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Tectonophysics 382 (2004) 51 – 84
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Slab behaviour and its surface expression: new insights from
gravity modelling in the SE-Carpathians
Blanka Sperner a,*, Dumitru Ioane b, Robert J. Lillie c
a
Geophysical Institute, Karlsruhe University, Hertzstr. 16, D-76187 Karlsruhe, Germany
b
Faculty of Geology and Geophysics, University of Bucharest, Bucharest, Romania
c
Department of Geosciences, Oregon State University, Corvallis, OR 97331-5506, USA
Received 21 March 2003; accepted 16 December 2003
Abstract
We use lithosphere-scale gravity models to calculate gravity anomalies resulting from oceanic subduction, continental
collision, slab steepening, delamination, and break-off. Local isostasy was assumed for determining vertical movements caused
by mass changes related to these tectonic processes. Our results show that subduction is accompanied by basin subsidence on
the upper plate caused by the heavy lithospheric root of the subducting slab. The basin evolution goes parallel with the slab
evolution and shows considerable modifications when the processes at depth change (slab steepening, delamination, break-off).
Characteristic gravity anomaly curves were acquired for the different tectonic scenarios. These curves together with other data
(e.g. basin evolution on the upper and the lower plate) were used for the reconstruction of the tectonic evolution of the SECarpathians which includes Tertiary subduction and collision followed by slab steepening and delamination.
D 2004 Elsevier B.V. All rights reserved.
Keywords: Gravity anomalies; Subduction; Isostasy; Vertical movements; Basins; Carpathians
1. Introduction
Subduction of oceanic lithosphere induces largescale mass transfers as heavy lithospheric mantle
replaces lighter asthenosphere. The density contrast
between the two materials is high enough to produce
a slab-pull force thought to be an important driving
mechanism for plate tectonics (e.g. Conrad and
Lithgow-Bertelloni, 2002). It can be expected that
* Corresponding author. Fax: +49-721-71173.
E-mail address: [email protected]
(B. Sperner).
0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2003.12.008
this heavy slab also has an effect on the surface
gravity field, not only at the plate boundary, but also
above the upper-plate region underlain by the heavy
lithospheric root of the subducted slab. Another
consequence is vertical movement due to isostatic
adjustments. We use gravity modelling to study
these effects in general and then to study an example
of young subduction in the SE-Carpathians. Our
type of gravity modelling differs from standard
gravity modelling in two ways: it includes the
tectonic evolution of the study area and it considers
isostatic adjustments.
The inversion of gravity data yields non-unique
solutions. This means that the integration of informa-
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
tion from other fields like geology and geophysics is
mandatory. By considering the tectonic evolution of a
region, gravity modelling can be used to test different
evolution models. This has the additional advantage
that the density distribution at depth is explained by
the tectonic evolution. No ‘‘mysterious’’ bodies with
density contrasts that best fit the gravity data are
presumed (as is sometimes the case in standard
gravity modelling). For isostatic balancing, local
(Airy) isostasy was assumed for the different stages
of tectonic evolution. The resulting vertical movements are assumed to be followed by erosion (of
uplifted areas) and sedimentation (in subsided areas).
Previous gravity modelling studies concentrated on
the crustal evolution of continental collision zones
(e.g. Lillie, 1991; Lillie et al., 1994). We consider
greater depths and thus include the influence of the
lithospheric root created by the subducted slab. We
first study processes related to the evolution of the
slab in general, namely subduction, steepening, delamination and break-off. Then, we apply this method
to the SE-Carpathians, test different models for the
tectonic evolution of the region, and compare the
results with measured data.
In the SE-Carpathians, Miocene oceanic subduction
was followed by continental collision at ca. 10 Ma.
Post-collisional shortening was small (soft collision),
so that the structures and basins which developed
during subduction are well preserved. Intermediatedepth seismicity (70 – 200 km) is concentrated in a
nearly vertical column (Bonjer et al., 2000; Oncescu
and Bonjer, 1997), suggesting that the subducted slab
is still hanging beneath the SE-Carpathians (Sperner et
al., 2001). Data from seismic experiments (refraction,
tomography) provide detailed information about the
lithospheric structure (Hauser et al., 2001; Martin et
al., 2001; Wortel and Spakman, 2000).
2. Modelling approach
Both free-air and Bouguer gravity anomaly data are
used in this paper.
2.1. Free-air anomaly (FAA)
The free-air correction accounts for gravitational
effects related to the latitude and elevation of gravity
stations with masses between the station and the
reference level being preserved. FAA data can give
some indication of the degree to which a region
approaches isostatic compensation.
2.2. Bouguer anomaly (BA)
The Bouguer correction accounts for the gravitational effects of masses between gravity stations and
the reference level. For stations on land, this is often
performed by subtracting the gravity effect of an
infinite slab with a specific density (typically 2670
kg m 3). For stations at sea, the opposite is done by
virtually filling the ocean with crustal material, i.e. by
adding the effects of an infinite slab with a density
equal to the difference between water and crustal
material. Therefore, oceans are always characterized
by positive BA values. On land, BA data often show a
strong correlation with the depth to the Moho because
the density contrast between crust and mantle is large
and topographic effects have been accounted for by
the Bouguer correction. Thus, undulations in the
Moho have large gravity effects at the surface.
2.3. Density contrasts
We use density contrasts because gravity anomalies, and especially the shape of the anomaly curves,
are the result of lateral density changes which depend
on density contrasts and not on absolute density
values. Additionally, the use of density contrasts has
the advantage that the depth-dependent increase of
density is inherent (density increases uniformly in all
units so that the contrast remains the same).
