GeolCarp_Vol67_No1_41

www.geologicacarpathica.com

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

, FEBRUARY 2016, 67, 1, 41—68 doi: 10.1515/geoca-2016

-0003

Evolution of the passive margin of the peripheral foreland

basin: an example from the

Lower

Miocene Carpathian

Foredeep (Czech Republic)

MICHAL FRANCÍREK and SLAVOMÍR NEHYBA

Institute of Geological Sciences, Faculty of Science, Masaryk University, Kotlářská 2, CZ- 611 37 Brno, Czech Republic;

francirekmichal@seznam.cz, slavek@sci.muni.cz

(Manuscript received April 17, 2015; accepted in revised form December 8, 2015)

Abstract: The Karpatian deposits of the central part of the Carpathian Foredeep in Moravia, which are deeply buried
under the Outer Western Carpathians, provide a unique opportunity to reconstruct the former evolutionary stages of this
peripheral foreland basin and its paleogeography. A succession of three depositional units characterized by a distinct
depositional environment, provenance, and partly also foreland basin depozone, have been identified. The first deposi-
tional unit represents a proximal forebulge depozone and consists of lagoon-estuary and barred coastline deposits.
The source from the “local” crystalline basement played here an important role. The second depositional unit consists
of coastline to shallow marine deposits and is interpreted as a forebulge depozone. Tidalites recognized within this unit
represent the only described tide-generated deposits of the Neogene infill of the Carpathian Foredeep basin in Moravia.
The source from the basin passive margin (the Bohemian Massif) has been proved. The third depositional unit is formed
by offshore deposits and represents a foredeep depozone. The provenance from both passive and active basin margin
(Silesian Unit of the Western Carpathian Flysch Zone) has been proved. Thus, both a stepwise migration of the foredeep
basin axis and shift of basin depozones outwards/cratonwards were documented, together with forebulge retreat.
The shift of the foreland basin depozones more than 50 km cratonward can be assumed. The renewed thrusting along
the basin’s active margin finally completely changed the basin shape and paleogeography. The upper part of the infill
was deformed outside the prograding thrust front of flysch nappes and the flysch rocks together with a strip of Miocene
sediments were superposed onto the inner part of the basin. The width and bathymetric gradient of the entire basin was
changed/reduced and the deposition continued toward the platform. The basin evolution and changes in its geometry are
interpreted as a consequence of the phases of the thrust-sheet stacking and sediment loading in combination with sea-
level change.

Key words: Carpathian Foredeep, Late Burdigalian—Karpatian, peripheral foreland basin, Flysch Thrust Wedge,
provenance.

Introduction

The  geometry  of  a  peripheral  foreland  basin  is  mainly
a  product  of  a  complex  dynamic  balance  between  the  oro-
genic loading, erosion and sedimentation, lithospheric flexu-
ral  response  to  these  processes,  and  sea-level  changes
(eustatic).  DeCelles  and  Giles  (1996)  subdivided  foreland
basins  into  four  distinct  depozones:  wedge-top,  foredeep,
forebulge, and backbulge. Several models were suggested to
explain  the  relations  between  thrust  loading,  sediment  sup-
ply,  and  basin  shape  (Flemings  &  Jordan  1989;  Jordan  &
Flemings  1991),  basin  character  and  depozone  migration
(Heller  et  al.  1988;  Catuneanu  et  al.  1998;  Yang  &  Miall
2010).  The  principal  data  about  the  alternating  phases  of
thrusting activity (orogenic loading) and tectonic quiescence
(sediment  loading)  are  provided  by  the  study  of  either  the
proximal foredeep or the forebulge depozones (Flemings &
Jordan 1989; Jordan & Flemings 1991; Sinclair et al. 1991;
Crampton & Allen 1995; Plint 2000; Plint et al. 2001; Yang
&  Miall  2010;  Leszczyński  &  Nemec  2014).  However,  as
thrust  movement  propagates  further  and  further  onto  the
foreland, the foreland basin structure is “continually” modi-

fied  as  inner  parts  of  the  basin  (proximal  to  active  margin)
are incorporated into the anticlinal thrust stack, and basin de-
pozones are shifted further onto the foreland. This situation
also means that foreland basins are continuously rearranged,
and  especially  the  most  proximal  parts  of  the  basin’s  pre-
vious  stages  are  poorly  preserved.  However  such  deposits
have  a  potential  to  provide  exclusive  information  about  the
basin’s evolution and paleogeography. The results of subsur-
face  exploration  in  eastern  Moravia,  where  several  deep
wells penetrated the deposits of the Outer Carpathian Flysch
Zone  into  the  submerged  part  of  the  Carpathian  Foredeep
basin,  offered  such  an  opportunity.  The  proposed  article
aims  to  improve  our  understanding  of  the  paleogeography
and  evolution  of  the  Carpathian  Foredeep  basin  during  the
Karpatian  (Lower  Miocene)  based  on  detailed  sedimentolo-
gical and sedimentary-petrographical studies using data from
16  deep  boreholes  and  31  cores.  “Alternative”  lithological,
sedimentological  and  paleontological  results  from  the  Car-
pathian  Foredeep  forebulge  depozone  have  been  published
(Cogan et al. 1993; Hladilová et al. 1999; Nehyba & Šikula
2007; Zágoršek et al. 2012; Hladilová et al. 2014; Holcová et
al. 2015).

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Geological setting

The studied Neogene deposits belong to the western part

of the Carpathian Foredeep, a peripheral foreland basin
formed due to the tectonic emplacement and crustal loading

of  the  Western  Carpathian  Thrust  Front  onto  the  passive
margin  of  the  Bohemian  Massif  (Nehyba  &  Šikula  2007)
(Fig. 1A). The infill and basin architecture varies throughout
the  Carpathian  Foredeep,  local  and  regional  unconformities
are  developed  due  to  varying  intensity  and  orientation  of
flexural loading and different geological and tectonic histo-
ries of the basement, along with a polyphase nature of the ac-
tive  basin  margin  and  gradual  change  of  its  position
(Brzobohatý & Cicha 1993; Eliáš & Pálenský 1998; Nehyba
2000; Nehyba & Petrová 2000; Krzywiec 2001; Kováč et al.
2003; 2004; Oszczypko et al. 2006; Nehyba & Šikula 2007).

The Moravian part of the Carpathian Foredeep was subdi-

vided into three segments (i.e. southern, central, and northern
ones),  with  partly  different  lithological  and  stratigraphical
contents  and  depositional  histories  (Brzobohatý  &  Cicha
1993). The sedimentary infill of the basin is composed of de-
posits  of  Egerian  to  Badenian  age;  however  only  the
Karpatian  deposits  (up  to  1200 m  thick)  were  identified  in
the  studied  area.  The  Karpatian  depositional  cycle  is  con-
nected with a shift of the basin axis to the northwest due to
continued  thrusting  of  the  Outer  Carpathian  Flysch  Wedge
(Brzobohatý & Cicha 1993), which also led to the overriding
of  the  significant  part  of  the  Carpathian  Foredeep  by  the
flysch nappes (Fig. 1B) and partial incorporation of the basin
infill into the nappes.

Seismic and borehole data showed that the pre-Miocene

basement  of  the  studied  area  is  formed  by  crystalline  rocks
of the Brunovistulicum, limestones of the Moravian-Silesian
Paleozoic  and  occasional  Carboniferous  clastic  deposits  of
Drahany Highland unit (Kalvoda et al. 2003, 2008; Zágoršek
et  al.  2012;  Hladilová  et  al.  2014).  The  basement  generally
dips southeastward, however the relief is very irregular. The
basement is covered by Karpatian deposits, which were sub-
divided  into  the  Mušov  Member,  the  Nový  Přerov  Member
and the Kroměříž Formation (Fig. 2). The Mušov Member is
represented by grey marine mudstones with rich microfauna
content  (called  “Schliers”).  The  Nový  Přerov  Member  is
formed of siltstones to mudstones with thin interbeds of fine-

to medium-grained sandstones (the so-called “Sandy Schlier
Formation”)  (Brzobohatý  &  Cicha  1993;  Adámek  et  al.
2003).  Mostly  chaotic  deposits  of  the  Kroměříž  Formation
represent the final phase of the Karpatian depositional cycle
(Benada & Kokolusová 1987; Adámek et al. 2003). Zádrapa
(1979), described five facies of these Karpatian deposits (i.e.
clastic,  psammitic-pelitic,  pelitic,  pelitic-psammitic,  and
variegated  ones).  Thonová  et  al.  (1987)  divided  them  into
three sections and Šikula & Nehyba (2006) described 6 well-
log  facies  of  these  deposits.  Detailed  paleontological  study
of the Karpatian deposits in the area under study is missing.

The studied deposits are superposed by up to 4 km thick

pile of the Western Carpathian Flysch Zone. In the Polish
Carpathian Foredeep, this part located beneath the Car-
pathian nappes is described as inner foredeep (Oszczypko &
Oszczypko-Clowes 2012, Waśkowska et al. 2014).

Methods

The presented results are based on the study of the data from

16 boreholes. These boreholes are Bařice 1 (Bar 1), Gottwal-
dov 1, 2, and 3 (G 1, 2, 3), Holešov 1 (Hol 1), Hulín 1, 2, and
3 (Hul 1, 2, 3), Jarohněvice 1 (Jar 1), Kroměříž 1 and 2 (Kro 1,
2), Rataje 1 and 2 (Rat 1, 2), Roštín 1 and 2 (Ros 1, 2), Slušo-
vice 1 (Slu 1), Tlumačov 1 and 2 (Tl 1, 2), and Vrbka 1 (Vr 1).
The positions of the wells are presented in Fig. 3.

The lithofacies analysis is based on the sedimentological

study of borehole cores, following the common rules of
Miall (1989), Walker & James (1992), Readi

ng (1996) and

Nemec  (2005).  The  quality  and  extent  of  the  cores  greatly
varies, however they were mostly about 1 m long, exceptio-
nally  reaching  up  to  3 m.  Therefore,  further  data  provided
evaluation  of  the  available  wire-line  logs;  “standard”  wire-
line  techniques  –  spontaneous  potential  (SP),  resistivity
(Rag 2,12) and gamma-ray (gamma-API) (Rider 1986).

Maps of the thicknesses of deposits were created in the

software Surfer 7 (gridding method). For the compilation of
maps, the data from a slightly broader area (boreholes
Kožušice 1, 4, Lubná 2, 4, Morkovice 1, 2, 4, Nítkovice 2, 4, 6,
Rusava 1, 3, 5, 6 and Stupava 1, 2) have been used.

Provenance analysis is based on the petrographical studies

of  core  samples.  The  framework  grains  were  point  counted
in  27  thin  sections  according  to  the  standard  method
(Dickinson & Suczek 1979; Dickinson et al. 1983; Ingersoll
1990; Zuffa 1980; 1985). The entire rock geochemistry was
evaluated  at  ACME  Laboratories  in  Vancouver,  Canada
(57 analyses).

The chemical index of alteration (CIA) is commonly

used (Nesbitt & Young 1982; Nesbitt & Young 1989; Fedo
et al. 1995; Nesbitt et al. 1996; von Eynatten et al. 2003; Li
& Yang 2010). Due to the absence of CO

data and different

contents of carbonate, a precise correction for the carbonate
CaO* was difficult. The correction is based on the indirect
method, in that it is necessary to deduct the mole fraction of
P

(apatite) from the mole fraction of CaO*. The value of

CaO* is accepted if the remaining mole fraction of
CaO<Na

O. However, if CaO>Na

O then CaO* corresponds

to CaO=Na

O (McLennan 1993).

Fig. 2. Regional stratigraphic scheme of the Neogene of the
Carpathian Foredeep in central Moravia (after Adámek et al. 2003).

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Heavy minerals were studied in the 0.063—0.125 mm

grain-size  fraction  in  33  samples.  The  data  of  the  heavy
minerals assemblages were obtained from the archives of the
MND  Group.  The  electron  microprobe  analysis  of  garnet
(336  grains  of  Karpatian  deposits  and  15  grains  of  the
crystalline basement) and rutile (103 grains) was performed
with  a  CAMECA  SX  electron  microprobe  analyser  in  the
Faculty  of  Science,  Masaryk  University,  Brno,  Czech  Re-
public.  The  following  analytical  conditions  were  used:
15 kV  accelerating  voltage,  20 nA  beam  current,  beam  dia-
meter  2—5

µm. Garnet and rutile grains were analysed in

their centres. Fe

and FeO were calculated from stoichio-

metry. The formula was standardized to 12 oxygens and
8 cations. Garnets and rutiles were studied in the sandstones
and siltstones of the boreholes G 1, G 2, Hol 1, Hul 2, Kro 1,
Kro 2, Ros 1, Ros 2, Slu 1, Tlu 1 and Tlu 2.

Results

Facies analysis

ndividual lithofacies were identified according to grain size

and sedimentary structures. Recognized lithofacies are briefly
described in Table 1 and can be followed in Fig. 4. Lithofacies
can be grouped into fine-grained, heterolithic, sands

tone, and

coarse-grained categories based on the dominant grain size.
Lithofacies have been combined, based on their spatial
grouping in cores into five facies associations (FA).

The FA1 is formed by the dominant lithofacies M1 and

less frequent M2. The FA1 is located on the base of the
depositional succession and is situated on the western and
central part of the studied area.

The FA2 is composed mostly of lithofacies S3, S1, and S4

and  less  commonly  S2,  S6,  H3,  and  G1.  Sandstone  facies
strongly  dominates  (forming  95.4 %),  whereas  gravel  and
heterolithic facies form only a few percent of the succession.
The  FA2  is  substituted  for  FA1  and  is  therefore  located  on
the  base  of  succession  or  above  FA1;  it  is  situated  in  the
western part of the area.

The FA3 is formed dominantly by sandstone lithofacies

S4, S2, S1, and S5, while fine-grained M1 and M2 and he-
terolithic H1 and H3 ones are less common. The occurrence
of individual lithofacies categories varies in individual wells
(sandstone  lithofacies  represent  47.1—51.1 %,  heterolithic
ones 20.4—39.0 %, and fine-grained ones 12.9—28.4 %). The
occurrence  of  herringbone  cross-lamination  (facies  H3)
within  the  upper  part  of  FA3  has  been  recognized  in  some
wells (G 1, G 2). The FA3 is superimposed on FA1 and FA2
and is located in the western part of the area.

