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Tectonophysics 430 (2007) 1 – 25
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Seismic crustal structure between the Transylvanian Basin
and the Black Sea, Romania
F. Hauser a,1 , V. Raileanu b , W. Fielitz c,⁎, C. Dinu d ,
M. Landes a , A. Bala b , C. Prodehl a
a
Geophysical Institute, University of Karlsruhe, Hertzstr. 16, D-76187 Karlsruhe, Germany
National Institute for Earth Physics, P.O.Box MG-2, RO-077125 Bucuresti-Magurele, Romania
c
Geological Institute, University of Karlsruhe, Hertzstr. 16, D-76187 Karlsruhe, Germany
Faculty of Geology and Geophysics, University of Bucharest, 6, Traian Vuia Street, sector 1, RO-70139 Bucuresti, Romania
b
d
Received 11 May 2005; received in revised form 17 October 2006; accepted 19 October 2006
Available online 11 December 2006
Abstract
In order to study the lithospheric structure in Romania a 450 km long WNW–ESE trending seismic refraction project was carried out
in August/September 2001. It runs from the Transylvanian Basin across the East Carpathian Orogen and the Vrancea seismic region to
the foreland areas with the very deep Neogene Focsani Basin and the North Dobrogea Orogen on the Black Sea. A total of ten shots with
charge sizes 300–1500 kg were recorded by over 700 geophones. The data quality of the experiment was variable, depending primarily
on charge size but also on local geological conditions. The data interpretation indicates a multi-layered structure with variable thicknesses
and velocities. The sedimentary stack comprises up to 7 layers with seismic velocities of 2.0–5.9 km/s. It reaches a maximum thickness
of about 22 km within the Focsani Basin area. The sedimentary succession is composed of (1) the Carpathian nappe pile, (2) the postcollisional Neogene Transylvanian Basin, which covers the local Late Cretaceous to Paleogene Tarnava Basin, (3) the Neogene Focsani
Basin in the foredeep area, which covers autochthonous Mesozoic and Palaeozoic sedimentary rocks as well as a probably PermoTriassic graben structure of the Moesian Platform, and (4) the Palaeozoic and Mesozoic rocks of the North Dobrogea Orogen. The
underlying crystalline crust shows considerable thickness variations in total as well as in its individual subdivisions, which correlate well
with the Tisza-Dacia, Moesian and North Dobrogea crustal blocks. The lateral velocity structure of these blocks along the seismic line
remains constant with about 6.0 km/s along the basement top and 7.0 km/s above the Moho. The Tisza-Dacia block is about 33 to 37 km
thick and shows low velocity zones in its uppermost 15 km, which are presumably due to basement thrusts imbricated with sedimentary
successions related to the Carpathian Orogen. The crystalline crust of Moesia does not exceed 25 km and is covered by up to 22 km of
sedimentary rocks. The North Dobrogea crust reaches a thickness of about 44 km and is probably composed of thick Eastern European
crust overthrusted by a thin 1–2 km thick wedge of the North Dobrogea Orogen.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Seismic refraction; Crustal velocity structure; Vrancea zone; Eastern Carpathians; Moesian Platform; Transylvanian Basin
⁎ Corresponding author. Tel.: +49 721 6082139; fax: +49 721
6082138.
E-mail address: [email protected]
(W. Fielitz).
1
Now at Geophysics Section, Dublin Institute for Advanced
Studies, Ireland.
0040-1951/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2006.10.005
1. Introduction
This study focuses on a crustal transect in Romania
with a complex geological history. It crosses from W to
E the Transylvanian Basin, the Carpathian Orogen and
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F. Hauser et al. / Tectonophysics 430 (2007) 1–25
the Carpathian Foredeep with the exceptionally deep
Focsani Basin, as well as the stable Moesian Platform
and the North Dobrogea Orogen in the foreland (Fig. 1).
Several plate or microplate boundaries (Tisza-Dacia in
the W, Moesia in the centre and North Dobrogea in the
E) are intersected, which are largely covered by the
nappes of the Carpathian Orogen or the Neogene
sedimentary basins. Whereas these plates were consolidated during complex pre-Mesozoic deformations, the
Carpathian Orogen and the younger sedimentary basins
are the result of plate convergence during Mesozoic and
Neogene times.
The youngest part of the orogen and basins in
Romania show still ongoing neotectonic activity. This is
manifested by near-surface crustal deformation resulting
in a very pronounced geomorphology (Fielitz and
Seghedi, 2005; Necea et al., 2005), subsidence in the
Focsani foreland basin (Tarapoanca et al., 2003;
Matenco et al., 2003) and a strong seismicity at
intermediate depths, which is concentrated in the
Vrancea area of the southeastern Carpathian bend
(Fig. 1; Fuchs et al., 1979; Oncescu et al., 1998).
While the shallow seismic activity scatters widely and
has moderate magnitudes (Mw b 5.6), the epicentral
region of the intermediate depth seismicity is confined
to an area of only about 40 × 80 km2 (Oncescu et al.,
1998). The epicenters of these earthquakes lie between
60 and 180 km depth within an almost vertically
elongated narrow zone and frequently have large
magnitudes (Mw = 6.9–7.4). Several major earthquakes
occurred there in the last century (1940, 1977, 1986 and
1990) causing many fatalities and an enormous
economical damage. In the 1977 earthquake alone,
1570 people died and more than 11,300 were injured.
In order to study this seismically high-risk area a joint
German–Romanian research program was initiated by
the Collaborative Research Center 461 (CRC 461)
“Strong Earthquakes — a Challenge for Geosciences
and Civil Engineering” at the University of Karlsruhe
(Germany) and the Romanian Group for Vrancea Strong
Earthquakes (RGVE) at the Romanian Academy in
Bucharest (Wenzel, 1997; Wenzel et al., 1998a). The
joint geoscientific and civil engineering research
activities of this project seek to better understand the
tectonic processes responsible for the strong intermediate-depth seismicity, and to reduce the risk by applying
appropriate techniques from civil engineering (e.g.
Wenzel et al., 1998a).
Two major active-source seismic refraction experiments were carried out in 1999 and 2001 as a
contribution to this research program (Hauser et al.,
2001, 2002). They were designed to study the crustal and
uppermost mantle structure to a depth of about 70 km
underneath the Vrancea epicentral region and were
jointly performed by the Geophysical and Geological
Institutes of the University of Karlsruhe (Germany), the
GeoForschungsZentrum in Potsdam (Germany), the
National Institute for Earth Physics in Bucharest
(Romania) and the University of Bucharest (Romania).
The obtained crustal velocity models provide additional important a priori information for passive
tomographic studies and help to calibrate the relative
velocity variations which are obtained by teleseismic
tomography (Wenzel et al., 1998b; Martin et al., 2005,
2006). The 1999 seismic refraction experiment was
published by Hauser et al. (2001). In this paper we
present the results of the VRANCEA2001 seismic
refraction project, which covers the 450 km long
segment from the town of Aiud in the Transylvanian
Basin to the town of Tulcea near the Black Sea (Fig. 1).
2. Geological and tectonic setting
The Eastern Carpathians are part of the Alpine–
Carpathian orogenic belt which resulted from the
convergence and collision of several microplates with
the Eurasian plate during the closure of the Tethys
Ocean (Sandulescu, 1984; Stampfli et al., 1998;
Neugebauer et al., 2001). In this context several tectonic
units have been accreted in the Carpathian area (Fig. 1).
The External (Flysch) Carpathians or Moldavides in the
east originated along the European (Moesian and East
European) margin (Sandulescu, 1988; see Fig. 1 for
location). Their convergence and the outline of the
present day geometry occurred mainly during the
Miocene (20–11 Ma) and climaxed in the early Late
Miocene (11–12 Ma, Sarmatian; Sandulescu, 1988;
Matenco and Bertotti, 2000; Matenco et al., 2003).
However, the Internal Carpathians further west (i.e.
Dacides and Transylvanides of Sandulescu, 1988; see
Fig. 1), which are part of the Tisza-Dacia microplate
(Csontos, 1995), converged already in mid-Cretaceous
time (Aptian/Albian), were then partially covered by
smaller basins with mid-Cretaceous (Aptian/Albian) to
Palaeogene sedimentary strata and were finally transported to their present position during the Sarmatian
accretion of the Moldavides against Europe (Sandulescu, 1988; Royden, 1988; Matenco and Bertotti,
2000). Their sediments accumulated on a thinned
continental (Sandulescu, 1988) or an oceanic domain
(Csontos, 1995; Stampfli et al., 1998; Neugebauer et al.,
2001) of Jurassic to Early Cretaceous age.
The geodynamic setting of the Sarmatian deformation is inadequately understood, but it appears that
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Fig. 1. Geological overview of the Eastern Carpathian bend area and its foreland with the main crustal units, nappe structures, faults and basins. The location of the VRANCEA'99 (N–S, small E–W
transverse) and VRANCEA2001 NE–SW seismic refraction profiles are shown with their shot points. Compiled from various sources given in the text.