2.4. Density contrast values
We used the crust as a reference level and assigned
it a relative density of zero (Table 1). The density
contrast between crust and lithospheric mantle was set
to 300 kg m 3 (Bielik, 1999; Lillie et al., 1994). For
the density contrast between lithospheric mantle and
asthenosphere, values in the literature range from 0 to
200 kg m 3 (e.g. Buiter et al., 1998; Jones et al.,
1996; Lachenbruch and Morgan, 1990; Lillie et al.,
1994; Royden and Karner, 1984). Estimations from
seismic tomography data, which use the velocity –
density relationships of, e.g. Birch (1964) or Cermak
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Table 1
Density contrasts used in the gravity models
Topography (young sediments)
Topography (crystalline crust)
Water
Young sediments
Continental crust
Oceanic crust, at depth < 30 km
Oceanic crust, subducted below 30 km
Lithospheric mantle
Asthenosphere
[kg m 3]
Symbol
2470
2670
1640
200
0
0
300
300
270
qTsed
qTc
qW
qSed
qCc
qOc
qEclo
qLM
qA
Values are relative to crustal material (qCc, qOc).
et al. (1990), lie in most cases between 10 and 50 kg
m 3. Therefore, we assume an average value of 30 kg
m 3 (i.e. 270 kg m 3 relative to the crust). The
density distribution in the down-going slab is depthand time-dependent due to the eclogitisation of the
oceanic crust at greater depths. Eclogitisation is,
amongst other things, controlled by temperature and
is thus dependent on the subduction velocity. For the
sake of simplicity, in our models we neglect the
influence of subduction velocity and we assume that
the oceanic crust undergoes eclogitisation at a depth
of 30 km. Thus, we set the density for the oceanic
crust at depths larger than 30 km to the same value as
the lithospheric mantle (i.e. a density increase of 300
kg m 3). Depth-dependent density increase in the
slab (crust and lithospheric mantle) is already taken
into account through the use of relative instead of
absolute densities. This means that density in the slab
increases in parallel with the density increase in the
surrounding asthenospheric material.
2.5. Local isostasy
For each step, the model was kept in local isostatic
equilibrium (Airy isostasy). This means that mass
changes are compensated directly beneath the mass
anomaly. For the resulting vertical mass movements,
the shear strength of the material is assumed to be zero
(no rigidity). This simplification makes the calculations much easier and faster in comparison with
flexural bending models and thus allows the testing
of a broader suite of models. The Airy approximation
is further justified because previous studies showed
that flexural rigidities in the study area are low (low
effective elastic thickness of the lithosphere: a few
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kilometres according to Artyushkov et al. (1996) and
14 km according to Matenco et al. (1997)).
Fig. 1 shows two examples for local isostasy, one
with mass deficit (crustal thickening; Fig. 1b) and one
with mass excess (lithospheric thickening; Fig. 1d) in
the subsurface. In both cases, Bouguer and free-air
anomalies are the same for the uncompensated stage
(Fig. 1b,d). The broad free-air anomalies indicate that
isostatic equilibrium is not realized and the Bouguer
anomalies reveal the uncompensated mass deficit and
mass excess for the two models. Note that the gravity
anomalies of the two uncompensated models are
different by an order of magnitude, the crustal thickening effect (Fig. 1b) being far greater than that of the
lithospheric thickening (Fig. 1d). Reasons for the
smaller anomalies in the lithospheric thickening model are the greater depth of the lithospheric root and the
lower density contrast between lithospheric mantle
and asthenosphere. Isostatic surface effects are uplift
in the case of mass deficit (Fig. 1c) and subsidence in
the case of mass excess (Fig. 1e). In the crustal
thickening model, the Bouguer anomaly is slightly
reduced (due to the rise of the Moho), while the freeair anomaly shows a typical ‘‘edge effect’’: the
uplifted area (topography) leads to a sharp increase
in the gravity anomaly, while the negative effect of the
crustal root is, due to its larger depth, of smaller
amplitude, but larger wavelength. Isostatic equilibrium is indicated when the areas under the FAA curve
(positive above zero line, negative below zero line)
sum to zero, which is the case in the example in Fig.
1c. In the lithospheric thickening model, isostatic
compensation leads to subsidence and thus the freeair anomaly flips to negative values in the middle part
of the model (Fig. 1e). The free-air anomaly is again
characterized by negative– positive couples typical of
the edge effect (see above). The Bouguer anomaly
curve shows only little changes due to the only
slightly deeper Moho.
2.6. Sedimentation
During the subduction and collision process, the
sediment volume on the two continental margins was
kept constant. Newly formed basins, like those on the
continental margin of the upper plate, were filled with
sediments when there was a nearby supply area with
topographic relief. Continental collision provides such
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Fig. 1. Effects of local isostasy due to crustal and lithospheric thickening respectively. (a) Starting model with no gravity anomalies. (b, c) Crustal thickening results in a mass deficit
in the lithospheric mantle (due to the lower density of the crust) which is compensated by uplift (heavy asthenospheric material is added at the bottom). The Bouguer anomaly is
negative, with a slight decrease after uplift. Without uplift, the free-air anomaly is the same as the Bouguer anomaly. The large negative free-air anomaly in (b) indicates that the model
is not in isostatic equilibrium. After isostatic adjustment, smaller free-air anomalies remain due to the edge effect caused by the different depths of the mass anomalies (see text for
details). (d,e) During lithospheric thickening, asthenosphere is replaced by higher-density lithospheric mantle. The resulting mass excess is compensated by subsidence. The Bouguer
anomaly is positive in both models. The initial free-air anomaly is also positive, indicating that the model is not in isostatic equilibrium. Subsidence leads in parts to negative free-air
anomalies, again showing an edge effect. Note the different scales of the gravity plots in (b,c) and (d,e).
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
uplifted areas, so that synorogenic sediment supply
was assumed. Basins were filled to zero-level, but
topographic heights were kept constant to provide
information about the maximum topography possible
for the respective tectonic evolution (i.e. no erosion
and thus no volume constancy).
2.7. Software
Modelling was done with the commercial software
package GM-SYS (v. 4.04 and 4.6) of Northwest
Geophysical Associates (NGA; http://www.nga.com/).
3. Models
The starting configuration for our models consisted of two continents with 30-km-thick crust
separated by an ocean with 8-km-thick crust (Fig.
2a). We assumed an old, thermally equilibrated
lithosphere with a horizontal lithosphere –asthenosphere boundary at 120-km depth. This implies a
thickness of the oceanic lithosphere of 116.6 km
which correlates with an age of 80 Ma (e.g.
Turcotte and Schubert, 2002). The width of the
transition zone between continent and ocean was
set to 100 km and the thickness of the water –
sediment package was limited to 10 km, with the
maximum sediment thickness at the continent –
ocean boundary. The compensation level was set
to a depth of 550 km, thus providing enough space
for the subducting slab. Models are shown on a
lithospheric scale (no vertical exaggeration) and on
a crustal scale (vertical exaggeration of 4:1). For
easier description, the orientation of the cross-sections is assumed to be west –east and the subduction direction is towards the west. All models have
a 2D-geometry, i.e. they extend to infinity in the
dimension perpendicular to the profiles shown.