The FA4 is represented dominantly by heterolithic litho-

facies H1, H2, and H3 (forming 46.1—66.2 %), whereas sand-
stone lithofacies S1, S2 and S5 (17.6—29.8 %) and fine-
grained lithofacies M1 and M2 (4.1—36.4 %) are less common.
The FA4 is situated above FA1 in the central part of the area.

The FA5 is formed predominantly by fine-grained litho-

facies M1 (50—87 %); heterolithic lithofacies H1 and H2 are
less common (13—38.8 %) and facies M2 is only occasional.
The FA5 is situated above FA1 or FA4 in the central part of

Fig. 3. Location of the studied boreholes and log cross-section (Fig. 5A-D).

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Table 1: Description and interpretation of lithofacies.

Symbol Description

Interpretation

Dark grey silty mudstone, massive to planar laminated with a occasional occurrence
of milimeters thick (max. 3 mm) mostly discontinuous laminas of very fine-grained
sand (with glauconite). Silt admixture is irregularly distributed. Abbundant content of
coalified plant detritus, which can be grouped into thin laminas. The bed thickness is
over 20 cm, sharp erosional base. Varied occurrence of light mica and pyrite.
Exceptional presence of foraminifera tests.

Dominant deposition from suspension,
occasionally disturbed by rapid
sediment delivery (upper flow regime,
storm activity, mass flows). The absence
of infauna, high content of plant debris
point to rapid deposition and/or oxygen
deficiency.

Dark grey massive to poorly visible planar laminated silty mudstone, bioturbated (the
bioturbation index varies). Typical is an occurrence of small (Ø max. 1cm)
subhorizontal tunnels with rounded or elliptical shape, filled with light grey very fine
sandstone. Admixture of coalified plant detritus and shell debris (molluscs,
foraminifera) is typical, similarly presence like presence of glauconite, white mica and
pyrite. The bed thickness ranges from 16 to 30 cm. The tops are sharp and erosional.

Deposition mostly from suspension,
suitable conditions (oxygen content,
sediment delivery, etc.) for the bottom
colonization, basinal environment.

Dark grey or dark brownish grey silty mudstone with occasional occurrence of laminas
(mostly discontinuous) or irregular thin lenses of light grey very fine to fine-grained
sandstone (max. about 1 cm thick) – lenticular bedding. The variable content of plant
debris, fine mica and shell fragments. The bed thickness ranges from 5 cm to 1 m. The
bases and tops are both flat/sharp.

The dominant deposition from
suspension occasionally alternates with
traction current deposition. Inner shelf
deposits (storm activity).

More or less regular alternation of laminas to thin beds (up to 2 cm thick) of dark grey
silty mudstone and beds (mostly 2–5 cm, max. 10 cm thick) of light grey very fine- to
fine-grained sandstone. Ripple cross lamination or planar lamination preserved in
thicker beds of sandstone. Thinner beds are typical by planar lamination. Wavy
bedding to rhytmites. Parting lineation. Very rare occurrence of foramineferal tests,
fish remnants, glauconite and plant fragments Abundant occurrence of synsedimentary
deformations (loading, pillow and flame structures, casps) and sole marks. Occurrence
of plant debris. Erosional bases and tops. Low bioturbation index.

The rapid (rhytmic) alternation of
deposition from traction currents and
deposition from suspension. Rapid
deposition, water escape deformation to
fluidization. Inner shelf deposits (storm
activity) and/or tidalites.

Light grey, micaceous, fine- to medium-grained sandstone with laminas and flasers of
dark silty mudstones (0.3 to 1.5 cm thick) – flasser bedding. Well preserved ripple
cross lamination in sandstones, sometime evident herringbone cross-lamination.
Mudstone laminas are often deformed (undulated, continuous lamina). The thickness
of beds ranges from 7–27 cm. The bases are flat and erosional. The tops are commonly
sharp and erosional. Parting lineation of whitish mica, rare occurrence of foraminefera
tests, spongies and glauconite.

Rapid alternation of sandstone
deposition from traction current and
mudstone deposition from suspension.
In some cases documented alternation of
flow directions (tidal activity), rapid
deposition.

Light grey, whitish grey, fine- or fine to medium grained sandstone, micaceous with
plane paralel lamination, well sorted. Bed thickness up to 20 cm. Sharp erosive base
(especially if superposed to fine grained M facies). Varied occurrence of plant
fragments (exceptionally up to 0.5 cm in diameter). Locally strings of small (about
2 cm) quartz pebbles (one grain thick).

Upper flow regime – upper shoreface
deposits, pebble strings point to storm
activity.

Light grey to whitish grey, fine to medium grained sandstone, micaceous. Structureless
due to intense bioturbation which complitely obliterated the primary structures. Rare
occurrence of small plant debris and fragments of organic rich mudstone.
Exceptionally 3 cm lense of quartzose gravellite (matrix supported – sand matrix).

Lower shoreface deposits – admixture of
coarser clasts points to storm activity.

Light grey locally green-grey, micaceous, fine or medium-grained sandstone, cross
stratified (ripple cross lamination), relative well sorted, exceptional occurrence of
small rounded quartz pebbles (up to 2 cm). The set thickness range between 15–20 cm,
thickness of cosets is about 40 cm. Sharp bases and tops. Sometime relative low degree
of bioturbation.

Traction current, lower flow regime,
unidirectional(?) flow, high energy
conditions, foreshore to upper shoreface
deposits.

Light grey locally green grey, fine grained sandstone, micaceous, ripple cross
laminated. Generally well sorted, but locally “nests” or scattered granules to small
pebbles (1–2 cm in diameter) of well rounded whitish quartzes. Sharp base locally (if
superimposed to M or H facies) loading structures. Varied occurrence of shell debris
and plant debris.

Lower flow regime, traction currents,
shoreface deposits.

Light grey localy green grey, micaceous, fine or medium-grained sandstone, undulated
lamination to hummocky cross-stratification (HCS) or its variations. Typically planar
lamination along the base, transition to undulated inclined/ripple lamination and finally
planar lamination. Upper interval of planar lamination is typical with alternation of
relative coarser and finer laminas. Irregularly inclined bases of sets. Relative common
was fining upward trend (both reduction of the grain size and admixture of silt towards
the top of sets). Irregular fragments (up to 2 mm) of mudstone sometime along the
base. Occurrence of light mica. The bed thickness usually several cm up to 15 cm.
Both the bases and tops are sharp and erosional. Varied but mostly low presence of
shell debris, pyrite, glauconite and plant fragments.

Combined action of current and wave,
result of storm activity. Deposition
below the fairweather wave base - lower
shoreace to inner shelf. Relative low
thickness and preservation of limited
part of the storm succession (Dott,
Bourgeois 1982) point to inner shelf
conditions.

Dark grey, fine to medium grained micaceous sandstone with abundant presence
several mm thick continuous laminas of coalified plant detritus or coal and mudstone
intraclasts. Typical is occurrence of pyrite. Facies recognized very exceptionally.

Intense and periodic delivery of plant
material – prohibited deposition
(backshore, lagoon, floodplain).

Light grey very coarse grained pebbly sandstone, gravellite or fine pebble
conglomerate. Subagular to angular clasts of deeply weathered and kaolinized
granitoids dominate within the pebbles. Well rounded quartzes (3 cm in diameter) are
less common. Only fragments of core preserved.

Partly short transport from adjacent
crystalline basement, proximity to
terrestrial conditions.

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

the area or on the base of the basin succession at the eastern
margin of the basin.

The lithofacies content of the FA1—5 can be compared

with the shape of well logs and combined with their areal po-
sition. Thus three depositional units (DU) can be identified.
Log cross-sections with spatial arrangement of these DU are
presented in Fig. 5A—D.

The DU I is located along the base of the basin fill. The

thickness  ranges  from  3  to  57  m  (Fig. 6)  and  is  greatest
around the borehole Vr 1. The whole unit reveals a broadly
lenticular shape, extended predominantly in the SW—NE di-
rection.  Two  subunits  (DSU)  with  different  facies  content
and  areal  position  can  be  recognized.  The  more  common
subunit IA is represented by the FA1 and is developed in the
central  part  of  the  studied  area.  The  subunit  IA  reveals
a  curved  funnel  shape  prolonged  in  the  NW—SE  direction
(Fig. 7A)  with  thickness  ranging  between  3  and  40 m.  The
subunit  IA  was  recognized  in  boreholes  Bar 1,  G 1,  2,  3,
Hul 1, Jar 1, Kro 1, 2, Rat 2, and Tl 1, 2 with a typical bell-
shape of the gamma and Sp logs (fining-upward trend).

The less common subunit IB was recognized in boreholes

Hol 1, Hul 2, 3, Ros 2, Slu 1, and Vr 1 and is represented by
the FA2. A generally blocky shape of gamma and Sp logs is
typical. However, the lower part of this subunit seems to be
less  monotonous  (slightly  irregular  and  funnel  shaped  of
logs), pointing to a coarsening-upward trend and alternation
of  sandstone  and  mudstone  beds.  The  upper  part  reveals
a  more  monotonous  blocky  shape  and  monotonous  litho-
logy. The subunit IB occurs along the southwest and north-
east  margins  of  subunit  IA  (Fig. 7B)  with  a  maximum
thickness of 54 m. The interfingering of subunits IA and IB
was recognized in wells G 3, Ros 2.

The DU II is situated in the southeast part of the area

(recognized  in  boreholes  G 1,  2,  3,  Ros 1,  2,  Slu 1,  Tl 1,  2,
and  Vr 1),  where  it  covers  either  subunit  IA  or  subunit  IB.
The  total  thickness  of  DU  II  ranges  from  20  to  426 m
(Fig. 8)  with  the  maximum  around  the  borehole  G 1.  The
unit  is  extended  in  the  SW—NE  direction,  parallel  with  the
thrust  front.  The  DU II  mostly  consists  of  FA3,  FA4,  and
FA5 and generally reveals a multiple repetition of a funnel-
shape (coarsening-upward trend) of well logs. Up to four of
such trends/cycles were recognized within the succession of
the  DU  II  with  thicknesses  of  each  one  ranging  from  12  to
55 m. Commonly a succession of FA5—FA4—FA3, were ob-
served from the base to the top of the cycle; however an in-
complete  succession  is  commonly  preserved.  The  lower
fine-grained  part  (FA5  and  FA4)  is  significantly  thicker
(about 15 m) than the sandier part (FA3). The lowermost cy-
cle usually reaches the greatest thickness.

The DU III is either superposed to DU II (southeast part of

the area) or directly overlaps the deposits of DU I (northwest
part). The DU III is described in all studied boreholes except
for borehole G 1. The unit occurred as a continuous tabular
belt prolonged in the SW—NE direction (parallel to the thrust
front) and inclined towards the southeast (where tectonic re-
moval of the unit is supposed). The thickness of DU III
ranges between 200 and 1000 m (Fig. 9) and incr

eases to-

wards the fronts of nappes. The log curves mostly have
a monotonous flat or slightly irregular shape, which, together

with  values  of  Rag,  Sp,  and  gamma,  indicates  dominantly
fine-grained content, while sandstone interbeds are rare and
thin.  The  interpretation  is  confirmed  by  cores,  where  FA5
was  recognized.  Strongly  tectonically  deformed  (cracks,
slickenslides,  etc.)  deposits  of  DU III  have  been  identified
mainly within the borehole Vr 1.

Interpretation

The FA1 is interpreted as lagoonal/estuarine deposits

which are faunistically poor and lacking the presence of fea-
tures  of  the  open  sea  (Dalrymple  et  al.  1992;  Sacchi  et  al.
2014).  The  FA2  is  determined  as  backshore-shoreline
(strandplain?)  deposits  (Wright  et  al.  1979;  Walker  &  Plint
1992).  The  FA3  is  interpreted  as  foreshore  to  lower  shore-
face deposits. Herringbone cross-lamination points to tidally
influenced deposition (tidal flats – Reineck & Singh 1973).
A  mudflat  and  mixed-flat  depositional  environment  is  sup-
posed for a part of the FA3 (FitzGerald et al. 2012; Reynand
&  Dalrymple  2012).  The  FA4  is  interpreted  as  deposits  of
the  lower  shoreface  and  the  transitional  zone  to  the  inner
shelf (Duke 1985; Duke et al. 1991). The FA5 is interpreted
as deposits of the inner and outer shelf.

The lithological infill of DU I can be compared to the

Mušov Mb. (Adámek et al. 2003). The DU I deposits are pa-
leontologically depleted. The remains of Teleostei and frag-
ments of skeletons and teeth of epipelagic species
(Lepidopus) were described (Thonová et al. 1987).

The areal extent and sedimentary infill of the subunit IA

point  to  the  important  role  of  basement  morphology  on  its
formation,  probably  due  to  reactivation  of  the  basement
faults in the compressive regime of the early stages of fore-
land  basin  evolution  (similarly  see  Gupta  1999;  Krzywiec
2001; Oszczypko et al. 2006; Nehyba & Roetzel 2010). Ma-
jor  faults  trend  NW—SE,  downthrowing  mainly  to  the  SW.
We  can  speculate  about  the  existence  of  a  paleovalley,
trending in the NW—SE direction. The role of basement tec-
tonics is also supported by the preservation of the Paleozoic
deposits  within  this  paleovalley  surrounded  by  crystalline
rocks on its margins. The depositional subunit IA could rep-
resent  an  “Early  transgressive  systems  tract”  (Koss  et  al.
1994;  Shanley  &  McCabe  1994)  due  to  localized  preserva-
tion, dominant vertical accretion, and its position below the
major transgressive surface.