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F. Hauser et al. / Tectonophysics 430 (2007) 1–25
subduction with slab break-off and lithospheric delamination during closure of the eastern prolongation of
the Penninic Ocean played a major role (Radulescu and
Sandulescu, 1973; Constantinescu et al., 1973; Airinei,
1977; Fuchs et al., 1979; Constantinescu and Enescu,
1984; Oncescu, 1984; Csontos, 1995; Girbacea and
Frisch, 1998; Mason et al., 1998; Seghedi et al., 1998).
Recently, Sperner et al. (2001, 2002) suggested a model
of Miocene subduction of oceanic lithosphere beneath
the Carpathian arc and subsequent “soft” continental
collision, which transported denser lithospheric material
into the mantle. Cloetingh et al. (2004) proposed postcollisional processes after subduction to explain the
actual geodynamic scenario, which are controlled by the
thermo-mechanical heterogeneities within the underthrusted lithosphere and lateral variations in the
interplay between the lithosphere and surface processes.
This subduction and collision connected the Tisza-Dacia
microplate to the Moesian and East-European plates and
a proposed steep Miocene suture zone, which is
concealed by the overthrusted Eastern Carpathian flysch
nappes, would separate both lithospheric domains from
each other (Sandulescu, 1988; Girbacea and Frisch,
1998; Sperner et al., 2001).
The External Eastern Carpathian fold- and thrust-belt
consists of a complex pile of nappes (Fig. 1; Sandulescu,
1984, 1988; Ellouz et al., 1994; Badescu, 1998; Zweigel
et al., 1998; Matenco and Bertotti, 2000). They are
made-up of Cretaceous marine basinal sediments and of
Paleogene to Neogene flysch and Neogene molasse
deposits. The outer nappes contain also Neogene
evaporitic formations with salt and/or gypsum. The
Internal East Carpathians are dominated by crystalline
rocks and a cover of Late Palaeozoic–Cretaceous
mostly marine sediments and Early Cretaceous flysch
deposits. The Moldavidian nappe pile has an estimated
thickness of 8–10 km (Stefanescu and Working Group,
1985; Ellouz et al., 1994; Morley, 1996; Matenco and
Fig. 2. Topographic map showing the VRANCEA2001 seismic line (thick line), as well as older parallel or transecting refraction and reflection lines
(thin and dashed lines). For detailed information and references see text.
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Bertotti, 2000). This thickness is constrained by surface
geology, seismic reflection and borehole data, mainly
from oil exploration, but for the inner nappes thicknesses are only estimated through balanced crosssections. Deeper structures are fairly unknown and
even the International Geotraverse XI (Figs. 2 and 3),
which crossed the Carpathians between Focsani and
Targu Secuiesc parallel to the VRANCEA2001 line,
gave only some vague indications (Radulescu et al.,
1976). Magnetotelluric data from Stanica and Stanica
(1998) indicate a depth of 8 km for the base of the
Moldavidian nappes and a depth of about 14–16 km for
the basement in the same area. Below the more internal
parts of the East Carpathians the magnetotelluric data
show the crystalline basement at 16–18 km and north of
the town of Focsani at approximately 10 km depth. The
crustal thickness reaches 50 km in both areas.
The Carpathian foreland is underlain by Precambrian
crust with a relatively undeformed Palaeozoic–Mesozoic
autochthonous platform cover and by deformed rocks of
the Triassic–Jurassic North Dobrogea Orogen (Fig. 1;
Sandulescu, 1984; Visarion et al., 1988; Tari et al., 1997;
Seghedi, 1998; Matenco et al., 2003). The crystalline
basement of the Moesian Platform is made-up of metamorphic and intrusive magmatic rocks, which sometimes
have a weak acoustic contrast to their older sedimentary
cover (Raileanu et al., 1994). Near the orogenic front of the
Carpathians the platform sediments are partly covered by
the foredeep sediments, which dip slightly towards the
central part of the foredeep (e.g. Tarapoanca et al., 2003).
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The Moesian Platform to the south and the Scythian
Platform to the north are characterised by distinct magnetic
anomalies, which result from petrological differences in
the crystalline basement, and by lithological differences of
the detritic and carbonaceous platform cover (Tari et al.,
1997; Seghedi, 1998 and references therein; Raileanu
et al., 2005). The major platform structures in the foreland
are Late Permian/Early Triassic rifts (Tari et al., 1997;
Seghedi, 1998; Landes et al., 2004; Raileanu et al., 2005;
Panea et al., 2005; Bocin et al., 2005). The deformed rocks
of the North Dobrogea Orogen include a complex polydeformed Variscan basement and a Permian–Cretaceous
sedimentary and volcanic cover (Seghedi, 1998). The
whole complex was overthrusted NNE-ward onto the
Scythian Platform between the Late Triassic and the Late
Jurassic.
The Scythian, Moesian and North Dobrogea crustal
blocks are thought to belong to the southeastern prolongation of the Trans-European suture zone (TESZ; e.g.
Pharao, 1999; Debacker et al., 2005). They are separated
by the Trotus (TF) and Peceneaga–Camena (PCF) faults
(Fig. 1). On the Moesian Platform the Capidava–Ovidiu
(COF) and Intramoesian (IMF) faults separate basements
of different composition. The COF separates a greenschist
basement to the north from a higher-grade metamorphic
basement to the south (Seghedi, 1998 and references
therein). The PCF and the COF are outcropping in the
Dobrogea area near the Black Sea, while sediments and
nappes cover the supposed NW-prolongation of the faults
as well as the IMF itself (Fig. 1; Visarion et al., 1988;
Fig. 3. International Geotraverse GT XI crustal seismic refraction line (Radulescu et al., 1976; Cornea et al., 1981; slightly modified; 1 — sediments,
2 — grantic layer, 3 — basaltic layer, 4 — boundary between sediments and basement, 5 — Conrad discontinuity, 6 — Moho, 7 — Eastern
Carpathian nappes, 8 — volcanic rocks, 9 — faults, 10 — crustal and subcrustal earthquake foci). For location see Fig. 2.
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F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Polonic, 1996; Ellouz et al., 1994; Seghedi, 1998). In the
autochthonous and overthrusted areas of the Moesian
Platform the seismic refraction line is expected to cross
two of these major crustal faults, the Capidava–Ovidiu
Fault (COF) and the Peceneaga–Camena Fault (PC).
Recent field studies show, that the Trotus/Peceneaga–
Fault system is a particularly active structure with a
pronounced tectonic geomorphology and local offsets of
200 m, which were achieved during the Quaternary
(Matenco et al., 2003; Tarapoanca et al., 2003). Based
mainly on reflection seismic data, these crustal faults are
thought to extend down to the Moho discontinuity and
possibly even deeper (e.g. Visarion et al., 1988;
Radulescu and Diaconescu, 1998).
Post-collisional (i.e. post-Sarmatian) sediments overlie
structures of the Carpathian Orogen and its foreland but
are also deformed by younger structures. The most
important of these structures is the Focsani Basin (Fig. 1).
The sediments of the foredeep and the Focsani depression
reach up to 8 km for the last 14 Ma (Matenco et al., 2003;
Tarapoanca et al., 2003). Late Pliocene to Early
Quaternary contractional deformation in the foreland
reaches as far as the Pericarpathian Front (Hippolyte and
Sandulescu, 1996; Hippolyte et al., 1999; Matenco and
Bertotti, 2000; Fig. 1). Along the western flank of the
Focsani Basin an eastward dipping monocline developed
from the Late Pliocene to the Holocene (Tarapoanca et al.,
2003; Cloetingh et al., 2003; Necea et al., 2005). Apatite
fission-track data show that strong uplift and exhumation
occurred in the Eastern Carpathians during the Pliocene
(2–7 Ma; Sanders et al., 1999). Ongoing vertical crustal
movements are revealed by geodetic measurements
(Popescu and Dragoescu, 1987; Radulescu et al., 1998;
Zugravescu et al., 1998).
The inner part of the Carpathians, which is partially
covered by sediments of the Tertiary Transylvanian
Basin (Fig. 1; Paleogene to Late Miocene in the central
and northern part and mainly Middle to Late Miocene in
the southern part) was also affected by some Neogene to
Quaternary compressional deformation (Ciulavu et al.,
2000). However, along its eastern and southeastern
margin the Transylvanian Basin is made-up of the postcollisional Late Miocene to Quaternary calc-alkaline
volcanic Calimani–Gurghiu–Harghita (CGH) mountains (Szakács and Seghedi, 1995, 1996; Seghedi
et al., 1998; Mason et al., 1998) and the Latest Pliocene
to Quaternary alkalic–basaltic volcanism in the adjacent
Persani mountains (Seghedi and Szakács, 1994;
Downes et al., 1995; Panaiotu et al., 2004), which are
structurally connected to the sinistral–transtensional
Gheorgheni–Ciuc–Brasov graben system (Fielitz and
Seghedi, 2005 and references therein; Fig. 1).
Fig. 4. Pre-2001 geological section along the main WNW–ESE VRANCEA2001 seismic-refraction line. The section is mostly transverse to the trend
of the main geological structures except for the basement structures of the foreland, which are highly oblique to the seismic line. For location see Fig. 1.