3.1. Oceanic subduction
3.1.1. Local isostasy
Oceanic subduction leads to the development of
a lithospheric root beneath the upper plate (Fig. 2).
The mass excess at depth results in basin subsidence on the upper plate. Due to the origin of this
basin, a subsurface load, we refer to it as a load
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basin. Continued subduction leads to an increase in
basin width; the basin depth also increases as long
as the lithospheric root continues to thicken. From
a certain point on, when the maximum thickness of
the lithospheric root is reached, the basin will only
widen without deepening (Fig. 3). The maximum
thickness of the lithospheric root (hLR) depends on
the thickness of the oceanic lithosphere (hOL) and
on the dip angle (a) of the slab (Fig. 3a):
hLR ¼ hOL ðcosðaÞÞ1
ð1Þ
The lithospheric root with density qLM replaces
asthenospheric material with density qA, so that the
mass excess DmLR can be calculated:
DmLR ¼ hLR ðqLM qA Þ
ð2Þ
For the mass calculations, the two other dimensions of the mass volume are constant and thus are
assumed to be 1 m each; they have no influence on
the result, so they are not shown in the formulas.
In the same way, the mass deficit of the water-filled
basin on the upper plate can be calculated (again,
asthenospheric material is replaced because it disappears out of the model at its lower boundary).
Subsidence hw of a water-filled basin causes a mass
deficit Dmw of
Dmw ¼ hw ðqw qA Þ
ð3Þ
For isostatic equilibrium, the sum of the mass
anomalies must equal zero:
hLR ðqLM qA Þ hw ðqA qw Þ ¼ 0
and thus the basin depth can be calculated as
hw ¼ hLR ðqLM qA ÞðqA qw Þ1
ð4Þ
The maximum basin depth then is given by substituting Eq. (1) into Eq. (4)
hw ¼ hOL ðcosðaÞÞ1 ðqLM qA ÞðqA qw Þ1
ð5Þ
Inserting the values in Table 1 for the density
contrasts, the formula can be simplified to
hw ¼ hOL ðcosðaÞÞ1 30=1910
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Fig. 2. Gravity anomalies (BA, Bouguer anomaly, FAA, free-air anomaly) and surface effects of oceanic subduction. The heavy lithospheric root
causes subsidence in the upper plate (load basin). Initial ocean width is 700 km, slab dip is 30j; density contrasts are given in Table 1 (V.E.,
vertical exaggeration).
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
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Fig. 3. Basin evolution on the upper plate during subduction (same configuration as in Fig. 2). (a) With ongoing subduction the basin increases
in width and depth until a certain stage is reached (stage 2). From then on the thickness of the lithospheric root does not increase any more
(hLR2 = hLR3, maximum thickness of the lithospheric root), so that the basin only widens while maintaining constant depth. (b) Uppermost part
of (a), with tenfold vertical exaggeration (V.E.). (c) Basin depth increases with increasing convergence until the maximum depth is reached.
Sedimentary fill produces a basin about four times deeper than with water fill (caused by the lower density contrast between sediment and
asthenosphere compared to the one between water and asthenosphere).
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A similar relationship exists for the maximum
depth of a sediment-filled basin (hSed) with density
qSed:
hSed ¼ hOL ðcosðaÞÞ1 ðqLM qA ÞðqA qSed Þ1
¼ hOL ðcosðaÞÞ1 30=470 ¼ 4:06hw
ð6Þ
Thus, sedimentary infill instead of water increases
the depth of the basin by a factor of about 4 due to
the lower density contrast between young sediments
and asthenosphere (470 kg m 3) in comparison to
the contrast between water and asthenosphere (1910
kg m 3).
Fig. 3b shows the increase of basin depth with
ongoing subduction until it reaches its maximum
depth, which in our model is 2.1 km for water fill
and 8.6 km for sediment fill for a slab dip a of 30j
and a thickness of the oceanic lithosphere hOL of
116.6 km. A steeper slab shows a larger maximum
basin depth due to its thicker lithospheric root (Fig.
4a). The limit for the maximum thickness of the
lithospheric root (and thus for the maximum basin
depth) is controlled by the down-dip length of the
subducted slab (otherwise, it would increase to infinity; Fig. 4b).
With increasing dip angle, the lithospheric root
not only increases in thickness, but it is also located
closer to the plate boundary (Fig. 5). In the same
way, the upper-plate basin shifts to a more eastern
position so that it lies on the continent – ocean
transition zone (Fig. 5b,c) or even on the downgoing plate (Fig. 5c,d). Due to the high-mass excess
in the subsurface, this region is pulled down to large
depths. In the most extreme case of a vertical slab,
the basin is completely situated on the down-going
plate and has a symmetric shape.
3.1.2. Free-air anomaly
The free-air anomaly is significantly affected by
the basin subsidence in the upper plate. Subsidence
results in a sharp and pronounced negative peak in the
FAA curve. The slab itself causes a broad, but only
small rise of the FAA data, leading to the typical ups
and downs near the basin margin (edge effect; see
above). Steeper slabs result in deeper, but narrower
basins with a corresponding effect on the negative
FAA peak (Fig. 5).
Fig. 4. Dependence of basin subsidence on the slab dip angle (same
configuration as in Fig. 2 except dip angle). (a) A larger slab dip
angle (a2>a1) results in a larger maximum basin depth. The amount
of convergence to reach this maximum depth is smaller than for
flatter slabs (c2 < c1). (b) Correlation between the slab dip angle and
the maximum thickness of the lithospheric root (left axes) and the
maximum basin depth (right axes). The upper limit of these two
parameters is given by the length of the subducted slab (400 km in
the example shown).
3.1.3. Bouguer anomaly
The influence of oceanic subduction on the Bouguer anomaly is only small (Figs. 2 and 5). The
lithospheric root causes a moderate rise of the BA
data, which is partly balanced by the down-warping of
the continental crust.
