface of Quaternary sheet) in the places were
the first approximation rose up to the recent surface.

Envelope surfaces were created for both, the AS and DS

from the highest ridge line points. Control by recent surface
was performed too (envelope surface cannot reach below re-
cent surface). The volume of preserved accumulation as well
as all retention space in the AS was calculated by subtraction
of the Quaternary base surface from the recent and envelope
surfaces. The volume of the DS was determined by analogy.

FT analysis

Sample preparation and FT analysis were carried out at the

Fission Track Laboratory of the Institute of Geology, Uni-

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versity of Innsbruck. Apatite and zircon separation was done
with  standard  heavy  mineral  separation  techniques  using
magnetic  and  heavy  liquid  separation.  The  apatites  were
mounted  in  epoxy  resin  and  zircons  in  PFA®  Teflon,  then
ground  and  polished.  Spontaneous  fission  tracks  were  re-
vealed by etching in 6.5% HNO

for 40 s at 20 °C. The zir-

cons were etched in a NaOH—KOH eutectic melt for 4—8 h at
235 °C. Irradiation of both apatite and zircon was carried out
at the FRM II research reactor in Garching, Germany. Neu-
tron flux was monitored using CN5 (for apatite) and CN1
(for zircon) dosimeter glasses. After irradiation, the induced
fission tracks in external detector muscovites were etched in
40% HF for 45 min at 20 °C. The fission tracks were counted
with 1250

× magnification with a dry objective using

a computer-controlled  Zeiss  Axioplan  microscope  equipped
with  an automated  AUTOSCAN  stage.  The  samples  were
analysed  using  the  external  detector  method  (EDM)  as  de-
scribed by Gleadow (1981). All ages are ‘central ages’ (Gal-
braith & Laslett 1993) with errors quoted as  ± l

σ. The central

ages were calculated using the zeta calibration method (Hur-
ford & Green 1983) with a zeta factor of 372 ± 35.8 year/cm

(apatite, CN5 glass) and 185.8 ± 11 year/cm

(zircon, CN1

glass)  (analyst:  S. Králiková).  Data  processing  was  carried
out  using  the  TRACKKEY  program,  version  4.2.f  (Dunkl
2002).  Only  grains  with  their  prism  planes  parallel  to  the
polished surface were used for age dating. The probability of
grains counted in a sample belonging to a single population
of  ages  was  assessed  by  a P(

) probability test (Galbraith

1981). Long axes of the FT etch-pits on polished apatite sur-
faces  (Dpar  method  –  Donelick  1993)  were  measured  as
a proxy for annealing properties. Horizontal confined fission
track  lengths  were  measured  in  apatite  samples  and  mean
horizontal  confined  track  lengths  (Mean  HCTL)  with  one-
sigma standard deviation (

σ) were determined.

Data used

Structural data

Structural data obtained from the Tatra Mts area were used

to  reconstruct  the  Late  Cretaceous  to  Quaternary  tectonic
evolution  of  the  study  area.  Four  main  deformation  phases
related  to  principal  burial  vs.  exhumation/denudation  pro-
cesses  were  identified  in  the  paleo-Alpine  nappe  units  and
one in the CCPB deposits. The last deformation phase affected
both  the  paleo-Alpine  nappe  units  as  well  as  the  CCPB  de-
posits.  The  Cenozoic  paleostress  fields  were  calculated  in
detail  and  published  in  the  works  of  Pešková  et  al.  (2009),
Vojtko et al. (2010), and Sůkalová et al. (2012). Analysis of
structural  measurements,  as  well  as  a geological  and  struc-
tural study of the Tatra region show relative clockwise rota-
tion  of  the  paleostress  field  during  the  Late  Cretaceous  to
Cenozoic (Figs. 3, 4 and Table 1).

Late Cretaceous—Paleocene orogenic collapse, exhuma-

tion/denudation,  post  dating  the  nappe  stacking,  is  docu-
mented  within  the  whole  CWC  territory  especially  in  the
Veporic  and  Gemeric  belt  (e.g.  Hók  et  al.  1993;  Plašienka
1993, 1999; Jeřábek et al. 2012). The compressional tectonic

regime was followed by extensional tectonics (Fig. 3) which
occurred predominantly on low angle normal faults.

From the Paleocene to Eocene a change from transpres-

sion to transtension together with a shift from WSW—ENE to
ENE—WNW compression can be observed (Fig. 3). The tim-
ing brackets for this stage are based on stratigraphic argu-
ments with the upper limit given by undeformed Upper
Eocene to Oligocene formations.

The Early—Middle Miocene was a time characterized by

reverse faulting and, to a lesser extent, by strike-slip faulting
with the maximum stress axis (

) oriented NW—SE (Fig. 3).

Especially, intense folding on all scales can be observed
throughout and the reverse sub-Tatra fault allowed for uplift
and exhumation of the Tatra Mts in the northern part of the
working area.

During the late Middle and Late Miocene a gradual

change in the paleostress orientation was observed. The
paleostress axis (

) rotated to the NE—SW position in the

Late  Miocene  (Fig. 3).  Additionally,  the  tectonic  regime
passed  from  transpression  through  transtension  to  tension.
The  Middle  to  Late  Miocene  was  characterized  by
a generally  N—S-trending  compression  in  a strike-slip  and
compressional tectonic regime.

The youngest tectonic regime (neotectonics) is character-

ized by the E—W extension, which has been documented in
the Orava—Nowy Targ Basin (Pešková et al. 2009). The Qua-
ternary tension is parallel to the Western Carpathian arc (S

and the S

of the stress field is generally N—S (Fig. 3). How-

ever, some disturbances can occur, for example in the Kozie
Chrbty Mts and the Hornád Depression in the south of the
study area (Sůkalová et al. 2012).

Sedimentological data

The aim of the sedimentological study was to better under-

stand the burial vs. exhumation story of the Tatra Mts region
during the Cenozoic. The main goal was to complement all
relevant previous sedimentary studies of the CCPB and Orava—
Nowy Targ Basin in combination with our data and to pro-
vide a contemporary view of the basins’ development.
Sedimentological data played a principal role in solving the
Cenozoic evolution of the Tatra region.

Formation of the CCPB started with the deposition of

transgressive  facies  from  the  Middle  to  Late  Eocene  (Lu-
tetian—Priabonian)  while  the  Oligocene  to  earliest  Early
Miocene is characterized by a thick complex of deep-sea fan
sediments (Fig. 3). Basin evolution was controlled by tecton-
ics, as documented by syn-sedimentary deformation (Starek
2001;  Pešková  et  al.  2009;  Vojtko  et  al.  2010),  by  climatic
and  by  sea-level  changes  (Janočko  &  Jacko  1998;  Soták
1998, 2010; Soták & Starek 2000; Soták et al. 2001).

Sedimentation of the Middle—Upper Eocene Borové For-

mation  started  with thick  alluvial,  deltaic  and  shallow  ma-
rine  clastics  overlain  by  shallow  marine  and  organodetrital
limestones.  The  uppermost  part  of  the  formation  is  com-
posed of bryozoan and Globigerina marls (Soták 2010). The
overall transgressive character of the Borové Formation was
interrupted  in  the  Priabonian  by  a short  regression  (Fig. 3),
namely the Pucov Member (cf. Starek et al. 2012).

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Table 1: Paleostress tensors from fault slip data. Explanations: Site – code of locality, n – number of fault used for stress tensor determina-
tion, n

– total number of fault data measured,

, and

– azimuth and plunge of principal stress axes,

Φ – stress ratio (σ

—

Method – direct inversion (DI), numerical dynamic analysis (NDA), Event – tectonic regime (flattening, strike-slip, extension).

Site

Formation

Method Event

CCP-H1 Huty

7 26

139/07 230/09

011/79

0.38

DI flattening

CCP-H2 Huty

6 26

096/71 226/13

319/14

0.48

DI extension

CCP-H3 Huty

4 26

306/83 120/07

210/01

0.5

DI extension

CCP-H4

Huty

3 26 195/22 298/28 072/53 – NDA flattening

CCP-H5 Huty

5 26

205/03 296/05

078/84

0.5

DI flattening

CCP-HBD1 Huty

6 308/15

217/04 113/74 0.44 DI

flattening

CCP-HBD2 Huty

6 063/22

153/01 244/68

–

NDA flattening

CCP-HBD3 Huty

6 215/76

010/12 101/06

–

NDA extension

CCP-HM1 Borové

9 46 153/14 350/75 244/04 0.61

strike-slip

CCP-HM2 Borové

6 46 210/20 065/65 305/13 0.38

strike-slip

CCP-HM3 Borové

15 46 237/87 001/02 091/02 0.29

extension

CCP-HM4 Borové

4 46 300/33 035/07 135/56 0.25

flattening

CCP-HM5 Borové

4 46 116/21 300/68 207/02 0.47

strike-slip

CCP-HNH1 Huty

4 270/06

167/64 003/25 –

NDA strike-slip

CCP-TDI.1 Zuberec

3 11 163/62 063/05 330/27 –

NDA extension

CCP-TDI.2 Zuberec

4 11 299/11 029/01 127/79 0.5 DI

flattening

CCP-TDI.3 Zuberec

2 11 359/25 190/65 097/01 – NDA strike-slip

CCP-TDII.1 Zuberec

33 208/82

036/08 306/01 0.5 DI

extension

CCP-TDII.2 Zuberec

33 131/18

222/02 317/72 0.61 DI

flattening

CCP-TDII.4

Zuberec

2 33 218/74 309/01 039/16 – NDA extension

CCP-Z1 Huty

255/30 032/52

152/22

0.5

DI strike-slip

CCP-Z2 Huty

163/21 026/63

260/17

0.46

DI strike-slip

CCP-Z3

Huty

3 25 157/59 055/07 312/30 – NDA extension

CCP-Z4 Huty

349/27 258/01

166/63

0.5

DI flattening

CCP-Z5 Huty

136/62 315/28

226/00

0.5

DI extension

PPTJA01A

Zuberec Fm

105/70

272/20

3/4

–

NDA extension

PPTJA01B

Zuberec Fm

4/16

273/2

176/74

–

NDA flattening

PPTJA02A

Huty Fm

109/38

206/8

306/51

–

NDA extension

PPTJA02B

Huty Fm

160/52

206/8

356/37

–

NDA extension

PPTJA04A

Zuberec Fm

216/82

44/7

314/1

0.43 DI

extension

PPTJA04B

Zuberec Fm

335/10

66/7

191/78

0.40 DI

flattening

PPTJA04C

Zuberec Fm

166/2

76/2

218/88

0.72 DI

flattening

PPTJA04D

Zuberec Fm

195/77

22/13 292/2

0.21 DI

extension

PPVTA01A

Gutenstein Lm.

91/0

1/83 181/7

0.20 DI

strike-slip

PPVTA01B

Gutenstein Lm.