Compiled from various sources. The near-surface geology is based on mapped exposures as published in geologic cross-sections by Stefanescu and
Working Group (1985) and Matenco and Bertotti (2000). The extensional structures of the inner Eastern Carpathians (Brasov basin system) are as
interpreted by the authors. The Moho depth is from Radulescu et al. (1976). The intracrustal thrust, alternative base of the Moho (dashed line) and the
low velocity zone (L.V.Z.) are from Hauser et al. (2001). The proposed steep Miocene suture zone separating the lithospheric domains of Tisza-Dacia
and Moesia is from Sandulescu (1988), Girbacea and Frisch (1998) and Sperner et al. (2001). Neogene to Quaternary andesites are tentatively projected
down-strike the East Carpathian Calimani–Gurghiu–Harghita volcanic chain into the cross-section to indicate the possibility of occurrence of such
rocks in the crust. The Neogene filling of the Transylvanian Basin is based on Ciulavu et al. (2000).
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Using published as well as unpublished geological
data, a crustal cross-section for the VRANCEA2001
refraction line was compiled (Fig. 4). The section is mostly
transverse to the trend of the main geological structures
except for the basement structures of the foreland, which
are highly oblique to the seismic line. The near-surface
geology is based on mapped exposures as published in
geological cross-sections by Stefanescu and Working
Group (1985) and by Matenco and Bertotti (2000). The
extensional structures of the inner Eastern Carpathians
(Brasov basin system) are as interpreted by the authors.
The Moho depth is from Radulescu et al. (1976). The
intracrustal thrust, alternative base of the Moho (dashed
line) and the low velocity zone (L.V.Z.) are from Hauser
et al. (2001). The proposed steep Miocene suture zone
separating the lithospheric domains of Tisza-Dacia and
Moesia is from Sandulescu (1988), Girbacea and Frisch
(1998) and Sperner et al. (2001). Neogene to Quaternary
andesites are tentatively projected down-strike the East
Carpathian Calimani–Gurghiu–Harghita volcanic chain
into the cross-section to indicate the possibility of
occurrence of such rocks in the crust. The Neogene filling
of the Transylvanian Basin is based on Ciulavu et al.
(2000). This cross-section, which is simplified and adapted
to the scale and resolution of the seismic line, forms the
base for a first geological interpretation of the velocitydepth model obtained from the seismic experiment as
presented in this paper. It can later be integrated into a
complex larger scale geodynamic model using additional
geological and geophysical (e.g. mantle tomography) data.
This is however not the focus of this paper.
3. Earlier seismic investigations
Parts of the study area have been investigated by
different geophysical methods during the last decades of
the past century. Besides gravity, magnetic, magnetotelluric and heat flow measurements, seismic reflection
and refraction data played an important role. Here we
summarize the results, which are relevant for the
VRANCEA2001 seismic refraction project.
On the Moesian Platform and in the Focsani and
Transylvanian basin areas the VRANCEA2001 seismic
line is either intersected by or sub-parallel to nearby seismic
reflection profiles recorded for oil and gas exploration.
These seismic lines mainly mapped the structure of the
Neogene cover and to a lesser extent the pre-Neogene
sedimentary succession and the crystalline basement.
A representative deep seismic reflection line across
the Moesian Platform extends from Ramnicu Sarat some
40 km towards the east (RmS in Fig. 2). It shows a rapid
thickening of the Neogene sediments from about 8 km at
7
the eastern end of the line to 12 km near Ramnicu Sarat
(Raileanu and Diaconescu, 1998). The interpretation of
deeper-crustal features defines the crystalline basement
at 16 km, an intra-crustal boundary at 32 km and the
Moho at about 42 km depth (Raileanu and Diaconescu,
1998).
Within the Transylvanian Basin two deep reflection
lines cross the basin from NNW to SSE and from W to E,
respectively and both lines end near the VRANCEA2001
profile (Tr1 and Tr2 in Fig. 2; Raileanu and Diaconescu,
1998; Raileanu, 1998). The sedimentary sequence is
composed of a pre-Cretaceous succession, a post-tectonic
Late Cretaceous to Palaeogene cover and a Neogene
succession with a total thickness of 9–10 km. The
crystalline crust is ‘transparent’ and no significant
reflections are observed except for some short, dipping
signals with weak coherency and most probably of
diffractive nature. Sometimes, and mainly within the
central part of the basin, the basement is marked by
reflections. Structures are observed between 9 and 10 s
two way time (TWT) or approximately 27–30 km depth
and 12–14 s TWT or approximately 36–42 km depth.
These reflections are interpreted as a transition zone from
crust to mantle, between approximately 27 km and 42 km
(Raileanu and Diaconescu, 1998; Raileanu, 1998).
Based on seismic reflection and well data, Polonic
(1998) produced a crystalline basement map of
Romania. This map predicts a dramatic deepening of
the basement along the VRANCEA2001 seismic line,
from 1–3 km in the Dobrogea area to 15 km in the
Focsani Basin, 8–10 km in the Carpathian Orogen area
and 5–8 km in the Transylvanian Basin.
Previous crustal refraction lines recorded within the
North Dobrogea (Pompilian et al., 1993), the Focsani
Basin (Enescu et al., 1972), the Moesian Platform and the
Focsani and Transylvanian basins (Radulescu et al., 1976;
Cornea et al., 1981) give further information on the deep
crustal structures. Within the North Dobrogea a 6 km thick
sedimentary cover is observed and the Moho reaches a
depth of about 42–43 km (Pompilian et al., 1993). The
crustal refraction line GT XI 1 (Fig. 2), recorded in the
1970's from Focsani southwards over a length of 60–
70 km, points out two sedimentary boundaries at 5 km and
10 km depth, respectively, a basement at approximately
17–18 km (K0 boundary), and an intra-crustal boundary
(K1) at 26 km (Enescu et al., 1972).
Another crustal seismic refraction line GT II (Fig. 2)
runs from the territory of Ukraine and the Moldavian
Republic through Galati with a SSW direction to
Calarasi near the Bulgarian border. It crosses the
Peceneaga–Camena Fault south of Braila (Fig. 1),
where a large offset at all crustal levels is recorded. The
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F. Hauser et al. / Tectonophysics 430 (2007) 1–25
interface K1 deepens from about 15 km south of the
fault to about 21 km to the north of it and the Moho
deepens from 40 km to 48 km (Radulescu et al., 1976;
Cornea et al., 1981).
The most important crustal seismic refraction line GT
XI (Figs. 2 and 3) was recorded in the first half of the
1970's (Radulescu et al., 1976). It runs nearly parallel to
the VRANCEA2001 profile at a distance of about 20 km
north of Braila and of about 50 km north of Medias. It
shows the following main features. In the Galati area the
crust has a total thickness of 42–43 km with an upper
crustal thickness of 20–22 km. In the Focsani Basin the
sedimentary cover reaches a depth of 20 km. A crustal
fault within the centre of the basin separates, at
basement level, an uplifted block to the east from a
subsided one to the west. An intra-crustal boundary (K1)
that delimits the upper and lower crust is located at
32 km depth to the west and approximately 27 km to the
east of the fault. The Moho is located at 47 km to the
west and at approximately 42 km to the east of the fault.
Further to the west the crustal boundaries are shallower.
In the center of the Eastern Carpathians the K0-boundary
lies at 7 km, K1 at 20 km and the Moho at 40 km depth.
In the Transylvanian Basin K0 is at 3 km, K1 at 12–
18 km and the Moho at 30 km (Radulescu et al., 1976).
Based on data collected before 1999 Radulescu
(1988) and Enescu et al. (1992) produced contour-maps
of the two major boundaries (Conrad and Moho
discontinuities) in the crust. Along the VRANCEA2001
line depth values of 22 km and 45 km are predicted for
the Dobrogea area, 26–28 km and 42 km for the Focsani
Basin, 24 km and 52 km for the Carpathian Orogen, and
14 km and 34 km for the Transylvanian Basin. A crosssection parallel to the VRANCEA2001 line with up to
scale projections of the Geotraverse XI, RmS and Tr1
lines is shown in Fig. 5 of Knapp et al. (2005). For the
Geotraverse XI see also Fig. 3.
The most recent refraction seismic line is the
VRANCEA'99 line, the first seismic line of the present
joint research project within the German–Romanian
program mentioned above. This VRANCEA'99 line
intersects the VRANCEA2001 profile in the East
Carpathian mountain area at the centre of the seismogenic “Vrancea area” (Figs. 1 and 4). Its interpretation
shows a multi-layered structure (Hauser et al., 2001) and
at the intersecting point it depicts 3 sedimentary layers,
an upper and a lower crustal layer. The base of the
sedimentary succession is at 11 km, the intra-crustal
boundary at 29–30 km and the Moho at about 41 km
depth. The P-wave velocities for the sedimentary
successions are 3.90–5.70 km/s, for the crystalline
crust 5.90–7.00 km/s and for the top of the upper mantle
7.90 km/s.
4. The VRANCEA2001 seismic experiment
In August/September 2001 a large seismic experiment, called VRANCEA2001, was performed. After the
Fig. 5. Trace normalised P-wave record section from shot point O with 6 km/s reduced time. The calculated travel times from the Vp model in Fig. 11
are superimposed on the data. Travel times are labelled as follows: P1–P5 = first arrival phases refracted within the sedimentary cover; Pg1 and Pg2 =
diving waves through the upper and middle crust, respectively; Pi1P = reflected waves from the top of the middle crust; Pi2P = reflected waves from
the top of the lower crust; P6P and P8P = reflected waves from the base of the low velocity layers L6 and L8; PmP = reflected wave from the crustmantle boundary (Moho); Pn = diving wave through the upper mantle; Diff. = supposedly diffracted waves.