3.2. Continental collision
3.2.1. Local isostasy
Continental collision is preceded by the collision of
the two sedimentary wedges lying on the margins of
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
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Fig. 5. Different slab dip angles (a) and their influence on gravity anomalies and on load basin location relative to the plate boundary (same
configuration as in Fig. 2, with 300 km convergence). (a, b) a = 30j, 45j: The basin is located on continental crust of the upper plate. The
transition zone underlies parts of the basin, but the basin axis lies on the continent. (c) a = 60j: The basin axis lies on the transition zone.
(d) a = 75j: The basin sits on the plate boundary. The high in the middle of the basin is caused by the down bended, not eclogised (i.e.
‘‘light’’) part of the oceanic crust. The location of the basin is also dependent on the thickness of the slab: a larger thickness shifts the basin
axis toward the ocean.
the two continents (Fig. 6a – c). Internal shortening of
this sediment wedge is assumed to be limited to 150
km which is about the shortening inside the Carpathian accretionary wedge (see below, Section 4.3).
When this limit is reached, the whole sediment
package starts to thrust over the down-going plate
(Fig. 6c – f). In our example, the maximum thickness
of the sedimentary wedge is 13.3 km and its maximum height above sea level reaches 2.1 km. Topography of the continental crust caused by crustal
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
thickening during collision reaches a maximum level
of 2.8 km.
The lithospheric root initially causes load basin
subsidence in the upper plate (Fig. 6a – c). However,
as soon as the continents collide, the crust thickens
and uplift begins, first in the accretionary wedge (Fig.
6d) and later in the eastern part of the basin (Fig. 6e).
With ongoing underthrusting, the uplifting region
propagates towards the west (Fig. 6e –f). At the same
time, the lithospheric root also shifts westward, forcing the basin to move in the same direction. Thus, the
westward-propagating basin is followed by an uplifting region, similar to a wave that first goes down and
then up again. This means that a basin with constant
width and geometry migrates westward, but is later
uplifted as part of an expanding mountain range. Note
that in our model, subsidence of this basin has nothing
to do with (back)thrusting and loading by the mountain range; it is a pure drag effect of the lithospheric
root.
3.2.2. Free-air anomaly
The load basin on the upper plate is characterized
by a negative free-air anomaly (Fig. 6c– f), while a
positive anomaly develops east of the basin in the
uplifted region (Fig. 6d –f).
3.2.3. Bouguer anomaly
Closure of the ocean basin and uplift of the sediments on the continental margins removes the oceanrelated positive Bouguer anomaly (Fig. 6a– c). A new
positive anomaly develops due to the lithospheric root
beneath the upper plate. Crustal thickening during
continental collision results in a negative Bouguer
anomaly which broadens with ongoing underthrusting
(Fig. 6d – f) (see also Lillie, 1991).
In the following sections, we will discuss the
evolution of the subducted slab after continental
collision. We assume a soft continental collision with
only 100 km of underthrusting and study the different
slab processes which have been proposed to take
place after collision (slab steepening, break-off and
delamination). A more detailed discussion about the
61
gravity effects of hard continental collision (i.e. collision accompanied by intense shortening and underthrusting) is given by Lillie (1991) and Lillie et al.
(1994).
3.3. Slab steepening
Numerical modelling by Dvorkin et al. (1993)
showed that the hydrodynamic suction between upper
and lower plate, induced by corner flow during
downwards movement, stabilizes the slab at a constant dip angle. However, several parameters and
processes can lead to a steepening of the slab. One
important factor is the along-strike length of the slab:
for narrow slabs, asthenospheric material can flow
around the corners of the slab, thus reducing the
hydrodynamic suction near the plate contact and
leading to slab steepening and rollback (Dvorkin et
al., 1993). Another factor is the end of down-dip
motion after continental collision. In this case, the
hydrodynamic suction is decreasing because corner
flow stops. Thus, the subducted lithosphere starts to
steepen under its own weight (for the same reason slab
break-off only takes place after continental collision;
Yoshioka and Wortel, 1995). To study the surface
effects of this important process, we modelled a slab
that first subducted at a 20j dip angle and then (after
soft continental collision) steepened to 75j. The
down-dip length of the slab is 300 km. The other
parameters are the same as in the previous models.
3.3.1. Local isostasy
As in the previous models, subduction leads to the
evolution of a basin on the upper plate. The maximum
basin depth is near the western margin of the basin
(Fig. 7a). We assumed that the basin was filled with
sediments from the surrounding uplifted regions created during continental collision (Fig. 7b). Steepening
of the slab (in our example from 20j to 75j) has two
effects (Fig. 7c): (a) the lithospheric root is now
narrower, but thicker, and its location is further to
the east. Thus, the basin axis also shifts towards the
east and the basin deepens; (b) the region west of the
Fig. 6. Collision of sedimentary wedges and of continents (slab dip: 20j). The sedimentary wedges thrust over each other until a predefined
maximum amount of internal shortening is reached (150 km in our models). Afterwards, the sedimentary wedge as a whole thrusts over the
down-going plate. Continental collision and underthrusting are accompanied by crustal thickening leading to surface uplift on the upper plate.
At the same time, the load basin on the upper plate moves westwards and its eastern parts are uplifted.
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Fig. 7. Effects of slab steepening after continental collision. (a) Subduction with an angle of 20j was followed by 100 km of continental
underthrusting, resulting in basin subsidence on the upper plate and shortening and uplift of the sediments on the continental margins. (b) The
basin was filled with sediments which were delivered from the uplifted surrounding regions (caused by continental collision). (c) Steepening of
the slab to 75j results in a new basin with an axis located east of the old one, while the region west of it rises above sea level. (d) Same as
(c), but without asthenospheric upwelling during slab steepening. Differences between (c) and (d) concerning isostatic movements and gravity
anomalies are minor.
new basin is now missing the formerly present lithospheric root so that it rises above sea level.
As shown in Fig. 7c, we assume that the steepening
slab leaves a gap in the lithospheric mantle which is
filled by upwelling asthenospheric material. The
opening of such a gap is dependent on the coupling
between the subducted slab and the overlying lithospheric mantle of the upper plate. Gvirtzman and Nur
(1999) assume a decoupling between slab and upper
plate beneath the Calabrian arc and the Kuriles. We
tested the gravitational effect by doing the same
model with and without a gap. The results show only
small differences in the calculated isostatic movements and gravity anomalies (Fig. 7d). Effects on
other parameters, like heat flow and volcanism, might
be more severe.
