New morphostructural subdivision of the Western Carpathians: An approach
integrating geodynamics into targeted morphometric analysis
Jozef Minár a,
⁎, Miroslav Bielik b,1
, Michal Kováč c,1
, Dušan Plašienka c,1
, Ivan Barka d,2
,
Miloš Stankoviansky a,1
, Hermann Zeyen e,3
a
Comenius University in Bratislava, Faculty of Natural Sciences, Department of Physical Geography and Geoecology, Bratislava, Slovakia
b
Comenius University in Bratislava, Faculty of Natural Sciences, Department of Applied and Environmental Geophysics, Bratislava, Slovakia
c
Comenius University in Bratislava, Faculty of Natural Sciences, Department of Geology and Paleontology, Bratislava, Slovakia
d
Forest Research Institute, National Forest Centre, Zvolen, Slovakia
e
Département des Sciences de la Terre, Université de Paris-Sud, Orsay, France
a b s t r a c ta r t i c l e i n f o
Article history:
Received 20 February 2009
Received in revised form 21 February 2010
Accepted 1 April 2010
Available online 10 April 2010
Keywords:
Western Carpathians
Morphostructures
Geodynamics
Targeted morphometric analysis
Geomorphic history
Neogene–Quaternary
A new quantitative method of morphostructure delimitation based on targeted morphometric analyses and
multivariate statistical methods was applied to the Western Carpathians. Nine specific morphostructural
regions and sixteen subregions were defined as an improvement on the preceding qualitative subdivision of
the area. The integration of geodynamics into the targeted morphometric analysis represents a prerequisite
for better interpretability of the delimited regions. The new subdivision of the Western Carpathians therefore
reflects first of all the Pliocene–Quaternary geodynamics that controls the development of the present-day
relief. The results also help to understand the timing of the basic dome-like morphostructural formation of
the Western Carpathians (which began 4–6 million years ago, with the main stage continuing until the Late
Pliocene and accelerated uplift taking place since the Middle Pleistocene), as well as the mechanism of its
formation. The importance of the Middle Miocene extension for the development of the basin-and-mountain
mosaic unique to the Western Carpathians is documented. The projection of the older structural boundaries
into the new morphostructural regions and the increasing prevalence of the young morpholineament
systems (N–S and W–E) on the southern and northern periphery of the Western Carpathians could be an
indication of the gradual spreading of the Western Carpathians into the surrounding lowlands during the last
stage of the morphotectonic development.
© 2010 Elsevier B.V. All rights reserved.
1. Introduction
The Western Carpathians are a discrete morphostructural unit of
the Alpine–Carpathian mountain chain. They differ from their
surroundings not only due to strong morphological individuality, but
also by their specific geological development and tectonic pattern.
Apart from the northern limit of the Western Carpathians, the margins
of this morphostructural unit are typically gradational (Fig. 1).
The existence of the first-order Western Carpathians megamorphostructure
was mentioned and described long ago by Mazúr
(1965) as a vague consequence of the ‘dictiogenetic’ (post-orogene)
evolution. Lower-order morphostructures were traditionally perceived
as geological (structural) regions. Morphologically-defined
second-order morphostructural regions were firstly defined by Lacika
and Urbánek (1998), but only for Slovakia and without exact
explanation and interpretation. The most distinctive basic morphostructures—a
mosaic of single mountain ranges and intramountain
basins—used to be interpreted simply as block structures without any
geodynamic explanation (Mazúr, 1979).
The new morphologically- and quantitatively-based approach
presented here comprehensively defines and characterizes the basic
second- and third-order morphostructures of the Western Carpathians
by taking into account geological and geophysical data and
geodynamic concepts, models and theories to build a synthetic
picture. We show that the morphological character of the mountains
can mirror the whole complex of young geological, geodynamic and
geophysical features that do not have to be visibly reflected in the
older geological structures. The suggested subdivision is an improvement
on earlier qualitative, morphologically-based subdivision
(Lacika and Urbánek, 1998; Urbánek and Lacika, 1998). This
Tectonophysics 502 (2011) 158–174
⁎ Corresponding author. Comenius University in Bratislava, Faculty of Natural
Sciences, Department of Physical Geography and Geoecology, Mlynská dolina, SK–
842 15 Bratislava 4, Slovakia. Tel.: +421 (2) 602 96 518; fax: +421 (2) 654 29 064.
E-mail addresses: minar@fns.uniba.sk (J. Minár), bielik@fns.uniba.sk (M. Bielik),
kovacm@fns.uniba.sk (M. Kováč), plasienka@fns.uniba.sk (D. Plašienka),
barka@nlcsk.org (I. Barka), stankoviansky@fns.uniba.sk (M. Stankoviansky),
hermann.zeyen@u-psud.fr (H. Zeyen).
1
Fax: +421 (2) 654 29 064.
2
Fax: +421 (45) 531 41 55.
3
Fax: +33 (1) 691 54 905.
0040-1951/$ – see front matter © 2010 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2010.04.003
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improvement comes from quantitative targeted morphometrical
analysis and cluster analysis that maximizes the internal morphological
homogeneity of the delimited morphostructural regions.
2. Geological setting
2.1. Geodynamic development and structural pattern of Western
Carpathians—an overview
The Western Carpathians form the northernmost, generally E–W
trending oroclinal segment of the European Alpides and are linked to
the Eastern Alps in the west and to the Eastern Carpathians in the east
(Fig. 2). The northern Carpathian foreland is formed by the North
European Platform that includes basement consolidated during the
Palaeozoic and platform cover of the Bohemian Massif and Polish
Platform in the NW and N. In the NE, the foreland is separated from
the Russian Platform by the Teisseyre–Tornquist Line. A large part of
the Central and most of the Internal Western Carpathians are covered
by remnants of Palaeogene sedimentary basins and thick Neogene
sedimentary and volcanic rock complexes that are related to the
hinterland Pannonian Basin System. The present structural pattern of
the Western Carpathians originated from the Late Jurassic–Cenozoic
subduction–collision orogenic processes in a mobile belt between the
stable North European Plate and drifting, Africa-related Apulian
continental fragments. One of the most characteristic features of
Western Carpathian evolution is a marked northward migration of
pre-orogenic and orogenic processes, including Mesozoic rifting and
extension of the epi-Variscan basement, compression and nappe
stacking of attenuated continental crust and subduction of longitudinal
oceanic basins (e.g., Maheľ, 1981; Birkenmajer, 1988; Plašienka
et al., 1997a, b; Froitzheim et al., 2008), as well as transpression and
transtension postdating the main shortening phases. The Western
Carpathian orogeny ended during the Late Cenozoic after slab
detachment terminated the southward subduction of the ocean–
crust substratum of the External Carpathian Flysch Belt (Tomek and
Hall, 1993; Bielik et al., 2004).
The Western Carpathians, like most of other Alpine collisional fold
belts, have been traditionally divided into outer (northern, dominated
by Tertiary deformation) and inner (southern, dominated by
Mesozoic deformation) structural zones. This subdivision partly
corresponds to the geographic–geomorphological subdivision of the
Western Carpathians into the Outer and Inner Western Carpathians.
Alternatively, a concept of triple division into the External, Central and
Internal Western Carpathians has been proposed (Maheľ, 1986;
Plašienka, 1999). The latter division emphasizes the presence of two
diachronously closed oceanic sutures. The Early Alpine (Cimmerian–
Late Jurassic) Meliata suture delimits the Internal and Central
Western Carpathians. However, being largely obliterated by Tertiary
cover, the precise position of this suture is not always clear. The
Central and External Western Carpathians are divided by the Pieniny
Klippen Belt, which is assumed to be one of the surface expressions of
the Late Cretaceous to Early Tertiary closure of a supposed Penninicrelated
oceanic domain along the northern Central Carpathian edge.