140/7

231/10

15/78

–

NDA

flattening

PPZDI01A

Huty Fm

107

317/5

204/78

48/11

0.14 DI

strike-slip

PPZDI01B

Huty Fm

107

106/85

262/4

352/3

0.36 DI

extension

PPZDI01C

Huty Fm

107

98/74

226/9

318/12

0.51 DI

extension

PPZDI01D

Huty Fm

107

172/4

81/6

298/82

0.39 DI

flattening

PPZDI01E

Huty Fm

107

351/6

122/82 261/7

0.12 DI

strike-slip

PPZDI01F

Huty Fm

107

332/0

62/16 241/7

0.25 DI

flattening

PPZDI01G

Huty Fm

107

190/78

62/7

331/9

0.38 DI

extension

PPZDI01H

Huty Fm

107

126/83

2340/6

250/4

0.41 DI

extension

PPZDI02A

Huty Fm

44/77

247/11 156/4

0.31 DI

extension

PPZDI02B

Huty Fm

278/74

145/11

53/11 0.58 DI

extension

PPZDI02C

Huty Fm

61/74

314/5

222/16

–

NDA extension

PPZDI02D

Huty Fm

257/25

153/27

9/58

–

NDA strike-slip

PPZDI02E

Huty Fm

351/70

253/3

162/20

0.52 DI

extension

PPZDI02F

Huty Fm

126/73

321/17 230/4

0.49 DI

extension

PPZDI03A

Carpathian Keuper

284/38

107/52 15/1

0.62 DI

strike-slip

PPZDI03B

Carpathian Keuper

246/74

118/10 26/13

0.41 DI

extension

PPZDI03C

Carpathian Keuper

265/88

21/1

111/2

0.37 DI

extension

PPZDI03D

Carpathian Keuper

181/20

294/47

76/36 0.36 DI

extension

PPZDI03E Carpathian

Keuper 4 37

1/22

96/12 212/65 0.35

flattening

PPZDI03F Carpathian

Keuper 5 37 203/17 324/60 105/25 0.22

strike-slip

PPZDI04A

Carpathian Keuper

329/11

236/13

98/73 0.47 DI

flattening

PPZDI04B

Carpathian Keuper

31/85

206/5

296/0

0.50 DI

extension

PPZDI06A

Borové and Huty Fm 14

150/6

241/11

29/77 0.47 DI

flattening

PPZDI08A Tokáreň Mb

168/23

19/64

263/12

0.51 DI

strike-slip

PPZDI08B Tokáreň Mb

160/78

349/12 259/2

–

NDA

extension

315

CRETACEOUS—QUATERNARY EVOLUTION OF THE TATRA MOUNTAINS (WESTERN CARPATHIANS)

GEOLOGICA CARPATHICA

GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

The thick Lower Oligocene Huty Formation comprises

mostly  fine-grained  deep  water  deposits  (Fig. 3).  Its  lower
regressive part is formed by various types of dark, organic-
rich  shales  and  thick  intraformational  conglomerates  of  the
Tokáreň  Member  (Sliva  2005).  In  the  eastern  part  of  the
study area, sedimentary facies suggest a rather proximal po-
sition  (Marschalko  &  Radomski  1970;  Janočko  &  Jacko
1998; Janočko et al. 2000; Sliva 2005). The N to NE directed
paleocurrents  (Fig. 3)  together  with  clasts  (CWC  source)
from fossil submarine canyons (Tokáreň Member), outcrop-
ping  on  the  northern  slopes  of  the  Tatra  Mts,  are  clear  evi-
dence for Gemeric and Veporic sources and non-existence of
the  Tatra  Mts.  Although  in  the  Podhale  and  Orava  parts  of
the  CCPB  the  paleotransport  is  not  unequivocally  deter-
mined,  meter-sized  Mesozoic  boulders  in  thick  unsorted
breccias and conglomerates favour short transport and prox-
imity to the basin margin and the CWC source.

which most probably have been redeposited in newly formed
Neogene basins in the hinterland.

The Neogene sedimentary record is preserved in the Orava—

Nowy Targ Basin, which originated during the Middle Mio-
cene  (Watycha  1976;  Nagy  et  al.  1996).  The  basin  was
previously  considered  to  be  a retro-arc  basin  (Roth  (Ed.)
1963). However, because of its position near the Periklippen
shear  zone,  some  authors  considered  the  basin  to  be  a pull-
apart type (Pospíšil 1990; Pomianowski 2003). The basin in-
fill  is  composed  of  Burdigalian,  Langhian,  and  Serravalian
coarse-grained  sandstone,  claystone,  and  intercalations  of
lignitic claystone to lignite. The sediments were deposited in
brackish,  limnic,  and  alluvial  environments.  The  basin  also
contains  Upper  Miocene  to  Lower  Pliocene  greenish-grey
claystone and siltstone with intercalations of sandstone lying
erosively on the Middle Miocene strata (Fig. 3).

Fig. 5. Paragenetic system of the Studený potok stream (Slavkovský štít and Lom-
nický štít peak area). DS – denudation part of the system (Malá and Ve ká studená
dolina – valleys in the High Tatras), AS – accumulation part of the system (Pod-
tatranská kotlina – basin). VT-4—VT-1, VTM – sample code locations. Contour in-
terval 10 m.

The Lower Oligocene transgressive part of

the  Huty  Formation  is  formed  by  typical
mud-rich  deep  marine  fan  deposits  without
thick  sandstones  and  conglomerates  (Soták
1998;  Soták  et  al.  2001).  Sediments  of  the
Huty  Formation  gradually  change  into  re-
gressive facies of the Zuberec and Biely potok
Formations – typical deep marine flysch de-
posits (Fig. 3). Paleocurrent analysis showed
reorientation from marginal to basin axis po-
sition,  indicating  an  increase  in  tectonic  ac-
tivity  and  change  of  sedimentary  sources
(Sliva  2005).  In  the  Orava—Spišská  Magura
Mts  region,  eastward  prograding  fans  were
deposited,  while  in  the  Šarišská  vrchovina—
Levočské  vrchy  Mts  fans  prograded  west-
wards  (e.g.  Marschalko  &  Radomski  1960;
Krysiak 1976; Starek 2001; Sliva 2005).

The youngest age (NP25—NN1 nanno-

plankton  zone)  of  the  CCPB  sediments  was
reported  from  the  Biely  Potok  Formation  in
many  sites  (Gedl  2000;  Soták  &  Starek
2000;  Starek  et  al.  2000;  Soták  et  al.  2001;
Starek  2001;  Garecka  2005;  Sliva  2005).
Sedimentological  research  in  these  deposits
confirmed their submarine fan origin (Janočko
&  Jacko  1998;  Soták  1998;  Janočko  et  al.
2000;  Soták  &  Starek  2000;  Starek  et  al.
2000;  Soták  et  al.  2001;  Starek  2001;  Sliva
2005).  Gradual  shallowing  and  decrease  of
salinity  is  indicated  by  microfossil  associa-
tions  (Starek  et  al.  2000;  Soták  et  al.  2001).
Regressive,  deltaic,  and  fluvial  facies  have
not been identified due to Miocene basin in-
version  followed  by  denudation.  The  upper
limit for the regressive deposits could be de-
termined  by  the  NN2  Zone  because  a new
transgressive sedimentary cycle started at the
NN2/NN3  boundary  in  the  East  Slovakian
Basin (e.g. Soták et al. 2001). Our estimation
points  to  several  hundred  meters  of  missing
regressive  sediments  (deltaic  and  alluvial)

Table 2: Denudation rates (DR) estimated from various possible ages of oldest Qua-
ternary sediments, DH – mean denudation levelling of the denudational segment.

Area

[km

]

Volume

[km

]

DH
[m]

DR 500 kyr

[mm/a]

DR 750 kyr

[mm/a]

DR 1000 kyr

[mm/a]

Denudation segment

14.05

2.07

Accumulation segment 69.51

Preserved sediments

0.83

0.12

0.08

0.06

Minimum (recent)
storage space

3.68

262

0.52

0.35

0.26

Construed maximum
storage space
(rising 0.03 mm/a)

3.96
4.10
4.24

282
292
302

0.56

0.39

0.32

316

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GEOLOGICA CARPATHICA

GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

Geomorphological data

The main goal of geomorphological analysis was to estimate a min-

imum  Quaternary  denudation  in  the  Studený  potok  paragenetic  system
(Fig. 5).  The  estimation  was  limited  by  the  following  vagueness:
(i) storage  sedimentary  space  on  the  mountain  foothill  is  open  so  only
a fragment of former accumulation is preserved; (ii) an unknown amount
of sediments was not deposited on the foothills (suspended, bed and dis-
solved load of the Studený potok (stream) shifted directly to the Poprad
River); (iii) the glacial and glaciofluvial sediments, preserved on the high-
est watersheds of the accumulation segment (AS), were never numerically
dated  and  their  age  is  only  roughly  estimated  as  Early  Pleistocene  to
Mindel (Elsterian) (Lukniš 1968; Nemčok et al. 1993); (iv) storage space
computed from the altitude of the highest recent remnants of Quaternary
sediments is underestimated because of their denudation; (v) glacial cy-
cles could infill and empty some parts of the storage space several times,
which again leads to underestimation of the resulting denudation rates.

Denudation rates from in situ produced cosmogenic nuclides from re-

golith profiles (Cockburn & Summerfield 2004) as well as denudation
rates computed for small flat river catchments from terrace deposits (e.g.
Schaller et al. 2002; Hidy et al. 2014) generally vary about 0.01—0.1 mm/a.
A set of estimations of minimum denudation rates was done with respect
to the age of the oldest sediments (500—1000 kyr) as well as the maximum
denudation rate of their remnants (0.03 mm/a – see Table 2).

Beside this, additional relevant morphometric data were obtained

from the morphometric analysis: (i) difference between mean altitude of
denudation  segment  (1880 m a.s.l.)  and  accumulation  segment
(779 m a.s.l.)  reached  1100  and  1200 m a.s.l.  in  medians,  respectively;
(ii) depths of valley bottoms below dividing ridges predominantly vary
from 300 to 600 m; (iii) volume of denudation segment – valley is ap-
proximately  half  the  volume  of  the  accumulation  segment  –  storage
space (denudation/accumulation volume ratio is 0.56 for minimum and
0.49 for maximum storage space). It means that about half the accumu-
lation has come from the space above the present valley and denudation
of ridges is about half the general denudation.

FT data

The sampling strategy was to quantify amplitude and timing of verti-

cal  movements  in  the  Tatra  Mts,  using  the  example  of  Slavkovský  štít
(peak). In order to determine exhumation rate, an elevation profile in the
facetted fault slope of Slavkovský štít (peak) was sampled every 400 m
of altitude. Because of the possibility to record rocks from above the re-
cent peaks in the mountains and to specify information about the denu-
dation  rate  in  the  Tatra  Mts  by  this  way,  the  oldest  moraine  material
below  Slavkovský  štít  (peak)  was  studied  as  well.  Four  samples  from
the Carboniferous granitic rocks of Slavkovský štít (peak) together with
one sample from the granite boulders of Middle Pleistocene moraine in
the foothills were dated by the FT method. All samples were taken from
surface outcrops (Figs. 5 and 6; Table 3).