The subunit IB is connected with the marine transgression

on the distal northwest margins of the basins. The sedimen-
tary infill might be connected with the distal part of a valley
(Cattaneo & Steel 2003) adjacent to the marine basin (to the
southeast) with a proximity to the terrestrial environment (to
the northwest?). A highly varied morphology of the coastal
area with a steep and irregular relief along the margins of the
paleovalley  and  the  flatter  and  more  depressed  relief  in  its
central part of valley is supposed. Such a situation could lead
to more rapid and further ingression of the marine flooding
into  the  foreland  through  the  valley.  Numerous  small  base-
ment  elevations  (forming  shoals,  small  islands  or  intrabasi-
nal highs) and a highly irregular pattern of the shore favour
significant differences in the slope and width of the coastline

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

for offshore sand delivery. The DU II is mostly interpreted as
a retrograding parasequence set, when general transgressive
trends  and  shifts  of  the  shoreline  to  the  foreland  are  inter-
rupted  by  short-term  regressions  (undulating  “zig-zag”
shoreline trajectory, Helland-Hansen & Gjelbeg 1994). Con-
tinuous  shoreline  migration  led  to  the  enlargement  of  the
shallow  marine  belt  and  a  “gradual”  reduction  of  the  sedi-
ment input into the basin, sediment reworking, and redeposi-
tion. We can partly speculate about “expanded backstepping
parasequences”  (Swift  et  al.  1991),  which  point  to  greater
formation of accommodation space than  sediment delivery.

Significant variations in the thickness of parasequences

and parasequence sets were recognized in the central part of
the  proposed  paleovalley  compared  to  its  marginal  parts
(similarly  see  Dalrymple  et  al.  1992;  Zaitlin  et  al.  1994),
and, similarly, differences in the thickness of the sandy beds
were recognized. Although variations in the shoreline trajec-
tory could reflect variations in sediment delivery, they most
probably  reflect  variation  in  the  basement  relief,  slope,  and
character  of  the  underlying  rocks  in  the  area  under  study.
A relatively slow shoreline retreat with a large role of wave
erosion and sediment reworking and so thicker and less ex-
tended transgressive deposits could be expected in the areas
with  higher  shoreline  slope.  However  in  the  flat  area,  very
rapid shoreline retreat and thinner and more extended trans-
gressive deposits can be supposed. Local structural elevation
could  also  serve  as  the  source  for  the  material  in  the  basin
and  further  affect  the  lateral  variability  of  the  transgressive
deposits.

The sedimentary infill of the DU III corresponds to the

Nový Přerov Mb. with almost faunistically sterile deposits,
where individual specimens of Haplophragmoides cf.

vasiceki  were  described  (Thonová  et  al.  1987).  The  DU III
represents  the  shallow  marine  deposits  of  the  inner  shelf,
where relatively quiet deposition was rarely disrupted by storm
activity (similarly Dott & Bourgeois 1982; Duke 1985). The
great  thickness  and  monotonous  character  of  deposits  point
to  an  aggrading  stacking  pattern  with  a  large  and  balanced
formation  of  accommodation  space  and  sediment  delivery.
However,  both  the  shape  and  the  thickness  of  the  DU III
were  probably  significantly  affected  by  later  thrusting.
Occurrence  of  both  deformed  and  undeformed  deposits  of
the DU III confirmed that the unit is partly composed of du-
plicated  sediment  packages  formed  by  deformation  or  en-
trainment of pre-existing basin sediments during the thrusting.

Provenance analyses

A wide spectrum of methods of provenance analysis

(petrography of sandstones, heavy mineral assemblages,
chemistry of garnet, rutile, major and trace elements) has
been used for better determination of source area and

its

evo-

lution.

Petrography of sandstones

The studied sandstones are in general very fine- to medium-

grained.  The  sandy  grains  have  angular  and  sub-angular
shapes.  The  sandstones  are  moderately  sorted  and  contain
a  large  amount  of  monomineral  quartz.  Aggregate  quartz
was detected less frequently. The subhedral grains of K-feld-
spars and plagioclases represent the common component and
probably originated from cataclased granitoids. The content

Fig. 9. Thickness map of the depositional unit III.

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

of  micas  (both  biotite  and  muscovite)  varied  slightly  and
they  were  probably  derived  from  mylonites  or  cataclased
granitoids.  Clasts  of  gneisses,  granulites,  and  feldspar  phe-
nocrysts were also identified. Accessory minerals are repre-
sented  by  garnet,  zircon,  and  rutile.  The  presence  of
authigenic glauconite was common. Matrix is mainly of the
coating  type.  Cement  is  formed  by  carbonates  and  Fe  oxy-
hydroxides.

The samples of the DU II are predominantly spread within

the arkoses field, whereas the samples of the DU III are
situated in the wack

e field (Fig. 10A – P

etránek 1963;

Kukal  1986).  Some  differences  in  the  provenance  of  sand-
stones of the DU II and the DU III are also detected from the
results of the discrimination diagram Qm—F—Lt (Fig. 10B –
Dickinson  &  Suczek  1979;  Dickinson  1985).  While  the
majority  of  the  samples  from  the  DU III  occupies  the  recy-
cled orogen field, the samples from the DU II are more wide-
spread and are located in the magmatic arc fields. Although
in  the  Q—F—L  diagram  (Fig. 10C  –  Dickinson  &  Suczek
1979; Dickinson 1985) the samples from both the DU II and

the DU III fall into the recycled field, the samples from the
DU II reveal a closer relation to the magmatic arc fields.

Studies of heavy minerals

Heavy minerals associations

The dominance of garnet is typical for the heavy minerals

association of the studied deposits. Garnet occurrence ranges
between  27.1  and  95.4 %  (average/AVG  70.4 %).  Further
commonly  identified  heavy  minerals  were  apatite  (0.9—
40.3 %, AVG 10 %), zircon (1.2—24.8 %, AVG 4.5 %), tour-
maline  (5.1—22.1 %,  AVG  4.5 %),  and  rutile  (0.7—41.3 %,
AVG 6.2 %). The occurrence of kyanite, sillimanite, stauro-
lite, monazite, anatase, titanite, epidote, amphibole, and py-
roxene  varies  greatly  and  reaches  a  maximum  of  1 %.  The
value of the ZTR (zircon+tourmaline+rutile index – Hubert
1962;  Morton  &  Hallsworth  1994)  varies  greatly,  ranging
between 8.4 and 64.9 (AVG 12.3 %). Any significant trends
were recognized in the heavy mineral assemblages or value
of the ZTR index within the studied succession.

The important occurrence of ultrastable heavy minerals is

generally  typical  for  deeply  buried  deposits  and  can  be  ex-
pressed  by  the  garnet/zircon  (GZi)  index  (Morton  &
Hallsworth  1994;  1999).  According  to  Morton  (1984),  gar-
nets underwent dissolution when buried at depths exceeding
3 km.  Milliken  (1988)  placed  this  dissolution  at  a  depth  of
4 km. However, the average values of the GZi index for G 2
(samples  from  depths  of  4297—4484 m)  and  Slu 1  (samples
from  depths  of  3203—3697 m)  vary  between  88  and  95 %,
which do not indicate such a burial effect.

Garnet

Garnet, as the most common heavy mineral, was further

evaluated by analysis of its chemistry, which is widely used
for the determination of provenance (Morton 1991; Morton
et al. 2004; Salata 2004, 2013a; b; Nehyba et al. 2012; Sug-
gate & Hall 2013). The composition of garnets was diverse,
and 10 garnet types (T1—T10) were determined. T1 is com-
posed of almandine garnets with low contents of pyrope,
grossular, and spessartine components and the usual compo-
sition is in the range of Alm

76—93

Prp

5—9

Grs

0—9

Sps

0—7

. T2 con-

sists of pyrope-almandine garnets with the composition
Alm

67—86

Prp

11—19

Grs

0—9

Sps

2—8

. T3 is composed of pyrope-

almandine garnets with an enriched pyrope component
and the typical composition is in the range of
Alm

45—77

Prp

21—46

Grs

0—6

Sps

0—3

. T4 is represented by grossu-

lar-almandine garnets with a composition in the range of
Alm

59—78

Grs

11—20

Prp

6—9

Sps

1—7

. T5 consists of grossular-

almandine garnets with increased content of grossular and the
composition Alm

53—68

Grs

21—32

Prp

5—9

Sps

1—9

. T6 is represented

by pyrope-almandine garnets with increased contents of gros-
sular and the composition Alm

55—75

Prp

12—27

Grs

11—18

Sps

0—3

. T7 is

formed by grossular-almandine garnets with increased contents
of pyrope and the composition Alm

40—76

Grs

12—28

Prp

11—26

Sps

1—6

T8 consists of spessartine-almandine garnets and a composi-
tion in the range of Alm

47—72

Sps

11—38

Prp

3—9

Grs

0—9

. T9 is com-

posed of almandine-spessartine garnets with the composition

Fig. 10. A – Classification ternary diagram (according to Petránek
1963;  Kukal  1986)  of  the  studied  sandstones.  M  =  matrix  (%),
F = plagioclase + K-feldspar (%), U = unstable rock fragments (%),
Q = quartz (%), S = stable rock fragments (%). B, C – Discrimi-
nation  ternary  diagrams  (according  to  Dickinson  1985)  of  the  stu-
died  sandstones.  (Q  =  Q

+ Q

, Q

– monocrystalline quartz,

– polycrystalline quartz; F = plagioclase + K-feldspar;

L = L

+ L

, L

– volcanic lithic fragments, L

– sedimen-

tary lithic fragments, L

– metamorphic lithic fragments;

= L + Q

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Sps

46—53

Alm

31—36

Prp

5—10

Grs

0—5

. Lastly, T10 is represented

by garnets with varied compositions and individual
occurrences. The representatives of this group are
Sps

41—45

Alm

36—39

Prp

13—16

(3 grains), Alm

55—61

Grs

17—18

Sps

11—15

(3 grains), Alm

37—59

Sps

17—37

Grs

(3 grains) Sps

Alm

Grs

Sps

Alm

, Grs

Alm

Sps

, and Prp

Alm

Grs

(one grain). Table 2 shows the relative abundance of these
garnet types in the studied samples.

The distribution of recognized garnet types T1—10 varies

within the identified DU. The subunit IB is characterized by
dominance  of  T1  (20.3 %),  T4  (16.9 %),  and  T2  (15.3 %).
The  most  comon  garnets  in  the  DU II  are  T1  (25.8 %),
T4  (21.1 %),  and  T2  (18.3 %).  The  DU III  is  characterized
by T2 (22.5 %), T4 (17.8 %), and T3 (14.6 %). The subunit
IB  is  characterized  by  a  significantly  higher  proportion
(10.2 %) of spessartine-almandine garnets (T8) compared to
the DU II (3.3 %) and the DU III (4.2 %). The detrital garnet
assemblages  of  the  DU II  contain  significantly  higher  num-
ber  of  (25.8 %)  of  almandine  garnets  (T1)  than  the  DU III
(8.0 %). The representation of specific types of garnets in the
individual DU is illustrated in Fig. 11.

The composition of the garnets of the studied Karpatian

deposits could be compared with the results of garnet analy-
ses from some potential source rocks, that is, the underlying
crystalline  rocks  of  the  Brunovistulicum,  sandstones  and
greywackes of the Moravo-Silesian Paleozoic deposits (Dra-
hany  and  Nízký  Jeseník  Culmian  facies),  sandstones  of  the
Magura Group (Rača Unit), sandstones of the Krosno-Meni-
lite  Group  (Silesian  Unit),  and  older  Eggenburgian-Ottnan-
gian infill of the Carpathian Foredeep Basin.

The pyrope-almandine-spessartine and pyrope-spessartine-

almandine garnets with compositions in the range of Sps

39—47

Alm

35—42

Prp

13—16

Grs

0—4

and Alm

38—46

Sps

35—40

Prp

13—15

Grs

0—2

are typical of the underlying crystalline rocks.

The detrital garnets of the older parts of the Moravian-

Silesian Paleozoic deposits (Protivanov and the lower part
of  the  Myslejovice  Formation)  are
clustered in the field of spessartine
in the ternary diagram Grs—Prp—Sps
(Fig.  12A).  The  upper  part  of  the
Myslejovice  Formation  contains
predominantly

pyrope-almandine

garnets  (Otava  et  al.  2000;  Čopja-
ková  et  al.  2002;  2005;  Čopjaková
2007)  clustered  in  the  field  of  py-
rope (see Fig. 12B).

The pyrope-almandine garnets

dominate in the Rača Unit (Otava et
al.  1997;  1998)  with  a  trend  of
Prp—Sps  in  the  ternary  diagram
Grs—Prp—Sps  (see  Fig.  12C).  The
pyrope-almandine  and  grossular—
almandine  garnets  predominate  in
the  Silesian  Unit  (Stráník  et  al.
2007)  and  form  a  linear  trend
Prp—Grs  in  the  ternary  diagram
Grs—Prp—Sps (see Fig. 12D).

Similarly, the pyrope-almandines

are the predominant garnets of

Table 2: Garnet types of the studied deposits (Alm – almandine, Grs – grossular,
Prp – pyrope, Sps – spessartine).

DEPOSITIONAL

SUBUNIT IB

[51 grains]

DEPOSITIONAL

UNIT II

[126 grains]

DEPOSITIONAL

UNIT III

[159 grains]

Alm

76–93

20.3% 25.8%

8.0%

Alm

67–86

Prp

11–19

15.3% 18.3% 22.5%

Alm

45–77

Prp

21–46

13.6%

6.6%

14.6%

Alm

59–78

Grs

11–20

16.9% 21.1% 17.8%

Alm

53–68

Grs

21–32

3.4%

4.2%

11.3%

Alm

55–75

Prp

12–27

Grs

11–18

6.8%

5.2%

6.1%

Alm

40–76

Grs

12–28

Prp

11–26

5.1%

12.2%

10.3%

Alm

41–72

Sps

11–42

10.2%

3.3%

4.2%

Sps

46–53

Alm

31–36

3.4%

1.4%

–

Sps

41–45

Alm

36–39

Prp

13–16

–

1.4%

Alm

55–61

Grs

17–18

Sps

11–15

3.4%

0.5%

0.9%

Alm

37–59

Sps

17–37

Grs

1.6%

1.4%

0.9%

Sps

Alm

Grs

–

0.5%

Grs

Sps

Alm

–

0.5%

Grs

Alm

Sps

–

0.5%

Prp

Alm

Grs

–

0.5%

the Eggenburgian and Ottnangian deposits (Nehyba &
Buriánek 2004).