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
3.3.2. Free-air anomaly
The FAA curves mainly reflect the surface movements. Negative anomalies exist where basins subside
and positive anomalies where uplift above sea level
occurs (Fig. 7a– d).
3.3.3. Bouguer anomaly
Filling the load basin with sediments results in a
negative Bouguer anomaly reflecting the downbended Moho (Fig. 7a,b). The crustal keel of the
underthrusted plate is responsible for a second negative BA farther east. After slab steepening, the old
load basin rises by 1.16 km and thus reduces the
negative Bouguer anomaly, but it cannot balance the
mass deficit caused by the loss of the lithospheric
root, so that the negative Bouguer anomaly is stronger
than before (Fig. 7c). The new load basin reduces the
local high in the BA curve (Fig. 7c).
3.4. Slab break-off
The fate of subducted lithosphere after continental
collision is still under discussion (e.g. Davies, 1980;
Davies and von Blanckenburg, 1995; Schott and
Schmeling, 1998). Most old subduction and collision
zones show no evidence for the existence of subducted material (e.g. no seismicity) so that the slab
somehow has to disappear. One possibility is slab
break-off followed by gravitational sinking of the
detached lithosphere into the deeper mantle (e.g.
Levin et al., 2002). Such a process has been proposed
for the Alps (von Blanckenburg and Davies, 1995);
Wortel and Spakman (1992) assume that break-off
beneath the Hellenic arc started at one lateral end of
the slab and now propagates laterally along the arc.
The slab gap is visible in seismic tomography and it
shows that break-off started in that region of the
Hellenic arc where continental collision already occurred. Numerical modelling by Yoshioka and Wortel
(1995) point in the same direction: slab break-off only
takes place when the down-dip movement stops, so
that extensional forces inside the slab can increase to
values large enough to initiate break-off. The region
where slab break-off will preferentially take place is
the transition zone between oceanic and continental
crust. Bending forces will be concentrated in this
region due to the buoyancy of the continental crust
which resists the downward pull of the slab. Addi-
63
tionally, this region is already weakened by extensional fault zones which developed during the
opening of the ocean. Thus, in our models we assume
slab break-off to occur at the continent –ocean boundary after soft continental collision took place. The
detached slab sinks into the deeper mantle so that no
isostatic or gravitational effect remains.
3.4.1. Local isostasy
Removal of the lithospheric root results in uplift of
the overlying region. Fig. 8 shows an example where
slab steepening took place before break-off, so that the
uplift occurred in two steps, first in the western part of
the load basin (due to slab steepening; Fig. 8a), later
in the eastern part (due to slab break-off; Fig. 8b).
3.4.2. Free-air anomaly
Slab break-off reduces the edge effects in the FAA
data produced during slab steepening and thus brings
the FAA curve closer to the zero line (Fig. 8a,b).
Small anomalies remain due to the thickened crust
(additional sediments) and the removal of parts of the
lithospheric mantle during steepening and break-off.
3.4.3. Bouguer anomaly
In comparison with the BA curve for the steep slab
(Fig. 8a), slab break-off shifts the curve to slightly
more negative values because the additional mass of
the subducted lithosphere has been removed (Fig. 8b).
Uplift cannot completely balance this effect because
the crust had been thickened by sediments, so that a
crustal root is left over.
3.5. Delamination
In contrast to slab break-off, which occurs along a
plane perpendicular to the surface of the lithosphere,
delamination takes place along a plane parallel to it.
One triggering mechanism for delamination of the
lithospheric mantle from the crust is the weight of an
overthickened lithospheric mantle caused by shortening in a continental collision zone (e.g. Tibet,
Alboran) (Bird, 1978; Seber et al., 1996), another is
the slab pull of subducted lithosphere (Girbacea and
Frisch, 1998; Sacks and Secor, 1990). In our model,
we assume 100-km delamination along the Moho
(Fig. 8c). The starting model and thus the preceding
evolution is the same as for the break-off model
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
(Fig. 8a,b), so that the effects of both processes can be
compared. The final position of the slab is vertical.
3.5.1. Local isostasy
Delamination removes the whole lithospheric mantle from the westernmost part of the lower plate, so
that this region experiences uplift (Fig. 8c). Further
eastward, the vertically hanging slab causes subsidence and basin formation. With ongoing delamination, the slab and thus the basin migrate towards the
east, while the previously down-warped region west
of the slab will rise. In the case that sediments had
been deposited into the basin, the topography after
delamination will be higher than before (due to the
larger crustal thickness).
3.5.2. Free-air anomaly
The FAA curve is similar to the steep –slab curve
with a shift of the minimum towards the east according
to the eastward shift of the slab during delamination
(Fig. 8a,c). Additionally, the steeper dip and the greater
vertical extent of the slab cause a more pronounced
minimum.
3.5.3. Bouguer anomaly
In the delaminated region, BA data are shifted to
more negative values due to the removal of the heavy
lithospheric slab (similar to slab break-off; Fig. 8b).
The region above the slab experiences subsidence
accompanied by down-bending of the Moho, which
also results in a shift to more negative values. The BA
curves for break-off and delamination are similar, but
the influence of the lithospheric slab is still visible in
the easternmost part where the BA data for delamination remain at a higher level (positive values) as
compared with the break-off data (Fig. 8e). If sediments were deposited into the basin above the slab,
the thickness of the crust increases resulting in an
additional down-bending of the Moho and thus a more
65
pronounced negative Bouguer anomaly; the FAA
curve changes only slightly (Fig. 8e).
4. Gravity modelling in the SE-Carpathians
(Vrancea region)
To date, gravity modelling of the Carpathians
has mainly covered the northern parts (i.e. Western
Carpathians) (Bielik, 1995; Bielik, 1999; Lillie and
Bielik, 1992; Lillie et al., 1994; Szafian et al.,
1999) or concentrated on large-scale structures
(Szafian et al., 1997). None of this work considered
the tectonic evolution of a subducting slab (including slab steepening, delamination and break-off)
and its influence on vertical movements at the
surface as well as on the gravity data. We have
chosen the SE-Carpathians as a study area because
intermediate-depth seismicity and seismic tomography indicate that cold subducted material is still
present in the upper mantle (Martin et al., 2001;
Oncescu, 1984; Sperner et al., 2001). Furthermore,
abundant data from different fields (e.g. structural
geology, sedimentology, paleomagnetics, seismics)
are available, allowing the reconstruction of the
Neogene tectonic evolution.