Zircon central ages were obtained from five samples (VT-1, VT-2,

VT-3, VT-4, and VTM) and range from 62.6 ± 6.0 to 76.8 ± 11 Ma (Ta-
ble 3, Fig. 2). All analyses passed the

-squared test (P(

) > 5 %; Gal-

braith 1981), indicating that all single grain ages statistically belong to
the same population. Unfortunately samples VT-1 and VTM yielded
a too low number of grains for further meaningful treatment although
the ages are in line with the other samples.

Apatite central ages are significantly younger than the aforementioned

zircon central ages, ranging between 9.3 ± 1.6 and 11.7 ± 1.8 Ma (Table 3,

Table 3:

ZFT

and

AFT

data

this

study.

–

number

counted

grains

per

sample,

ρρρρρ

(ρρρρρ

)

–

density

spontaneous

(induced)

tracks

(×

tr/cm

Ns,

(Ni)

–

number

counted

spontaneous

(induced)

tracks,

ρρρρρ

–

density

dosimeter

tracks

(×

tr/cm

–

number

counted

dosimeter

tracks,

χχχχχ

22222

)

–

probability

obtaining

values

for

degrees

freedom

where

number

crystals-1;

central

age

(Ma)

error

(Galbraith

Laslett

1993),

Dpar

–

measured

long

axis

etch-pits

which

are

parallel

the

ystallographic

c-axis

(Donelick

1993),

–

mean

horizontal

confined

track

lengths,

–

standard

deviation

the

sample,

(L)

–

number

horizontal

confined

tracks

measured,

–

standard

error

the

mean.

318

KRÁLIKOVÁ, VOJTKO, SLIVA, MINÁR, FÜGENSCHUH, KOVÁČ and HÓK

GEOLOGICA CARPATHICA

GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

Fig. 2). All apatite samples (VT-1, VT-2, VT-3, VT-4, and
VTM) passed the

-squared test (P(

) > 5 %; Galbraith

1981). Only sample VT-1 provided too few grains for mean-
ingful interpretation. Despite the fact that not too many fis-
sion  track  lengths  could  be  measured,  the  track  length
distributions  (TLD)  of  horizontally  confined  fission  tracks
are unimodal, negatively skewed, with fairly broad standard
deviations [SD between 1.4 (VTM sample) and 1.7 µm (VT-2
sample) Table 3] and short mean horizontally confined track
lengths  [Mean  HCTL  between  13.4  (VTM  sample)  and
12.1 µm  (VT-2  sample)  Table 3].  Such  TLD’s  are  common
in basement rocks with slow protracted cooling through the
apatite  partial  annealing  zone  (APAZ;  ~ 60—120 °C;  e.g.
Wagner & Van den Haute 1992). In all of the analysed apa-
tites, Dpar values do not vary much, ranging between 1.1 and
1.6 µm (Table 3), indicating fairly similar chemical composi-
tions and relatively fluorine rich apatites (Burtner et al. 1994)
with a low resistance to annealing (Ketcham et al. 1999).

Interpretation and discussion

The first tectonic stage (TS-1; ~ 95—80 Ma)

TS-1 reflects mid-Cretaceous nappe stacking originated

during the northwards progressing paleo-Alpine Cretaceous
convergence (Fig. 7). Within the working area detached cover
nappes formed in the northern frontal part of the accreted Car-
pathian system (e.g. Plašienka et al. 1997; Jurewicz 2005).

Generally,  NW  directed  thrusting  propagated  from  the  inner
CWC outwards into the foreland lower plate (Plašienka 2003;
Prokešová  et  al.  2012).  Deformation  occurred  under  brittle-
ductile  to  brittle  conditions  with  consistent  top-to-the-north-
west shear (Fig. 8a,b). During this stage the Tatric Unit took
a lower  position  in  the  nappe  stack  and  was  metamorphosed
under  anchizonal  and/or  lower  greenschist  facies  conditions
(Fig. 8b). The fully annealed zircon samples indicate that the
metamorphic  temperature  was  in  excess  of  ~ 320 °C  (pro-
posed upper temperature limit of the zircon partial annealing
zone  –  ZPAZ;  Tagami et al. 1998). The peak metamorphic
temperature  had  to  be  less  than  ~ 350 °C  within  the  Tatric
crystalline  basement,  because  the

Ar/

Ar and Rb/Sr on

muscovite and biotite methods yielded isotopic age values for
the granitic rocks in the range of 330—280 Ma. These ages
most probably record magmatic cooling followed by exhuma-
tion of the granite pluton during the Late Variscan orogenesis
and they are not affected by the Alpine tectogenesis (Janák
& Onstott 1993; Maluski et al. 1993; Janák 1994).

The second tectonic stage (TS-2; ~ 80—45 Ma)

After the tectonic burial beneath the paleo-Alpine nappe

pile, an exhumation by unroofing of the Tatric Unit occurred
during the Late Cretaceous—Paleocene (Fig. 7). Exhumation
processes occurred predominantly on the low angle normal
faults where the principal tension axis operated in the gener-
ally W—E direction with general top-to-the-east shear (see
Fig. 3). In the western part of the mountains, tectonic contact

Fig. 7. Summarizing time-temperature evolution of the Tatra Mts crystalline basement drawn
from regional geological considerations (main tectonic events and tectonic stages recorded in the
study area are depicted at the top of the diagram), from results of ZFT and AFT data, and addi-
tional published thermochronological data. ZPAZ – zircon partial annealing zone, APAZ – ap-
atite partial annealing zone; double dot-dashed line represents idealized fit for the tectono-thermal
evolution of the Tatric crystalline basement.

of  the  Hronic  and  extremely  re-
duced  Fatric  units  vs.  the  Tatric
crystalline basement along the low-
angle  faults  is  considered  to  be
a consequence  of  this  exhumation
process  but  the  kinematics  are  al-
most opposite.

The Late Cretaceous—Early Paleo-

cene cooling of the Tatric basement
from the depth of ca.  ~ 10 km was
also revealed by the presented ZFT
data  with  the  ages  scattered  be-
tween  ~ 60  and  70 Ma.  Due  to  the
late-Eocene transgression it is obvi-
ous  that  at  least  the  western  Tatric
basement  cooled  from  Cretaceous
T

max

to surface temperatures. Only

because  of  reheating  in  the  course
of formation of the CCPB, the AFT
samples  do  not  record  this  early
phase  of  cooling  (e.g.  Krá   1977;
Anczkiewicz 2005 and data herein)
except at least two detrital samples
in  the  Podhale  Basin  (5/02  and
22/02)  with  AFT  peak  ages  of
55—60 Ma (cf. Anczkiewicz 2005).
Formation  of  the  CCPB  caused
substantial  burial  of  the  Tatric
basement and total resetting of the
AFT  system  in  the  east.  However,

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GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

partial resetting and thus less overburden may probably be
indicated by older (mixed?) AFT ages in the Western Tatra
Mts (Burchart 1972; Struzik et al. 2002 and Anckiewicz
2005 – see samples T3 and 9/01).

Basically, these exhumation and denudation processes

were recorded in the whole of the Tatra Belt as well as in the
Veporic  and  Gemeric  units.  Published  ZFT  data  from  the
Tatra  Belt  yielded  ages  of  ~ 70 Ma  in  the  Považský  Inovec
Mts (Kováč et al. 1994),  ~ 53 Ma in the Tribeč Mts (Kováč
et al. 1994), and  ~ 45 to 70 Ma in the Malá Fatra Mts and the
basement of the Turiec Basin (Danišík et al. 2010; Králiková
2013;  Králiková  et  al.  2014).  In  the  southern  zone  of  the
CWC,  the  Veporic  Unit  yielded  ZFT  age  of  ~ 65—75 Ma
(Plašienka  et  al.  2007;  Králiková  2013;  Vojtko  et  al.  2013)
and the Gemeric Unit provided ZFT ages between  ~ 62 and
88 Ma (Plašienka et al. 2007). According to the ZFT dating
results, the first conspicuous exhumation processes after the
Alpine  nappe  stacking  definitely  took  place  mainly  during
the Late Cretaceous—Paleocene. It is evident, that this exhu-
mation,  well-known  in  the  Veporic  Unit  (e.g.  Hók  et  al.
1993;  Plašienka  1993,  1999;  Jeřábek  et  al.  2007,  2008,
2012), prograded to the external, Tatra Unit.

After the exhumation the hanging wall leftovers were mark-

edly eroded and a last remnant of the Hronic Unit is preserved
in  the  north-western  edge  of  the  Tatra  Mts.  In  the  Western
Tatra  Mts  (Košariská  and  Mačacie  diery  sites),  erosion  was
even more efficient and Upper Eocene strata transgressed onto
the pre-Mesozoic Tatric basement (cf. Passendorfer 1958; An-
drusov 1959; Gorek & Scheibner 1966; Nemčok et al. 1994).
It  should  be  stressed  that  some  authors  interpret  the  contact
between Upper Eocene strata and granite in the Košariská site
as tectonic origin (e.g. Uhlig 1897; Lugeon 1903). However,
this contact lacks any tectonic reworking. Scarce paleo-current
markers  (towards  west  to  north-west),  the  Borové  Formation
clasts  and  granite  pebbles  in  the  stratotype  site  of  the  Pucov
Member conglomerates (Priabonian) point to a source area in
the  Western  Tatra  Mts  during  the  Priabonian.  In  the  central
and  eastern  portion  of  the  Tatra  Mts  such  direct  evidence  of
the Tatric crystalline basement erosional level has never been
recorded (Fig. 8c). Moreover, there are no indications of Tat-
ric  crystalline  pebbles  in  the  Borové  Formation  and  Tokáreň
Member so far (e.g. Janočko et al. 2000; Sliva 2005).