Rutile

The concentration of the main diagnostic elements of

rutile (Fe, Nb, Cr, and Zr) varies significantly in the studied
samples.  The  concentration  of  Fe  ranges  between  450  and
9390  ppm  (average  3184  ppm),  the  concentration  of  Nb
ranges  between  70  and  9630  ppm  (average  2636  ppm),  the
concentration  of  Cr  ranges  between  0  and  2680  ppm
(average  643  ppm),  and  the  concentration  of  Zr  varies  be-
tween 20 and 5390 ppm (average 438 ppm).

The content of Fe is used an as indicator to distinguish the

igneous  and  metamorphic  origin  of  rutiles  (Zack  et  al.
2004a;  b).  Provenance  from  igneous  rocks  was  very  minor
and such rutiles represent about 9.7 % of the studied spectra.
The  absolute  majority  of  the  investigated  rutiles  originated
from  metamorphic  rocks.  Rutiles  derived  from  metapelitic
rocks  dominate  in  the  DU  III  (87.5  %)  over  rutiles  from
metamafites  (12.5  %).  Similarly,  in  the  DU  II  metapelitic
(75.6 %) rutile predominates over metamafic (24.4 %) ones
(see Fig. 13). The rutiles from the DU III (47 grains) reveal
higher  contents  of  Nb  and  Cr  compared  to  the  rutiles  from
the  DU  II  (43  grains).  It  is  supposed  that  felsic  granulites
and  paragneisses  dominate  and  garnet  amphibolites,  and
eclogites  have  a  subordinate  role  in  the  source  area  (see
Meinhold  et  al.  2008).  The  contents  of  Zr  in  rutile  can
be  used  as  a  thermometer  (Zack  et  al.  2004a;  Watson
et al. 2006; Meinhold et al. 2008; Meinhold 2010). The cal-
culated  temperatures  range  between  372—1088  °C  (equation
Zack  et  al.  2004)  and  464—958  °C  (equation  Watson  et  al.
2006) respectively (Fig. 14). The Figure 15 shows the possible
relative  abundance  of  different  metamorphic  facies  in
the  source

area of the studied deposits with dominance of

medium- to high-temperature amphibolite/eclogite facies.

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Fig. 13. Discrimination plot Cr vs. Nb of investigated rutiles
(according to Meinhold et al. 2008).

Fig. 14. Histogram of calculated temperatures for rutiles from
DU II and DU III. Temperatures were calculated based on the equa-
tion of Zack et al. (2004a) and equation of Watson et al. (2006).

Fig. 15. Diagrams showing the percentage of the different metamor-
phic facies for the source rocks according to calculated temperatures.

cluster (sandstones), whereas higher values of Al

and

TiO

are typical for the second one (mudstones). While posi-

tive correlation of Al

and TiO

is typical for the data from

the DU I and the DU II, the results for the DU III shows sig-
nificant differences in the Al

content and a relatively

stable content of TiO

. Such a horizontal distribution reflects

the different weathering and sorting of the sand fraction.

The TiO

/Zr versus Zr/Al

diagram (Fig. 16B) reveals

an almost linear distribution and negative correlation.
Generally similar arrangements for results from individual
depositional units can be followed in this diagram as in the
Al

versus TiO

diagram.

Distinct trends in the Al

versus TiO

and TiO

/Zr ver-

sus Zr/Al

diagrams are significantly influenced by grain-

size sorting. The main TiO

origin probably come from

minerals such as biotite, amphibole, and pyroxene or olivine
(Nesbitt 1979; Taylor & McLennan 1985). The low Al

and TiO

contents of the studied deposits could point to

a source from granulites and granitoids (Passchier & White-
head 2006), and the position of all samples in an almost curvi-
linear line suggests a uniform/similar source (Fralick &
Kronberg 1997; Passchier & Whitehead 2006).

The studied samples are sedimentary rocks from heteroge-

neous  source  rocks  which  might  have  undergone  different
weathering  processes  and  intensities.  The  values  of  CIA
range from 56.84 to 80.89 for the studied samples. The CIA
of sediments is, in general, about 50 in the case of first cycle
sediments  and  tends  to  increase  as  chemical  weathering  in-
tensifies (Nesbitt & Young 1982). The obtained values cor-
respond  to  intermediate  chemical  weathering  and  an
important source from recycled material (Fedo et al. 1995).
The variations in CIA reflect differences in the proportion of
the  content  of  weathered/recycled  material.  The  data  were
plotted  in  the  A—CN—K  (Al

—(CaO+Na

O)—K

O diagram

(Fig. 16C). All the studied samples are distributed parallel to
the A—CN axis and follow a trend of increasing Al

with

decreasing CaO+Na

O. The elongated distribution reflects

the varied role of the weathering trend/clay minerals and can
be associated with grain-size variations (Corcoran 2005).
The A—CN—K triangle is also used to determine the composi-
tion of the parent rocks (Fedo et al. 1995) and the results in-
dicate a similar source.

SiO

/Al

is known as the index of chemical maturity of

sediments (Roser et al. 1996; Roser & Korsch 1999). The
SiO

/Al

ratio ranges from 2.54 to 13.82 (AVG 5.10). Ac-

cording to Zhang (2004), the low SiO

/Al

ratios indicate

a low sediment recycling and deposition from a nearby
source. Negative correlation between SiO

and Al

for the

studied samples can be followed in the Fig. 16D and is a re-
sult of the varied presence of mud material. A similar distri-
bution of samples in Fig. 16A—E points to the principal role
of grain size (Ross & Bustin 2009; Adegoke et al. 2014).

The degree of chemical maturity of sediments is also ex-

pressed by the diagram of Al

O+Na

O versus SiO

(Suttner & Dutta 1986; Sen et al. 2012). The studied samples
have values of chemical maturity in the range from 0.10 to
0.51 (AVG 0.3). The results in the arid field are derived from
mudstones while the results in the humid field are derived
from sandstones (Fig. 16E).

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

With no extra input of detritus, the sediment recycling re-

sults in a negative correlation of SiO

and TiO

(Gu et al.

2002). Such a trend can generally be followed for the studied
samples (Fig. 16F). In the ideal case (Cox & Lowe 1995;
Corcoran 2005) the overlying sequence/formation should

contain more quartz (i.e. SiO

) and less feldspar and clays

(lower contents of TiO

, Al

, and MgO). The studied

cases show the increase of the sediment recycling for the
succession of the DU, but the results are also influenced by
grain-size sorting.

Fig. 16. Discrimination plots of major element geochemistry. A – TiO

vs. Al

, B – TiO

/Zr vs. Zr/Al

, C – ternary diagram

—(CaO+Na

O)—K

O, D – Al

vs. SiO

, E – SiO

vs. Al

O+Na

O, F – SiO

vs. TiO

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Ali et al. 2014). The studied samples have an extremely low
Cr/V ratio and low Y/Ni ratios, indicating predominantly fel-
sic source rocks.

Interpretation

The poor development of the burial effect on the associa-

tion of the heavy minerals and garnet chemistry could be ex-
plained by the relatively short duration of diagenetic
processes. These results signify that the studied heavy mi-
nerals associations were dominantly influenced by the com-
position of the source rocks.

The rutile analyses proved that the primary source was

crystalline rocks of the Bohemian Massif (Moldanubian and
Moravian  Zones).  The  rocks  of  the  Bohemian  Massif  were
recognized  in  the  Račice  and  Luleč  conglomerates  of  the
Moravo-Silesian  Paleozoic  deposits  (Čopjaková  2007)  and
these conglomerates are probably supposed to represent the
source of the studied deposits.

Although similar heavy mineral associations are known

from both the Rača Unit (Gilíková et al. 2002) and Myslejov-
ice Formation (Otava 1998; Čopjaková 2007), the chemistry
of garnets discriminates between the sources more precisely.

Occurrence of spessartine-almandine garnets (T8) proved

that  the  crystalline  rocks  of  the  local  basement  can  be  con-
sidered as an important source for the subunit IB. The source
of  almandine  garnets  (T1)  can  be  traced  well  to  the  Luleč
conglomerates  (the  young  part  of  Myslejovice  Formation)
(Čopjaková  2007).  These  garnets  are  missing  in  the  Rača
Unit  and  the  Silesian  Unit  that  contain  pyrope-almandine
(T2, 3) and grossular-almandine garnets (T4, 5). It can be as-
sumed that the early phases of the sedimentation (DU I and II
—  the  forebulge  depozone)  were  supplied  from  the  passive/
foreland  margin  of  the  basin,  that  is,  from  the  Bohemian
Massif.  The  slight  differences  in  provenance  between  units
I  and  II  (the  variations  in  the  role  of  the  local  crystalline
basement  and  the  more  distant  Moravo-Silesian  Paleozoic
deposits)  can  be  explained  by  the  basin  and  source  area
enlargement.

Both the small content of garnet T1 and the abundance of

T3 point to a change of the source area for the later phases of
deposition (i.e. the DU III). The significant role of the source
from the Silesian Unit (active margin — the Western Car-
pathian Flych Zone) and its mixing with material from the
Bohemian Massif is proposed. This points to shift of the de-
pozones towards the orogenic wedge, that is, to the foredeep
depozone.

The cannibalization of the older basin infill (Eggenburgian

and Ottnangian deposits) and its role as a partial source for
the studied Karpatian deposits cannot be excluded.

The results of garnets are partly supported by petrography

(see Fig. 10B).

Discussion

The stacking pattern of facies associations of the Karpa-

tian basin infill in the studied well cores, the temporal and

spatial  evolution  of  the  depozones,  and  the  results  of  the
provenance  analyses  can  all  be  summarized  and  evaluated
according to the principles of the sequence stratigraphy and
proposed  models  of  peripheral  foreland  basin  evolution
(Flemings  &  Jordan  1990;  Jordan  &  Flemings  1991;  Beau-
mont et al. 1993; Crampton & Allen 1995; DeCelles & Giles
1996). The recognized paleo-environmental changes and pa-
leogeographic  evolution  of  the  area  point  to  significant  re-
construction  of  the  basin  shape  and  geometry,  lateral  shifts
of

the depozones and an important role of basement mor-

phology. A model of the evolution of the basin in the area
under study with several depositional stages can be proposed
(see Fig. 18):

Stage 1: Increased crustal loading by a thickened orogenic

wedge results in subsidence of the inner part of the basin and
a coeval uplift of the forebulge (Fig. 18A). The fault reacti-
vation (Holešov

faults) facilitated the varied/predisposed cre-

ation of an accommodation space in the forebulge depozone.
Eastward to southeastward flexural dip of the passive basin
margin  is  supposed,  with  an  irregular,  generally  SW—NE,
prolonged  shoreline  disturbed  by  a  perpendicular  paleoval-
ley  oriented  in  the  NW—SE  direction  (along  the  basement
faults,  see  Fig. 5A,  B)  where  started  sedimentation  of  the
depositional subunit IA (Pre- to Early transgressive deposits).
The basin depozones continued towards the E-SE.

The stage corresponds to a major pulse of the Carpathian

nappe-pile contraction by the Early Styrian phase of
thrusting (Ottnangian—Early Karpatian). The stage can be
compared with sequence II of Nehyba & Šikula (2007).

Stage 2: The subsequent major marine transgression

drowned  the  embayed  margins  of  the  basin.  This  transgres-
sion  can  most  probably  be  correlated  with  the  Karpatian
TB  2.2.  sea-level  cycle  (Haq  et  al.  1988;  Nehyba  &  Šikula
2007; Hohenegger et al. 2009; 2014). Basin facies belts had
been  shifted  cratonward  (i.e.  generally  towards  the  north-
west)  and  the  adjacent  pre-Neogene  basement  had  been
drowned. The position of the shoreline (peripheral bulge de-
pozone) was located in the area under study and is represen-
ted by the depositional subunit IB. The lagoonal depositional
environment,  represented  by  depositional  subunit  IA,  was
developed behind the barred coast. This situation is reflected
by  the  NE—SW  distribution  of  depositional  subunit  IA  de-
posits  northwestward  to  the  depositional  subunit  IB  (see
Fig. 7A). The continued cratonward shift of the facies belts
led  to  the  transition  of  the  shoreline  condition  to  a  shallow
marine one (transition of depositional subunit IB to the DU II)
with more rapid shift towards the northwest within the paleo-
valley (Fig. 18B). This situation probably mainly reflects the
gradual sea-level rise, although early phases of the forebulge
retreat might also play a role (Plint et al., 1993; Lesczyński
&  Nemec  2014).  The  interplay  between  sediment  supply,
basement  relief,  eustatic  sea-level  changes,  and  tectonics
affected the frequent bathymetric changes and shoreline tra-
jectory (dynamic stratigraphy of foredeep peripheral uncon-
formity - see Lesczyński & Nemec 2014).

The Neogene tidalites within the Central Paratethys sedi-

mentary basins are known from the Pannonian Basin
(Eggenburgian) (Sztanó & de Boer 1995), the North Alpine
Foreland Basin (Ottnangian) (Bieg et al. 2007, 2008),

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

the Korneuburg Basin (Karpatian) (Zuschin et al. 2014), and
the  Polish  and  Ukrainian  parts  of  the  Carpathian  Foredeep
(Badenian)  (Lis  &  Wysocka  2012).  The  tidal  flat  deposits
were  not  recognized  in  the  sedimentary  succession  of  the
Moravian  part  of  the  Carpathian  Foredeep  and  their  excep-
tional  occurrence  is  described  here  only  from  the  spatially
restricted area. Such a situation points to an embayed coast.
A  typical  environment  for  the  development  of  the  tidal  flat
deposits is a tide-dominated estuary (Dalrymple et al. 1992),
where  its  funnel  shape  tends  to  increase  the  tidal  current
strength  (Dalrymple  &  Choi  2007).  The  tidal  flat    deposits
probably originated in micro- to meso-tidal conditions.