The External Western Carpathians define the northerly convex,
arcuate shape of the West Carpathian orocline. The External Western
Carpathians comprise molassic sediments of the Late Cenozoic
Carpathian Foredeep deposited on the margins of the North European
Plate in Moravia and southern Poland, and the broad Flysch Belt
composed of numerous thrust units. The Flysch Belt that forms the
Cenozoic accretionary wedge of the Carpathian orogen is generally
divided into two groups: (1) the Silesian-Krosno nappes in the north
and (2) the Magura thrust system in the south. Layers of mostly
siliciclastic Flysch sediments, which reach several thousand metres in
thickness, are of Jurassic to Lower Miocene age and were detached
from strongly-attenuated continental and, in some parts, probably
also from truly oceanic crust (prolongation of the North Penninic).
This oceanic crust was produced by Jurassic–Tertiary rifting and sea
floor spreading. The basement substratum of the Flysch Belt was
shortened and underthrust below the Central Western Carpathians.
The boundary between the External Western Carpathian Flysch
Belt and the Central Western Carpathian basement-cover thrust units
is formed by a steep, narrow zone with an intricate structure, the
Pieniny Klippen Belt. This belt is formed exclusively of post-Triassic
sedimentary formations sheared off their unknown basement
substratum, which is interpreted to be a continental ribbon rifted
off the North European Platform during the Jurassic and overridden by
the Central Carpathian nappes during the Late Cretaceous. Geometrically,
the basement to the Pieniny Klippen Belt corresponds to the
Fig. 1. Position and gradational nature of boundaries in the Western Carpathians.
159J. Minár et al. / Tectonophysics 502 (2011) 158–174
Briançonnais high that divides the Northern and Southern Penninics,
i.e. it occupied a Middle Penninic position (e.g., Tomek, 1993;
Plašienka, 2003). This basement probably forms the lower crust of
the northern Central Western Carpathian zones (Tomek, 1993; Bielik
et al., 2004).
The Central Western Carpathians consist of three principal crustalscale
super-units (from N to S): the Tatricum, Veporicum and
Gemericum, and three detached cover nappe systems (Fatric, Hronic
and Silicic; Fig. 2). Crustal units embody a pre-Alpine crystalline
basement and its Late Palaeozoic–Mesozoic sedimentary cover. From
the point of view of regional tectonics, the Central Western
Carpathians consist of two belts: the Tatra-Fatra Belt of “core
mountains” and the Vepor-Gemer Belt.
The Tatra-Fatra Belt is characterized by pre-Alpine basementcored
mountainous areas (the so-called “core mountains”), which are
Late Cenozoic transtensional horst structures. From bottom to top, the
following units are exposed in the core mountains: (1) the Tatric preAlpine
crystalline basement and its Late Palaeozoic–Mesozoic sedimentary
cover; (2) the Fatric (Krížna) Mesozoic cover nappe system;
(3) the Hronic (Choč) cover nappe system; (4) Upper Cretaceous and
Cenozoic post-nappe sedimentary and volcanic cover. The Tatra-Fatra
Belt is separated from the Vepor-Gemer Belt by the so-called
Čertovica Line, which is a moderately south-dipping, crustal-scale
fault between the Tatric footwall and Veporic hanging wall basementinvolved
thrust sheets. Based on low-temperature thermochronology,
the core mountains (from west to east: Malé Karpaty, Považský
Inovec, Tribeč, Strážovské vrchy, Žiar, Malá Fatra, Veľká Fatra,
Ďumbierske Tatry, Tatra) were uplifted during the Neogene in
conjunction with subsidence of intervening intramontane basins
(Kováč et al., 1994; Danišík et al., 2004).
The Vepor-Gemer Belt comprises: (1) the Veporic basement/cover,
thick-skinned thrust sheet; (2) the Gemeric thick-skinned thrust
sheet composed dominantly of low-grade Palaeozoic volcano-sedimentary
complexes; (3) nappe outliers of Triassic–Jurassic volcanosedimentary
complexes of the oceanic Meliata Unit; (4) superficial
nappes of the Silicic system; (5) Late Cretaceous–Cenozoic sedimentary
and volcanic cover rocks. The Veporic basement and cover is
characterized by palaeo-Alpine (Cretaceous) metamorphism that
reaches amphibolite facies (e.g., Janák et al., 2001). Geochronological
cooling data indicate uplift and exhumation of the Veporic metamorphic
complexes during the latest Cretaceous (Kráľ, 1977; Kováč et al.,
1994; Dallmeyer et al., 1996; Koroknai et al., 2001; Plašienka et al.,
2007). Though partly rejuvenated during the Cenozoic, the VeporGemer
Belt represents the oldest exposed part of the Western
Carpathians with remnants of the earliest relief (Lukniš, 1972).
The Internal Western Carpathians encompass Palaeozoic and
Mesozoic volcanic and sedimentary, mostly carbonate, complexes in
southern Slovakia and northern Hungary. The Internal Western
Carpathians are represented by the so-called Pelso mega-unit that is
exposed mostly in the north-Hungarian inselbergs (Transdanubian
Range, the Bükk, Uppony, Szendrő and the Aggtelek–Rudabánya
Mountains; c.f. Kovács et al., 2000). The supposed Central/Internal
Western Carpathian boundary is indicated by the oceanic complexes
of the Meliata Unit, but this suture is mostly obliterated by
superimposed nappe units and overstep rocks. The Internal Western
Carpathians are mostly composed of un-metamorphosed or lowgrade
Palaeozoic and Mesozoic complexes that form a south-directed
thrust–fold belt. From the point of view of palaeogeography, the Pelso
mega-unit exhibits close similarities to the South Alpine facies realm
(Transdanubian Range), or even to the Dinarides (Bükk Mountains)—
see e.g., Schmid et al., 2008.
2.2. Neogene birth of the Carpathian Chain and Pannonian Basin System
Formation of the present Carpathian Chain and Pannonian Basin
System began at the end of the Early and beginning of the Middle
Miocene (Mazúr, 1965; Kováč, 2000; Popov et al., 2004; Rasser et al.,
2008). During this time, the Early Miocene microplates (Alcapa and
Tisza-Dacia, forming a large part of the internal zones of the
Carpathians and the whole basement of the Pannonian Basin) reached
more or less their present position adjacent to the mid-Hungarian
Fig. 2. Schematic tectonic map of the Western Carpathians: 1. Alpine–Carpathian foreland, 2. foredeep: Neogene sediments (unfolded molasse), 3. Silesian-Krosno units of the Flysch
Belt, 4. Magura units of the Flysch Belt, 5. Pieniny Klippen Belt, 6. Neogene to Quaternary sediments of the Pannonian Basin system, 7. Neogene to Quaternary volcanic rocks, 8.
Eocene to Early Miocene sedimentary rocks of the Buda Basin, 9. sediments of the Central Carpathian Paleogene Basin, 10. sediments of the Gosau Group (Late Cretaceous to Eocene),
11. Tatricum, 12. Veporicum and Fatricum, 13. Hronicum, 14. Gemericum, 15. Meliaticum, 16. Turnaicum, 17. Silicicum, 18. Transdanubicum and Bükkicum, 19. Uppony-Szendro
Palaeozoic.
160 J. Minár et al. / Tectonophysics 502 (2011) 158–174
tectonic zone (Csontos et al., 1992; Csontos, 1995; Kováč et al., 1993,
1994, 1998).
The Early Miocene movement of the Western Carpathian
lithosphere north-eastwards (due to tectonic extrusion of the Alcapa
microplate from the East Alpine domain) led to oblique collision
with the European platform. In front of the colliding palaeo-Alpine
part of the Western Carpathians, the neo-Alpine accretionary wedge
of the Outer Western Carpathians was formed. Flexure at the
European platform margin, in front of the Carpathian overthrust, was
flooded by foredeep basins. In contrast to the compressional
tectonics in the accretionary wedge zone, the intra-Carpathian
domains were subjected to stretching caused by the “slab-pull
effect” of the subducting plate in front of the orogen (Konečný et al.,
2002; Kováč et al., 1998; Royden, 1993a, b). Back-arc extension led
to initial rifting and a large-scale back-arc basin system started to
develop in the Pannonian domain (Horváth, 1993; Lankreijer et al.,
1995).