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
9
Fig. 6. Trace normalised P-wave record section from shot point S. For further explanations see Fig. 5.
completion of the first north–south running VRANCEA'99 line, briefly described above, it was the second
experiment to target the deep structure of the seismogenic Vrancea area with the aim of testing new and
existing geodynamic models.
The experiment recorded refraction/wide-angle reflection seismic data along a 700 km line from east of
Tulcea through the Vrancea zone to Karcag in Hungary
(see Figs. 1 and 2). Ten large drill hole shots (300–
1500 kg charge) were fired in Romania between Aiud at
the western margin of the Transylvanian Basin and the
Black Sea which resulted in an average shot point
spacing of 40 km. To the west an additional shot
(500 kg) was fired in Hungary. These shots generated 11
seismogram sections recorded by almost 800 geophones. The spacing of the geophones was variable. It
was around 1 km from the eastern end to Aiud (ca.
450 km length), 6 km from Aiud to Oradea (Romanian–
Hungarian border) and about 2 km on the Hungarian
territory. Between shot points T and U the geophones
were deployed at a spacing of 100 m. In total, 790
recording instruments were available. The 640 onecomponent geophones (TEXAN type) were mostly
deployed in the open field outside of localities, while for
security reasons the 150 three-components geophones
(REFTEK and PDAS type) were deployed in guarded
Fig. 7. Trace normalised P-wave record section from shot point T. For further explanations see Fig. 5.
10
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Fig. 8. Trace normalised P-wave record section from shot point W. For further explanations see Fig. 5.
properties within towns and villages. The experiment
was jointly performed by research institutes and
universities from Germany, Romania, the Netherlands,
Hungary and the United States. They included the
University of Karlsruhe and the GeoForschungszentrum
Potsdam, Germany, the Free University of Amsterdam,
Netherlands, the National Institute for Earth Physics and
the University of Bucharest, Romania, the Eötvös
Loránd Geophysical Institute in Budapest, Hungary
and the Universities of Texas at El Paso and South
Carolina, USA. Field recording instruments were
provided by University of Texas at El Paso and the
IRIS/PASSCAL instrument pool, USA (TEXAN 1component stations), and the GeoForschungsZentrum,
Germany (3-component REFTEK and PDAS stations).
In this paper we only deal with the 450 km long main
line extending from the Transylvanian Basin to the
Black Sea.
a point along the refraction model with respect to shot
point Z.
Due the complex geological structures of the
Carpathian Orogen, the data quality of the experiment
is quite variable (Tables 1 and 2) but seems to depend
primarily on the charge size. It was also influenced by
local geological conditions around the shot points, the
receivers, as well as by propagation conditions between
the source and the receivers. Nevertheless, the shots
Table 1
Summary of data quality as a function of offset for each shot point
No.
Shot
point
Branch
1
Z
2
Y
3
X
4
W
5
U
6
T
7
S
8
R
9
P
10
O
5. Seismic sections
The seismic record sections were compiled and
plotted using the SeismicHandler program package of
Stammler (1994). All seismic sections presented in this
paper are displayed with a reduction velocity of 6 km/s
(Figs. 5–8). Seismograms are normalized with respect to
the maximum amplitude per trace. A general bandpass
filter between 3 and 12 Hz was applied to the data to
improve the signal-to-noise ratio.
For this paper the term “travel time” refers to the
reduced travel time of the seismic section. The term
“offset” is used for the distance between the source and
the receivers, while the term “distance” is used to locate
Good
Acceptable
Poor
Distance
(km)
Distance
(km)
Distance
(km)
West
East
West
East
West
East
West
East
West
East
–
0–140
0–70
0–95
0–40
0–30
0–100
0–90
0–50
0–40
–
West
East
West
East
West
East
West
East
West
East
0–95
0–40
0–105
0–50
0–105
0–120
0–70
0–70
0–95
0–30
–
–
105–215
50–130
–
–
–
–
95–220
–
–
140–220
–
95–175
–
30–130
100–145
90–130
50–150
40–100 very poor;
170–220 poor PmP
95–135
40–85; 120–200
215–290
130–155
105–205
–
70–235
–
220–285
–
–
–
–
–
–
–
–
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
11
Table 2
Summary of the P-wave phase correlation and their relative quality
O–e
O–w
P–e
P–w
R–e
R–w
S–e
S–w
T–e
T–w
U–e
U–w
W–e
W–w
X–e
X–w
Y–e
Y–w
Z–e
P1
P2
P3
P4
P5
P6P
Pg1
Pi1P
P8P
Pg2
Pi2P
PmP
Pn
–
–
–
–
3
3
3
3
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
3
3
3
3
3
3
3
3
3
3
3
3
3
–
–
–
–
–
–
3
3
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
3
3
3
3
3–2
3
3
3–2
3
3
3
3
3
3
3
3
3
3
3
2–1
2
–
2
2
2
2–3
3–2
2
2
3
2–3
3
–
–
–
–
–
–
–
3–2
2
2–3
3
2–3
2–3
3
2
3
2
–
–
3
3
3
3–2
3
–
2
–
–
2–3
–
3–2
3
3
–
2
3–2
2
3
2–3
2–3
2–3
2
2
2–1
2
2–3
2
2–1
1
2
2–3
3–2
1
2
2
2
2–3
–
–
–
–
–
–
–
3–2
1
2–1
1
1–2
1–2
2
1
2
2
2
2
–
2
–
–
–
–
–
2
–
–
–
–
–
2
–
–
1–2
–
3
1
2
2
2
2
1–2
2
2
1
2–1
–
2–1
2
2–1
–
2–1
2
1
2
1
2–3
1
1
1
–
2
3
1
2–3
1
2–1
2
1
1
1
1
1
2–1
–
1
–
–
–
–
–
3
1
–
–
–
–
–
–
–
–
–
2
Correlation ranks: 3—easy correlation, 2—difficult correlation, 1—very difficult correlation. Combination of two numbers means intermediate
assessments between two ranks. (–): phase is not present or the high noise did not allow any correlation.
from the eastern half of the line were more effective then
the shots from the western half. Especially shot point S
(300 kg charge) generated coherent signals up to the
western end of the line, although some secondary
arrivals are suspected to be local diffractions. Some
seismic sections show segments of good or acceptable
data (i.e. clean and easy to correlate traces), followed by a
segment of noisy traces (Figs. 5–9). This is probably due
to the local noise levels during the recordings. Higher
quality data seem to be located below the Carpathians
and lower quality data below the Transylvania Basin
and partially the Focsani Basin and North Dobrogea
Orogen. Table 1 summarizes the data quality as a function
of the offset.
Not all phases could be identified on all record sections, sometimes because the phase is not present, but
sometimes also due to low signal-to-noise ratio (Table 2).
We picked only those phases, which are coherent over
several traces. Up to five first-arrival refracted phases
with apparent velocities b 6 km/s could be identified.
Fig. 9. Trace normalised P-wave record section from shot point Z. For further explanations see Fig. 5.
12
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
find the model that fits the seismic data best we used
several major steps in the modelling procedure:
They have been labelled P1 to P5 and are attributed to
waves travelling through different sedimentary layers or
intercalations of sedimentary, metamorphic, and volcanic layers (Figs. 5–9). Most of the data show a clear first
arrival Pg1 phase (a diving wave through the upper crust)
from 20 km to over 100 km offset (Figs. 5 and 7–9). This
phase is characterised by strong undulations and sometimes very small amplitudes, making picking generally
difficult beyond 80 km offset. Especially in Transylvania
(shot points Z–W) this phase shows very high and
variable apparent velocities (6.0–6.2 km/s) and appears
to be split into two phases Pg1 and Pg2 (e.g. Fig. 9),
separating a low velocity layer which generated the
reflected phase P6P on its base (Figs. 6–8). A Pg2 phase
is also observed in the eastern part, but only for shot
points O and S (Figs. 5 and 6). A reflected Pi1P phase
from the top of a low velocity layer in the western half
and from an intracrustal boundary in the eastern half of
the seismic line is correlated almost along the whole line
(Figs. 5–9). Another reflected phase P8P was generated
at the base of a second low velocity layer within the
western half of the seismic line (Figs. 6–9).
Within the middle crust we could identify secondary
Pi2P arrivals (the reflection from the top of the lower
crust (Figs. 5–9), between 30 and 140 km offset. The
phase is prominent and laterally coherent beyond
100 km offsets, but weak and hardly visible at subcritical offsets. Critical offsets for Pi2P ranges from
120–130 km for shots that recorded data at both ends of
the line to 90–110 km within the central part. The
reflection from the crust-mantle boundary (PmP) is well
observed on several record sections, especially from the
larger shots at both ends of the profile and from shot
point S (Figs. 5–9). The critical offset for this phase
varies from ca. 90 km in the west to 120 km in the east
and an apparent velocity of 6.8–6.9 km/s is observed at
larger offsets. A Pn diving wave through the upper
mantle can only be observed for a few shot points (Figs.