Stage 3: The rapid forebulge flexural retreat and sediment

delivery  from  the  active  margin  are  both  reflected  by  the
DU III (Fig. 18B), which was deposited within the foredeep
depozone (i.e. the inner part of the foreland basin). The fore-
bulge  disconformity/shoreline  was  rapidly  shifted  craton-
ward (northwestward) out of the area under study (Fig. 5A).
The  position  of  the  coastline/peripheral  bulge  depozone
was probably located more than 50 km northwest. The closest
recognized  coastline  deposits  of  Karpatian  age  were  docu-
mented  in  the  northwest  of  Prostějov  in  the  borehole
Slatinky  MH-10  (Bubík  &  Dvořák  1996).  The  basin  was
significantly enlarged and its bathymetry changed. A similar
cratonward  shift  of  the  foreland  basin  depozones  was
described  in  the  North  Alpine  Foreland  Basin  in  Switzer-
land,  where  the  forebulge  was  shifted  100  km  towards
the  northwest  (Kempf  &  Pfiffner  2004).  This  stage  is  con-
nected  with  orogenic  quiescence  and  increased  sediment
supply  from  the  active  margin  (Flemings  &  Jordan  1989;
Johnson  &  Beaumont  1995).  Stages  2  and  3  represent  the
time  period  between  the  Early  and  Late  Styrian  phases  of
thrusting and are compared with Sequence III of Nehyba &
Šikula (2007).

Stage 4: Renewed loading by orogenic wedge results in

subsidence of the inner part of the basin and a coeval uplift
of the forebulge. Continuous thrusting led to tectonic defor-
mation of the basin infill and its partial incorporation in the
thrust  belt  in  the  most  proximal/inner  part  of  the  basin.
Generally westward advance of the nappes led to shift of the
lithosphere flexure cratonward and the basin subsidence axis
northwestward of the area under study. The inner part of the
basin  was  buried  under  the  nappes  or  destroyed.  Here  the
original  thickness  of  individual  depositional  units  was  tec-
tonically  reduced  on  one  side  and  enlarged  on  the  other
(Fig. 18C).  The  thicknesses  of  the  DU  II  and  the  DU  III
in  the  studied  area  are  relatively  large.  These  thicknesses
are  explained  by  the  duplicated  sediment  packages  by  so-
called overlap of the allochthonous and para-autochthonous
Karpatian deposits before the front of nappes. The southeast
part of the area under study (around the boreholes G 1, G 2,
and G 3) is characterized by complete removal of the DU III
and the replacement of the DU II (para-autochthon). The tec-
tonically  removed  deposits  of  the  DU  III  were  shifted  to-
wards  the  northwest.  Benada  (1986)  recognized  the
Eggenburgian  deposits  (Ždánice  Unit)  between  the  Karpa-
tian  deposits  in  the  borehole  Rat  1  and  determined  the
spreading of allochthonous deposits of the Carpathian Fore-
deep  mainly  in  the  northern  part  of  the  studied  area  (bore-

holes Bar 1, Kro 1, 2, and Rat 1, 2). The basin was limited on
its  cratonward  margin  by  the  uplifted  forebulge,  which  fur-
ther  affected  basin  extent  and  bathymetry,  so  the  resulting
foreland basin was narrower and deeper. In short, the basin
underwent  a  complete  reconstruction  of  its  shape,  position,
and lateral extent. This stage can be connected with the Late
Styrian phase of thrusting (Late Karpatian) and is compared
with Sequence IV of Nehyba & Šikula (2007).

Stage 5: The most northwestern part of the area under

study  belonged  to  the  proximal  part  of  the  foreland  basin
along  its  active  margin.  Sediment  delivery  was  from  the
stacked  thrust  zone  and  was  highly  sensitive  to  individual
tectonic  events.  This  situation  is  connected  with  mass  flow
deposits. The position of the active margin of the basin was
probably generally similar to the present state. Further depo-
sitional zones of the basin were located more to the W-NW,
where they are also preserved (Fig. 18D).

More precise timing of these stages by biostratigraphic

data are of little help, because the whole Karpatian is within
a single nannoplankton zone NN 4.

Similarly, the inner part of the Polish Carpathians, located

beneath the Carpathian nappes, is more than 50 km wide and
is  composed  of  Early  to  Middle  Miocene  deposits,  up  to
1500 m  thick.  However,  these  Lower  Miocene  strata  are
mainly terrestrial in origin (coarse grained alluvial deposits)
with a significant portion of material derived from the Car-
pathian  nappes  (Oszczypko  &  Oszczypko-Clowes  2012).
Differences in the depositional environments along the basin
active  margin  and  in  the  recognized  role  of  the  Carpathian
nappes  in  the  provenance  (Moravian  vs.  Polish  part  of  the
basin)  point  to  significant  variation  in  the  basin  geometry,
paleobathymetry  and  sedimentary  budget  (underfilled  stage
vs. overfilled conditions).

Conclusions

The Karpatian deposits of the central part of the Car-

pathian Foredeep in Moravia, are known only from the deep
boreholes  since  they  are  deeply  buried  under  the  Flysch
Thrust  Wedge  of  the  Outer  Western  Carpathians.  These
rocks  have  been  studied  by  a  complex  of  sedimentological
and sedimentary-petrographical methods with the aim of im-
proving our understanding of the paleogeography and evolu-
tion of the Carpathian Foredeep basin and reconstructing the
former  evolutionary  stages  of  this  peripheral  foreland  basin
and its paleogeography. The chemistry of detrital garnets has
proved to be an important indicator of changes in the prove-
nance of the deposits.

Three depositional units were determined. They differ in

their  depositional  environment,  basin  depozones  and  reflect
successive  stages  of  the  basin’s  evolution.  Depositional
unit I  is  represented  by  lagoon-estuary  and  barred  coastline
deposits  (backbulge  and  forebulge  depozones)  and  reflects
the Pre- to Early transgressive phase over the pre-Cenozoic
bedrock.  The  deposition  was  influenced  by  reactivation  of
the basement faults due to the Early Styrian phase of thrusting
(Ottnangian—Early  Karpatian).  The  source  from  the  “local”
crystalline  basement  formed  by  crystalline  rocks  of  the

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Brunovistulicum played an important role here. The second
depositional  unit  is  represented  by  the  coastline  to  shallow
marine deposits (forebulge depozone), reflecting the Karpa-
tian transgression (TB 2.2. sea-level cycle), which drowned
the embayed passive margins of the basin. Continued craton-
ward shift of the facies belts is typical for this unit. Tidalites
(tidal flat deposits) recognized within this unit represent the
only  described  tide-generated  deposits  of  the  Carpathian
Foredeep  basin  in  Moravia.  The  source  from  the  basin  pas-
sive margin (especially from greywackes and conglomerates
of the Moravian—Silesian Paleozoic deposits/Culm unit) has
been  proved.  The  third  depositional  unit  is  formed  by  off-
shore  deposits  and  represents  a  foredeep  depozone.  The
provenance from both passive and active basin margin (Sile-
sian Unit of the Western Carpathian Flysch Zone) has been
proved. The cratonward forebulge flexural retreat continued.
The succession of depositional units reflects both a stepwise
migration of the foredeep basin axis and shift of basin depo-
zones  outwards/cratonwards,  together  with  the  forebulge
retreat. The shift of the foreland basin depozones more than
50  km  cratonward  can  be  assumed.  The  basin  was  signifi-
cantly enlarged and its bathymetry changed.

Renewed thrusting (Late Styrian phase of thrusting—Late

Karpatian) led, in the most proximal/inner part of the basin,
to tectonic deformation of the basin infill and its partial in-
corporation into the thrusts. Generally westward advance of
the nappes led to shift of the lithosphere flexure cratonward
and the basin subsidence axis northwestward of the area
under study. The inner part of the basin was buried under the
nappes or destroyed. The deposition in the foreland basin
continued only in the flexed periphery in front of the nappes,
so the basin was narrower and deeper.

The complete reconstruction of the basin’s shape, position,

and lateral extent are all interpreted as a consequence of the
phases of the thrust-sheet stacking and sediment loading in
combination with eustatic sea-level change.

Acknowledgements:  The  study  is  the  result  of  specific
research  at  Institute  of  Geological  Sciences  of  Masaryk
University.  The  authors  thank  Moravian  Oil  Mines  Inc.
which  provided  the  primary  data.  Proof-Reading-Service
kindly  improved  the  written  English  of  the  manuscript.
We also give thanks to three anonymous reviewers and edi-
tor  Michal  Kováč  for  their  helpful  comments  and  sugges-
tions.

References

Adámek J., Brzobohatý R., Pálenský P. & Šikula J. 2003: The Kar-

patian  in  the  Carpathian  Foredeep  (Moravia).  In:  Brzobohatý
R.,  Cicha  I.,  Kováč  M.  &  Rögl  F.  (Eds.):  The  Karpatian,
a  Lower  Miocene  Stage  of  the  Central  Paratethys.  Masaryk
University, Brno 75—92.

Adegoke A.K., Abdullah W.H., Hakimi M.H. & Yandoka B.M.S.

2014: Geochemical characterisation of Fika Formation in the
Chad (Bornu) Basin, northeastern Nigeria: Implications for
depositional environment and tectonic setting. Appl. Geochem.
43, 1—12.

Ali S., Stattegger K., Garbe-Schönberg D., Frank M., Kraft S. &

Kuhnt W. 2014: The provenance of Cretaceous to Quaternary
sediments in the Tarfaya basin, SW Moroco: Evidence from
trace element geochemistry and radiogenic Nd-Sr isotopes.
J. Afr. Earth Sci. 90, 64—76.

Andersson P.O.D., Worden R.H., Hodgson D.M. & Flint S. 2004:

Provenance evolution and chemostratigraphy of a Palaeozoic
submarine fan-complex: Tanqua Karoo Basin, South Africa.
Mar. Petrol. Geol. 21, 555—577.

Bahlburg H. 1998: The geochemistry and provenance of Ordovician

turbidites in the Argentinian Puna. In: Pankhurst R.J. & Rapela
C.W. (Eds.): The Proto-Andean Margin of Gondwana. Geol.
Soc. London, Spec. Publ. 142, 127—142.

Beaumont C., Quinlan G.M. & Stockmal G.S. 1993: The evolution

of  the  Western  Interior  Basin:  causes,  consequences  and  un-
solved  problems.  In:  Caldwell  W.G.E.  &  Kauffman  E.G.
(Eds.):  Evolution  of  the  Western  Interior  Basin.  Geol.  Assoc.
Can. Spec. Paper 39, 97—117.

Benada S. 1986: New findings on the extension of the paraauthoch-

tonous deposits of Karpatian in the central part of the Car-
pathian Neogene Foredeep in Moravia. Zemní Plyn Nafta 31, 4,
485—492 (in Czech).

Benada S. & Kokolusová A. 1987: The new knowledge about the

geological positions of the coarse-grained clastic deposits of
the Karpatian in the Central part of the Carpathian Foredeep.
Zemní Plyn Nafta 32, 1, 1—15 (in Czech).

Bergman K.M. & Walker R.G. 1988: Formation of Cardium Ero-

sion  surfaces  E5,  and  associated  deposition  of  conglomerate:
Carrot  Creek  field,  Cretaceous  Western  Interior  Seaway,  Al-
berta.  In:  James  D.P.&  Leckie  D.A.  (Eds.):  Sequences,  Sedi-
mentology, Surface and Subsurface. Canad. Soc. Petrol. Geol.,
Memoir 15, 15—24.

Bhatia M.R. & Crook A.W. 1986: Trace element characteristics of

graywackes and tectonic setting discrimination of sedimentary
basins. Contr. Mineral. Petrology 92, 181—193.

Bieg U., Nebelsick J.H. & Rasser M. 2007: North Alpine Foreland

Basin (Upper Marine Molasse) of Southwest Germany: Sedi-
mentology, Stratigraphy and Palaeontology. Geol. Alp 4,
149—158.

Bieg U., Süss M.P. & Kuhlemann J. 2008: Simulation of tidal flow

and circulation patterns in the Early Miocene (Upper Marine
Molasse) of the Alpine foreland basin. Spec. Publ. Int. Ass.
Sediment. 40, 145—169.

Borges J. & Huh Y. 2007: Petrography and chemistry of the bed

sediments of the Red River in China and Vietnam: Provenance
and chemical weathering. Sed. Geol. 194, 155—168.

Brzobohatý R. & Cicha I. 1993: The Carpathian Foredeep. In:

Přichystal A. et al. (Eds.): The Geology of the Moravia and
Silesia. MZM a PřF MU, Brno, 123—128 (in Czech).

Bubík M. & Dvořák J. 1996: About the finding of the Karpatian

(Miocene) and other results of the Slatinky MH-10 borehole.
Zpr. geol. výzk. v r. 1995 20—21 (in Czech).

Cattaneo A. & Steel R.J. 2003. Transgressive deposits: a review of

their variability. Earth-Sci. Rev. 62, 187—228.

Catuneanu O., WilliS A.J. & Miall A.D. 1998: Temporal signifi-

cance of sequence boundaries. Sed. Geol. 121, 157—178.

Clifton H.E. 1976: Wave-formed sedimentary structures - a concep-

tual model. In: Davis R.A. & Ethington R.L. (Eds.): Beach and
Nearshore processes. SEPM Spec. Publ. 24, 126—148.

Cogan J., Lerche I., Dorman J.T. & Kanes W. 1993: Flexural plate

inversion: application to the Carpathian Foredeep, Czechoslo-
vakia. Mod. Geol. 17, 355-392.

Corcoran P.L. 2005: Recycling and chemical weathering in tectoni-

cally controlled Mesozoic-Cenozoic basins of New Zealand.
Sedimentology 52, 757—774.

Cox R. & Lowe D.R. 1995: A conceptual review of regional-scale

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

controls on the composition of clastic sediment and the co-evo-
lution of continental blocks and their sediment cover. J. Sed.
Res. A65, 1, 1—12.