The Middle Miocene geodynamic evolution of the Western
Carpathians was still controlled by subduction at the orogenic front
(Tomek and Hall, 1993). Termination of subduction was followed by
gradual uplift of the accretionary wedge and palaeo-Alpine part of the
Western Carpathians to the end of this period (Konečný et al., 2002;
Kováč, 2000; Kováč et al., 2007; Kvaček et al., 2006). Sedimentation in
the Carpathian foredeep retreated from its western part and
continued in the eastern regions where the deepest basin depocentres
developed during this time (Meulenkamp et al., 1996). In the
Pannonian back-arc area, basin formation was influenced by extension.
The synrift stage of basin development resulted in subsidence of
individual depocentres that were filled by sediments transported by
rivers flowing from the surrounding uplifting mountains (Horváth,
1993; Kováč et al., 1997, 1998). In the western and central part of the
Pannonian back-arc basin, upheaval of an asthenospheric mantle
diapir caused block tilting along deep-seated detachment surfaces
(see Lankreijer et al., 1995; Kováč, 2000). The opening of the Danube
Basin was associated with structural unroofing of the lowermost
nappe units of the palaeo-Alpine structural pattern (Tari et al., 1992).
The asthenospheric mantle uplift was followed by voluminous acid
and calc-alkaline volcanism. Whilst this volcanic activity has been
documented in the basement of the Danube Basin, it is mainly evident
in the Central Slovakian Neovolcanic field at this time (Kováč, 2000;
Konečný et al., 2002; Lexa and Konečný, 1998; Pécskay et al., 1995,
2006). In the east, back-arc basin formation was dominantly
influenced by the pull of the submerging plate. This phenomenon
was especially manifested by extension along the hinterland of the
Eastern Carpathian accretionary wedge. The Transcarpathian Basin
opened parallel to the orogen (Kováč et al., 1995, 1998; Kováč, 2000)
and basin formation was initially associated with voluminous acid
volcanic activity, and later with calc-alkaline to basaltic volcanism of
island arc type, with a direct connection to subduction (Konečný et al.,
2002; Pécskay et al., 1995, 2006).
The Late Miocene period represented a time of subduction retreat
from the Western Carpathians eastward and south–eastward to the
front of the Eastern Carpathians (Matenco et al., 1997; Linzer, 1996,
Linzer et al., 1998). The end of subduction in the northern part of the
Carpathian orogen and isostatic rebound led to accelerated uplift of
the Western Carpathians at the end of this period. In the Pannonian
Basin domain, the extensive rifting that initially took place was driven
by the Eastern Carpathian subduction pull. After the Early Pannonian,
moderate post-rift thermal subsidence was controlled by the
development of individual basins that were mostly filled with deltaic
deposits (Horváth, 1993; Lankreijer et al., 1995; Konečný et al., 2002).
The Late Miocene basin formation was associated with basalt
volcanism (Kováč, 2000; Konečný et al., 2002; Pécskay et al., 1995,
2006).
Tectonic inversion of the Pannonian Basin began in the Pliocene.
This fact is well documented in the development of the Western
Carpathians by the uplift of mountains and gradual termination of
subsidence in the Neogene basins.
2.3. Deep-seated lithospheric structure of the Western Carpathians
versus topography
An analysis and critical evaluation of studies related to the deepseated
lithospheric structure of the Carpathians and Pannonian Basin
System showed clearly that the present topography in the region
(Fig. 3a) is influenced significantly by recent structure, composition
and geodynamics of the lithosphere (Zeyen and Bielik, 2000; Zeyen
et al., 2002; Dérerová et al., 2006).
Lillie et al. (1994) showed changes in the topography, Moho depth,
lithospheric thickness and gravity anomalies that relate to the current
structure and tectonic history of the region. Their gravity models
illustrated that the changes in the degree of continental collision in
the Western Carpathian region, the result of plate convergence and
stretching in the Pannonian back-arc basin, led to differences in
lithospheric thickness. Thicker crust and higher topography in the
Western Carpathians, along with the wavelength (about 50 km) and
amplitude (about −60 mGal) of observed gravity anomalies, are
consistent with about 50 km of continental crustal shortening and
4 km of isostatic rebound. Preservation of the thick Outer Carpathians
Flysch Belt deposits and little isostatic rebound are attributable to the
high-density, shallow mantle of the intact continent–ocean transition
zone. Lillie et al. (1994) also concluded that the effect of the
asthenosphere must be taken into account in modelling longwavelength
gravity anomalies in the Carpathian–Pannonian region.
The first consideration of the lithosphere–asthenosphere boundary
for interpreting the gravity field in this region was made by Šefara
(1986).
Geophysical study of the lithosphere was based on 2D integrated
lithospheric modelling that combined the interpretation of surface
heat flow, gravity, geoid and topography (Zeyen and Fernandez,
1994). The new model of lithospheric thickness resulting from this
modelling (Fig. 3b, c) indicates that the differences in recent
topography of the Carpathian–Pannonian region can be explained
by variations in the structure and thickness of both the crust and the
lithosphere. These differences are also reflected in gravity and surface
heat flow. The 2D modelling showed that the relationship between
recent topography and geological and geophysical parameters
changes significantly not only across the Carpathians, but also along
the mountain chain (Fig. 3a, b, c).
From the theory of isostasy, it is well known that the high
topography of mountain areas represents a topographic mass excess
that is often compensated at depth by a mass deficiency (typically
thick crust). In contrast, the relationships between topography,
crustal thickness, topography and lithospheric thickness in the
Western Carpathian region are different. While crustal thickening
results in uplifted topography, lithospheric thinning leads to lowering
of the topography (Dérerová et al., 2006).
Unlike the eastern part of the Western Carpathians and the Eastern
Carpathians, no lithospheric thickening was observed in the western
part of the Western Carpathians (the junction between the Eastern
Alps, Western Carpathians and Bohemian Massiff). The lower
topography (only about 300–400 m on average) is associated here
with relatively thinner lithosphere (∼110 km on average). This
phenomenon might be explained by the oblique collision (dominated
by strike–slips) of the Western Carpathians (Alcapa microplate) with
the European Platform (the Bohemian Massif) margin (Ratschbacher,
1991a, b; Konečný et al., 2002).
The thin and warm crust (∼26–28 km) and lithosphere (∼60–
90 km) in the Pannonian Basin System reflects stretching (extension)
of the crust and associated uplift of the asthenospheric mantle. These
low crust and lithosphere thicknesses are consistent with the low
topography (150–200 m on average) in the Pannonian Basin region
161J. Minár et al. / Tectonophysics 502 (2011) 158–174
(Fig. 3c). Here, the elevated Moho, which represents the mass excess,
is isostatically compensated by the mass deficiencies of the Pannonian
Basin infill (sediment thickness varies from 0 to 10 km) and the
elevated asthenosphere beneath the Pannonian Basin. Regional
geophysical cross sections (e.g., Lillie et al., 1994; Szafián et al.,
1997; Zeyen et al., 2002 and Dérerová et al., 2006) indicate that
lithospheric and crustal thinning is confined only to terranes of the
Alcapa and Tisza-Dacia microplates (the region of the Pannonian
Basin System and/or intra-Carpathian region). The elevation of the
lithosphere–asthenosphere boundary and the Moho in the Pannonian
Basin is thermally controlled (e.g., Čermák, 1982; Zeyen et al., 2002
and Dérerová et al., 2006).
3. Western Carpathian morphostructures
3.1. State of the art
The Western Carpathians comprise multiple hierarchically-ordered
morphostructures. At the first level, they are an extensive,
relatively flat and elliptical elevated area (dome). Its rise Mazúr
(1965) attributed to repeated vertical movements conditioned
probably by sub-crustal magma movement or lateral pressure in
combination with an isostatic response.
Along its entire periphery, this elevated dome is bordered by the
depressions of the Pannonian basin systems and Vienna basin,
Carpathian Foredeep and the Transcarpathian depression. The surface
of the Western Carpathian megaform is irregular, but a mosaic of
discrete mountains (mainly horst and dome structures) and basins
(mainly graben and flexures) creates distinctive patterns (Fig. 4).