The third tectonic stage (TS-3; ~ 45—20 Ma)

TS-3 is linked to the formation of the CCPB. It shows

a fore-arc  basin  position  developed  on  the  destructive  plate-
margin of the ALCAPA microplate and behind the EWC ac-
cretionary wedge (Tari et al. 1993; Soták et al. 2001; Kázmer
et al. 2003). During the sedimentation of the CCPB sequences,
the former area of the Tatra Mts was buried under the relatively
thick  Middle  Eocene—lowermost  Miocene  strata  (Figs. 7  and
8d). The known thickness of the CCPB sediments is  ~ 3.5 km
in  the  Levočské  vrchy  Mts  (Gross  et  al.  1999),  ~2.25 km  in
the Spišská Magura Mts (Janočko et al. 2000),  ~ 3.0 km in the
Podhale Basin (Kępińska 1997),  ~ 1.5 km in the Liptov Basin
(e.g.  Gross  et  al.  1980),  and  ~ 2.5 km  in  the  Orava  region
(Fusán et al. 1987). At present, the Liptov Basin has the mini-
mum thickness of the CCPB strata among the portions of the

CCPB. This is apparently caused by its more intense uplift fol-
lowed by erosion during the Neogene and Quaternary. In addi-
tion, the known paleotransports documented in the Oligocene
sequences (in the Orava, Liptov, Podhale, and Spišská Magura
Mts) point to eastward transport of sedimentary material fed
by  gravity  flows  (Sliva  2005).  From  the  sedimentological
standpoint,  the  Oligocene  sequences  of  the  Orava  region,
Liptov  and  Podhale  basins,  and  Spišská  Magura  Mts  repre-
sent  one  submarine  fan  system  (Fig. 8d).  It  means  that  the
Tatra  region  was  completely  overburdened  by  the  CCPB
sediment pile. The inferred thickness of the eroded and pre-
served section varies from  < 3.1—3.9 km in the Orava region
to  > 5.5—6.9 km  in  the  Spišská  Magura  Mts  close  to  the
Ružbachy fault, calculated for 20 and 25 °C/km paleogradi-
ents (e.g. Środoń et al. 2006). On the basis of these data, the
total  thickness  of  the  CCPB  sedimentary  sequences  in  the
Tatra  region,  eroded  and  preserved,  can  be  estimated  as
~

4.0—5.0 km. Thus, a combined tectonic and sedimentary to-

tal overburden of the Tatric basement on the order of  ~ 4.0 km
in the west and up to 7.0 km in the eastern portion of the study
area can be assumed after the CCPB formation, as was depicted
also by K-Ar dating on bentonite at  ~ 16—19 Ma (Środoń et al.
2006).  With  an  assumed  geothermal  gradient  of  ~ 25 °C/km
this leads to metamorphic temperatures in the Tatric basement
(except  the  west  with  slightly  lower  temperatures)  between
160  and  220 °C.  Published  zircon  U-Th/He  (ZHe)  data
(Śmigielski et al. 2012) and the presented AFT data are well in
line  with  this  estimate.  Predominantly  there  are  two  pre-
ferred models: (i) on the basis of X-ray diffraction study of
illite combined with K-Ar dating of bentonites, Środoń et al.
(2006)  postulate  burial  beneath  variable  thickness  of  Ceno-
zoic sediment pile ( ~ 4 km in the west to  ~ 7 km in the east of
the study area) at the stable geothermal gradient of 21 ± 2 °C/km
as the only heat source. The authors evaluation of the burial
depth  is  independently  supported  by  fluid-inclusion  data
published by Hurai et al. (2002) that refer to burial depth of
5.3—6.5 km  (130—205 °C)  close  to  the  Ružbachy  fault,  and
2—3 km (95—100 °C) in the Orava territory. Besides, Środoń
et  al.  (2006)  published  additional  evidence  of  burial  as  the
only heat source that was revealed by measurements of dif-
ferent  grain  densities  and  porosities  in  the  Bukovina  and
Chocholow wells in the Podhale Basin. According to the au-
thors,  these  differences  respond  primarily  to  the  lithostatic
load and not to the temperature; (ii) according to AFT dating
study from the Podhale—Spišská Magura Basin, Anczkiewicz
et  al.  (2013)  assume  both  heating  associated  with  mid-Mio-
cene volcanism and variable thickness of Oligocene- and po-
tentially  also  Miocene  sediments  as  the  responsible  heat
source.  Their  interpretation  shifts  the  extent  of  the  mid-Mio-
cene  thermal  episode  to  the  northern  edge  of  the  Pannonian
block and thus provides an extra heat source. According to the
authors, the additional heat source could be linked with the el-
evated  paleogeothermal  gradient  ( > 25 °C/km)  estimated  by
Marynowski  &  Gaweda  (2005)  and  Kepińska  (2006),  on  the
basis of vitrinite reflectance and illite-smectite data.

On the basis of our research together with known geologi-

cal  knowledge  we  inclined  to  the  first  model  presented  by
Środoń  et  al.  (2006).  For  many  reasons,  it  is  disputable  to
consider  the  interpretation  of  ‘the  mid-Miocene  thermal

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event as the extra heat source that prograded to the northern
edge  of  the  Pannonian  block’  proposed  by  Anczkiewicz  et
al. (2013). One of the arguments is the lack of evidence that
this  thermal  event  affected  the  Tatra  Mts  crystalline  base-
ment. This is documented by a big scatter in AFT ages that
were  not  completely  reset  in  the  Middle  Miocene,  or  even
later. Another important fact is that in the areas close to the
huge volcanic edifices, such as the Tribeč, Žiar, and Vepor-
ské  vrchy  Mts,  the  AFT  system  was  not  influenced  by  this
thermal event, providing the ages older than Miocene or Oli-
gocene  (see  AFT  data  in  the  works  of  Danišík  et  al.  2004,
2008;  Plašienka  et  al.  2007;  Králiková  2013;  Vojtko  et  al.
2013).  Therefore,  it  is  very  unlikely  that  the  mid-Miocene
thermal  event  would  affect  the  very  remote  and  relatively
cold periphery of the CWC without influencing immediately
surrounding  areas  with  the  main  volcanic  activity  (Central
Slovak Neovolcanic Field region). Additionally, our new ZFT
data of Late Cretaceous to Paleocene age, that were not influ-
enced  by  an  elevated  paleogeothermal  gradient  ( > 25 °C/km)
in the Middle Miocene, may be regarded as further circum-
stantial evidence. If we assume  ~ 7 km of total overburden of
the Tatra Mts paleo-Alpine basement after the formation of
the CCPB in the Early Miocene and the Mesozoic strata both
nappe and cover sequences, then a paleogeothermal gradient
of 30 °C/km, or even more, should at least partially or com-
pletely  reset  the  ZFT  system  in  the  eastern  part.  However,
we missed this evidence. The same argument applies to the
ZHe data published by Śmigielski et al. (2012). On the basis
of these arguments it is necessary to look for another inter-
pretation because the paleogeothermal gradient had to be in
the range of 20—25 °C/km. For these reasons, the most likely
interpretation  appears  to  be  the  combination  of  burial  be-
neath  the  CCPB  strata  and  subsequent  exhumation  in  the
Miocene as was described in detail in the work of Králiková
et al. (2014).

The fourth tectonic stage (TS-4; ~ 20—7 Ma)

TS-4 is characterized by exhumation of the Tatric crystal-

line basement after the deposition of the Oligocene—lower-
most Miocene strata followed by a quiet period predominantly
during the Late Miocene to Early Pliocene (Fig. 7). The mea-
sured brittle structures in the CCPB sediments indicate that the
exhumation process was controlled by a compressional tec-
tonic regime (Early Miocene), followed by a compressional to
strike-slip tectonic regime (Middle to Late Miocene) and fi-
nally by an extensional tectonic regime (Pliocene to Quater-
nary). The paleostress field progressively changed, where the
S

Hmax

relatively rotated from NW—SE through N—S to NE—SW

directions (Fig. 3; for further information and detailed paleo-
stress  data  see  Pešková  et  al.  2009;  Vojtko  et  al.  2010;
Sůkalová et al. 2012). The obtained data helped us to under-
stand  the  tectonic  processes  during  the  exhumation,  denuda-
tion, and formation of the modern Tatra Mts relief. The Tatra
Mts are an asymmetrical horst where the Mesozoic sedimen-
tary cover and nappe units occur on the northern slopes with
the moderate to steep inclination of bedding planes northward
(Fig. 8e).  The  principal  role  during  the  exhumation  of  the
mountains is ascribed to the W—E striking sub-Tatra fault with

its several cross-cutting and arching fault segments such as the
Ružbachy  fault  on  the  east  and  Prosečné  (a.k.a.  Choč-
Prosečné-Krowiarski)  fault  on  the  west.  The  results  of  fault-
slip analysis and paleostress reconstruction make it possible to
interpret  the  kinematics  of  the  main  fault  zones  participating
in the Tatra Mts structure.

Previous models of the Tatra Mts exhumation/uplift were

predominantly based on map-scale geological data. Generally,
five models were used in the past: (i) exhumation happened on
the W—E sub-Tatra reverse fault with north inclination as was
accepted predominantly during the first half on the 20

cen-

tury  (e.g.  Sokolowski  1948;  Gorek  1956;  Andrusov  1958,
1968;  Mahe   et  al.  1967;  Biely  &  Fusán  1967;  Piotrowski
1978);  (ii) the  exhumation  occurred  on  the  sub-Tatra  normal
fault. This interpretation was supported by technical and geo-
physical works (Gross 1973; Gross et al. 1980) where the fault
was interpreted as subvertical, or even a steeply south-dipping
fault  plane.  However,  this  fault  plane  orientation  was  recog-
nized only on its accompanying subsidiary faults. This model
was widely accepted especially in the second half of the 20

century  (e.g.  Mahe   1986;  Nemčok  et  al.  1993);  (iii) a  new
conceptual and kinematically consistent model was proposed
by Sperner (1996) and Sperner et al. (2002). The model was
based  on  measurement  of  brittle  deformation  in  the  broader
area of the Tatra Mts. The main displacement on the sub-Tatra
fault was considered to be reverse southward movement with
strike-slip  combination  on  the  Ružbachy  and  Prosečné  faults
and  the  Tatra  massif  was  uplifted  as  a compressional  horst
structure with a combination of strike-slip duplex (Ružbachy
horst); (iv) the sub-Tatra fault is a part of planar normal faults
which were produced by tilt and spin rotations of domino-type
prismatic  upper  crustal  blocks  that  formed  due  to  horizontal
top-to-the  north  simple  shear  of  the  CWC  crust  triggered  by
underthrusting  of  the  Northern  European  plate  (Grecula  &
Roth  1978;  Marko  1995);  (v) the  last  proposed  model  of  the
Tatra massif exhumation as a  compressive structure involving
the basement was published by Janák et al. (2001) according
to  whom  the  sub-Tatra  fault  operated  as  a frontal  subvertical
fault  and  the  fault  bend  syncline  was  located  in/along  the
northern foot of the Tatra Mts.

Interpretation of fault-slip analysis and paleostress recon-

struction  can  help  us  to  understand  the  kinematics  of  main
fault shear zones in the Tatra region (Figs. 3 and 8). The Early
Miocene paleostress was caused by a compressional tectonic
regime with the NW—SE principal compression, and exhuma-
tion of the Tatra Mts started most probably along the reverse
Ružbachy fault and transpressive to oblique reverse sub-Tatra
fault.  The  uplift  started  as  a  compressive  structure  involving
the basement as was proposed by Janák et al. (2001). During
the  Middle  Miocene,  the  compressional  axis  rotated  to  the
generally N-S direction and the Ružbachy fault operated as a
sinistral  strike-slip  fault  (transform  fault)  and  the  sub-Tatra
fault  as  a  reverse  fault.  Exhumation  was  carried  out  by
a combination  of  strike-slip  movement  and  reverse  faulting
(Sperner 1996; Sperner et al. 2002). The Late Miocene paleo-
stress field can be characterized by the NE-SW oriented prin-
cipal  compressional  axis  (e.g.  Sperner  1996;  Pešková  et  al.
2009; Vojtko et al. 2010; Sůkalová et al. 2012) and the kine-
matics  of  the  Ružbachy  and  Prosečné  faults  successively

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changed  from  transpression  via  transtension  up  to  normal
faulting. Along the Ružbachy fault, the hanging wall was lo-
cated  in  the  SE  side  with  down-throw  movement  and  along
the Prosečné fault, the hanging wall down-throw north-west-
ward.  The  Tatra  Mountains  started  to  be  an  individualized
horst structure in the northern portion of the CWC. During this
stage, the reverse movement of the sub-Tatra fault changed to
oblique  transpressive  sinistral  strike-slip  up  to  transtensive
sinistral  strike-slip  at  the  end  of  the  Late  Miocene.  The
Pliocene  to  Quaternary  period  is  characterized  by  an  exten-
sional  tectonic  regime  with  general  principal  tension  in  the
NW-SE  direction.  The  orientation  of  the  paleostress  field
caused  predominant  normal  faulting  along  the  sub-Tatra  and
Ružbachy faults where the south block subsided. Most authors
agree  that  the  Prosečné,  sub-Tatra,  and  Ružbachy  faults,
which  restricted  the  asymmetric  horst-like  structure  of  the
Tatra  Mts,  operated  as  normal  faults  during  the  neotectonic
phase (e.g. Mahe  1986; Nemčok et al. 1993).