5, 6 and 9; Table 2). If present, Pn can be seen between
140 and 230 km offset. The apparent velocities lie
between 7.9 and 8.1 km/s. Table 2 shows the correlated
phases and their relative quality on a scale from 1 (low)
to 3 (high).
In some cases certain strong and coherent signals are
displayed on seismograms which might be interpreted as
diffractions (Fig. 5, 6, 8, 9). Sometimes those diffracted
waves overlap with useful signals.
An initial 2D velocity model was obtained by
Hauser et al. (2002) by picking first arrival times only
and inverting them using the non-linear tomographic
technique of Hole (1992). The most prominent result of
this inversion is seen in the center of the model
extending to a depth of some 20 km. It is associated
with the Carpathian Orogen and the Focsani Basin (see
Fig. 3 in Hauser et al., 2002). Since reflections from the
crust-mantle boundary and from an intra-crustal layer
are visible on most sections, this information was used
to construct initial 1D models for individual shot
points.
Both informations were then used to construct a 2D
starting model for ray-tracing, making due allowance for
the offsets in the different phases. Next, the data were
interpreted by 2D forward ray-tracing. For this we used
the RAYINVR program, which is based on the method
of Zelt and Smith (1992) and which allows the inclusion
of reflected phases, therefore using more of the information present in the seismic wave field.
In the ray-tracing procedure, travel times are
successively re-calculated for a sequence of models
which are constructed by the interpreter. In the initial
modelling, the uppermost velocity structure is determined using the appropriate seismic phases. This is
followed by stepwise modelling of phases from deeper
layers, usually in a sequence in which the velocity is
determined from the refracted phase, after which the
thickness is determined by the relevant reflected phase.
Usually only super-critical reflections are strong enough
to be identified in the seismic record sections. This
procedure is continued until a reasonable fit is achieved,
as judged from the assessed uncertainty of each
modelled seismic phase.
6. Interpretation techniques and velocity model
6.1. Uncertainties
For the interpretation of the P-wave seismic data only
vertical component seismograms were used. In order to
A total of 13 seismic phases were correlated and
picked (see Table 2), which already indicates the
(1) Travel times and associated errors were picked for
each of the seismic phases described above. The
integrity of the picks and the consistency of the
phase identification were checked by comparing
reciprocal travel times where possible.
(2) One-dimensional (1D) and two-dimensional (2D)
velocity-depth functions were calculated using
tomographic inversion methods as well as raytracing techniques.
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
complexity of the area. The pick uncertainties of the
seismic phase arrivals were estimated to ± 0.1 s for all
phases, except PmP and Pn which were estimated as
13
± 0.15 s. At the end of the forward modelling, when a
reasonable agreement between observed and calculated
travel times was reached, an inversion of all 5142 picked
Fig. 10. (a) Ray coverage through the final model connecting all shot and receiver pairs for all phases (upper figure) and only for Pi2P, PmP and Pn
(lower figure). Small black dots indicate the shot points named as in Figs. 1, 11 and 12. (b) Comparison of observed and calculated travel times for all
shots and all phases. Vertical bars indicate observed data with height representing pick uncertainties: ±0.10 s for all phases except PmP and Pn with
±0.15 s. Solid lines indicate calculated travel times.
14
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
data points was carried out. The final model gave a rootmean-square (RMS) travel time residual of ± 0.117 s and
a normalized χ2 = 1.244. The largest deviations from
χ2 = 1 were observed for P4, P5, Pg1 and Pn with values
between 2 and 5. A ray coverage through the final model
for all shot and receiver pairs is presented in Fig. 10a,
and a comparison of the observed and calculated travel
times for all shots and all correlated phases in Fig. 10b.
From this we estimate the uncertainties for the P-wave
velocity from ± 0.1 km/s for the upper and middle crust
to ± 0.15–0.20 km/s for the lower crust and upper
mantle. For the interfaces we estimate depth uncertainties
from ±0.5–1.0 km for shallow layers of the model to
±1.0–1.5 km for deeper interfaces. Based on all above
data the model interfaces were drawn with thick solid
lines for well resolved regions, and with dashed lines for
less well resolved regions (Fig. 11).
6.2. The 2D velocity model
The final 2D velocity model derived using the
methods described above is shown in Fig. 11. It has a
multi-layered character and reflects the different tectonic
units (Transylvanian Basin, Eastern Carpathian Orogen,
Moesian Platform including the Focsani Basin and
North Dobrogea Orogen), which are crossed by the
seismic line. The main horizontal structures can be
separated into two different groups : (1) the sedimentary
cover with imbricated crystalline and volcanic rocks
showing velocities b 6 km/s; (2) the crystalline crust
down to the crust-mantle boundary (Moho).
The sedimentary succession with imbricated crystalline and volcanic rocks along the profile consists of up
to seven layers (L1–L6 and L8 in Fig. 11) with
velocities ranging from 2.0 to 5.9 km/s. Within the
Fig. 11. VRANCEA2001–2D velocity-depth model. Labelled dots at the top of the model indicate the shot points from O to Z and the decimal
numbers show the P-wave velocities in km/s. The VR 99 arrow marks the intersection with the VRANCEA'99 seismic line and PCF the Peceneaga–
Camena Fault. L1 to L11 indicate the seismic layers used in the interpretation. Thick solid lines indicate areas which are well constrained by
reflections and/or refractions. Dashed lines indicate less well constrained areas, while thin lines are extrapolations (see also Table 2).
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Eastern Carpathian Orogen we find four layers (L2, L4,
L5 and L6) above the crystalline basement (L7 with
6.0 km/s) with variable thicknesses and increasing
velocities from 3.60 km/s at the surface to 5.70 km/s at
the base of the succession. The lowest layer (L6) is a
low-velocity zone with 5.30 km/s sandwiched between
higher-velocity layers. Further to the west and into the
Transylvanian Basin we find three layers (L2, L4 and
L5) with variable thicknesses and seismic velocities
from 3.00 km/s near the top to 5.80 km/s at the
sediment-basement interface.
To the east and into the Focsani Basin the sediment
thickness increases drastically to about 20–22 km. The
seismic velocities cover a wide range and increase from
2 km/s at the surface to 5.6 km/s in layer L5. The deepest
part of the Focsani Basin shows three layers (L6–L8)
with velocities between 5.1 and 5.9 km/s, which are
attributed to a succession of different sedimentary rocks
characterised by a high-velocity layer (L7) wedged in
between layers with lower velocities (L6, L8). The North
Dobrogea Orogen is covered by a thin wedge of
sediments, volcanics and imbricated basement (L5; 1–
3 km thick) with rather high velocities (5.00–5.90 km/s).
The seismic basement (L7–L10) coincides with a
depth where velocities exceed 5.9 km/s. The upper
crustal velocities are very heterogeneous and seem to
reflect different tectonic units. In addition, we find a
low-velocity layer (L8 with 5.50–6.00 km/s) within the
upper crystalline crust, which extends from the western
end of the seismic line to the Focsani Basin. A distinct
intra-crustal boundary separates the middle crust (L9;
6.1–6.5 km/s) from the lower crust (L10; 6.7–7.1 km/s)
with varying depth from 27 km at the western end to
29 km below the Focsani Basin and 27 km at the eastern
end of the seismic line.
Wide-angle Moho reflections (PmP) indicate the
existence of a first-order crust-mantle boundary (between L10 and L11). The Moho topography shows a
thickening of the crust from 37 km in the west to 42–
46 km below the Focsani Basin and 44 km in the North
Dobrogea Orogen area in the east. No pronounced
crustal root below the Carpathian Orogen is recognizable. Some constraints on upper mantle seismic velocities are provided by Pn arrival times picked from
several shot points.
7. Discussion and interpretation
The VRANCEA2001 seismic refraction model in
Figs. 11 and 12 demonstrates considerable lateral
thickness and velocity variations within the sedimentary
succession as well as in the deeper crust. Three different
15
crustal blocks characterised by clearly distinct geometries and velocity structures were identified (Fig. 12,
from west to east): (1) the Tisza-Dacia crustal block,
which underlies the Transylvanian Basin and most of the
Eastern Carpathian Orogen, (2) the Moesian Platform
crustal block, which underlies the Focsani Basin, and (3)
the North Dobrogea Orogen crustal block. The last two
tectonic units are separated by the Peceneaga–Camena
Fault while the contact between the first two units is less
well defined and concealed by the East Carpathian
nappes. A proposed steep Miocene suture zone is,
however, thought to separate both lithospheric domains
(Sandulescu, 1988; Girbacea and Frisch, 1998; Sperner
et al., 2001) and we therefore attribute the complete
crust with its crystalline and sediment-derived nappes to
the individual plates (Tisza-Dacia and Moesian crustal
blocks; details in Section 7.3.). Additional constraints
for the interpretation of the seismic model were
provided by seismic reflection data, borehole data for
which published velocity logs are available, and other
geophysical data.
7.1. The sedimentary sequence with imbricated
crystalline rocks
The sedimentary sequence is composed of the East
Carpathian flysch nappes, the Neogene infill of the
Transylvanian Basin in the hinterland and the Focsani
Basin in the foredeep, the autochthonous Mesozoic and
Palaeozoic sedimentary rocks of the Moesian Platform
and the North Dobrogea Orogen, and the deeper
sedimentary basins below the Transylvanian and the
Focsani basins (Fig. 12). The different outcropping
geological units are reflected by the laterally variable
velocity structure in the seismic model (Fig. 11). The
sedimentary sequence along the line, which also
comprises some imbricated crystalline rocks as specified
below, was subdivided into several geological units
(shown in Fig. 12 with different colours) made up of a
single or several seismic layers as identified in the
velocity model.