The compactness of the megaform decreases from the NE to SW
and, in the same direction, the area covered by depressions also
increases. The SW margin is open to the Pannonian Basin where the
extent of the elliptical morphostructure is only indicated by spurs of
narrow, low mountains and slightly-elevated hilly lands in the
Pannonian Basin. The average altitude of the megaform ranges from
300 to 1500 m, with the highest ranges located near the NE focus of
the ellipse. In contrast, the altitudinal minimum lies near the SW focus
on the gradational boundary with the Pannonian Basin.
While the highest-order unit (Western Carpathian dome) and the
lowest-order units (discrete mountains and basins—basic geomorphic
units) seem to be clear, the situation for intermediate units is
ambiguous. Geomorphological subdivisions and morphostructural
sketches (Pécsi and Somogyi, 1969; Balatka et al., 1971; Mazúr and
Lukniš, 1978; Kondracki, 1978; Mazúr, 1979; Demek et al., 2007) tend
only to reflect geological structures with small morphological
modifications (Fig. 5). Therefore, geomorphological subprovinces
and regions clearly correspond with geological structures (Fig. 5b)
but their correspondence with topography alone is weaker (Fig 5a).
Neglecting the neotectonic Pliocene–Quaternary development, they
largely reflect the influence of pre-Neogene and Miocene lithospheric
evolution.
However, various geological and geomorphological markers
indicate the crucial role of Late Miocene, Pliocene and Quaternary
evolution in the creation of recent relief. Fission track and radiometric
data confirm the very young (since Middle Miocene) denudational
history of many individual mountain ranges (e.g., Kováč et al., 1994;
Baumgart-Kotarba and Kráľ, 2002; Struzik et al., 2002; Bíl et al., 2004;
Danišík et al., 2004, 2008). Increasing tectonic activity in the
Quaternary can be deduced from the height of river terraces
(Mazúrová, 1978). The neotectonic map of Slovakia (Maglay et al.,
1999) stresses Pliocene–Quaternary tectonic activity. In the Outer
Western Carpathians, the formation of planation surfaces and river
terraces shows a number of narrow (15–20 km) but long (50–
150 km) elevated and subsided structures. These structures originated
in response to Pliocene–Quaternary relaxation of remnant
horizontal movement within the Flysch Belt (Zuchiewicz, 1998).
Young, active morphostructures are also described from various parts
of the Western Carpathians (e.g., Mazúr, 1965; Činčura, 1969; Lukniš,
1973; Harčár, 1983; Lacika, 1990; Stankoviansky, 1993, 1994; Minár
et al., 2003; Bizubová et al., 2005; Beták and Vojtko, 2009).
The most recent morphostructural subdivision of Slovakia (Lacika
and Urbánek, 1998; Urbánek and Lacika, 1998) reflects the dominant
role of the Late Miocene–Quaternary arched uplift of the Western
Carpathians. The main morphostructural units are defined as parts of
the Western Carpathians dome morphostructure (Central, Transitional
and Marginal), complemented by southern depressed and
elevated morphostructures. Although the link to the main morphostructural
units of relief is evident, the purely expert-based approach
without exact definition of regionalization criteria and methods raises
doubts about detailed results (boundary location, delimitation of 3rdorder
morphostructural units).
3.2. New method of morphostructural subdivision
Delimitation of morphostructures can be done in two principal
ways. Looking for the morphological expression of geological
structures is the traditional, more static and very simple approach
(Fig. 6-a). However, this approach is limited in the case of complex
tectonic evolution. The morphologically-based method used
here arises from the application of targeted morphometric
analysis and multivariate statistical methods to the delimitation of
the most morphologically distinct and most clearly interpretable units
(Fig. 6-b).
Targeted morphometric analysis can be considered as a part of the
specific geomorphometry (Evans, 1972, 1986), where definition of
analyzed map units, selection of generally used morphometric
variables and definition of specific variables as well as procedures of
their analyses and using are dependent on given target.
Therefore we used the basic geomorphological units that are
morphologically defined but they also represent basic morphostructures—mostly
individual horsts, grabens or flexures. Larger morphostructures
can be delimited by examining the clustering of these basic
units.
The basic geomorphological units (170 units, mostly 200–
1000 km2
in size) were characterized by mean, median and extreme
values of three standard morphometric variables (altitude, slope and
available relief—the difference between maximum and minimum
altitude in a given area). A good morphostructural interpretation of all
used morphometric variables is the first aspect of targeting of
morphometric analyses. Altitude reflects the intensity of vertical
tectonic movements. Uniform tectonic blocks tend to homogeneity of
average altitude; fault boundaries manifest themselves by step
changes in average altitude. Slope, together with available relief,
defines the relief roughness that is influenced by rock resistance
(passive morphostructure) and age of the morphostructure. Of course,
altitudinal position (e.g. height of the tectonic uplift) also influences
on roughness.
As exogenic processes considerably influence slope aspect,
morpholineaments were used instead of aspect to express the spatial
orientation of morphostructures. Moreover, the morpholineaments
correlate with important structural boundaries and can be used for
morphostructure delimitation. Morpholineaments were independently
selected on DEM at two different scales: 1:3 000 000 (basic
and oldest morpholineaments) and 1:1 000 000 (also reflecting
younger development). To highlight regularities in the network of
lineaments and to ensure compatibility with fault systems, parallel
and orthogonal morpholineaments were preferred (Fig. 7).
Comprehensive knowledge of geological, geophysical and geomorphological
models and concepts is a prerequisite for effective
targeted morphometric analysis. This knowledge leads to more
clearly interpretable morphostructural units. After considering
previous work on the Western Carpathians (see above), we
162 J. Minár et al. / Tectonophysics 502 (2011) 158–174
Fig. 3. (a) Topography (m) of the Carpathian–Pannonian Basin System region (from the GTOPO30 data set Gesch et al., 1999). (b) Map of lithospheric thickness (km) in the
Carpathian–Pannonian Basin System region (modified after Dérerová et al., 2006). Legend: VI—location of the transect shown in c. (c) Lithospheric model for profile VI (modified
after Dérerová et al., 2006). The dots correspond to measured data with uncertainty bars and solid lines to calculated values.
163J. Minár et al. / Tectonophysics 502 (2011) 158–174
identified three crucial stages in the morphostructural evolution of
the Western Carpathians important for targeted morphometric
analysis:
a) Pre-Neogene processes that lead to the formation of a passive
morphostructure, but also influenced the formation of morpholineaments
that, after neotectonic reactivation, could later become
the boundaries to active morphostructures.
b) Mostly extensional Miocene tectonics leading to creation of a
mosaic of elevated and depressed areas that resulted from
stretching of the overriding plate by the subduction pull effect.
c) Structural inversion connected with intensive Late Miocene–
Quaternary uplift and creation of the recent Western Carpathian
domal mega-morphostructure.
While the first two geotectonic stages are distinctly expressed in
the mosaic of basic geomorphological units (lowest-order morphostructures),
the influence of the last stage (c) is overprinted by the
former (a, b). To identify a morphological expression in (c) we have to
eliminate especially the strong morphological influence of (b).
Separation of positive (elevations) and negative (depressions)
basic morphostructures and subsequent separate cluster analysis
Fig. 3 (continued).
164 J. Minár et al. / Tectonophysics 502 (2011) 158–174
proved to be the solution. To achieve it a specific morphometric
variable, relative height, was defined. Relative height along a
boundary segment is defined as the difference in mean altitude of
the units on either side (Fig. 8a). The affinity of a basic unit to a
depression or elevation is calculated by averaging the positive and
negative differences along each boundary segment, weighted by the
length of each segment (Fig. 8b—the degree of affinity to elevation is
expressed by positive values and affinity to a depression is expressed
by negative values).