The timing of the Middle to early Late Miocene exhumation

processes of the Tatra Mts crystalline basement was revealed
by two independent thermochronological methods, using AFT
and  apatite  U-Th/He  (AHe)  analysis  (Burchart  1972;  Krá
1977; Struzik et al. 2002; Anczkiewicz 2005; Anczkiewicz et
al. 2005; Śmigielski et al. 2012 – Fig. 2). The presented new
AFT ages in the range of 9.3 ± 1.6 and 11.7 ± 1.8 Ma (Figs. 2
and 6) from the Slavkovský štít (peak) area are in good agree-
ment  with  the  previous  published  AFT  data  from  the  Tatra
crystalline basement (Krá  1977; Anczkiewicz et al. 2005), re-
vealing  the  late  Middle  to  early  Late  Miocene  exhumation
from the depth of at least 5 km (recalculated by the geother-
mal gradient at 20—25 °C/km). The elevation profile in Slav-
kovský  štít  (peak)  allowed  us  to  determine  an  approximate
exhumation rate of ca. 0.35±0.1 mm/a. The result is consistent
with  obtained  track  length  measurements  of  identical  short
Mean HCTL and fairly broad SD (VT-2—VT-4 samples), indi-
cating simple continuous moderate cooling of the Tatric crys-
talline  rocks  through  the  APAZ  at  9.3±1.6  to  11.7 ± 1.8 Ma
(Fig. 6). It is clear that no particular event occurred in this time
span.  A  similar  exhumation  rate  (0.2 ± 0.1 mm/a)  was  deter-
mined  in  the  Mount  Rysy,  High  Tatra  Mts,  during  the  Mio-
cene  (Anckiewicz  et  al.  2005).  The  revealed  early  Late
Miocene slow exhumation rate can be related to the initial pla-
nation  surface  formation  of  the  modern  relief  (intramountain
level)  in  the  Western  Carpathians  (e.g.  Mazúr  1965;  Starkel
1969; Minár et al. 2011; Zuchiewicz 2011). This period was
replaced by accelerated uplift during the latest Late Miocene
to  Pleistocene  (Fig. 8f).  According  to  Baumgart-Kotarba  &
Krá   (2002),  the  uplift  from  the  ~ 2 km  depth  ( ~ 60°)  most
probably started in the time span of  ~ 7—2 Ma. Moreover, the
authors postulate an uplift rate of up to 1.0 mm/a for this time
range because of a short time to uncover the Tatric crystalline
basement of the High Tatra Mts. Additionally, sedimentologi-
cal  data  confirm  the  outlined  scenario.  After  Tokarski  et  al.
(2012) all the Neogene gravels of the Miocene—Pliocene sedi-
ments of the Orawa—Nowy Targ Basin are devoid of clasts de-
rived from the Tatra Mts where no prominent relief existed at
that time. To the contrary, the Pieniny Klippen Belt was sub-
ject to denudation and had more considerable relief then now.
The Late Miocene relatively quiet tectonic period and low re-

lief in the Tatra Mts is documented by the oldest fine-grained
cave sediments in the Belianska cave, dated to ~6.15—4.18 Ma
(Bella et al. 2011). Similar results were obtained from speleo-
them  dating  in  the  Polish  Tatra  caves,  indicating  their  oldest
denudation  surface  as  latest  Miocene  or  younger  in  age
(Głazek 1996). The upper limit of denudation surface forma-
tion  ( > 1.2 Ma)  was  gained  by  study  of  the  evolution  of  the
phreatic stage in the Tatra caves (Gradziński et al. 2009).

The neotectonic stage ( ~ 7—0 Ma)

The impressive morphology of the Tatra relief points to

a new  acceleration  in  tectonic  activity  during  the  latest  Late
Miocene  to  Pleistocene.  Recent  research  on  the  sedimento-
logical record from the Orava—Nowy Targ Intramontane Basin
supports this view. The research showed a considerable differ-
ence between the Neogene (fine-grained) and the Quaternary
(coarse-grained)  deposits  (Tokarski  et  al.  2012).  Neotectonic
fault  activity  is  also  proved  by  significant  evolution  of  cal-
careous tufa and travertine mounds in the Polish (e.g. Mastella
& Rybak-Ostrowska  2012)  and  also  in  the  Slovak  (e.g.
Sůkalová et al. 2012) parts of the Tatra Mts area. The young-
est  story  was  indirectly  documented  by  granite  boulders  of
Middle  Pleistocene  moraine  located  at  the  foot  of  the  High
Tatra  Mts.  The  granite  boulders  revealed  significantly  older
AFT age ( ~ 11.0 ± 1.4 Ma, VTM sample; Table 3) than the age
of  deposition.  This  AFT  age  refers  to  the  simple  moderate
cooling event of Slavkovský štít (peak) during the latest Mid-
dle to early Late Miocene, documented by the same age and
track length measurement records as in the VT-2 to VT-4 sam-
ples.  In  addition,  the  lag  time  between  the  AFT  age
(11.0 ± 1.4 Ma) inherited from the Tatra massif and the time of
deposition of the VTM sample (Middle Pleistocene) indicates
an acceleration of exhumation rate during this time span. This
premise  indicates  that  an  accelerated  exhumation  rate  of  the
Tatra  massif  must  have  occurred  ~ 6.5—1.0 Ma.  Beside  this,
slip  rate  analysis  of  the  Vikartovce  fault  which  is  parallel  to
the  sub-Tatra  fault  close  to  the  Tatra  Mts  provided  dip-slip
movement  along  the  fault  ~ 1 mm/a  during  the  Late  Pleis-
tocene (Vojtko et al. 2011).

Holocene (postglacial) denudation of the High Tatra Mts

ridges  was  estimated  from  the  Holocene  sedimentary  infill
volume  of  17 valleys.  Statistically  consistent  data  lead  to  a
mean lowering of the ridges by  ~ 5 m for the Studený potok
valley (Lukniš 1968). It is equivalent to a mean denudation
rate  of  ~ 0.5 mm/a.  From  the  denudation/accumulation  vol-
ume  ratio  (0.56  –  see  chapter  4.3)  of  the  Studený  Potok
paragenetic system and considering that a significant part of
the removed rocks is not reflected in the volume of accumu-
lation storage space, an estimate of the Holocene denudation
rate of  ~ 1 mm/a can be assumed.

U-series dating of cave speleothems (Gradziński et al. 2009)

points  to  a  mean  rate  of  valley  deepening  of  0.2—0.3 mm/a
during the last 200 kyr in the northern part of the Tatra Mts.
However,  this  significant  discordance  can  be  eliminated  by
several  factors:  (i) N—S  asymmetry  of  the  Tatra  Mts  uplift
could lead to smaller deepening of the valleys located on the
north side; (ii) a fundamental part of the denuded ridge mate-
rial remained in the valleys during the interglacial times, but

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GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

was  later  removed  by  glaciers  during  the  next  glacial);
(iii) the  intensity  of  interglacial  ridge  denudation  was  in-
creased  by  the  effect  of  rapid  denudation  in  the  paraglacial
stage; (iv) more intensive widening of valleys in comparison
with  their  deepening  in  the  last  200 kyr  could  play
a significant  role.  On  the  basis  of  the  aforementioned  data,
a minimum  denudation  rate  ~ 0.5 mm/a  is  considered  to  be
most presumable since the Middle Pleistocene.

Conclusions

The Tatra Mts and their surroundings underwent

a complex Alpine tectonic evolution, which can be divided
into several tectonic stages, based on structural, sedimentary,
geomorphic, and FT data:

A first tectonic stage (TS-1; ~ 95—80 Ma) can be dated

back  to  the  late  Early  Cretaceous  and  is  coeval  with  nappe
stacking  (Fig. 7).  At  this  time,  the  Tatric  crystalline  base-
ment was buried below the paleo-Alpine – Fatric and Hronic
units  (Fig. 8b).  The  fully  annealed  zircon  samples  suggest
that the metamorphic temperature was in excess of  ~ 320 °C
(Fig. 7);

The principal compressional stage of the Alpine orogene

was  replaced  by  the  Late  Cretaceous  to  Paleogene  orogene
collapse followed by an orogen-parallel extension (Fig. 8c),
revealed by structural and ZFT data of  ~ 60—70 Ma (Figs. 2
and 7). The extensional tectonics were replaced by transpres-
sion  to  transtension  (Fig. 3)  during  the  Late  Paleocene  to
Eocene period (TS-2;  ~ 80—45 Ma);

The Late Eocene to Earliest Miocene can be character-

ized by formation of a marginal sea of the Peri-Tethyan Basin
represented by the CCPB (Fig. 8d). It shows a fore-arc basin
position developed on the destructive plate-margin and be-
hind the EWC accretionary wedge. During the sedimentation
of the CCPB sequences, the crystalline basement of the Tat-
ric Unit was buried in between the APAZ and ZPAZ (TS-3;
~ 45—20

Ma; Fig. 7);

The final cooling of the Tatra massif was a result of

asymmetric  neotectonic  exhumation  that  could  be  linked
with the sub-Tatra faulting in the southern edge of the massif
during  the  Neogene  (TS-4;  ~ 20—7 Ma;  Figs. 6,  7  and  8e).
The exhumation processes of the Tatric crystalline basement
from  the  depth  of  ~ 5 km  were  dated  to  the  Middle/early
Late  Miocene,  according  to  AFT  and  AHe  data  (Burchart
1972;  Krá   1977;  Struzik  et  al.  2002;  Anczkiewicz  2005;
Anczkiewicz  et  al.  2005;  Śmigielski  et  al.  2012;  and  data
herein). Additionally, the early Late Miocene evolution was
characterized by formation of the basic Western Carpathian
planation surface – intramountain level;

It can be stated that the final appearance of the moun-

tains  range  in  morphology  above  the  surrounding  foreland
was a consequence of a new acceleration in tectonic activity
during  the  latest  Late  Miocene  to  Pleistocene  (neotectonic
stage – TS-5;  ~ 7—0 Ma; Figs. 7 and 8f). On the basis of the
geomorphological  data,  a denudation  rate  of  ~1 mm  can  be
considered  for  the  whole  Quaternary,  whereas  a minimum
denudation  rate  of  ~ 0.5 mm/a  can  be  assumed  since  the
Middle Pleistocene (Table 2).