The first geological unit represents the western part
of seismic layer L2 and the upper part of seismic layer
L4 and extends from west of shot point Z to east of shot
point X (yellow unit in Fig. 12). The thickness of the
upper seismic layer L2 is nearly constant, at about 1 km,
and the velocities within this layer range from 3.0 to
3.3 km/s. In seismic layer L4 the upper approximately
up to 2 km of this geological unit with velocities of
3.9 km/s are not well separated from the next second
geological unit probably because of similar composition
and/or consolidation of the sedimentary rocks. Because
16
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Fig. 12. Interpreted geological cross-section (top: 4.5 × vertical exaggeration, bottom: without vertical exaggeration) from the 2D seismic model of
Fig. 11 along the main VRANCEA2001 seismic refraction line between the Transylvanian Basin and the Black Sea. The upper crustal geological
structures of the Tisza-Dacia and the Moesian crustal blocks are transverse to the section. The proposed out-of-sequence thrusting in the crystalline
basement (labeled with number 1) and the geologic structures of the North Dobrogea crustal block in the foreland (labeled with number 2) are oblique
to the seismic line. For location of the section and for location of the major geological structures compare with Fig. 1.
of their proximity to the surface and from reflection
seismic lines and boreholes from gas exploration
(Ciulavu, 1999; Ciulavu et al., 2000) they are however
known to exist. This geological unit with in total up to
3 km of thickness represents the Neogene cover of the
Transylvanian Basin (Fig. 12).
The second geological unit (seismic layer L4 between west of shot point Z and shot point Y in Fig. 11)
underlies geological unit 1 within the Transylvanian
Basin (green unit in Fig. 12). It has an asymmetric shape
with a steeper eastern flank and its thickness increases
from 0 km in the west to about 2 km in the east. The
velocity within this layer increases from 3.9 km/s at the
top to 4.2 km/s at the bottom. It represents the Tarnava
Basin, which has a half-graben geometry and is filled
with Late Cretaceous to Paleogene sediments and is
underlain by Early Cretaceous sediments and Jurassic
volcanics. It is known from reflection seismic lines and
boreholes from gas exploration (Ciulavu, 1999; Ciulavu
et al., 2000).
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
The third geological unit comprises seismic layers
L2 and L4 from west of shot point W to shot point T
(Fig. 11; brown unit in Fig. 12) and has a thickness of 4–
6 km. Its velocities range between 3.6 and 5.1 km/s with
strong lateral as well as vertical variations. This unit
represents the presumed unrooted sedimentary nappe
pile of the Carpathian Orogen made up mainly of
Triassic to Neogene rocks (mostly Cretaceous and
Tertiary flysch of the Moldavide and Outer Dacide
nappes; see Figs. 1 and 4).
The fourth geological unit (seismic layer L4 from
shot point Y to west of shot point W and layer L5 from
west of shot point Z to halfway between shot points U
and T in Fig. 11) underlies geological units 1, 2 and 3
(Transylvanian and Tarnava Basins and sedimentary
nappe pile of the Carpathian Orogen; uppermost part of
violet unit in Fig. 12). It reaches the surface between
shot points X and W, where it can be correlated to
surface outcrops. The thickness of this unit ranges
between 2 and 5 km with highly variable velocities
between 4.0 and 5.8 km/s reaching the maximum
velocity (5.4 to 5.8 km/s) at the bottom of layer L5 at a
depth of about 7 km. This unit represents imbricated
basement nappes of the Median Dacides (mainly
Bucovinian nappes; Sandulescu, 1984) of the Internal
and External Carpathian Orogen. It is composed of
metamorphic and other crystalline rocks. A heterogeneous composition with participation of lower-grade
metamorphic, possibly Palaeozoic rocks, and a localized
thin cover of autochthonous Palaeo and Mesozoic rocks
could explain the velocity variations, especially the low
velocities (b 4.9 km/s) within seismic layer L4. This
geological unit may, west of shot point Y, contain thin
thrust sheets of ophiolithic rocks from the Transylvanide
nappes as proposed in the section of Fig. 4. But
velocities point to a more generally crystalline composition of the crust. There is also no indication of the
proposed Miocene suture zone below the third geological unit (Eastern Carpathian flysch nappes) since
velocities seem to be laterally and horizontally relatively
continuous and lower than expected for mafic or
ultramafic ophiolitic rocks.
The fifth geological unit (seismic layer L6 in Fig. 11)
is an about 2 to 4 km thick low velocity zone with
velocities of 5.3 km/s (blue unit in Fig. 12). It underlies
geological unit 4 (basement nappes of the Carpathian
Orogen) and its western boundary near shot point Y is in
western down-dip prolongation of the overthrusted
basement nappes at the surface (see preview crosssection in Fig. 4, east of shot point X). Because of this
relationship it can be interpreted as the autochthonous
Palaeozoic and/or Mesozoic cover of the basement
17
overthrusted by the deeper, higher velocity metamorphic
crystalline basement of geological unit 4 (Fig. 12;
Sandulescu, 1984).
The sixth geological unit (seismic layers L1–L3 east
of shot point T and L4 below and east of shot point T in
Fig. 11; eastern yellow unit in Fig. 12) partially
underlies geological unit 3 (sedimentary nappe pile of
the External Carpathians). It has its greatest thickness of
about 10 km east of shot point T at the front of the
Carpathian nappes and thins continuously from shot
point S towards shot point R to only 1 km and
disappears further east. In its lower part (seismic layer
L4) it displays an asymmetric shape with a steep western
flank, which clearly separates this unit from the western
area, and a relatively gentle dipping eastern flank. In the
upper parts (seismic layers L1 to L3) its more
symmetrical shape correlates with data known from
surface geology and reflection seismic lines (Tarapoanca
et al., 2003). Velocities increase from 2.0 to 4.8 km/s at
the base of the unit and are laterally continuous. This
geological unit corresponds to the Middle Miocene to
Quaternary sedimentary fill of the Focsani Basin.
The seventh geological unit (seismic layer 5 in Fig. 11)
begins about halfway between shot points U and T and
can be followed along the base of geological unit 6
(Focsani Basin) until it reaches the surface between shot
points R and P (dark-blue unit in Fig. 12), from where it
stays at the surface to the eastern end of the section
(easternmost light blue unit in Fig. 12). Its thickness is
about 2–3 km, while slightly decreasing to about 1–
1.5 km east of shot point R. Velocities are in the range of
5.4–5.9 km/s with only minor lateral variations. The
deeper part of this unit (dark blue segment) probably
represents the autochthonous Mesozoic and maybe the
very thin Cenozoic sedimentary cover rocks of the
Moesian Platform below the Neogene Focsani Basin.
The high velocities would correlate with widely occurring
carbonate rocks in these layers (Tari et al., 1997).
Observations of Poisson's ratio along the VRANCEA'99
seismic refraction line seem to confirm this interpretation
(Raileanu et al., 2005). In the Dobrogea area, this unit is
also composed of Triassic volcanic rocks and imbricated
Palaeozoic sedimentary, magmatic or metamorphic rocks.
The whole layer in this crustal block probably represents
the NNE-ward overthrusted North Dobrogea Orogen
(light blue segment in Fig. 12).
The eighth geological unit (seismic layer 6 between
east of shot point T and shot point S in Fig. 11; orange
unit in Fig. 12) is a homogeneous low-velocity layer with
a thickness of about 5 km and velocities of 5.1 km/s. Its
isolated appearance with sharp lateral boundaries point
to a graben-like structure, while the lower velocities
18
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
indicate a more clastic sedimentary succession. Since it
is located below the proposed autochthonous Mesozoic
cover rocks of the Moesian Platform, we propose a
Permo-Triassic graben structure (Fig. 12). Similar
geological structures of the Moesian Platform in this
or nearby areas have already been described or
proposed by Tari et al. (1997), Landes et al. (2004),
Panea et al. (2005), Bocin et al. (2005) and Raileanu
et al. (2005).
The ninth geological unit (seismic layers L7 from
halfway between shot points U and T until shot point R
and seismic layer L8 between shot points W and S in
Fig. 11) has a relatively constant thickness of about 6 km
while its velocities range from 5.5 to 5.9 km/s (light blue
unit in Fig. 12). Its western part is located below the
Carpathian nappes at 14 km depth, its central part at
greater depth below the proposed Permo-Triassic
graben, and the continuation to the east would be
covered by Mesozoic platform rocks. Because this unit
underlies the above graben structure, we interpret it for
the central and eastern part as the autochthonous
Palaeozoic cover rocks of the Moesian Platform
(Fig. 12). The western part (west of halfway between
shot points U and T) could be made up of similar rocks
covering the eastern margin of the Tisza-Dacia block
with a slightly thinned middle and lower crust, which
were later overthrusted by crystalline Carpathian nappes
with higher velocities. The high velocities can, again, be
correlated in the Moesian domain with widely occurring
carbonate rocks in these layers (Tari et al., 1997) and
Poisson's ratio observations along the VRANCEA'99
seismic refraction line seem to confirm this as well
(Raileanu et al., 2005).