The Neogene to recent tectonic evolution of the
SE-Carpathians has for a long time been debated
and discussions are still going on, especially
concerning the existence and the type of Miocene
subduction (e.g. Girbacea and Frisch, 1998; Girbacea and Frisch, 1999; Linzer, 1996a; Linzer, 1996b;
Pana and Erdmer, 1996; Pana and Morris, 1999).
Here we present our preferred model for the tectonic evolution, test it using gravity modelling, and
compare it with other models for the recent platetectonic configuration. The time scale used in this
paper is based on the Central Paratethys stages as
compiled by Rögl (1996) (Table 2).
Fig. 8. Effects of slab break-off and delamination. (a, b) The subducted and steepened slab (a) breaks off at the ocean – continent boundary and
sinks into the deeper mantle (b) so that no gravitational effect remains. As a result, the region above the removed slab raises. The BA curve is
shifted to slightly more negative values because lithospheric mantle is replaced by less-dense asthenosphere. The FAA curve approaches the
zero line because the high-density slab is removed. (c) Delamination along the Moho causes uplift above the region of the removed slab, while
subsidence occurs above the new position of the slab. The BA curve is similar to that for slab break-off, while the FAA curve shows a
pronounced minimum at the slab location (strong edge effect). (d) Infill of sediments into the basin above the slab changes the FAA curve only
slightly, but results in a stronger negative Bouguer anomaly due to the deepening of the Moho. (e) For better comparison, the BA and FAA
curves for the different tectonic situations are plotted in one diagram.
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Table 2
Correlation between Mediterranean and Paratethys stages after Rögl (1996)
4.1. Neogene tectonic evolution of the Carpathian
region
Subduction beneath the Carpathians had begun by
late Cretaceous when the intra-Carpathian blocks
(North Pannonian block and Tisia – Dacia block in
Fig. 9) and the Adriatic microplate moved north(west)ward. In the Alps, continental collision had already
started at the end of the Cretaceous (Stampfli et al.,
2002), while subduction continued beneath the Carpa-
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
67
Fig. 9. Tectonic map of the Carpathian – Pannonian region showing Tertiary – Quaternary structures and the location of intermediate-depth
earthquakes in the SE-Carpathians (Vrancea region; geology based on Horvath, 1993). The blow-up (lower map) shows the relationship between
topography and main basins in the SE-Carpathians.
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
thians. An oceanic embayment in the European foreland, which was still open when collision took place in
the Alps, provided additional space for the northeast-
and eastwards movement of the intra-Carpathian
blocks (Fig. 10). The Miocene movements of these
blocks can be reconstructed by taking into consider-
Fig. 10. Geodynamic evolution of the Carpathian – Pannonian region during the Miocene as derived from kinematic, sedimentary, and
paleomagnetic data (Sperner et al., 1999). Opposite rotation of two crustal blocks (a) resulted in an oblique collision with the European foreland
that started in the north and successively propagated towards the southeast and south (b). TDB, Tisia – Dacia block.
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
ation data from structural geology, paleomagnetism,
and sedimentology (e.g. Balla, 1985; Csontos et al.,
1992; Fodor et al., 1999; Linzer et al., 1998; Royden et
al., 1983). Continental collision in the Carpathians first
took place in the western part (18 Ma), then shifted to
the north (13 Ma) and finally reached the SE-Carpathians (9 Ma; Fig. 10). Intermediate-depth seismicity is
only recorded from the southeastern corner of the
Carpathians (Vrancea region; Fig. 9), so that we assume
that the subducted slab is only preserved in this region
and is still hanging beneath the collisional orogen. In all
other regions of the Carpathians, the slab has already
broken off and sunk into the deeper mantle (e.g.
Sperner, 1996; Sperner et al., 2001). The fact that the
subducted slab, as defined by the intermediate-depth
seismicity in the Vrancea region, is not hanging beneath the upper plate or the suture zone, but is located
beneath the accretionary wedge and the foredeep
(Fig. 11), has led to a number of different interpretations (e.g. Csontos, 1995), which we have tested by
means of gravity modelling.
4.2. Available data
The Romanian gravity maps were compiled at the
Institute of Geology and Geophysics (Bucharest) and
69
are based on 78,000 point observations (mean coverage 0.33 data points per km2) (Ioane and Radu, 1995).
The normal gravity was calculated with the International Gravity Formula (Cassinis, 1930) using a reduction density of 2670 kg m 3. The Bouguer
correction was done with an infinite slab; for the
terrain correction the Schleusener method was used
(Schleusener, 1940). Both the Bouguer and the
free-air gravity anomaly map are shown in Fig. 12.
For topography, the GTOPO30 data were used which
are available online at http://edcwww.cr.usgs.gov/
landdaac/gtopo30/gtopo30.html. Sediment thickness
and other geological information come from maps
and profiles published by the Geological Institute in
Bucharest (e.g. Dumitrescu and Sandulescu, 1970;
Sandulescu et al., 1978; Stefanescu, 1985). Vertical
movement rates have been studied for several decades
by using repeated levelling data (Popescu and Lazarescu, 1988; Zugravescu et al., 1998). They show that
the SE-Carpathians are uplifting, while the foreland
subsides. Fission-track data can be used to estimate the
beginning of uplift, but they have to be handled with
care because they identify the time of exhumation
(erosion) and not of uplift, i.e. the acquired ages are
in many cases younger than the beginning of uplift.
For the northern and central part of the Eastern
Fig. 11. Profile through the Vrancea region in the SE-Carpathians (modified after Radulescu et al., 1976; Stefanescu, 1985; for location, see
Fig. 9). Note that the earthquake epicentres are located beneath the accretionary wedge and the foredeep, and not beneath the Miocene suture.
Shown on the left side is the estimation of the slab dip angle (a) from the distance between the plate boundary and subduction-related volcanism
(240 km) and the depth of the magma generating window (90 km): a = arctan (90/240) = 21j.
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Fig. 12. Free-air (a) and Bouguer (b) anomaly maps of Romania (Ioane and Atanasiu, 1998). From 78,000 point observations, mean values were calculated for blocks of 5V 7.5Vsize.