A limiting value for relative height of ±50 m was applied to
eliminate ambiguous units (with low affinity to both elevations and
depressions). A set of statistical tests was separately applied to the
clear elevations (63 units) and depressions (45 units). Input data are
given in Table 1. The factor analysis used Varimax rotation and
concentrated the morphometric information into a small number of
independent and representative variables, or factors. Pairs of similar
variables (mean and median altitude, total relief within a unit and 2km
relief or two boundaries characteristics) were selected in order to
eliminate any statistical imperfection of the particular variables. More
than 99% variance was contained in the first four factors (F1 to F4) for
both elevation and depression sets (Table 2). Factors F3 and F4 mainly
represent affinity to depressions/elevations and the total relief that we
would like to eliminate as an expression of the older Neogene
evolution. Moreover, only 7.6% (elevations) or 18.4% (depressions)
of total variability is contained in these two factors, so we used only the
first two factors in the subsequent analysis. Despite this, factors F3 and
F4 show the relative significance of membership of the units in the sets
of elevations or depressions (the same relative height is less important
in a highly-dissected area than in a flat area) and can be used for the
explanation of some abnormalities in the final regionalization.
To find the most easily interpreted results, various sets of input
data and different factor and cluster-analysis procedures and settings
Fig. 4. Elliptical morphostructural dome of the Western Carpathians (1) and axes of positive (2—in the Carpathians, 4—in the Pannonian Basin) and negative (3—in the Carpathians,
5—in the Pannonian Basin) morphostructural units of the lowest-order (after Mazúr, 1976).
Fig. 5. Higher-order units of traditional geomorphological subdivision (synthesis of national geomorphological subdivisions: Mazúr and Lukniš, 1978; Balatka et al., 1971; Kondracki,
1978; Pécsi and Somogyi, 1969) in comparison with (a) topography and (b) geological structure (see Fig. 2 for details).
165J. Minár et al. / Tectonophysics 502 (2011) 158–174
were tested. The most commonly used methods (Nearest Neighbour,
Furthest Neighbour, Centroid, Median, Group Average, Ward's and Kmeans)
and three distance metrics (Squared Euclidean, Euclidean, and
Simple Manhattan) were applied. The results obtained were evaluated
from the point of view of the stability and spatial continuity of
the subdivision, as well as morphostructural interpretability. The best
results were achieved using the Squared Euclidian metric as it most
clearly singled out interpretable morphological domains. Ward's
method was selected for displaying the clustering results shown in
Fig. 9 and Fig. 10 because these results were required to minimal
change in subsequent steps.
The clusters obtained ensure that there is maximum statistical
difference between clusters in attribute space, but no maximal
contrast between neighbouring units in geographical space. In
ambiguous areas where various cluster methods offered different
subdivisions, we tested the separation of geographically neighbouring
units in the attribute space (F1, F2). Subsequently, we reclassified
some boundary units so that we achieved maximum contrast on the
geographical boundary of a morphostructure (Fig. 11a). A few units in
the elevation set remain surrounded (in geographical space) by units
that are classified differently. In all cases, we found reasons for these
differences and reclassified them (Fig. 11b, see also next chapter).
Both groups of reclassified units are depicted in Fig. 9.
An overlay of elevation and depression clusters (Fig. 10) correlates
well in geographical space and was the basis for the final delimitation of
morphostructural units. We selected five elevation clusters but only four
depression clusters because the depressions do not include a unit like
the Tatra Mountains with extreme elevations. A small shift between the
character of elevation and depression clusters can have a systematic or
local character and sometimes this was a reason for the definition of
morphostructural subregions. Subregions were also defined on the basis
of evident local morphological contrasts (distance in F1, F2 space) that
were generally reflected in the alternative clustering (other methods,
distance metrics or number of clusters).
The last step in our analysis was to make a schematic generalization
of morphostructure boundaries. For better morphotectonic
interpretation of delimited regions, the location of the boundaries
was constrained by the nearest morpholineaments (Fig. 7).
3.3. New morphostructural subdivision
The delimited morphostructural units (Fig. 12) have significantly
higher morphological homogeneity then traditional geomorphological
units of a similar dimension. When compared to the 8 new
regions (compounded from more than one elevation and one
depression unit), the average homogeneity of the 15 traditional
Fig. 6. Procedures for morphostructure identification.
Fig. 7. Topographic lineaments derived at two different scales: a) 1:3000 000, b) 1:1000000.
166 J. Minár et al. / Tectonophysics 502 (2011) 158–174
regions (Fig. 5) is almost three times less for elevation units and
more than two times less for depression units. This increase in
homogeneity is doubled for 16 newly delimited subregions (Table 3).
The following description of the new regions outlines their basic
character and serves as a basis for more detailed morphostructural
interpretation.
1– The Tatra region is morphologically the most extreme region and is
represented by a single elevation unit. Despite having a very small
area, we had to separate it at the region level because of its
extreme morphological contrast. The morphological distance from
the nearest unit (Nízke Tatry Mountains) is greater than the
distance range in all other regions and the region was distinct in all
clustering experiments. The distinct topographic anomaly cannot
be explained simply as a culmination of dome elevation—other
concepts for explaining the creation of the Tatra Mountains horst
must be considered (e.g., Bac-Moszaszwili, 1993; Marko et al.,
1995; Sperner et al., 2002). In summary, these concepts can be
described in terms of an upper-crustal, basement-involved
compressive structure developed in response to late-Tertiary
collision events along the northern Carpathian boundary. Related
shortening was estimated to be 30% and the maximum vertical
uplift of the Tatra Mountains is about 8 km (Janák et al., 2001).
2– The Central Region consists of 7 elevation and 6 depression units.
Three elevations (Kozie chrbty Mountains, Skorušinské vrchy
Mountains, Gubałówka Mountains) have poorly-fitting (significantly
lower) relief (see Fig. 9a) that can be explained by younger
uplift of the original basin areas. Lukniš (1973) and Jakál (1992)
confirm changes in Quaternary river networks and uplift of the
first unit, while Kondracki (1978) describes Late Miocene–
Quaternary uplift of the other two units that were once part of
the large Obniżenie Orawsko-Podhalańskie Depression. We will
term them “delayed elevations” because they were raised as
morphological elevations later than the surrounding mountains.
Without them, the Central Region is quite homogeneous—all
experiments separated it (or its two subregions) at the level of 4–6
clusters. The region is the second highest part of the Western
Carpathian dome and is divided along a distinct N–S morpholineament
(hereafter the Central morpholineament) into two
morphologically different parts: Tatra Subregion (2a) and Fatra
Subregion (2b). Their morphological distance is close to the
distance between regions and the trend of the morpholineament
system changes sharply from W–E (2a) to SW–NE (2b). Although
the Chočské vrchy Mountains share morphological similarities
with the Fatra Subregion, they are probably only a later-activated
part (delayed elevation) of the Tatra Subregion, because of their
exceptional structure (no outcrop of the core), morphology (many
transverse disequilibrium canyons) and their position in relation
to the surrounding depression units.
3– The Transitional Region encloses Central region and consists of 23
elevation and 11 depression units. Despite the highest number of
units, a good separation of the region was achieved in the majority
of the clustering experiments. Only five classified elevation units
are specific (without clear membership to the transitional cluster).
Four of these (Žiar Mountains, Dzialy Orawskie Mountains,
Bachureň Mountains, Súľovské vrchy Mountains) are probably
delayed elevations (c.f. Klimaszewski, 1972; Minár et al., 2004),
while the last (Javorníky Mountains) have a transitional character
on the region boundary. Six selected subregions are parts of an
annulus that is lower to the west and east and that is broken up in
a N–S direction by the Central morpholineament that also divides
the Central Region. Subregions westwards of the Central morpholineament
(3d, e, f) are on average 200 m lower, have a radial
morpholineament system and the basic morphostructures have a
fragmented pattern. The morpholineament system of the eastern
subregions is more sparse and homogeneous and the subregions
are more massive and compact. Excepting boundaries coinciding
with the Central morpholineament, the separation of subregions is
less distinct.
4– The West Marginal Region consists of 7 elevation and 5 depression
units. The region is well separated from its neighbours except for
the boundary with the Javorníky subregion (3e) that is a transition
between the Transitional Region and West Marginal Region (some
clustering experiments classified it as a part of the West Marginal
Region). The Javorie Mountains (4b), a small subregion east of the
Central morpholineament, is significantly higher, but local
morphological distances expressly place it into this region. A
major western subregion (4a) represents a low periphery of the
fragmented western part of the Western Carpathians dome. The
region has a character (morpholineament system and broken,
small basic morphostructures) that is partially similar to the
western part of the Transitional region.