Acknowledgment: This publication is the result of the pro-
ject  implementation:  Comenius  University  in  the  Bratislava
Science  Park  supported  by  the  Research  and  Development
Operational  Programme  funded  by  the  ERDF  Grant  No.:
ITMS 26240220086. The work was financially supported by
the Slovak Research and Development Agency under the con-
tracts Nos. APVV-0625-11, APVV-0099-11, APVV-0315-12,
APVV ESFEC-0006-07, by the Grant Agency of the Czech
Republic  No.  GAČR:13-15123S  and  by  the  VEGA  agency
under contracts No. 1/0193/13. B. Fügenschuh acknowledges
financial support by the Austrian Science Fund (FWF) Grant
No. I 138-N19. Thanks are due to reviewers for their accu-
rate and constructive reviews.

References

Anczkiewicz A. 2005: Verification of maximum paleo-temperatures on

the basis of smectite illitization estimated for the Tatra Mts.,
Podhale Basin and adjacent area of the External Carpathians using
fission track method. PhD. Thesis, Inst. Nauk Geol. Pol. Akad.
Nauk, Osrodek Badawczy, Krakow, 123 (in Polish).

Anczkiewicz A., Zattin M. & Środoń J. 2005: Cenozoic uplift of the

Tatras and Podhale basin from the perspective of the apatite fis-
sion track analyses. Pol. Towarzystwo Mineralogiczne – Prace
Specjalne 25, 261—264.

Anczkiewicz A.A., Środoń J. & Zattin M. 2013: Thermal history of

the Podhale Basin in the internal Western Carpathians from the
perspective of apatite fission track analyses. Geol. Carpathica
64, 2, 141—151.

Andrusov D. 1958: Geology of the Czechoslovakian Carpathians. I. 1

edition. Vydav. SAV, Bratislava, 304 (in Slovak).

Andrusov D. 1959: Geology of the Czechoslovakian Carpathians. II.

edition. SAV Publ., Bratislava, 375 (in Slovak).

Andrusov D. 1968: Grundriss der Tektonik der Nördlichen Karpaten.

Vydav. SAV, Bratislava, 188.

Andrusov D., Bystrický J. & Fusán O. 1973: Outline of the structure

of the West Carpathians. Guide book, X Congress CBGA, D. Štúr
Institute of Geology, Bratislava, 5—45.

Angelier J. 1984: Tectonic analysis of fault slip data sets. J. Geophys.

Res. 89, B7, 5835—5848.

Angelier J. 1990: Inversion of field data in fault tectonics to obtain the

regional stress. III. A new rapid direct inversion method by ana-
lytical means. Geophys. J. Int. 103, 2, 363—376.

Angelier J. 1994: Fault slip analysis and paleostress reconstruction. In:

Hancock P.L. (Ed.): Continental deformation. Pergamon Press,
University of Bristol (U.K.), London, 53—100.

Baumgart-Kotarba M. & Krá J. 2002: Young tectonic uplift of the

Tatra Mts (Fission track data and geomorphological arguments).
Proceedings of XVII. Congress of Carpathian-Balkan Geological
Association. Geol. Carpathica, Spec. Issue, 53 (CD version).

Bella P., Bosák P., Pruner P., Głazek J. & Hercman H. 2011: The de-

velopment of the River Biela Valley in relation to the genesis of
the Belianska Cave. Geogr. Čas. 63, 4, 369—387.

Bezák V., Broska I., Ivanička J., Reichwalder P., Vozár J., Polák M.,

Havrila M., Mello J., Biely A., Plašienka D., Potfaj M., Konečný
V., Lexa J., Kaličiak M., Žec B., Vass D., Elečko M., Janočko J.,
Pereszlényi M., Marko F., Maglay J. & Pristaš J. 2004: Tectonic
map of Slovak Republic. MŽP SR – ŠGÚDŠ, Bratislava.

Biely A. & Fusán O. 1967: Zum Problem der Wurzelzonen der subta-

trischen Decken. Geol. Práce, Spr. 42, 51—64.

Burchart J. 1972: Fission-track age determination of accessory apatite

from the Tatra mountains, Poland. Earth Planet. Sci. Lett. 15,
418—422.

Burtner R.L., Nigrini A. & Donelick R.A. 1994: Thermochronology of

Lower Cretaceous source rocks in the Idaho-Wyoming thrust belt.
Amer. Assoc. Petrol. Geol. Bull. 78, 10, 1613—1636.

324

KRÁLIKOVÁ, VOJTKO, SLIVA, MINÁR, FÜGENSCHUH, KOVÁČ and HÓK

GEOLOGICA CARPATHICA

GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

Cockburn H.A.P. & Summerfield M.A. 2004: Geomorphological ap-

plications of cosmogenic isotope analysis. Prog. Phys. Geogr.
28, 1—42.

Danišík M., Kohút M., Dunkl I. & Frisch W. 2008: Thermal evolution

of the Žiar Mountains basement (Inner Western Carpathians, Slo-
vakia) constrained by fission track data. Geol. Carpathica 59,
19—30.

Danišík M., Kohút M., Broska I. & Frisch W. 2010: Thermal evolu-

tion of the Malá Fatra Mountains (Central Western Carpathians):
insights from zircon and apatite fission track thermochronology.
Geol. Carpathica 61, 1, 19—27.

Danišík M., Dunkl I., Putiš M., Frisch W. & Krá J. 2004: Tertiary

burial and exhumation history of basement highs along the NW
margin of the Pannonian Basin – an apatite fission track study.
Austrian J. Earth Sci. 95/96, 60—70.

Delvaux D. & Sperner B. 2003: New aspects of tectonic stress inver-

sion with reference to the TENSOR program. In: Nieuwland
D.A. (Ed.): New insights into structural interpretation and model-
ling. Geol. Soc. London, Spec. Publ. 212, 75—100.

Delvaux D., Moeys R., Stapel G., Petit C., Levi K., Miroshnichenko

A., Ruznich V. & San’kov V. 1997: Paleostress reconstructions
and geodynamics of the Baikal region, Central Asia. Part 2. Cen-
ozoic rifting. Tectonophysics 282, 1—38.

Donelick R.A. 1993: Apatite etching characteristics versus chemical

composition. Nucl. Track Radiat. Meas. 21, 4, 604.

Dunkl I. 2002: TRACKKEY: a Windows program for calculation and

graphical presentation of fission track data. Comput. and Geosci.
28, 2, 3—12.

Etchecopar A., Vasseur G. & Daignieres M. 1981: An inverse problem

in microtectonics for the determination of stress tensor from fault
striation analysis. J. Struct. Geol. 3, 1, 51—65.

Froitzheim N., Plašienka D. & Schuster R. 2008: Alpine tectonics of the

Alps and Western Carpathians. In: McCann T. (Ed.): The geology
of Central Europe. Geol. Soc. London 1141—1232 (Abb. 2—8).

Fusán O., Biely A., Ibrmajer J., Plančár J. & Rozložník L. 1987: Base-

ment of the Tertiary of the Inner West Carpathians. GÚDŠ, Bra-
tislava, 1—123 (in Slovak).

Galbraith R.F. 1981: On statistical models for fission track counts.

Math. Geol. 13, 6, 471—478.

Galbraith R.F. & Laslett G.M. 1993: Statistical models for mixed fis-

sion track ages. Nuclear Tracks 21, 459—70.

Garecka M. 2005: Calcareous nannoplankton from the Podhale Flysch

(Oligocene—Miocene, Inner Carpathians, Poland). Stud. Geol. Pol.
124, 353—369.

Gedl P. 2000: Biostratigraphy and paleoenvironment of the Podhale

paleogene (Inner Carpathians, Poland) in the light of palynologi-
cal studies. Part I. Stud Geol. Pol. 117, 69—154.

Gleadow A.J.W. 1981: Fission track dating methods: what are the real

alternatives? Nuclear Tracks 5, 3—14.

Głazek J. 1996: Karst and caves of the Polish Tatra Mountains, state of

knowledge and perspectives. In: Kotarba A. (Ed.): The Nature of
the Tatra National Park and Man. Vol. 1. Nauki o Ziemi, Tatrzan-
ski Park Narodowy, Polskie Towarzystwo Przyjaciół Nauk o Ziemi
Oddział w Krakowie, Zakopane, Kraków, 31—44 (in Polish).

Golab J. 1959: On the geology of the western Podhale Flysch area. Biul.

(Inst. Geol.) 149, 225—240 (in Polish with English summary).

Gorek A. 1956: Geological structure of the Western Tatra Mts. Geol.

Zbor. SAV, Bratislava 7, 1, 125—130 (in Slovak).

Gorek A. & Scheibner E. 1966: Final report. Basic geological maps,

Nižná and Bystrá sheets (1 : 50,000). Manuscript, Archive Geo-
fond, Bratislava, 342 (in Slovak).

Gradziński M., Hercman H., Kicińska D., Barczyk G., Bella P. &

Holúbek P. 2009: Karst of the Tatra Mountains – the develop-
ment of knowledge in the last thirty years. Przegl. Geol. 57,
674—684 (in Polish).

Grecula P. & Roth Z. 1978: Kinematic model of the West Carpathians.

Sbor. Geol. Věd, Ř.G., 49—73.

Gross P. 1973: On the character of the Choč-sub-Tatra fault. Geol.

Práce, Spr. 61, 315—319 (in Slovak).

Gross P., Köhler E. & Samuel O. 1984: A new lithostratigraphic sub-

division of the Central Carpathian Paleogene. Geol. Práce, Spr.
81, 103—117 (in Slovak).

Gross P., Köhler E., Mello J., Halouzka R., Haško J. & Nagy A. 1993:

Geology of Southern and Eastern Orava. GÚDŠ, Bratislava, 292
(in Slovak).

Gross P., Köhler E. (Eds.), Biely A., Franko O., Hanzel V., Hricko J.,

Kupčo G., Papšová J., Priechodská Z., Szalaiová V., Snopková P.,
Stránska M., Vaškovský I. & Zbořil . 1980: Geology of the Lip-
tovská kotlina Basin. GÚDŠ, Bratislava, 242 (in Slovak).

Gross P., Buček S., Ďurkovič T., Filo I., Maglay J., Halouzka R.,

Karoli  S.,  Nagy  A.,  Spišák  Z.,  Žec  B.,  Vozár  J.,  Borza  V.,
Lukáčik  E.,  Janočko  J.,  Jetel  J.,  Kubeš  P.,  Kováčik  M.,  Žáková
E., Mello J., Polák M., Siráňová Z., Samuel O., Snopková P., Ra-
ková J., Zlinská A., Vozárová A. & Žecová K. 1999: Explanation
to Geological map of Popradská kotlina Basin, Hornádska kotli-
na Basin, Levočské vrchy Mts., Spišsko-šarišské medzihorie de-
pression,  Bachureň  Mts.  and  Šarišská  vrchovina  highland
(1 : 50,000).  Ministerstvo  životného  prostredia,  Geologická
služba Slovenskej republiky, Bratislava, 239.