7.2. The structure of the crystalline crust
The tenth geological unit (seismic layer L7 west of
halfway between shot points U and T and east of shot
point R; seismic layer L8 west of halfway between shot
points W and U and seismic layers L9 and L10 in
Fig. 11) makes up the crystalline crust of the TiszaDacia, Moesian and North Dobrogea crustal blocks
(violet, pink and gray–blue units in Fig. 12). It shows
thickness variations related to the different crustal
blocks, while there are only small lateral velocity variations (especially in the middle crust) along the entire
seismic line. The velocities increase from 6.0 km/s at the
top of basement to approximately 7.0 km/s at the Moho.
The total thickness of the crystalline crust lies between 30 and 34 km for the western part of the model,
which corresponds to the Tisza-Dacia crustal block. This
block is characterised by basement thrusts in its upper
crustal layers down to 12–15 km depth as described in
the previous section. A low-velocity layer (L8 with
6.0 km/s in Fig. 11) between 11 and 15 km depth is
interpreted as being related to another intra-crustal
basement thrust connected to the Carpathian Orogen,
where a higher-velocity deeper crustal unit (L7 with 6.0–
6.2 km/s in Fig. 11) was thrusted over a lower-velocity
shallower crustal unit (Fig. 12). Between shot point W
and east of shot point U, the crystalline crust of the
seismic layer L7 and the western part of layer L8 cover
lower velocity rocks of the ninth geological unit
(Fig. 12). We interpret this again as crystalline basement
overthrusted on top of Palaeozoic sedimentary rocks
inside the Tisza-Dacia block. This structure is also seen
in the 3D crustal tomography model of Landes et al.
(2004) and has been correlated with a SSE-ward directed
Late Pliocene/Early Pleistocene out-of-sequence basement thrust (Landes et al., 2003; Fielitz and Seghedi,
2005). The middle crust comprises velocities of 6.3–
6.5 km/s at depth between 15 and 29 km, whereas the
lower crust with velocities of 6.8–7.0 km/s is thinner and
has a thickness of only between 8 and 10 km (Fig. 11). As
discussed later on (Section 7.3.), the middle and lower
crust of the Tisza-Dacia block have their eastern
boundary between shot points U and T (Fig. 12). And
like for the upper crust, there is also no indication of the
proposed Miocene suture zone in the middle and lower
crust, since velocities seem to be laterally and horizontally relatively continuous and lower than to be expected
for mafic or ultramafic ophiolitic rocks. The velocity
model gives also no indication of voluminous Neogene
to Quaternary volcanic rocks (basalts and andesites) as
tentatively proposed in the geological section of Fig. 4
for this part of the seismic profile. Small-volume dikes or
sub-volcanic intrusions are however still possible, but
cannot be resolved.
To the east, until shot point R, a clearly distinct
structure in the central part of the model is associated
with the Moesian crustal block. The total thickness of
the Moesian crystalline crust is 19 to 25 km. The only
7–9 km thick middle crustal layer with the top between
approximately 20 and 22 km depth shows velocities of
6.1–6.3 km/s, while the 9–16 km thick lower crustal
layer has velocities of 6.7–7.1 km/s. The eastern
boundary of this Moesian crustal block correlates well
with the down-dip prolongation of the Peceneaga–
Camena crustal fault (Figs. 2 and 12). As described in
Section 7.1, above the thinned middle to lower crust, the
Moesian block is characterised by alternating high- and
low-velocity layers (layers L1–L8 in Fig. 11), which we
interpret from bottom to top as a thick Palaeozoic autochthonous sedimentary cover, a Permo-Triassic
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
graben structure, a Mesozoic autochthonous sedimentary cover and a deep Middle Miocene to Quaternary
sedimentary basin. The superposition of these different
structures suggests a repeated reactivation of possibly
Palaeozoic or older major crustal discontiniuties within
the Moesian block. The verification of this geological
model, however, has to take into account the 3Dorientation of the individual structures, which cannot be
distinguished on this two-dimensional section. In
addition, especially the deeper structures, which do
not outcrop at the surface, are not yet clearly established.
East of the Peceneaga–Camena Fault, the North
Dobrogea crustal block shows a very distinct threelayered crystalline crust with a total thickness of 44 km.
The upper crust has velocities of 6.0–6.2 km/s (seismic
layer L7), while middle and lower crustal velocities
(seismic layer L9 and L10) range between 6.3–6.4 km/s
and 6.7–7.1 km/s, respectively. The thick North
Dobrogea crystalline crust is connected with the
Scythian Platform, which is a continuation of the East
European Platform further to the east and northeast, as
shown by large-scale tomographic data from this area
(Wortel and Spakman, 2000). We therefore suggest, that
the North Dobrogea crustal block is mainly composed of
crust from the Scythian Platform (see Fig. 1) and that the
uppermost layer of 1 to 2 km thickness represents the
central or frontal parts of the overthrusted wedge of the
North Dobrogea Orogen. Layers L7, L9 and L10 show
no lateral velocity variations between shot points R and
O. Therefore, an important deep reaching ultramafic
nappe, as proposed in the geological section of Fig. 4,
seems improbable. Additionally the velocity model in
this area is sub-parallel to the geological structures. For
this reason such a nappe would not be a steeply dipping
structure but a near-surface horizontal to shallow dipping
body. Velocities of 5.8 and 5.9 km/s in the thrust wedge
of seismic layer L5 could relate to such an ultramafic
nappe, whose down-dip continuation must however be
searched south and parallel to the actual seismic section.
7.3. Plate boundaries
As discussed in the introduction, the geodynamic
setting of the region covered by the VRANCEA2001
seismic line is thought to relate to the final stages of a
subduction process (e.g. Sandulescu, 1988; Csontos,
1995; Girbacea and Frisch, 1998; Sperner et al., 2001;
Cloetingh et al., 2004). This subduction involved the
upper Tisza-Dacia lithospheric plate, which was already
affected by important contractional deformation
(thrusts, nappes and a palaeosuture) related to an earlier
Early Cretaceous subduction and collisional event. The
19
composite lower plate originated by the accretion of
different lithospheric domains, which involved the
relatively undeformed but compositionally distinct
Moesian, Scythian and East European platform areas
as well as the Late Triassic to Late Jurassic North
Dobrogea Orogen with its Variscan basement, all
separated by important crustal faults. The collision and
climax of deformation between both plates took place in
the Middle to Late Miocene (Sarmatian) and resulted in
a steep suture zone thought to be concealed by the
overthrusted Eastern Carpathian flysch nappes, which
would represent the unrooted accretionary prism. This
model is represented in the geological section of Fig. 4.
The VRANCEA2001 profile can be subdivided into
three crustal domains with distinct characteristics
concerning thickness, composition, structuring and
geometry of the different crustal layers (Fig. 13).
The western domain, which we relate to the TiszaDacia plate, has the thinnest crust with a Moho depth of
37–33 km. The middle crust is significantly thicker than
the lower crust. The upper crust shows an alternation of
high and low velocity zones, which we interpret as
largely related to imbricated thrust sheets with alternating sedimentary and crystalline rocks. Much of this
deformation could already have been emplaced during
the earlier Early Cretaceous subduction and collisional
event with some reactivation and new deformation
during the later Sarmatian event. In this domain
Bouguer anomalies show negative values decreasing
across the Transylvanian Basin with the lowest values
due to thick sediments around shot point U (− 50 mgal)
and shot point Y (− 60 mgal; Visarion, 1998). Positive
values of the magnetic anomaly component ΔZ of up to
300 gamma dominate the centre of the Transylvanian
Basin, while negative values of up to −100 gamma were
observed between shot points X and R (Airinei et al.,
1985). These anomalies are related to different basement
compositions (greenschist basement with negative
anomalies and other basements with positive anomalies;
Airinei et al., 1985).
The central domain, which we relate to the Moesian
plate, has a thick crust with a Moho depth down to
45 km. Here the middle crust (with slightly lower
velocities than the adjacent plates) is thinner than the
lower crust. The upper crust is well layered, wholly
sedimentary and shows in its lower part also an
alternation of high and low velocity zones. Deformation
seems to be limited to Permo-Triassic extension and
Miocene to Quaternary subsidence in the Carpathian
foreland. In this domain the Bouguer anomalies show
negative values with a small minimum of − 85 mgal
around shot point T and an absolute minimum of
20
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
− 100 mgal between shot points S and T in the Focsani
Basin, where the sedimentary basin is deepest (Visarion,
1998). Variations of the magnetic anomaly component
ΔZ are similar to the western domain with negative
values of up to − 100 gamma (Airinei et al., 1985).
The eastern domain, which we relate to the North
Dobrogea Orogen, has a thick crust with a Moho depth of
44 km. Here the middle and lower crust are relatively
thick, the latter being somewhat thicker. The upper crust
is relatively thick, very homogeneous and crystalline,
except for the very thin uppermost thrust wedge of mixed
composition related to the Triassic to Jurassic deformation of the orogen. The thick crystalline crust would be
the continuation of the Scythian/East European platform.