5– The North–East Marginal Region consists of 3 elevation and 4
depression units with good separation of elevations from the
surrounding regions in all clustering experiments. However, 3 of
the depression units have a specific character (relatively low, but
highly-dissected marginal hilly lands) that placed them into the
cluster of the Transitional regions in many experiments (see
Fig. 9b). The narrower western part (5a) represents the transition
between the Transitional Region (3) and the North Foreland (6).
Fig. 8. Determination of the affinity of geomorphic units to elevations and depressions. Affinity (A) is expressed by the summary relative height (Rh) of its boundary segments
A = ∑
n
i=1
Rhi⋅Li, where n is a number and L length of boundary segments.
167J. Minár et al. / Tectonophysics 502 (2011) 158–174
This subregion consists of a series of dispersed smaller positive
morphostructures and one significant tectonic depression (Kotlina
Sądecka Basin). The broader and larger eastern part of the region
(5b) represents a transverse depression through the Flysch
Carpathians between the Sandomierz and Pannonian Basins (the
morphostructure of intervening belt according to Mazúr, 1979).
The region neighbours the Eastern Carpathians and has a
transitional character. It is characterised by a classification that
is different to the national geomorphological subdivisions (Mazúr
and Lukniš, 1978 versus Kondracki, 1978), a morpholineament
system typical of the Eastern Carpathians (the majority of the
morpholineaments trend NW–SE) and reduced crustal thickness
in the Transcarpathian depression (Hók et al., 2000; Lenkey et al.,
2002).
6– The North Foreland consists of 7 only depression units. It represents
the transition between the Carpathian foredeep on one side and,
on the other side, the North–East (5), partially also the West (4)
Marginal Regions and, in one portion, even a subregion of the
Transitional Region (3f). The region was separated on the basis of
basin unit properties and a lack of clarity in the subdivision was
recorded only on the boundary with the West Marginal Region
(depression units of both regions are morphologically similar and
fall within one cluster—see Fig. 9b). The unit is orientated in a W–E
direction and morpholineaments also dominantly trend in this
direction. However, N–S and SW–NE trends are more evident for
second-order morpholineaments. Though its southern boundary
is mostly structurally and lithologically controlled (Krosno versus
Magura nappe groups of the Flysch belt), the morphological
difference with southern units also results from young tectonics
(resulting in marked denudation of the neighbouring Moravian–
Silesian Beskids).
7– The South–East Region has a specific character. It consists of
8 elevation units, 8 depression units and, in particular, a high
number (9) of transitional units. The region is divided into
two subregions: a higher South–East Marginal Subregion (7a)
and a lower South–East Foreland (7b) that could be considered
also as two independent regions. The subregions are distinct in
the presented cluster classification (7a—Marginal regions, 7b—
Forelands, c.f. Fig. 9), but in some clustering experiments they
fitted into one cluster. Because division on two independent
regions did not significantly improve the mean homogeneity of
regions (see Table 3) we vote their distinguishing only on the
subregion level. Moreover, the character of the Slanské vrchy
Mountains is similar to the Transitional Region (3), but local
morphological distances support our final classification. Another
elevation unit of the marginal subregion 7a (Cerová vrchovina
Mountains) fitted into the foreland cluster, but it is a clear
delayed elevation (Lacika, 1990). Overall, this region shows the
lowest individuality (mainly because its boundaries with
regions 5 and 8 are ambiguous) and its morphological
neutrality (i.e. many transitional units) connects it with the
Transdanubian Mountains. Values of F1, F2 generally decrease
from the NE to SW that confirms ambiguous gradient character
of the region. The dominant morpholineament directions
clearly change with their order (SW–NE for first-order
morpholineaments; N–S for second-order morpholineaments).
This change could reflect very young tectonic changes in the
region.
8– The South–West Foreland consists of 8 elevation (and 3 parts of
elevations) and 2 big depression units. The elevations create
small spurs into young (Quaternary) uplifted parts of the
Pannonian and Vienna Basins. Elevations are significantly lower
than in the adjacent West Marginal Region (4). The southern
parts of three elevation units (Považský Inovec Mountains,
Tribeč Mountains, Štiavnické vrchy Mountains) were separated
from the West Marginal Region and integrated into South–West
8Foreland. This solved an ambiguous classification mainly
related to the Tribeč Mountains. The separation from other
surrounding regions is also clear. Depression units of the region
are morphologically less distinctive than those of the South–East
Region and elevation units are more distinctive than those of the
Transdanubian Mountains. The westernmost part was separated
as a special subregion (8a) and is characterised by lower
morphological factor-values and the dominance of transitional
units. The absolute dominance of NE–SW first-order morpholineaments
(higher in the 8a subregion) is replaced by a
significant increase in the abundance of the N–S second-order
morpholineaments (dominating in the 8b subregion).
9– The Transdanubian Mountains consist of only 3 clear elevation
units, but these were expressly separated in all clustering
experiments. Despite this, transitional units dominate both by
frequency (10 units) and area. The transitional units are similar
to the foreland regions by nature of their morphological
neutrality and low relief, but they differ in the absence of
depression units. A distinct morphological individuality results
from the specific morphotectonic development of this region.
This morphostructure comprises a series of indistinct horsts. In
the Tertiary, the Transdanubian Mountains were subjected to
differentiated faulting, burial and renewed uplift. The individual
small horsts and grabens acquired different altitudes. In a
geomorphological sense, the mountain range was only created
by epeirogenic uplift that began in the Upper Miocene and
intensified in the Pliocene and Quaternary. Thus, the range of
planated horsts with graben-like basins is morphologically not a
block mountain, but a young Alpine structure (Pécsi, 1996).
Distribution of the first- and second-order morpholineaments is
very similar to the neighbouring regions 7 and 8, but the
inversion is more distinct and W–E directions dominate the
second-order morpholineaments.
Table 1
Input variables to factor analyses. A digital elevation model (DEM) derived from SRTM
data (Jarvis et al., 2008) with 90 m resolution and GRASS GIS (Neteler and Mitasova,
2008) was used for computation. Explanation of symbols see also in the text.
Variable Definition Used for:
A1 Arithmetic mean of altitude Elevations+
depressions
A2 Median of altitude Elevations+
depressions
Amax Maximum of altitude Elevations
Amin Minimum of altitude Depressions
TR Total relief within a unit=Amax −Amin Elevations+
depressions
S Arithmetic mean of slope Elevations+
depressions
AR Arithmetic mean of 2 km available relief (2 km window) Elevations+
depressions
Rh Arithmetic mean of relative height weighting by length of
boundary segment
Elevations+
depressions
% Rh+ % of boundary length with positive Rh Elevations
% Rh− % of boundary length with negative Rh Depressions
Table 2
Summary of the factor analysis. STATGRAPHICS Plus for Windows was used for
computation. See text and Table 1 for detailed explanation.
Elevations Depressions
Factor Variance Maximal loads Variance Maximal loads
F1 79.0% A1 (0.90), A2 (0.89) 67.8% Amin (0.95), A1 (0.90),
A2 (0.89)
F2 13.4% S (0.76), AR (0.72) 13.9% S (0.90), AR (0.85)
F3 5.0% TR (0.78), Amax (0.61) 12.0% Rh (0.69), % Rh+(0.62)
F4 2.6% % Rh+(0.77), Rh (0.63) 6.4% TR (0.84)
168 J. Minár et al. / Tectonophysics 502 (2011) 158–174
3.4. Discussion of geomorphic history and new morphostructural
subdivision
The geomorphic history of the Western Carpathians, including
timing, character and reasons for tectonic uplift leading directly to
their present (modern) relief, are all reflected in the new morphostructural
subdivision. The most important features include:
a) The previously-documented dome-like character of the Western
Carpathians supra-region with its very young features.
b) The good conformity of river terraces and planation surfaces with the
delimited morphostructural regions, but not with fission-track data
(thescattered character of exhumation in individual regionsindicates
uplift these regions prior to the young morphostructures formation).
c) The cascade arrangement of the defined regions and the saddle
shape of the Transitional region (3) that indicates the character of
the young differential vertical tectonic movements.
d) The growth of altitudinal differences between depression and
elevation units for the central parts of the supra-region.
e) The projection of the older structural boundaries into some newlydefined
morphostructural regions, mainly in the northern and
southern periphery.
f) The increased number of young (N–S and W–E trend) morpholineament
systems, mainly on the southern, but also partially on
the north periphery.
g) The distinctive change in character of the regions and subregions
along the central morpholineament boundary.