Haq B.U., Hardenbol J. & Vail P.R. 1988: Mesozoic and Cenozoic

chronostratigraphy and cycles of sea level changes. In: Wilgus
C.K., Hastings B.S., Kendall C.G.St.C., Posamentier H.W., Ross
C.A. & Van Wagoner J.C. (Eds.): Sea level changes – An inte-
grated approach. SEPM Spec. Publ. 42, 71—180.

Havrila M. 2011: Hronicum: paleogeography and stratigraphy (Upper

Pelson – Tuvalian), tectonic individualization and structure.
Geol. Práce, ŠGÚDŠ, Bratislava, 117, 7—103 (in Slovak).

Hidy A.J., Gosse J.C., Blum M.D. & Gibling M.R. 2014: Glacial—in-

terglacial variation in denudation rates from interior Texas, USA,
established with cosmogenic nuclides. Earth Planet. Sci. Lett.
390, 209—221.

Hók J., Kováč P. & Madarás J. 1993: Extensional tectonics of the

western part of the contact area between Veporicum and Gemeri-
cum (Western Carpathians). Miner. Slovaca 25, 172—176 (in Slo-
vak, with English summary).

Hurai V., Kihle J., Kotúlová J., Marko F. & Świerczewska A. 2002:

Origin of methane in quartz crystals from the Tertiary accretion-
ary wedge and fore-arc basin of the Western Carpathians. Ap-
plied Geochemistry 17, 1259—1271.

Hurford A.J. & Green P.F. 1983: The zeta age calibration of fission-

track dating. Chem. Geol. 1, 285—317.

Janák M. 1994: Variscan uplift of the crystalline basement, Tatra Mts,

Central Western Carpathians: evidence from

Ar/

Ar laser

probe dating of biotite and P-T-t paths. Geol. Carpathica 45,
293—300.

Janák M. & Onstott T.C. 1993: Pre-Alpine tectono-thermal evolution

of metamorphism in the Tatra Mts., Western Carpathians: P-T
paths and

Ar/

Ar laser probe dating. Terra Abstr., Suppl. 1 to

Terra Nova 5, 238.

Janák M., Plašienka D. & Petrík I. 2001: Excursion to the Tatra Moun-

tains, Central Western Carpathians: Tectonometamorphic record of
Variscan and Alpine orogeny. In: Excursion guide. 6

meeting of

the Czech Tectonic Studies Group, Donovaly – Nízke Tatry, May
3—6, 2001. Geolines 13, 141—148.

Janák M., O’Brien P.J., Hurai V. & Reuter C. 1996: Metamorphic evo-

lution and fluid composition of garnet-clinopyroxene amphibolites
from the Tatra Mountains, Western Carpathians. Lithos 39, 57—79.

Janák M., Hurai V., Ludhová L., O’Brien P.J. & Horn E. 1999: Dehy-

dration melting and devolatilization during exhumation of high-
grade metapelites: the Tatra Mts, Western Carpathians. J.
Metamorph. Geology 17, 379—395.

Janočko J. & Jacko S. 1998: Marginal and deep-sea deposits of Central-

Carpathian Paleogene Basin, Spišská Magura region, Slovakia:
Implication for basin history. Slovak. Geol. Mag. 4, 281—292.

Janočko J., Gross P., Buček S., Karoli S., Žec B., Rakús M., Potfaj M.

& Halouzka R. 2000: Geological map of the Spišská Magura
Mts. 1 : 50,000. ŠGÚDŠ, Bratislava.

Jeřábek P., Lexa O., Schulmann K. & Plašienka D. 2012: Inverse duc-

325

CRETACEOUS—QUATERNARY EVOLUTION OF THE TATRA MOUNTAINS (WESTERN CARPATHIANS)

GEOLOGICA CARPATHICA

GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

tile  thinning  via  lower  crustal  flow  and  fold-induced  doming  in
the West Carpathian Eo-Alpine collisional wedge.  Tectonics  31,
TC5002.  Doi:10.1029/2012TC003097

Jeřábek P., Stünitz H., Heilbronner R., Lexa O. & Schulmann K.

2007: Microstructural-deformation record of an orogen-parallel
extension in the Vepor Unit, West Carpathians. J. Struct. Geol.
29, 1722—1743.

Jeřábek P., Faryad W.S., Schulmann K., Lexa O. & Tajčmanová L.

2008: Alpine burial and heterogeneous exhumation of Variscan
crust in the West Carpathians: insight from thermodynamic and
argon diffusion modelling. J. Geol. Soc. 165, 479—498.

Jurewicz E. 2005: Geodynamic evolution of the Tatra Mts. and the

Pieniny Klippen Belt (Western Carpathians): problems and com-
ments. Acta Geol. Pol. 55, 3, 295—338.

Kahan Š. 1969: Eine neue Ansicht über den geologischen Aufbau des

Kristallinikums der West Tatra. Acta Geol. Geograph. Univ.
Com. 12, 115—122.

Kázmer M., Dunkl I., Frisch W., Kuhlemann J. & Ozsvárt P. 2003:

The Paleogene forearc basin of the Eastern Alps and Western
Carpathians: subduction erosion and basin evolution. J. Geol.
Soc., London 160, 413—428.

Ketcham R.A., Donelick R.A. & Carlson W.D. 1999: Variability of

apatite fission-track annealing kinetics. III. Extrapolation to geo-
logical time scale. Amer. Mineralogist 84, 1235—1255.

Kępińska B. 1997: Geologic and geothermal model of the Podhale Ba-

sin. Studia, Rozprawy, Monografie 48, Polska Akademia Nauk,
Krakow, 1—105 (in Polish).

Kępińska B. 2006: Thermal and hydrothermal conditions of the Podhale

geothermal system. Studia, Rozprawy, Monografie IGSMiE PAN w
Krakowie 135, 1—112 (in Polish, English summary).

Kováč P. & Hók J. 1996: Tertiary development of the western part of

the Pieniny Klippen Belt. Slovak Geol. Mag. 2, 137—149.

Kováč M., Krá J., Márton E., Plašienka D. & Uher P. 1994: Alpine

uplift history of the Central Western Carpathians: geochronologi-
cal, paleomagnetic, sedimentary and structural data. Geol. Car-
pathica 45, 83—96.

Kozur H. & Mock R. 1996: New paleogeographic and tectonic inter-

pretations in the Slovakian Carpathians and their implications for
correlations with the Eastern Alps. Part I. Central Western Car-
pathians. Miner. Slovaca 28, 151—174.

Kozur H. & Mock R. 1997: New paleographic and tectonic interpreta-

tions in the Slovakian Carpathians and their implications for cor-
relations with the Eastern Alps. Part II. Inner Western Carpathians.
Miner. Slovaca 29, 164—209.

Králiková S. 2013: Low-thermal evolution of the Central Western

Carpathian rock complexes during the Alpine tectogenesis. PhD.
Thesis, Comenius University, Faculty of Natural Sciences, Bra-
tislava, 130.

Králiková S., Vojtko R., Andriessen P., Kováč M., Fügenschuh B.,

Hók J. & Minár J. 2014: Late Cretaceous—Cenozoic thermal evo-
lution of the northern part of the Central Western Carpathians
(Slovakia): revealed by zircon and apatite fission track thermo-
chronology. Tectonophysics 615—616, 142—153.

Krá J. 1977: Fission track ages of apatites from some granitoid rocks

in West Carpathians. Geol. Zbor. Geol. Carpath. 28, 269—276.

Krysiak Z. 1976: Transport directions of the Podhalie Flysch material

based on the data from the Lesnica brook catchment area. Kwart.
Geol. 20, 2, 323—330 (in Polish).

Lugeon M. 1903: Les nappes de recouvrement de la Tatra et l’origine

des Klippes des Carpathes. Bull. Lab. Géol. Minéral. Géophys.
Mus. Géol. Univ., Laussane 4, 17—63.

Lukniš M. 1968: Geomorphological Map of the Vysoké Tatry Mts.

(High Tatra Mts.) and their Foreland 1 : 50,000. Geologický ústav
Dionýza Štúra, Bratislava.

Lukniš M. 1973: Relief of the High Tatra Mts. and their foreland. Vyd.

SAV, Bratislava, 375 (in Slovak).

Mahe M. 1986: Geological structure of Czechoslovak Carpathians.

Part 1. Paleo-Alpine Units. Veda, Bratislava, 1—503 (in Slovak).

Mahe M., Kamenický J., Fusán O. & Matějka A. 1967: Regional geo-

logy of the ČSSR. Part II. Western Carpathians. Volume 1. Vyd.
ČSAV, Praha, 407 (in Czech).

Maluski H., Rajlich P. & Matte P. 1993:

Ar—

Ar dating of the Inner

Carpathians Variscan basement and Alpine mylonitic overprint.
Tectonophysics 223, 313—337.

Marko F. 1995: Dynamic analysis of fault deformation in the Central

Carpathian Paleogene Basin based upon structural observations
from NW and S periphery of the Levočské vrchy Mts. MS GIÚ
SAV, Bratislava, 1—24 (in Slovak).

Marschalko R. & Radomski A. 1960: Preliminary results of investiga-

tion of current direction in the flysh basin of the Central Car-
pathians. Rocznik Polskiego Towarzystva Geologicnego, Krakow
XXX, 3.

Marschalko R. & Radomski A. 1970: Sedimentary structures and de-

velopment of marginal facies of the Eocene flysch near Ždiar.
Geol. Práce, Spr. 53, 85—99 (in Slovak).

Marynowski L. & Gawęda A. 2005: Correlation between biomarkers

and thermal maturity of the organic matter from the Paleogene
sedimentary rocks of the Podhale trough. Mineral. Soc. Pol.,
Spec. Pap. 25, 329—332.

Mastella L. & Rybak-Ostrowska B. 2012: Tectonic control of tufa oc-

currences in the Podhale Synclinorium (Central Western Car-
pathians, southern Poland). Geol. Quart. 56, 733—744.

Mazúr E. 1965: Major features of the West Carpathians in Slovakia as

a result of young tectonic movements. In: Mazúr E. & Stehlik O.
(Eds.): Geomorphological problems of Carpathians. 1. Vyd. SAV,
Bratislava, 9—53.

Midriak R. 1983: Morphogenetic of the high-mountain surface. Veda,

Bratislava, 516 (in Slovak).

Minár J., Bielik M., Kováč M., Plašienka D., Barka I., Stankoviansky

M. & Zeyen H. 2011: New morphostructural subdivision of the
Western Carpathians: An approach integrating geodynamics into
targeted morphometric analysis. Tectonophysics 502, 158—174.

Nagy A., Vass D., Petrík R. & Pereszlényi M. 1996: Tectonogenesis

of the Orava Depression in the light of latest biostratigraphic in-
vestigations and organic matter alteration study. Slovak Geol.
Mag. 2, 49—58.

Nemčok M. & Nemčok J. 1994: Late Cretaceous deformation of the

Pieniny Klippen Belt, West Carpathians. Tectonophysics 239,
81—109.

Nemčok J., Bezák V., Janák M., Kahan Š., Ryka W., Kohút M.,

Lehotský I., Wieczorek J., Zelman J., Mello J., Halouzka R., Racz-
kowski W. & Reichwalder P. 1993: Explanation to geological map
of the Tatra Mts. 1 : 50,000. GÚDŠ, Bratislava (in Slovak).