In this domain the Bouguer anomalies show positive
values (Visarion, 1998) and the magnetic anomaly
component ΔZ rises to almost 200 gamma between
shot points R and P and to approximately 150 gamma at
the eastern end of the line (Airinei et al., 1985).
The boundaries between the three plates are steep and
relatively sharp (Figs. 12 and 13). The Peceneaga–
Camena Fault between the North Dobrogea and the
Moesian plate is a well known crustal discontinuity and
well defined from surface geology and geophysical data
(Visarion et al., 1988; Radulescu and Diaconescu, 1998;
Seghedi, 1998; Matenco et al., 2003; Tarapoanca et al.,
2003). The boundary between the Moesian and TiszaDacia plates is generally poorly constrained, because it is
concealed by the Sarmatian overthrusted flysch nappes
of the Eastern Carpathians (Figs. 4, 12 and 13). The
proposed Miocene suture (Sandulescu, 1988; Girbacea
and Frisch, 1998; Sperner et al., 2001) cannot be
identified in the VRANCEA2001 profile, neither in the
location shown in Fig. 4 nor further to the east. The only
sharp boundary separating two distinct crustal domains
is found between shot points U and T (Fig. 12).
Therefore, we tentatively interpret this crustal discontinuity to be the boundary between the Tisza-Dacia and the
Moesian plates. There is, however, no indication of a
suture zone, since crustal velocities do not point to mafic
or ultramafic rocks. The nature and detailed geometry of
this contact is not known, but an alternative could be a
crustal fault with lateral displacement, eventually a
transfer zone reactivating older (Permo-Triassic ?)
crustal structures. From the conventionally proposed
subduction models important horizontal displacements
between the Tisza-Dacia and Moesian plate would
generally be expected, also for deeper parts of the
21
crust. This cannot be confirmed from the presented plate
characteristics (Fig. 13). It has however to be taken into
account for the correlation, geometry and interpretation
of the presented data that the orientation and age of the
crustal structures change considerably between and
inside the involved plates (Tisza-Dacia with Early
Cretaceous and Sarmation deformation, Moesia with
Triassic–Permian and Miocene–Quaternary deformation, North Dobrogea with Variscan and Triassic–
Jurassic deformation, Scythian/East European platform
with Precambrian deformation) and that the 3Dorientation of the individual structures can be difficult
to distinguish in a two-dimensional section. Additionally, structures in the probably mostly Precambrian
crystalline crust might be concealed because possible
compositional and therefore structural differences might
not be shown by differences in the seismic velocities.
Also no obvious relation to the steep Vrancea seismic
body can be seen. This could be because of decoupling of
crustal and mantle processes.
The three plates are generally thought to belong to the
southeastern prolongation of the Trans-European suture
zone (TESZ; e.g. Pharao, 1999; Debacker et al., 2005) and
therefore the VRANCEA2001 profile also crosses this
major plate boundary. The overall geometry of the
presented velocity model shows a high degree of
similarity to the velocity models and seismic profiles
across the TESZ in Poland further to the northwest (e.g.
Jensen et al., 2002; Janik et al., 2002; Grad et al., 2002).
These similarities consist mainly in a thick three-layered
crust of the Precambrian Craton (42–45 km in Poland,
44 km in North Dobrogea), a thinner crust with a thin
(∼8 km) lower crust in the areas to the southwest (29–
32 km in the Palaeozoic terranes of Poland, 37–33 km in
the Tisza-Dacia terran) and the TESZ itself is covered by a
deep sedimentary basin with Permian origins or precursors (20 km thick Polish Basin, 22 km thick Focsani
Basin area). This suggests strongly a southeastward
prolongation of the TESZ structure into Romania along
the southwestern margin of the East European Precambrian craton. However, there are also several differences,
which have to be considered: The POLONAISE profiles
cross the TESZ perpendicular to their overall structures,
whereas the VRANCEA2001 profile is highly oblique to
it. In Poland the southwestern terranes belong to Avalonia,
which experienced Caledonian and Variscan deformation.
In Romania they belong to the Moesian terran with its still
poorly understood Palaeozoic evolution and to the North
Fig. 13. Geological characteristics and crustal thicknesses of the main crustal domains (plates) along the VRANCEA2001 seismic refraction line.
Steep boundaries between the deduced Tisza-Dacia, Moesia and North Dobrogea plates seem to be recognizable although important horizontal
displacements between the Tisza-Dacia and Moesian plate would generally be expected.
22
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
Dobrogea orogen, which experienced Variscan and
Triassic–Jurassic deformation. The Tisza-Dacia terran,
which makes up the whole western half of the
VRANCEA2001 profile and already experienced an
Early Cretaceous collisional event, is generally thought
to have collided with the TESZ only during the Miocene
and, separated by an oceanic domain, was formerly
located much farther to the west. Therefore its crustal
geometry cannot be compared easily with the Avalonia
terran. Also the younger basins show marked differences.
The Polish Basin is a Carboniferous–Permian and
Mesozoic structure that was inverted during the Late
Cretaceous and Early Tertiary. The Focsani Basin is
mainly a Late Cenozoic structure marginally affected by
Late Alpine deformation. It possibly had a Permo-Triassic
precursor basin, but its geometry and relation to the
overlying Cenozoic basin is only very poorly constrained.
In summary, the POLONAISE and VRANCEA2001
profils show globally many similarities, expecially due to
the contrast between the exceptionally thick East
European Precambrian crust and the thinner southwestern
accreted terranes. Timing of accretion and deformation of
these terranes might, however, be very different and
crustal and basinal similarities partly result only from the
mechanical differences between both crustal domains.
A more in-depth interpretation of the VRANCEA2001 is only possible in the context of a complex
larger scale geodynamic model using additional geological and geophysical (e.g. mantle tomography;
Wenzel et al., 1998b; Martin et al., 2005, 2006) data.
This is however not the focus of this paper.
8. Conclusions
A 700 km long WNW–ESE trending seismic refraction
line was carried out in Romania in order to study the
lithospheric structure. Here we present results from a subsection between the Transylvanian Basin across the SECarpathians to the Carpathian foreland areas. The
geophysical and geologic interpretation of the data by
forward and inverse modeling gave the following results:
The sedimentary succession can be subdivided into 7
layers with a total thickness of up to 22 km. It is
composed of (1) the Carpathian nappe pile, (2) the postcollisional (post-Early Cretaceous) Paleo to Neogene
Transylvanian Basin, which covers the local Late
Cretaceous to Paleogene Tarnava Basin, (3) the Neogene
Focsani Basin in the foredeep area, which covers
autochthonous Mesozoic and Palaeozoic sedimentary
rocks as well as a proposed Permo-Triassic graben
structure of the Moesian Platform, and (4) the Palaeo and
Mesozoic rocks of the North Dobrogea Orogen.
The underlying crystalline crust shows considerable
thickness variations, in total as well as in its individual
subdivisions, which correlate well with the Tisza-Dacia,
Moesian and North Dobrogea crustal blocks, respectively. Only minor lateral changes in velocity structure
of these blocks were observed. The Tisza-Dacia block is
about 35 km thick and low velocity zones in its
uppermost 15 km are presumably basement thrusts
imbricated with sedimentary successions related to the
Carpathian Orogen. The crystalline crust of Moesia does
not exceed 23 km and is covered by up to 22 km of
sedimentary rocks. The North Dobrogea crust reaches a
thickness of about 44 km including an up to 2 km thick
mixed sedimentary-volcanic-crystalline cover, which is
mainly composed of a thin overthrusted wedge of the
North Dobrogea Orogen.
The presented velocity model intersects the TransEuropean suture zone (TESZ) and shows a high degree
of similarity in its overall geometry and velocities to the
velocity models and seismic profiles across the TESZ in
Poland further to the northwest, although the specific
crustal evolution of both areas appears to have clear
differences.
Acknowledgements
This investigation was only possible by the continuous effort of many volunteers, in particular students
from the Universities of Amsterdam, Bucharest and
Karlsruhe. The National Institute for Earth Physics
(NIEP) and the University of Bucharest (Geology and
Geophysics Department) provided the logistics for the
fieldwork in Romania. The Romanian Exploration
Company PROSPECTIUNI S.A., Bucharest, was responsible for the environmental study as well as the
drilling and shooting operations. Data were collected
using the seismic equipment of the geophysical
instrument pool of the GeoForschungsZentrum Potsdam
(150 units) as well as the joint pool of IRIS /PASSCAL at
Socorro, New Mexico and the University of Texas at El
Paso (640 units). The Deutsche Forschungsgemeinschaft
(German Science Foundation) funded the project
through the Collaborative Research Centre 461 (CRC
461) at the University of Karlsruhe, Germany: “Strong
Earthquakes — a Challenge for Geosciences and Civil
Engineering”. The Romanian Ministry for Education and
Research funded the Romanian researchers in this
project via the CERES program (CERES 1 no. 34/
2001 and CERES 4 no. 38/2004). The NATO Science
Collaborative Research Linkage Grant no. EST.CLG
974792 assisted the project by additional travel funding.
Laszlo Csontos and an anonymous reviewer are
F. Hauser et al. / Tectonophysics 430 (2007) 1–25
thankend for their careful and constructive reviews,
which helped to improve and clarify parts of the paper.
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