Crampton S.L. & Allen P.A. 1995: Recognition of forebulge uncon-

formities associated with early stag foreland basin develop-
ment: example from the North Alpine foreland basin. AAPG
Bulletin 79, 1495—1514.

Čopjaková R. 2007: The reflection of provenance changes in the

psefitic and psamitic sedimentary fraction of the Myslejovice
Formation (heavy mineral analysis). PhD thesis, Masaryk
University, Brno, 1—137 (in Czech).

Čopjaková R., Sulovský P. & Otava J. 2002: Comparison of the

chemistry  of  detritic  pyrope-almandine  garnets  of  the  Luleč
Conglomerates  with  the  chemistry  of  granulite  garnets  from
the  Czech  Massif.  Geol.  výzk.  Mor.  Slez.  v r.  2001  44—47
(in Czech).

Čopjaková R., Sulovský P. & Paterson B.A. 2005: Major and trace

elements in pyrope—almandine garnets as sediment provenance
indicators of the Lower Carboniferous Culm sediments,
Drahany Uplands, Bohemian Massif. Lithos 82, 51—70.

Dalrymple R.W. & Choi K. 2007: Morphologic and facies trends

through the fluvial—marine transition in tide-dominated deposi-
tional systems: A schematic Framework for environmental and
sequence-stratigraphic interpretation. Earth Sci. Rev. 81,
135—174.

Dalrymple R.W., Zaitlin B.A. & Boyd R. 1992: Estuarine facies

models: conceptual basis and stratigraphic implications. J. Sed.
Petrology 62, 1130—1146.

DeCelles P.G. & Giles K.A. 1996: Foreland basin systems. Basin

Res. 8, 105—123.

Dickinson W.R. 1985: Interpreting provenance relations from detri-

tal modes of sandstones. In: Zuffa G.G. (Ed.): Provenance of
Arenites. D. Reidel Publication Co., 333—361.

Dickinson W.R., Beard L.S., Brakenridge G.R., Erjavec J.L.,

Ferguson R.C., Inman K.F., Knepp R.A., Lindberg F.A. &
Ryberg P.T. 1983: Provenance of North American Phanerozoic
sandstones in relation to tectonic setting. Geol. Soc. Amer.
Bull. 94, 222-235.

Dickinson W.R. & Suczek C.A 1979: Plate tectonics and sandstone

composition. Amer. Assoc. Petrol. Geol. Bull. 63, 2164—2182.

Dott R.H. & Bourgeois J. 1982: Hummocky stratification: Signifi-

cance of its variable bedding sequences. Geol. Soc. Amer. Bull.
93, 663—680.

Drake D.E. & Cacchione D.A. 1985: Seasonal variation in sediment

transport on the Russian River shelf, California. Cont. Shelf
Res. 4, 495—514.

Duke W.L. 1985: Hummocky cross-stratification, tropical hurri-

canes, and intense winter storms (palaeogeography). Sedimen-
tology 32,167—194.

Duke W.L., Arnott R.W.C. & Cheel R.J. 1991: Shelf sandstones

and hummocky cross-stratification: new insights on a stormy
debate. Geology 19, 625—628.

Eliáš M. & Pálenský P. 1998: The model for the development of the

Miocene foredeeps in the Ostrava area. Zpr. geol. výzk. v
r. 1997, 65—66 (in Czech).

Fedo C.M., Nesbitt H.W. & Young G.M. 1995: Unraveling the

effects of potassium metasomatism in sedimentary rocks and
paleosols, with implications for paleoweathering conditions
and provenance. Geology 23, 921—924.

FitzGerald D., Buynevich I. & Hein C. 2012: Morphodynamics and

facies architecture of tidal inlets and tidal deltas. In: Davis J.A.
Jr. & Dalrymple R.W. (Eds.): Priciples of Tidal Sedimento-
logy. Springer, New York, 301—333.

Flemings P.B. & Jordan T.E. 1989: A synthetic stratigraphic model

of foreland basin development. J. Geophys. Res. 94, B4,
3851—3866.

Flemings P.B. & Jordan T.E. 1990: Stratigraphic modelling of fore-

land basins: Interpreting thrust deformation and lithosphere
rheology. Geology 18, 430—434.

Fralick P.W. & Kronberg B.I. 1997: Geochemical discrimination of

clastic sedimentary rock sources. Sed. Geol. 113, 111—124.

Galloway W.E. & Hobday D.K. 1996: Terrigenous Clastic Deposi-

tional Systems: Applications to Fossil Fuel and Groundwater
Resources, 2nd edition. New York, Springer, 1—489.

Gilíková H., Otava J. & Stráník Z. 2002: Petrological characteris-

tics of sediments of the Magura Flysch at the map-sheet
Holešov. Geol. výzk. Mor. Slez. v r. 2001, 26—29 (in Czech).

Gu X.X, Liu J.M., Zheng M.H., Tang J.X. & Qi L. 2002: Prove-

nance and tectonic setting of the Proterozoic turbidites in
Hunan, South China: geochemical evidence. J. Sed. Res. 72,
393—407.

Gupta S. 1999: Controls on sedimentation in distal margin pallaeo-

valleys in the Early Tertiary Alpine foreland basin, south-
eastern France. Sedimentology 46, 357—384.

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

chronostratigraphy and cycles of sea-level change. In: Wilgius
C.K. et al. (Eds.): Sea-Level Changes. SEMP Spec. Publ. 42,
71—108.

Helland-Hansen W. & Gjelberg J.G. 1994: Conceptual basis and

variability in sequence stratigraphy: a different perspective.
Sed. Geol. 92, 31—52.

Heller P.L., Angevinw C.L., Winslow N.S. & Paola C. 1988: Two-

phases stratigraphic model of forelad-basin sequences. Geology
16, 501—504.

Hiscott R.N. 1984: Ophiolitic Source Rocks for Taconic-Age

Flysch: Trace-Element Evidence. Geol. Soc. Amer. Bull. 95,
1261—1267.

Hladilová S., Nehyba S., Dolákova N. & Hladíková J. 1999: Com-

parison of some relics of Miocene sediments on the eastern
margin of the Bohemian Massif. Geol. Carpathica 50, Spec.
Iss., 31—33.

Hladilová Š., Nehyba S., Zágoršek K., Tomanová Petrová P., Bitner

M.A. & Demeny A. 2014: Early Badenian transgression on the
outer flank of Western Carpathian Foredeep, Hluchov area,
Czech Republic. Ann. Soc. Geol. Pol. 84, 259—279.

Hohenegger J., Ćorić S. & Wagreich M. 2014: Timing of the Mid-

dle Miocene Badenian Stage of the Central Paratethys. Geol.
Carpathica 65, 1, 55—66.

Hohenegger J., Rögl F., Ćorić S., Pervesler P., Lirer F., Roetzel R.,

Scholger R. & Stingl K. 2009: The Styrian Basin: A key to the
Middle Miocene (Badenian/Langhian) Central Paratethys
transgressions. Austrian J. Earth Sci. 102, 103—132.

Holcová K., Brzobohatý R., Kopecká J. & Nehyba S. 2015: Recon-

struction of the unusual Middle Miocene (Badenian) palaeoen-
vironment of the Carpathian Foredeep (Lomnice/Tišnov
denudational relict, Czech Republic) Geol. Quarterly 59,
654—678.

Hubert J.F. 1962: A zircon—tourmaline—rutile maturity index and

the interdependance of the composition of heavy mineral
assemblages with the gross composition and texture of sand-
stones. J. Sed. Petrol. 32, 440—450.

Ingersoll R.V. 1990: Actualistic sandstones petrofacies: discrimi-

nating modern and ancient source rocks. Geology 18, 733—736.

Johnson D.D. & Beaumont C. 1995: Preliminary results from

a planform kinematic model of orogen evolution, surface pro-
cesses and the development of clastic foreland basin strati-
graphy. In: Dorobek S.L. & Ross G.M. (Eds.): Stratigraphic
Evolution of Foreland Basins. Spec. Publ. Soc. Econ. Paleont.
Mineral. 52, 3—24.

Jordan T.E. & Flemings P.B. 1991: Large-scale stratigraphic archi-

tecture, eustatic variation, and unsteady tectonism: a theoreti-
cal evaluation. J. Geophys. Res. 96, B4, 6681—6699.

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Kalvoda J., Bábek O., Fatka O., Leichmann J., Melichar R., Nehyba

S. & Špaček P. 2008: Brunovistulian terrane (Bohemian Mas-
sif, Central Europe) from late Proterozoic to late Paleozoic:
a review. Int. J. Earth Sci. 97, 3, 497—517.

Kalvoda J., Leichmann J., Bábek O. & Melichar R. 2003: Brunovis-

tulian  terrane  (Central  Europe)  and  Istanbul  zone  (NW  Tur-
key):  Late  proterozoic  and  paleozoic  tectonostratigraphic
development  and  paleogeography.  Geol.  Carpathica  54,  3,
139—152.

Kempf O. & Pfiffner A. 2004: Early Tertiary evolution of the North

Alpine Foreland Basin of the Swiss Alps and adjoining areas.
Basin Res. 16, 549—567.

Koss J.E., Ethridge F.G. & Schumm S.A. 1994: An experimental

study of the effect of base-level change on fluvial, coastal and
shelf systems. J. Sed. Res. B64, 90—98.

Kováč M. 2000: Geodynamic, paleogeographical and structural de-

velopment of the Miocene Carpatho-Pannonian region. New
view on the Slovak Neogene basins. Veda, Bratislava, 1—176
(in Slovak).

Kováč M., Sergeyevna A.-G., Brzobohatý R., Fodor L., Harzhauser

M., Oszczypko N., Pavelić D., Rögl F., Saftić B., Sliva  . &
Stráník  Z.  2003:  Karpatian  paleogeography,  tectonics  and
eustatic  changes.  In:  Brzobohatý  R.  et  al.  (Eds.):  The  Karpa-
tian,  A  Lower  Miocene  Stage  of  the  Central  Paratethys.
Masaryk University, Brno, 49—72.

Kováč M., Grygorovich A.A., Bajraktarević Z., Brzobohatý R.,

Filipescu S., Fodor L., Harzhauser M., Oszczypko N., Pavelić D.,
Rögl  F.,  Saftić  B.,  Sliva  .,  Kvaček  Z.,  Hudáčková  N.  &
Slamková  M.  2004:  Paleogeography  of  Central  Paratethys
during  the  Karpatian  and  Badenian.  Scripta  Fac.  Sci.  Nat.
Univ. Masaryk Brunensis, Geology 31—32, 7—17.

Krzywiec P. 2001: Contrasting tectonic and sedimentary history of

the central and eastern parts of the Polish Carpathian foredeep
basin – results of seismic data interpretation. Mar. Petrol.
Geol. 18, 13—38.

Kukal Z. 1986: Manual of practical sediment nomenclature and

classification. Czech Geol. Surv., Praha, 1—80 (in Czech).

Leszczyński S. & Nemec W. 2014: Dynamic stratigraphy of com-

posite peripheral unconformity in a foredeep basin. Sedimen-
tology 62, 3, 645—680.

Li C. & Yang S. 2010: Is chemical index of alteration (CIA) a reliable

proxy for chemical weathering in global drainage basins?
Amer. J. Sci. 310, 111—127.

Lis P. & Wysocka A. 2012: Middle Miocene deposits in Carpathian

Foredeep: facies analysis and implications for hydrocarbon
reservoir prospecting. Ann. Soc. Geol. Pol. 82, 239—253.

McLennan S.M. 1993: Weathering and global denudation. J. Geology

101, 295—303.

McLennan S.M., Heming S.R., McDaniel D.K. & Hanson G.N.

1993: Geochemical approaches to sedimentation, provenance,
and tectonics. In: Johnsson M.J. & Basu A. (Eds.): Processes
controlling the composition of clastic sediments. Geol. Soc.
Amer., Spec. Pap. 284, 1—19.

Meinhold G., Anders B., Kostopoulos D. & Reischmann T. 2008:

Rutile chemistry and thermometry as provenance indicator:
An example from Chios Island, Greece. Sed. Geol. 203,
98—111.

Meinhold G. 2010: Rutile and its applications in earth sciences.

Earth Sci Rev. 102, 1—28.

Miall A.D. 1989: Architectural elements and bounding surfaces in

channelized clastics deposits: notes on comparisons between
fluvial and turbidite systems. In: Taira A. & Masuda F. (Eds.):
Sedimentary Facies in the Active Plate Margin. Terra Sci.
Publ. Company 3—15.

Milliken K.L. 1988: Loss of provenance information through sub-

surface diagenesis in Plio-Pleistocene sediments, northern

Gulf of Mexico. J. Sed. Petrology 58, 992—1002.

Morton A.C. 1984: Stability of detrital heavy minerals in Tertiary

sandstones from the North Sea Basin. Clay Miner. 19,
287—308.

Morton A.C. 1991: Geochemical studies of detrital heavy minerals

and their application to provenance studies. In: Morton A.C. et
al. (Eds.): Developments in Sedimentary Provenance Studies.
Geol. Soc. London, Spec. Publ. 57, 31—45.

Morton A.C. & Hallsworth C.R. 1994: Identifying provenance-spe-

cific features of detrital heavy mineral assemblages in sand-
stones. Sed. Geol. 90, 241—256.

Morton A.C. & Hallsworth C.R. 1999: Processes controlling the

composition of heavy mineral assemblages in sandstones. Sed.
Geol. 124, 3—29.

Morton A.C., Hallsworth C.R. & Chalton B. 2004: Garnet composi-

tions in Scottish and Norwegian basement terrains: a frame-
work for interpretation of North Sea sandstone provenance.
Mar. Petrol. Geol. 21, 393—410.

Nehyba S. 2000: The cyclicity of Lower Miocene deposits of the

SW part of the Carpathian Foredeep as the depositional
response to sediment supply and sea-level changes. Geol.
Carpathica 51, 1, 7—17.

Nehyba S. & Buriánek D. 2004: Chemistry of garnet and tourmaline

– contribution to provenance studies of fine grained Neogene
deposits of the Carpathian Foredeep. Acta Mus. Moraviae, Sci.
Geol. 89, 149—159 (in Czech).