The explanation for the dome-like character of the Western
Carpathians supra-region with its very young features was
mentioned only briefly in older literature. However, the reasons
for this phenomenon were either not dealt with or the tectonic
explanation was limited by the level of knowledge at that time (c.f.
Mazúr, 1965; Klimaszewski, 1981). We suggest that the source of
this uplift could be very similar to the Alps, naturally with some
peculiarities.
In the Alps there are principally two concepts for explaining uplift.
Some researchers argue that climate is/was important for the young
uplift and the increase in erosion (e.g., Schlunegger, 1999; Kuhlemann
et al., 2006). The second school argues for tectonically-driven uplift. In
this group, a variety of mechanisms have been proposed to explain the
uplift. Genser et al. (2007) assumes that the late-stage widespread
uplift of the entire Alpine realm (at about 6 Ma), including the
northern peripheral foreland basin, is most likely explained by
distributed delamination and/or convective removal of over-thickened
lithosphere and the northward spread of the mechanical
decoupling between the mantle and crust of the subducted lithosphere.
These processes are thought to be enhanced by erosional
unloading. The experiments described by Willingshofer and Sokoutis
(2009) depict such a situation during the late Miocene–Pliocene,
where the foreland plate was thrust under the orogenic wedge,
leading to crustal thickening and uplift of the internal part of the
orogen and subsidence of the foreland plate. They argue that
collisional mountain belts in general are expected to evolve from a
decoupled to a coupled system, leading to a change in the dominant
Fig. 9. Basic clusters of elevational (a) and depressional (b) units derived by using Ward's method and Squared Euclidean distance. Factors F1 and F2 are detailed in Table 1.
Ambiguous (neutral) units are not depictured.
169J. Minár et al. / Tectonophysics 502 (2011) 158–174
deformation mechanism as portrayed in the vertical motions of the
orogenic wedge and the foreland.
The neotectonic rise of the Western Carpathian dome clearly
and continuously deforms the initial planation surface (Lukniš,
1962)—‘Mid-mountain level’ (Mazúr, 1963) and the uplift should,
therefore, be younger. In contrast, the exhumation age derived from
fission-track data is very variable within the delimited regions (e.g.
Kováč et al., 1994; Kováč, 2000; Baumgart-Kotarba and Kráľ, 2002;
Danišík et al., 2004, 2008; Struzik et al., 2002). This indicates that the
detected exhumation history is older than the development of the
dome. The youngest fission-track data from about 10 Ma determines
the maximum age of the ‘Mid-mountain level’ that itself requires a
Fig. 10. Map view of the separate clusters (4 depression clusters, 5 elevation clusters) and their synthesis.
Fig. 11. Two reasons of unit reclassification: a) Ambiguous boundary unit classified by various experiments to various clusters; finally classified on the basis of local contrast.
b) Indistinctive elevation unit surrounded by uniformly classified units; classified into cluster of neighbouring units. DFab—distance of units a and b in the attribute space (F1, F2).
170 J. Minár et al. / Tectonophysics 502 (2011) 158–174
few million years to form. Generally, fine-grained Late Pannonian and
Pontian correlative sediments in the Pannonian Basin and Western
Carpathians intramountain basins are also indicative of the formation
of the ‘Mid-mountain level’. The general character of sedimentation
was demonstrably changed in the Pliocene (coarse sediments
substituted fine sediments). Consequently, the dome probably first
rose sometime during the last 4–6 million years. However, while the
altitude of the ‘Mid-mountain level’ corresponds with the mean
altitudes of the delimited regions, the younger ‘River Level’ (Upper
Pliocene–Early Quaternary pediment after Mazúr, 1963) and Quaternary
river terraces differ far less between individual regions. This
indicates that the main stage of the dome formation occurred in the
Pliocene and that both the ‘River level’ and the river terraces were
formed in the existing dome.
While the isostatic component of the accepted model described
above is clearly confirmed by lithospheric thickness modelling, the
annular form of the regions and the saddle shape of the Transitional
region are well explained by the influence of pressure from the
encroaching Adriatic plate. This influence began in the Western
Carpathians after relative cooling and strengthening of the Pannonian
Basin lithosphere. Moreover, the time elapsed between the end of the
extensional tectonic phase, with an elastic character reflecting
overheated lithosphere (Sarmatian–Early Pannonian), and the relative
strengthening of the Pannonian Basin lithosphere in a postrift/
tectonic inversion state (Pontian–Early Pliocene), offers sufficient
time for the formation of the ‘Mid-mountain level’ (its character,
timing and extent represent an important part of regional geomorphology
that has been intensively discussed during the last decade—
see e.g. Činčura, 1998; Urbánek, 2001; Minár et al., 2004; Bíl et al.,
2004; Danišík et al., 2008; Beták and Vojtko, 2009).
Middle Miocene extension in the Central and Internal Western
Carpathians was connected with back–arc extension and active
elongation of the Western Carpathians orogen in response to
subduction roll back in front of the Eastern Carpathians (Csontos,
Fig. 12. Morphostructural subdivision of the Western Carpathians (numbers correspond with labelling on the map), regions and (in brackets) subregions: 1. Tatra Region, 2. Central
Region (2a. Tatra Subregion, 2b. Fatra Subregion), 3. Transitional Region (3a. Beskid Żywiecki—Gorce Subregion, 3b. Beskid Sądecki—Levočské vrchy Subregion, 3c. Slovenské
rudohorie Subregion, 3d. Strážovské vrchy—Kremnické vrchy Subregion, 3e. Javorníky Subregion, 3f. Moravian–Silesian Beskids Subregion), 4. West Marginal Region (4a. Biele
Karpaty–Štiavnické vrchy Subregion, 4b. Javorie Subregion), 5. North–East Marginal Region (5a. Beskid Wyspowy Subregion, 5b. Low Beskids Subregion), 6. North Foreland, 7.
South–East Region (7a. South–East Marginal Subregion, 7b. South–East Foreland), 8. South–West Foreland (8a. Vienna Basin Subregion, 8b. Danube Basin Subregion), 9.
Transdanubian Mountains. Geomorphological units mentioned in the text: A Tatra Mountains, B Nízke Tatry Mountains, C Kozie chrbty Mountains, D Skorušinské vrchy Mountains, E
Gubałówka Mountains, F Obniżenie Orawsko-Podhalańskie Depression, G Chočské vrchy Mountains, H Žiar Mountains, I Dzialy Orawskie Mountains, J Bachureň Mountains, K Javorie
Mountains, L Javorníky Mountains, M Kotlina Sądecka Basin, N Moravian–Silesian Beskids, O Slanské vrchy Mountains, P Cerová vrchovina Mountains, R Považský Inovec Mountains,
S Tribeč Mountains, T Štiavnické vrchy Mountains.
Table 3
Mean distance of elevation (E) and depression units (D) of new and traditional
morphostructural subdivision in the attribute space F1, F2. Distance was not computed
if only one clear elevation or depression unit create a region.
New regions and
subregions
E D Traditional subprovinces
and regions
E D
1 Tatry Region – – Inner Western
Carpathians
5.51 3.73
2 Central Region 0.95 2.45 Fatra-Tatra Region 10.11 1.59
2a – 0.20 Slovak Middle Mountains 1.44 0.82
2b 0.16 0.79 Slovak Ore Mountains 1.51 –
3 Transitional Region 0.40 0.75 Lučenec–Košice
depression
– 0.55
3a 0.10 0.25 Matra–Slansky Region 0.35 0.40
3b 0.08 1.01 Outer Western
Carpathians
2.72 2.95
3c 0.09 0.28 Austrian–Moravian
Carpath.