For details about the data processing, see Ioane and Atanasiu (1998). Profiles run through the seismically active Vrancea region (white star) where the existence of a subducted slab is
postulated. For map location, see Fig. 9.
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Carpathians, Sanders et al. (1999) conclude from
fission-track data that erosion rates increased in Late
Badenian –Sarmatian time (11– 14 Ma) and that this
increase is related to the thrusting of the accretionary
wedge onto the European foreland. In the southern
part, a major erosion phase took place during the
Pliocene (1.8 – 5.3 Ma), while at the same time rapid
subsidence and sedimentation occurred in the foreland
(Sanders et al., 1999).
Moho maps have been published in different
versions, some based on gravity data with consideration of isostasy and geology (Socolescu et al.,
1964), others with additional information from refraction seismic data (Radulescu and Diaconescu,
1998). Using these maps for our gravity modelling
would result in circular arguments (gravity data were
used to construct a Moho map which was then used
to do gravity modelling). We thus rely on the new
refraction seismic profiles of Hauser et al. (2001,
2002), which give Moho information that is independent of gravity data, but have the disadvantage
that they do not give map-scale information for all of
Romania. The depth of the lithosphere – asthenosphere (L/A) boundary has been estimated from heat
flow (Radulescu and Diaconescu, 1998), refraction
seismic (Enescu, 1992) or seismological and magnetotelluric data (Praus et al., 1990). The results show
large variations, e.g. ranging between 100 and 200
km for the East European platform. In our models,
we assumed an initial depth of the L/A boundary of
90 km for the whole model. This depth is reported
for the Transylvanian basin (western part of our
modelling profile) (Radulescu and Diaconescu,
1998) and we assume that the lithosphere of the East
European platform had about the same thickness near
the continental margin; thickening towards the old
and cold, 200-km-thick lithosphere in the inner parts
of the East European platform is not taken into
account. Creation of oceanic lithosphere started at
ca. 160 Ma when the opening of the Alpine Tethys
was initiated (Stampfli and Borel, 2002). Parts of this
ocean were already subducted during Cretaceous
time, but the Carpathian part survived until Oligocene time (ca. 30 Ma). We can thus assume an age of
the oceanic lithosphere of 100 Ma or more and an
associated thickness of ca. 90 km according to the
age –thickness relationship of, for example, Forsyth
(1977).
71
Seismic tomography results are available at different scales ranging from regional (Mediterranean
region; Wortel and Spakman, 2000) to local (Vrancea
region; Martin et al., 2001). They give information
about the dimension and orientation of the subducted
lithosphere beneath Vrancea. Together with data
from structural geology and other fields, this information can be used to reconstruct the (plate-)tectonic
evolution of the Carpathian – Pannonian region
(Csontos et al., 1992; Fodor et al., 1999; Linzer et
al., 1998; Sperner and the CRC 461 team, in press).
This evolution serves as the basis for our gravity
modelling.
4.3. Parameters and boundary conditions for the
gravity models
For the Carpathian gravity models, we used the
same density contrasts as in the previously discussed
models (Table 1). For the density contrast between
lithospheric mantle and asthenosphere, velocity data
from seismic tomography can be used for an independent estimation. Results of seismic tomography show
a high-velocity body that encloses most of the intermediate-depth earthquakes and can be interpreted as
the subducted slab (Fan et al., 1998; Martin et al.,
2001; Wenzel et al., 1998a). Therefore, the velocity
difference between high-velocity body and background model (representing the surrounding asthenosphere) can be used to estimate the density contrast
between lithospheric mantle and asthenosphere. The
mean velocity contrast for the high-velocity body is
about 2% (Martin et al., 2001) and the velocity of the
background model ranges between 7.9 and 8.3 m s 1
for depths of 70 to 300 km. Thus, the velocity
difference Dvp ranges between 0.158 and 0.166 m
s 1. Using the velocity –density relationship of Cermak et al. (1990) for ultra-basic rocks (i.e., for
lower crust and upper mantle; q = 170vp + 1850;
Dq = 170Dvp) the resulting density contrast is ca. 27
kg m 3 which is about the same density contrast that
was used in the models described previously (30 kg
m 3; Table 1).
For the starting model, we assumed a distance of
200 km between the European continent and the
Tisia– Dacia block. This distance is based on three
observations. First, minimum amounts of Tertiary
shortening, as recorded by cross-section balancing of
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
the Eastern Carpathian accretionary wedge, range
from 100 km (Burchfiel and Bleahu, 1976) through
130 km (Roure et al., 1993) to 150 –160 km (Morley,
1996). Roure et al. (1993) and Morley (1996) restored
only parts of the accretionary wedge, so that the total
amount of shortening must be larger. Secondly, paleomagnetic data indicate an early to middle Miocene
clockwise rotation of the Tisia – Dacia block by about
60j (Panaiotu, 1998; Patrascu et al., 1994). For a
reconstruction of the early Miocene position, the
Tisia – Dacia block has to be rotated backwards around
the northwestern edge of Moesia (Fig. 10; note that
the early Miocene (late Egerian, 20 – 22 Ma) position
of the Tisia – Dacia block was even more southeastwards than the Ottnangian to Karpatian position
shown; Sperner et al., 1999). Thus, the distance
between the Tisia – Dacia block and the European
foreland (i.e. the width of the ocean) was about the
length of the northern margin of Moesia, i.e. about
200 km. Thirdly, seismic tomography shows that a
high-velocity body exists beneath the Vrancea region
that reaches depths of about 350 km (Wortel and
Spakman, 2000). This high-velocity body can be
interpreted as the subducted slab indicating that the
assumed 200 km oceanic subduction is a minimum
value. Post-collisional processes like lengthening due
to slab pull or mantle delamination along the Moho
could have increased the originally shorter slab to its
present-day length of 350 km.
The thickness of continental crust and lithospheric
mantle were defined to be 30 and 60 km, respectively.
The thickness of the oceanic crust was assumed to be
8 km and the depth to the L/A boundary was fixed to
90 km (see above). The maximum depth of water plus
sediments on the continental margin was limited to 6
km in the starting model.
The dip angle of the down-going plate can be
estimated from the distance between the plate boundary and subduction-related calc-alkaline volcanism.