Nehyba S. & Petrová P. 2000: Karpatian sandy deposits in the

southern part of the Carpathian Foredeep in Moravia. Věstník
Čes. Geol. Úst. 75, 1, 53—66.

Nehyba S. & Roetzel R. 2010: Fluvial deposits of the St. Marein-

Freischling Formation – insights into initial depositional pro-
cesses on the distal external margin of the Alpine-Carpathian
Foredeep in Lower Austria. Austrian J. Earth Sci. 103, 2,
50—80.

Nehyba S. & Šikula J. 2007: Depositional architecture, sequence

stratigraphy and geodynamic development of the Carpathian
Foredeep (Czech Republic). Geol. Carpathica 58, 1, 53—69.

Nehyba S., Roetzel R. & Maštera L. 2012: Provenance analysis of

the Permo—Carboniferous fluvial sandstones of the southern part
of the Boskovice Basin and the Zöbing Area (Czech Republic,
Austria): implications for paleogeographical reconstructions of
the post-Variscan collapse basins. Geol. Carpathica 63, 5,
365—382.

Nemec W. 2005: Principles of lithostratigraphic logging and facies

analyses. Institut for geovitenskap, Univ. Bergen, 1—28.

Nesbitt H.W. 1979: Mobility and fractionation of rare earth ele-

ments during weathering of granodiorite. Nature 279,
206—210.

Nesbitt H.W. & Young G.M. 1982: Early Proterozoic climates and

plate motions inferred from major element chemistry of lutites.
Nature 299, 715—717.

Nesbitt H.W. & Young G.M. 1989: Formation and diagenesis of

weathering profiles. J. Geology 97, 129—147.

Nesbitt H.W. & Young G.M. 1998: Processes controlling the distri-

bution of Ti and Al in weathering profiles, siliciclastic sedi-
ments and sedimentary rocks. J. Sed. Res. 68, 3, 448—455.

Nesbitt H.W., Young G.M., McLennan S.M. & Keays R.R. 1996:

Effects of chemical weathering and sorting on the petrogenesis
of siliciclastic sediments, with implications for provenance
studies. J. Geology 104, 525—542.

Nittrouer C.A., DeMaster D.J., Kuehl S.A. & McKee B.A. 1986:

Association of sand with mud deposits accumulating on continen-
tal shelves. In: Knight R.J. & McLean J.R. (Eds.): Shelf Sands
and Sandstones. Canad. Soc. Petrol. Geol., Memoir 11, 17—25.

Oszczypko N., Krzywiec P., Popadyuk I. & Peryt T. 2006: Car-

pathian Foredeep Basin (Polish and Ukraine): Its sedimentary,

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

structural, and geodynamic evolution. In: Golonka J. & Picha
F.J. (Eds.): The Carpathians and Their Foreland: Geology and
Hydrocarbon Resources. AAPG Memoir 84, 293—350.

Oszczypko N. & Oszczypko-Clowes M. 2012: Stages of develop-

ment in the Polish Carpathian Foredeep Basin. Central Eur. J.
Geosci. 4, 1, 138—162.

Otava J. 1998: Trends of changes in the composition of siliciclastics

of the Lower Carboniferous at Drahany Upland (Moravia) and
their geotectonic interpretation. Geol. výzk. Mor. Slez. v r.
1997, 62—64 (in Czech).

Otava J., Krejčí O. & Sulovský P. 1997: First results of electron

microprobe analyses of detrital garnets from the Rača Unit of
the Magura Group. Geol. výzk. Mor. Slez. v r. 1996, 39—42
(in Czech).

Otava J., Sulovský P. & Kejčí O. 1998: The results of study of the

detrital garnets from the Cretaceous sediments of the Rača
Unit of the Magura Group (Outer Carpathians). Geol. výzk.
Mor. Slez. v r. 1997, 29—31 (in Czech).

Otava J., Sulovský P. & Čopjaková O. 2000: Provenance changes of

the Drahany Culm greywackes: statistical evaluation. Geol.
výzk. Mor. Slez. v r. 1999, 94—98 (in Czech).

Passchier S. & Whitehead J.M. 2006: Anomalous geochemical

provenance and weathering history of Plio—Pleistocene glacio-
marine fjord strata, Bardin Bluffs Formation, East Antarctica.
Sedimentology 53, 929—942.

Petránek J. 1963: Sedimentary rocks. Czech Academy of Sciences,

Praha, 1—717 (in Czech).

Plint A.G. 2000: Sequence stratigraphy and paleogeography of

a Cenomanian deltaic complex: The dunvegan and lower
kaskapau formations in subsurface and outcrop, Alberta and
British Columbia, Canada. Bull. Canad. Petrol. Geol. 48, 43—79.

Plint A.G., Hart B. & Donaldson W.S. 1993: Lithospheric flexure

as a control on stratal geometry and facies distribution in
Upper Cretaceous rocks of the Alberta Foreland Basin. Basin
Res. 5, 69—77.

Plint A.G., McCarthy P.J. & Faccini U.F. 2001: Nonmarine

sequence  stratigraphy:  Updid  expression  of  sequence  boun-
daries  and  systems  tracts  in  a  high-resolution  framework,
Cenomanian  Dunvegan  Formation,  Alberta  foreland  basin,
Canada. AAPG Bulletin 85, 1967—2001.

Posamentier H.W., Allen G.P., James D.P. & Tesson M. 1992:

Forced regressions in a sequence stratigraphic framework:
concept, examples, and exploration significance. AAPG Bulle-
tin 76, 1687—1709.

Reading H.G. 1996: Sedimentary environments: processes, facies

and stratigraphy. Blackwell Science, Oxford, 1—704.

Reineck H.E. & Singh I.B. 1973: Depositional sedimentary environ-

ments: with reference to terrigenous clastics. Springer-Verlag,
Berlin, 1—439.

Reynand J.-Y. & Dalrymple R.W. 2012: Shallow-marine tidal

deposits. In: Davis J.A. Jr. & Dalrymple R.W. (Eds.): Priciples
of Tidal Sedimentology. Springer, New York, 335—369.

Rider M.H. 1986: The geological interpretation of well logs. John

Wiley & Sons, New York, 1—175.

Roser B.P. & Korsch R.J 1999: Geochemical characterization,

evolution and source of Mesozoic accretionary wedge:
the Torlesse terrane, New Zealand. Geol. Mag. 136, 493—512.

Roser B.P., Cooper R.A., Nathan S. & Tulloch A.J. 1996: Recon-

naissance sandstone geochemistry, provenance, and tectonic
setting of the lower Paleozoic terranes of the West Coast and
Nelson, New Zealand. New Zeal. J. Geol. Geop. 39, 1—16.

Ross D.J.K. & Bustin R.M. 2009: Investigating the use of sedimen-

tary geochemical proxies for paleoenvironment interpretation
of thermally mature organic-rich strata: Examples from the
Devonian—Mississippian shales, Western Canadian Sedimentary
Basin. Chem. Geol. 260, 1—19.

Sacchi M., Molisso F., Pacifico A., Vigliotti M., Sabbarese C. &

Ruberti  D.  2014:  Late-Holocene  to  recent  evolution  of  Lake
Patria,  South  Italy:  An  example  of  a  coastal  laggon  within
a  Mediterranean  delta  system.  Global  Planet  Change  117,
9—27.

Salata D. 2004: Garnet provenance in mixed first-cycle and poly-

cycle heavy-mineral assamblages of the Ropianka and Menilite
Formations (Skole Nappe, Polish Flysch Carpathians): con-
straints from chemical composition and grain morphology.
Ann. Soc. Geol. Pol. 83, 161—177.

Salata D. 2013a: Detrital garnets from the Upper Cretaceous—

Palaeocene sandstones of the Polish part of the Magura nappe
and the Pieniny Klippen Belt: chemical constraints. Ann. Soc.
Geol. Pol. 74, 351—364.

Salata D. 2013b: Source rocks for heavy minerals in lower part of

Menilite Formation of Skole Nappe (Polish Flysch Car-
pathians), based on study of detrital garnet and tourmaline.
Ann. Soc. Geol. Pol. 83, 1—17.

Sen S., Das P.K., Bhagaboty B & Singha L.J.C. 2012: Geochemistry

of shales of Barail group occurring in and around Mandardisa,
North Catchar Hills, Assam; India: Its implications. Int. J.
Chem. Appl. (IJCA) 4, 1, 25—37.

Shanley K.W. & McCabe P.J. 1994: Perspectives on the sequence

stratigraphy of continental strata. AAPG Bulletin 78, 544—568.

Sinclair H.D., Coakley B.J., Allen P.A. & Watts A.B. 1991: Simu-

lation of foreland basin stratigraphy using a diffusion model of
mountain belt uplift and erosion: an example from the central
Alps, Switzerland. Tectonics 10, 3, 599—620.

Stráník Z., Hrouda F., Otava J., Gilíková H. & Švábenická L. 2007:

The  Upper  Oligocene—Lower  Miocene  Krosno  lithofacies  in
the  Carpathian  Flysch  Belt  (Czech  Republic):  sedimentology,
provenance  and  magnetic  fabrics.  Geol.  Carpathica  58,  4,
321—332.

Suggate S.M. & Hall R. 2013: Using detrital garnet compositions to

determine provenance: a new compositional database and pro-
cedure. In: Scott. R.A. et al. (Eds.): Sediment Provenance
Studies in Hydrocarbon Exploration and Production. Geol.
Soc. London, Spec. Publ. 386, 373—393.

Suttner L.J. & Dutta P.K. 1986: Alluvial sandstone composition and

paleoclimate, I. framework mineralogy. J. Sed. Petrology 56,
3, 329—345.

Swift D.J.P., Philips S. & Thorne J.A. (1991): Sedimentation on

continental  margins:  V.  Parasequences.  In:  Swift  D.J.P.  et  al.
(Eds.):  Shelf  Sand  and  Sandstone  Bodies:  Geometry,  Facies
and  Sequence  Stratigraphy.  Spec.  Publ.  Int.  Assoc.  Sediment.
14, 153—187.

Sztanó O. & de Boer P.L. 1995: Basin dimensions and morphology

as controls an amplification of tidal motions (the Early
Miocene North Hungarian Bay). Sedimentology 42, 665—682.

Šikula J. & Nehyba S. 2006: Interpretation of the Neogene deposits

from the area Vizovice Hills according the subsurface data.
Geol. výzk. Mor. Slez. v r. 2005, 54—57 (in Czech).

Taylor S.R. & McLennan S.M. 1985: The Continental Crust: Its

Composition and Evolution. Blackwell, London, 1—312.

Thonová H., Holzknecht M., Krystek I. & Brzobohatý R. 1987:

New insights into the development of Karpatian under flysch
nappes in the section „Center“. Zemní Plyn Nafta 32, 3,
383—389 (in Czech).

Tomkins H.S., Powell R. & Ellis D.J. 2007: The pressure depen-

dence of the zirkonium-in-rutile thermometer. J. Metamorph.
Geology 25, 703—713.

von Eynatten H., Barceló-Vidal C. & Pawlowsky-Glahn V. 2003:

Modelling compositional changes: the example of chemical
weathering of granitoid rocks. Math. Geology 35, 3, 231—251.

Walker R.G. & James N.P. 1992: Facies models: Response to sea

level change. Geol. Assoc. Canada, 1—380.

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67, 1, 41—68

Walker R.G. & Plint A.G. 1992: Wave- and storm-dominated shal-

low marine systems. In: Walker R.G. & James N.P. (Eds.):
Facies Models: Response to Sea Level Change. Geol. Assoc.
Canada, 219—238.

Waśkowska A., Cieszkowski M., Golonka J. & Kowal-Kasprzyk J.

2014: Paleocene sedimentary record of ridge geodynamics in
Outer Carpathian basin (Subsilesian Unit). Geol. Carpathica
65, 1, 35—54.

Watson E.B., Wark D.A. & Thomas J.B. 2006: Crystallization ther-

mometers for zircon and rutile. Contrib. Mineral. Petrol. 151,
413—433.

Wright L.D., Chappel J., Bradshaw M.P. & Cowell P. 1979: Mor-

phodynamics of reflective and dissipative beach and nearshore
systems, southeastern Australia. Mar. Geol. 32, 105—140.

Yang Y. & Miall A.D. 2010: Migration and stratigraphic fill of an

underfilled foreland basin: Middle—Late Cenomanian Belle
Fourche Formation in southern Alberta, Canada. Sed. Geol.
227, 1—4, 51—64.

Zack T., Moraes R. & Kronz A. 2004a: Temperature dependence of

Zr in rutile: empirical calibration of rutile thermometer.
Contrib. Mineral. Petrol. 148, 471—488.

Zack T., von Eynatten H. & Kronz A. 2004b: Rutile geochemistry

and its potential use in quantitative provenance studies. Sed.
Geol. 171, 37—58.

Zádrapa M. 1979: Heavy minerals in the Karpatian deposits in the

Central and Southwestern parts of the Carpathian Foredeep in

the Moravia. Zemní Plyn Nafta 24, 3, 447—451 (in Czech).

Zágoršek K., Nehyba S., Tomanová Petrová P., Hladilová Š.,

Bitner M.A., Doláková N., Hrabovský J. & Jašková V. 2012.
Local catastrophe near Přemyslovice (Moravia, Czech Republic)
during Middle Miocene due to the tephra input. Geol. Quarterly
56, 269—284.

Zaitlin B.A., Dalrymple R.W. & Boyd R. 1994: The stratigraphic

organization of incised-valley systems associated with relative
sea-level change. In: Dalrymple R.W. et al. (Eds.): Icised-
Valley Systems: Origin and Sedimentary Sequences. SEPM
Spec. Publ. 51, 45—60.

Zhang K.-J. 2004: Secular geochemical varitions of the Lower

Cretaceous siliciclastic rocks from central Tibet (China) indi-
cate a tectonic transition from continental collision to back-arc
rifting. Earth Planet. Sci. Lett. 229, 73—89.