0.02 –
3d 0.10 – Middle Moravian Carpath. – –
3e 0.17 – Slovak–Moravian Carpath. 0.34 0.05
3f 0.12 0.16 Western Beskids 0.76 0.16
4 West Marginal Region 0.17 0.16 Middle Beskids 1.22 3.89
4a 0.11 0.05 Eastern Beskids 0.11 5.08
4b – – Low Beskids – 1.61
5 North–East marginal
region
0.10 0.52 Western Beskids Piedmont – 0.29
5a 0.08 – Eastern Beskids Piedmont – 0.96
5b – 0.20 Podhôľno–Magura Region – 3.66
6 North Foreland – 0.41 Transdanubian Mountains 0.35 –
7 South–East Region 1.63 0.51 Small Danube basin – –
7a 0.17 0.17 Danube Lowland – –
7b 2.78 0.22 Vienna Basin – –
8 South–West Foreland 0.58 0.21 Záhorská Lowland – –
8a 0.23 – Moravian Basins – –
8b 0.05 –
9 Transdanubian
Mountains
0.20 –
Mean regions 0.58 0.72 Mean subprovinces 2.86 3.34
Mean subregions 0.33 0.33 Mean regions 1.62 1.59
171J. Minár et al. / Tectonophysics 502 (2011) 158–174
1995; Kováč, 2000; Konečný et al., 2002). This extension played
an important role in the development of the basins–mountains
(depressions–elevations) mosaic specific to the Western Carpathians.
This mosaic was later amplified on the younger rising (uplifting)
domal mega-morphostructure with an extensional regime on the
surface. The influence of different unloading of the basins and
mountains created conditions suitable for the observed Pliocene–
Quaternary growth of the altitudinal differences between the
depression and elevation units, mainly in the central parts of the
Western Carpathians (c.f., Genser et al., 2007; Willingshofer and
Sokoutis, 2009).
The projection of the older structural boundaries into our newlydefined
morphostructural regions and the increased abundance of
young morpholineament systems (N–S and W–E), mainly on the
southern and northern periphery, also confirm and complete the
concept outlined above. The first factor suggests that while the
influence of the older structure in the central part of the dome was
markedly reduced by dome uplift (because uplift is more intensive
and probably older), the periphery preserves the influence of the
older structural elements because of less distinct and younger uplift.
The younger character of the periphery is confirmed by the
abundance of the young morpholineament system, but also by the
change in ‘River level’ character from the pediment (in the central
part of the Western Carpathians) into a pediplain that truncates the
youngest Neogene (Pliocene) sediments. However, a high abundance
of the young morpholineaments could also reflect: (a) the
presence of thin-skinned crustal bodies (thin-skinned tectonics) in
the peripheral regions that are more sensitive to the most recent
tectonic pulses; or (b) southward fading of the oroclinal shape
influence on the recent stress field (the recent stress field indicates
compression orthogonal to the Western Carpathians arc and
extension parallel to it—c.f. Hók et al., 2000; Cloetingh et al., 2002;
Vojtko et al., 2008).
The distinctive splitting of the Western Carpathians dome along
the N–S oriented Central morpholineament (Central Slovak Fault
System) can be explained by geophysical data that suggest rapid
thinning of the crust and lithosphere to the west that led to lowering
and disintegration of morphostructural units. However, the influence
of crust and mantle thinning diminishes in the south (where the
lithosphere is thinner to the east) and the young dynamics related to
the recent approach of the Adriatic plate provides an alternative
explanation (the Danube River has followed the Central morpholineament
only since the Middle Pleistocene). Together with the
abundance of young morpholineaments found frequently in the
youngest Neogene sediments in the south, increasing river activity
since the Middle Pleistocene (based on the height of the river
terraces) and possibly also young elevated and subsided structures of
the Outer Western Carpathians (c.f. Zuchiewicz, 1998) could indicate
that the most recent and more active stage of the morphotectonic
development started in the Middle Pleistocene.
4. Conclusions
• Original targeted morphometric analysis enabled the delimitation
and basic description of nine specific morphostructural units
(regions) creating parts of the Late Miocene–Quaternary uplifted
dome-like mega-morphostructure of the Western Carpathian
supraregion: 1. Tatra Region, 2. Central Region, 3. Transitional
Region, 4. West Marginal Region, 5. North–East Marginal Region, 6.
North Foreland, 7. South–East Region, 8. South–West Foreland, 9.
Transdanubian Mountains.
• The morphologically-oriented morphostructural subdivision of
the Western Carpathians is an alternative to older structurallybased
subdivisions (Pécsi and Somogyi, 1969; Balatka et al., 1971;
Mazúr and Lukniš, 1978; Kondracki, 1978; Demek et al., 2007) and
improves and completes the first morphologically-oriented
subdivision (Lacika and Urbánek, 1998) in a quantitative manner.
The morphostructural subdivision is established not only from
information contained in relief characteristics, but also on the
basis of the geodynamic evolution of the orogen. The use of
consistent regional elements of geomorphological, geological and
geophysical theories and models, as well as the quantitative basis
for regionalization, all offer sound data for a morphostructural
interpretation that is more objective than the preceding qualitative
subdivisions.
• The targeted morphometric analysis (separation of the elevation
and depression units) and subsequent cluster analysis used here
(maximizing the internal morphological homogeneity and external
differentiation of regions) lead to delimited regions whose
morphological homogeneity is several times higher than for
traditional divisions (Mazúr and Lukniš, 1978).
• The morphostructural plan of the Western Carpathians presented
here shows a certain degree of correlation with geological patterns,
but mainly with the geodynamic evolution of the Western
Carpathians. While the Palaeozoic and Mesozoic geological pattern
rarely influenced the morphostructural regions (it can only be
detected in some boundary morpholineaments), the Neogene to
Quaternary geodynamic processes expressed in the young active
fault structures determined the limits and character of the delimited
morphostructures. Therefore, all delimited Western Carpathian
morphostructural units can be regarded as active Neogene–
Quaternary morphostructures (endogenous landform elements in
terms of Gerasimov and Meščerjakov, 1967). Passive (older)
morphostructures (reflecting the resistance of the rocks affected
by exogenous processes) do not significantly manifest themselves in
the morphometric variables used.
• The modern synthetic hypothesis of Western Carpathian relief
formation was outlined in connection with the presented results.
Particular characteristics of the delineated morphostructural units
can contribute to the determination of the main stages of modern
Western Carpathian relief formation: the planation of a large part
of the Western Carpathians at the Miocene/Pliocene boundary;
the beginning of the dome morphostructure formation before 4–
6 million years; the main stage of the uplift to the Late Pliocene;
and the last, more active stage starting in the Middle Pleistocene.
• The new morphostructural subdivision is also in good agreement
with the probable mechanism of the Western Carpathian uplift
(distributed delamination, convective removal of over-thickened
lithosphere, spreading of the mechanical decoupling between the
mantle and crust and erosional unloading preceded by an
extensional roll-back effect). The subdivision enabled us to define
various spatial specifications and their consequences (e.g. young
spreading of the Western Carpathians into the Pannonian Basin
realm and the specific character of morphostructural units that are
divided by the N–S Central morpholineament).
• All the outlined interpretations create a new holistic framework for
subsequent detailed regional studies focused on the consistent
geomorphodynamic explanation of all morphological features
within the delineated morphostructural regions.
Acknowledgements
This work was supported by the Slovak Research and Development
Agency under contracts ESF-EC-0006-07, APVV-0158-06,
APVT-51-002804 and APVV LPP-0120-06, as well as by the Scientific
Grant Agency of the Ministry of Education of the Slovak Republic
and Slovak Academy of Science (projects Nos. 1/3066/27, 1/0461/
09, 2/0107/09, 2/0072/08, 2/6019/6, 1/3066/06, 1/4042/07 and 1/
4044/07).
The authors thank to R. Hackney for assistance with English and
anonymous reviewers and I. S. Evans for inspiring remarks and
comments.
172 J. Minár et al. / Tectonophysics 502 (2011) 158–174
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