Nemčok J., Bezák V., Biely A., Gorek A., Gross P., Halouzka R., Janák

M., Kahan Š., Kotański Z., Lefeld J., Mello J., Reichwalder P.,
Raczkowski W., Roniewicz P., Ryka W., Wieczorek J. & Zelman
J. 1994: Geological map of the Tatra Mountains. MŽP SR,
GÚDŠ, Bratislava.

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.

Passendorfer E. 1958: About sedimentation of the Eocene in the Tatra

Mts. [W sprawie sedimentacji eocenu tatrzamskiego.] Acta Geol.
Pol. 8, 451—476 (in Polish).

Pešková I., Vojtko R., Starek D. & Sliva . 2009: Late Eocene to Qua-

ternary deformation and stress field evolution of the Orava region
(Western Carpathians). Acta Geol. Pol. 59, 1, 73—91.

Piotrowski J. 1978: Mesostructural analysis of the main tectonic units

of the Tatra Mountains along the Kościeliska Valley. Stud. Geol.
Pol. 55, 90 (in Polish with English summary).

Plašienka D. 1993: Structural pattern and partitioning of deformation

in  the  Veporic  Federata  cover  unit  (Central  Western  Car-
pathians).  In:  Rakús  M.  &  Vozár  J.  (Eds.):  Geodynamic  model
and  deep  structure  of  the  Western  Carpathians.  Konf.,  Symp.,
Sem., GÚDŠ, Bratislava, 269—277.

Plašienka D. 1999: Tectonochronology and paleotectonic model of the

Jurassic—Cretaceous evolution of the Central Western Car-
pathians. Veda, Vyd. SAV, Bratislava, 125 (in Slovak).

326

KRÁLIKOVÁ, VOJTKO, SLIVA, MINÁR, FÜGENSCHUH, KOVÁČ and HÓK

GEOLOGICA CARPATHICA

GEOLOGICA CARPATHICA, 2014, 65, 4, 307—326

Plašienka D. 2003: Development of basement-involved fold and thrust

structures exemplified by the Tatric-Fatric-Veporic nappe system
of the Western Carpathians (Slovakia). Geodynamica Acta 16,
21—38.

Plašienka D., Broska I., Kissová D. & Dunkl I. 2007: Zircon fission-

track dating of granites from the Vepor – Gemer Belt (Western
Capathians): constraints for the Early Alpine exhumation history.
J. Geosci. 52, 113—123.

Plašienka D., Grecula P., Putiš M., Kováč M. & Hovorka D. 1997:

Evolution and structure of the Western Carpathians: an overview.
In: Grecula P., Hovorka D. & Putiš M. (Eds.): Geological evolu-
tion of the Western Carpathians. Miner. Slovaca, Monograph,
Bratislava, 1—24.

Pomianowski P. 2003: Tectonics of the Orava-Nowy Targ Basin –

results of the combined analysis of the gravity and geoelectrical
data. Przegl. Geol. 51, 498—506.

Pospíšil L. 1990: The geophysical gravity models of the Orava Neo-

gene basin. Zemný Plyn a Nafta 35, 301—310 (in Czech).

Prokešová R., Plašienka D. & Milovský R. 2012: Structural pattern

and emplacement mechanisms of the Krížna cover nappe (Cen-
tral Western Carpathians). Geol. Carpathica 63, 1, 13—32.

Ratschbacher L., Frisch W., Linzer H., Sperner B., Meschede M.,

Decker K., Nemčok M., Nemčok J. & Grygar R. 1993: The Pie-
niny Klippen Belt in the Western Carpathians of northeastern
Slovakia: structural evidence for transpression. Tectonophysics
226, 471—483.

Roniewicz P. 1969: Sedimentation of the Eocene Nummulite in the

Tatra Mts). [Sedimentacja eocenu numulitowiego Tatr.] Acta
Geol. Pol. 19, 3, 503—608 (in Polish).

Roth Z. (Ed.) 1963: Explanation to geological map at a scale of

1 : 200,000, sheet M-34-XX (Trstená). Archive-Geofond, Bratislava
(in Slovak).

Royden L.H. & Báldi T. 1988: Early Cenozoic tectonics and paleo-

geography of the Pannonian Basin and surrounding regions.
Amer. Assoc. Petrol. Geol. Mem. 45, 1—16.

Schaller M., von Blanckenburg F., Veldkamp A., Tebbens L.A., Hovius

N. & Kubik P.W.A. 2002: A 30,000 yr record of erosion rates
from cosmogenic 10 Be in Middle European river terraces. Earth
Planet. Sci. Lett. 204, 307—320.

Sliva . 2005: Sedimentary facies of the Central Carpathians Paleo-

gene Basin in the area of Spišská Magura Mts. PhD. Thesis,
Comenius University, Bratislava, 137 (in Slovak).

Sokolowski S. 1948: Belianske Tatry Mts. Geology of the southern

slopes. [Tatry Bielskie. Geologia zboczy poludniowych.] Pr.
Panstw. Inst. Geol., Warszawa 4, 3—47 (in Polish).

Soták J. 1998: Sequence stratigraphy approach to the Central Carpathian

Paleogene (Eastern Slovakia): eustasy and tectonics as controls of
deep-sea-fan deposition. Slovak Geol. Mag. 4, 3, 185—190.

Soták J. 2010: Paleoenvironmental changes across the Eocene-Oli-

gocene boundary: insights from the Central-Carpathian Paleo-
gene Basin. Geol. Carpathica 61, 5, 393—418.

Soták J. & Starek D. 2000: Synorogenic deposition of turbiditic fans

in the Central Carpathian Paleogene Basin: evidence for and
against sea-level and climatic changes. Slovak Geol. Mag. 6, 2—3,
191—193.

Soták J., Pereszlényi M., Marschalko R., Milička J. & Starek D. 2001:

Sedimentology and hydrocarbon habitat of the submarine-fan de-
posits of the Central Carpathian Paleogene Basin (NE Slovakia).
Mar. Petrol. Geol. 18, 87—114.

Sperner B. 1996: Computer programs for the kinematic analysis of brit-

tle deformation structures and the Tertiary tectonic evolution of the
Western Carpathians (Slovakia). Tübinger Geowiss. Arbeiten,
Reihe A, Band 27, 120.

Sperner B., Ratschbacher L. & Nemčok M. 2002: Interplay between

subduction retreat and lateral extrusion: Tectonics of theWestern
Carpathians. Tectonics 21, 1028—1051.

Starek D. 2001: Sedimentology and paleodynamics of the Paleogene

formations in the Central Western Carpathians (Orava region).
Manuscript, PhD. Thesis, Geol. Inst. Slovak Acad. Sci., Bratislava,
152 (in Slovak).

Starek D., Andreyeva-Grigorovich A.S. & Soták J. 2000: Suprafan de-

posits of the Biely Potok Fm. in the Orava region: sedimentary
facies and nannoplankton distribution. Slovak Geol. Mag. 6, 2—3,
188—190.

Starek D., Sliva . & Vojtko R. 2012: Eustatic and tectonic control on

late Eocene fan delta development (Orava Basin, Central West-
ern Carpathians). Geol. Quart. 56, 1, 67—84.

Starkel L. 1969: The age of the stages of development of the relief of

the Polish Carpathians in the light of the most recent geological
investigations. Studia Geomorphologica Carpatho-Balcanica 3.
33—44.

Struzik A.A., Zattin M. & Anczkiewicz R. 2002: Apatite fission track

analyses from the Polish Western Carpathians. Geolines 14, 87—89.

Sůkalová ., Vojtko R. & Pešková I. 2012: Cenozoic deformation and

stress field evolution of the Kozie chrbty Mountains and the
western part of Hornád Depression (Central Western Car-
pathians). Acta Geol. Slovaca 4, 1, 53—64.

Śmigielski M., Stuart F.M., Krzywiec P., Persano C., Sinclair H.D.,

Pisaniec K. & Sobien K. 2012: Neogene exhumation of the
Northern Carpathians revealed by low temperature thermochro-
nology. Geophys. Res. Abstr., 14, EGU 2012—12063.

Środoń J., Kotarba M., Biroň A., Such P., Clauer N. & Wójtowicz A.

2006: Diagenetic history of the Podhale-Orava Basin and the un-
derlaying Tatra sedimentary structural units (Western Car-
pathians): evidence from XDR and K-Ar of illite smectite. Clay
Minerals 41, 75—774.

Tagami T., Galbraith R.F., Yamada R. & Laslett G.M. 1998: Revised

annealing kinetics of fission tracks in zircon and geological im-
plications. In: Van den Haute P. & De Corte F. (Eds.): Advances
in fission-track geochronology. Kluwer Academic Publishers,
Dordrecht, 99—112.

Tari G., Báldi T. & Báldi-Béke M. 1993: Paleogene retroarc flexural

basin beneath the Neogene Pannonian Basin: a geodynamic model.
Tectonophysics 226, 433—456.

Tokarski A.K., Świerczewska A., Zuchiewicz W., Starek D. & Fodor L.

2012: Quaternary exhumation of the Carpathians: a record from
the Orava-Nowy Targ Intramontane Basin, Western Carpathians
(Poland and Slovakia). Geol. Carpathica 63, 4, 257—266.

Uhlig V. 1897—1898: Die Geologie des Tatragebirges I. Denkschr. (Kais.

Akad. Wiss.), Math.-Naturwiss. Kl., Wien 64, 68, 643—684.

Vojtko R., Tokárová E., Sliva . & Pešková I. 2010: Reconstruction

of Cenozoic paleostress fields and revised tectonic history in the
northern part of the Central Western Carpathians (the Spišská
Magura and Východné Tatry Mountains). Geol. Carpathica 61,
3, 211—225.

Vojtko R., Králiková S., Minár J. & Fügenschuh B. 2013: Low ther-

mal evolution of the Southern Veporic Unit crystalline basement
(Central Western Carpathians) constrained by new fission track
data. Berichte der Geologischen Bundesanstalt, Band 99, Work-
shop on Alpine Geological Studies and European Symposium on
Fossil Algae, Schladming 7—14, 9, 96—97.

Vojtko R., Marko F., Preusser F., Madarás J. & Kováčová M. 2011:

Late Quaternary fault activity in the Western Carpathians: evi-
dence from the Vikartovce fault (Slovakia). Geol. Carpathica 62,
563—574.

Wagner M. 2011: Petrologic studies and diagenetic history of coaly

matter in the Podhale flysh sediments, southern Poland. Ann.
Soc. Geol. Pol. 81, 173—183.

Wagner G.A. & Van den Haute P. 1992: Fission-track dating. Kluwer

Academic Publishers, Dordrecht, 285.

Watycha L. 1976: The Neogene of the Orava—Nowy Targ Basin.

Kwart. Geol. 20, 575—585 (in Polish with English summary).

Zuchiewicz W. 2011: Planation surfaces in the polish Carpathians: myth

or reality? Part 2. Geographia Pol., Spec. Issue 84, 155—178.