Zimmermann U. & Bahlburg H. 2003: Provenance analysis and tec-

tonic setting of the Ordovician clastic deposits in the southern
Puna Basin, NW Argentina. Sedimentology 50, 1079—1104.

Zuffa G.G. 1980: Hydrid arenites: Their composition and classifica-

tion. J. Sed. Petrology 50, 21—29.

Zuffa G.G. 1985: Optical analyse of arenites: influence of

methodology on compositional results. In: Zuffa G.G. (Ed.):
Provenance of Arenites. D. Reidel Publication Co., 165—189.

Zuschin M., Harzhauser M., Hengst B., Mandic O. & Roetzel R.

2014: Long-term ecosystem stability in an Early Miocene
estuary. Geology 42, 1—4.

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67,1

Electronic supplement

Appendix 1: The major element composition (wt. %) of the studied samples.

Sample
(borehole)

SiO

MgO

CaO

TiO

MnO

LOI

BAR1

53.05

11.73

4.56

3.45

8.12

1.17

2.61

0.57

0.13

0.08

0.014

14.3

BAR1

51.77

12.25 4.81 3.60 8.27 1.03 2.52 0.56 0.11 0.13 0.013 14.7

BAR1

54.16

15.70 5.11 2.06 4.08 0.91 2.80 0.74 0.13 0.03 0.017 14.1

BAR1

53.27

11.97 4.62 3.50 8.15 1.14 2.43 0.57 0.14 0.08 0.014 13.9

70.88 5.13 2.15 2.53 6.96 0.90 1.11 0.27 0.09 0.05

0.004 9.9

52.61

16.22 5.68 3.55 5.03 1.31 3.44 0.80 0.15 0.06 0.020 10.9

68.16

6.57

2.71

2.29

8.23

1.25

1.16

0.50

0.09

0.08

0.007

8.9

65.85 5.62 2.78 2.82 8.87 1.25 0.86 0.50 0.09 0.09

0.007 11.2

54.44

14.90 4.98 3.29 5.88 1.19 3.17 0.73 0.14 0.06 0.016 11.0

74.03 5.51 2.01 2.20 6.00 1.18 1.04 0.38 0.09 0.06

0.006 7.4

49.97

14.47 5.17 3.10 7.82 0.89 2.99 0.57 0.16 0.04 0.017 14.6

45.08

14.37

4.84

2.50

10.81

0.89

2.93

0.60

0.21

0.04

0.016

17.5

50.88

17.57 5.54 3.74 4.91 1.28 3.84 0.80 0.14 0.06 0.019 11.0

52.71

16.05 5.55 3.36 4.64 0.98 3.36 0.61 0.18 0.04 0.018 12.2

48.87

16.15 6.06 3.58 6.62 0.92 3.49 0.70 0.12 0.06 0.018 13.2

44.27

13.97 4.65 2.44

11.88 0.75 2.79 0.59 0.18 0.04 0.015 18.2

49.39

19.39 5.67 4.09 3.21 1.06 4.85 0.85 0.18 0.04 0.022 11.0

48.90

15.46

5.70

3.20

6.35

0.90

3.27

0.64

0.24

0.04

0.018

15.0

49.44

16.56 5.99 3.68 5.55 1.07 3.82 0.77 0.12 0.06 0.020 12.7

52.73

14.99 5.62 3.43 6.19 1.32 3.30 0.74 0.15 0.06 0.017 11.3

HOL1

48.06

15.05 5.19 2.41 8.22 0.88 2.82 0.64 0.15 0.04 0.018 16.3

HOL1

50.26

14.02 5.32 3.75 6.96 1.02 2.84 0.62 0.12 0.08 0.016 14.8

HOL1

60.02 9.41 3.53 3.15 8.39 1.21 1.98 0.52 0.12 0.08

0.010 11.4

HOL1

56.60

11.69 4.52 2.85 7.30 1.07 2.45 0.57 0.11 0.07 0.011 12.6

HOL1

56.61

12.10 4.36 2.62 7.20 1.00 2.51 0.57 0.09 0.06 0.012 12.7

HOL1

58.04

11.72 4.22 2.67 6.88 1.06 2.49 0.57 0.11 0.06 0.009 12.0

HOL1

57.22

14.33 5.32 2.95 3.59 1.08 2.76 065 0.10 0.06 0.013 11.8

HOL1

59.98 7.16 3.09 2.72

11.05 1.14 1.59 0.41 0.12 0.14

0.007 12.3

HOL1

63.32

9.32

3.34

2.20

6.89

1.07

2.26

0.46

0.12

0.06

0.006

10.8

HOL1

56.30

15.12 5.67 3.00 2.50 1.09 2.89 0.68 0.11 0.05 0.013 12.4

HOL1

59.84 9.25 3.72 3.08 8.40 1.20 1.93 0.53 0.16 0.11

0.009 11.4

HUL2

57.31

12.13 4.50 3.08 6.66 1.17 2.56 0.59 0.13 0.07 0.014 11.6

HUL2

45.34

14.09 4.92 2.88

10.04 0.84 2.88 0.60 0.12 0.04 0.017 18.0

HUL2

64.07 9.89 3.86 2.27 5.48 1.07 2.37 0.50 0.16 0.06

0.011 10.1

HUL2

63.37

10.17

3.76

2.25

5.98

1.06

2.43

0.46

0.12

0.06

0.009

10.2

HUL3

54.51

14.35 5.17 2.92 4.51 1.03 2.76 0.63 0.13 0.05 0.015 13.8

HUL3

38.95

11.06 3.98 6.53

15.02 0.87 2.28 0.49 0.10 0.05 0.011 20.4

JAR1

51.71

13.18 4.97 3.54 6.86 1.04 2.71 0.58 0.12 0.10 0.015 15.0

JAR1

50.71

13.55 4.98 3.30 7.33 1.05 2.86 0.61 0.12 0.07 0.016 15.2

KRO1

54.44

13.89 5.12 3.11 5.91 1.06 2.77 0.68 0.11 0.06 0.015 12.6

KRO1

65.51

11.99 4.51 1.31 1.06 1.27 2.27 0.64 0.15 0.04 0.011 11.1

KRO1

50.07

12.11 5.03 3.33 7.68 0.88 2.47 0.52 0.14 1.71 0.014 15.9

KRO1

60.32 9.76 3.64 3.11 7.43 1.27 2.04 0.62 0.15 0.07

0.011 11.4

KRO2

50.16

13.31 4.87 3.56 8.31 1.08 2.65 0.60 0.11 0.06 0.016 15.1

KRO2

53.27

13.34 4.96 3.40 6.54 1.02 2.68 0.60 0.11 0.05 0.014 13.8

KRO2

66.89

7.32

2.56

2.73

7.15

1.07

1.68

0.46

0.10

0.05

0.009

9.8

KRO2

60.45

12.57 4.52 2.85 3.97 1.09 2.70 0.58 0.11 0.06 0.014 10.9

SLU1

52.45

13.28 4.80 3.70 7.10 1.03 2.95 0.60 0.14 0.07 0.015 13.7

SLU1

53.34

11.68 4.39 3.13 9.74 1.27 2.26 0.64 0.14 0.07 0.013 13.2

SLU1

61.20 9.59 3.78 2.38 8.85 1.34 1.76 0.50 0.14 0.07

0.009 10.2

SLU1

50.03

15.48 5.83 3.52 7.23 1.25 3.16 0.72 0.14 0.07 0.016 12.3

TL1

59.70

7.54

3.44

3.40

9.58

1.16

1.53

0.58

0.12

0.07

0.010

12.7

TL1

61.80 6.87 3.13 3.21 9.43 1.14 1.47 0.46 0.11 0.06

0.010 12.2

TL1

52.58

12.91 4.78 3.65 7.40 1.13 2.69 0.58 0.14 0.07 0.013 13.9

TL2

52.09

12.71 4.69 3.00 8.04 0.99 2.54 0.57 0.13 0.06 0.015 15.0

TL2

56.52

10.97 4.14 3.12 7.36 1.13 2.26 0.53 0.13 0.07 0.013 13.6

TL2

53.61

12.01

4.60

3.37

8.36

1.12

2.45

0.59

0.11

0.08

0.015

13.5

EARLY MIOCENE PASSIVE MARGIN EVOLUTION OF THE CARPATHIAN FOREDEEP

Electronic supplement

Appendix 2: The trace element composition (ppm) of the studied samples.

Sample
(borehole)

BAR1

319

9.2

107.2

8.5

2.5

109

133.1

27.1

BAR1

59 11 308 9.6

10.6

112.5 9.4 2.3 117

107.2

18.9 26

BAR1

69 14 449

10.4

10.2

136.3 11 5.5

185

144.4

24.3 31

BAR1

49 11 308 9.6 9.3 99.1 9.1 2.6 101

124.5

18.1

26.5

14.7 4 203 4.2 7 40 3.5 1.8 32

68.9

12.6

11.3

52.8 15 465 15.2 16.4 154.3 12.5 3.7 151

187.2 26.1 34.6

15.3 6 210 4.9 8.2 43.5 4.4 1.8 37

123.3

16.9

17.2

12.7 4 281 4.4 8.6 32.1 4.8 1.7 30

148.7

15.9

17.8

46.9 13 479 14.6 14.6 145.6 11.1 3.4 150

164.9 24.6 35.5

10.5 5 194 4.2 7.4 37.5 3.3 1.2 29

124.7

13.3

12.5

58 14 565 13.1 11.8 153.7 11.1 6.4 233

103.2 22.9 33.1

65 13 577 11.6 11.7 152.4 10.8 12.9 226

108.4 23.3 33.9

58 17 474 17.6 16.8 178.7 13.7 3.7 181

164.6 26.8 43.5

671

15.8

13.7

169.2

12.4

6.7

227

117.5

25.8

36.8

56 15 471 18.1 14.8 173.1 13.1 4.4 196

136.1 26.7 40.2

59 13 597 11.7 12.2 148.9 10.3 12.3 217

111.5 24.2 35.3

74 18 488 16.4 16.1 234.1 14.6 4.8 205

151.2 29.7 46.2

102 15 578 21.9 12.5 169.3 11.3 15.2 341

110.2 25.3 36.2

63 16 466 15.9 16.6 182.6 13.8 3.7 171

144.5 27.4 38.9

422

14.8

13.7

144.8

10.9

2.9

134

169.3

22.5

31.9

HOL1

83 14 416 13.8 11.7 134.1 10.2 6.6 244

104.3 19.9 27.6

HOL1

57 13 386 11.7 15.1 147.1 11.3 3.2 146

120.7 22.8 32.1

HOL1

38 7 313 7.2 7.5

178.3 6.1 1.8 69

131.8

18.6

21.3

HOL1

44 10 354 9 8.4

106.3

10.2 2.8

105

172.7

19.7 27

HOL1

48 10 374 9 7.8 109 9.6 2.9

110

160.1

18.3

27.8

HOL1

351

8.6

9.6

101.5

9.4

3.1

182.9

19.5

27.3

HOL1

59 12 368

10.9 9.8

125.1 11 3.2

123

144.4

20.9

29.7

HOL1

25 6

1505

4.8

4.8 55 4

1.4

108

16.7

21.5

HOL1

35 7 362 8.1 6.5 78.3 8.5 2.9 71

191.8

24.1

24.4

HOL1

62 13 355

11.3

11.9

130.9

10.3 3.1

139

120.7 19 28

HOL1

35 7

2044 9

6.3 70

5.5

2.0

134.7

17.9

20.5

HUL2

44 10 438 9.4

11.2 113 9.3 2.9 100 186 21

26.2

HUL2

74 13 422 10.4 11.4 137.9 9.5 6.3 296 93.9 18.4 25.7

HUL2

45 8 458 9.2

10.4 95.1

11.5 3.3 77

262.1

24.5

33.7

HUL2

42 8

458 9

8.3

94.1

9.1

2.7

174.4

18.3

24.9

HUL3

54 13 317 11.8 12.7 136.8 10.1 3.6 138

116.7 18.7 28

HUL3

48 10 287 8.7 8.4

107.1 9.4 3.7 106

109.1

18.8

25.7

JAR1

48 12 324 11.9 10.8 131.5 9.5 2.9 128

115.1 17.9 25.4

JAR1

333

11.2

134.6

9.3

2.8

138

105

19.5

26.7

KRO1

56 12 396 11.4 12.5 136.5 10.7 3.3 138

164.8 23.3 31.6

KRO1

54 10 387

12.7

10.7

104.5 8.8 9.1

114

179.7 20

24.7

KRO1

59 11 339

20.3 9.5

115.1 9.5 2.4

124

100.6 18

26.5

KRO1

34 8 384 9.2 9.8 78.6 9.8 2.9 76

265.3

21.6

28.1

KRO2

58 11 366 12.5 13.2 136.9 10.1

3 127

123.2 22.7 29.1

KRO2

361

10.3

11.2

130.5

10.6

126

132.5

19.1

KRO2

31 7

289

6.5

8.1

66.5 6

1.9

194.1

17.2

19.5

KRO2

65 10 401 12.2 10.4 118.6 10.5 2.9 117

159.5 21.9 28.2

SLU1

36 12 383

11.8

143.2 10 3.2

131

113.4

20.7

31.7

SLU1

37 10 355 12 12.7 104.8

9 2.8 105

174.7 23.6 30.4

SLU1

24 8

312 8

9.6

74.1

6.9

2.0

162.7

24.5

26.1

SLU1

468

15.7

15.6

154.4

12.9

3.8

173

160.3

25.8

37.9

TL1

27 7

312 7

9.9

61.1

5.5

1.9

144.3

16.1

18.4

TL1

24 6 309 6.4 8.1 56.9 4.6 1.5 50

119.7

13.7

15.4

TL1

46 12 318

11.1 11 127 8.8 3

117

111.8

19.6

25.5

TL2

59 12 311 10.9 11.7 126 9.2 3.9 154

106.2 19.6 25.1

TL2

41 10 312

3.12

10.1 105 8 2.2 96

113.9

18.6

23.5

TL2

51 11 314 9.1

10.1

102.6 8.6 2.3 101 124

18.9

24.3

FRANCÍREK and NEHYBA

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2016, 67,1