This distance is about 240 km (Fig. 11). Assuming
the depth of the magma-generating window to be
about 90 km, the mid-Miocene dip angle of the
subducting slab must have been about 20j. All other
parameters, like the width of the transition zone
between oceanic and continental crust (100 km) or
the maximum shortening in the sediments on the
continental margin (150 km), are the same as in the
models discussed in Section 3. The assumed maxi-
mum shortening is in agreement with the minimum
amount of Tertiary shortening estimated from balanced cross sections (see above).
4.4. Subduction
Miocene oceanic subduction beneath the Tisia –
Dacia block lead to the development of a basin on the
upper plate with a maximum water depth of 1.28 km.
This depth corresponds to a sediment thickness of
5.20 km at the final stage of subduction (Fig. 13a).
The equivalent to this basin in reality is the Transylvanian basin (for location, see Fig. 9), which is known
to be a cold basin with normal crustal thickness (ca.
30 km) and without large-scale extensional structures
despite its up to 5-km-thick Paleogene to Upper
Miocene sediment fill (Fig. 14). These features led
to the development of various, sometimes complicated
models for the evolution of the Transylvanian basin
(e.g. Ciulavu and Dinu, 1998; Nielsen et al., 1999).
Royden et al. (1982) offered a simple solution by
assuming that the steepening of the slab resulted in a
suction force on the upper plate. Our solution is even
simpler: the pure existence of the slab is a reason
enough to cause a downward pull of the upper plate
(Fig. 13a). Slab steepening occurs later in our model
and has a different effect (see below).
Sediments for the basin fill were sourced in the
early Miocene from the north and northwest (De
Broucker et al., 1998). Later, sediment supply also
came from the Southern Carpathians in the south and
from the Apuseni mountains in the west. Both of these
mountain ranges show uplift and erosion from late
Badenian time (15 Ma; Sanders et al., 1999).
4.5. Continental collision
The youngest significant thrusting movements in
the SE-Carpathians, marking the end of the convergence between Tisia – Dacia block and European foreland, took place during the late Sarmatian (Roure et
al., 1993). Younger shortening is related to local
events (Girbacea and Frisch, 1998). Continental collision of the Tisia – Dacia block with the European
foreland started earlier (Badenian) when the transition
zone entered the subduction zone (Zweigel et al.,
1998). Crustal thickening in the SE-Carpathians only
reaches values of about 40 km (Hauser et al., 2001)
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
73
Fig. 13. Gravity model for the SE-Carpathians. We started the model with 200 km distance between the two continents. Thicknesses are 8 km
for oceanic crust, 30 km for continental crust, and 60 km for continental lithospheric mantle. (a) Oceanic subduction resulted in the evolution of
a basin on the upper plate, the Transylvanian basin. The basin was filled during the Paleogene and lower Miocene with sediments from the
Apuseni mountains in the west and the Southern Carpathians in the south. (b) Slab steepening led to a southeastwards migration of the
depocentre in the Transylvanian basin, so that middle and upper Miocene sediments have their maximum thickness in the southeastern part of
the basin (see also Fig. 14). (c) Delamination caused a shift of subsidence further southeastward into the foredeep and uplift of the Eastern
Carpathian flysch belt. Note that sedimentary filling of the subsiding area results in a 15-km-deep basin and a Moho at 45-km depth
accompanied by gravity anomalies much larger than the ones measured in the SE-Carpathians. (d) Adaptation of Moho (40 km) and basin depth
(ca. 10 km Neogene and Quaternary sediments) in combination with a more realistic (smoothed) topography in the foreland results in gravity
anomalies which are close to the measured values. FAA, free-air anomaly (calculated: dashed line; measured: filled circles); BA, Bouguer
anomaly (calculated: solid line; measured: open circles).
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Fig. 14. Profile trough the Transylvanian basin (after Ciupagea et al., 1970) (for location, see Fig. 9). Between the Lower and Middle Miocene,
the axis of maximum basin depth migrated southeastward, consistent with the pattern that results from steepening of the subducted slab
(Fig. 13a,b).
indicating a soft collision where only the two transition zones thrusted over each other, but where no
continental subduction took place. Thus, in our model
convergence stops after 100 km of underthrusting
(Fig. 13a). Fig. 13b shows the steepening of the
subducted slab after continental collision stopped.
The steepening process might already have been
initiated during an earlier stage when subduction
velocity was slowed due to the entrance of transition-zone crust into the subduction zone. In our
model, slab steepening results in a shift of the basin
axis towards the southeast (Fig. 13b). The same
evolution can be seen in the sediments of the Transylvanian basin: the depocentre, as indicated by the
thickest sediment pile, migrates towards the southeast
(Fig. 14).
4.6. Post-collisional evolution
Strong intermediate-depth seismicity in the SECarpathians indicates that the subducted slab is still
hanging beneath the Vrancea region (Fig. 11). Earthquake epicentres are located in a narrow, nearly
vertical column extending from 70 to 200 km depth
(Bonjer et al., 2000; Oncescu and Bonjer, 1997).
Focal mechanisms show, in most cases, a downward
extension, while the compression axes are subhorizontal with no preferred orientation (Oncescu, 1987).
These features, as well as the high deformation rate of
the seismogenic volume (6.3 10 15 s 1; Wenzel et
al., 1998b), demonstrate that the slab is still mechanically coupled with the overlying lithosphere (Sperner
et al., 2001). Slab break-off, as sometimes proposed
for the SE-Carpathians (e.g. Fuchs et al., 1979), would
result in a sinking of the slab into the deeper mantle
without producing strong earthquakes.
One of the most interesting points is the location of
the earthquakes relative to the surface structures: they
are located beneath the accretionary wedge and the
foredeep (Fig. 11). For a Wadati – Benioff zone with
the earthquakes near the upper surface of the slab and
a slab being located at the suture zone, one would
expect the earthquakes beneath the upper plate or the
suture zone. Thus, different plate-tectonic configurations, called ‘‘model 1’’ to ‘‘model 4’’, can be discussed to explain this situation (Fig. 15). Model 1
assumes a subduction zone beneath the suture zone
between Tisia – Dacia block and European platform
(Fig. 15a). As already mentioned, this model can not
explain the location of the earthquakes because they
are situated ca. 100 km further east than expected. A
subduction zone beneath the foredeep (model 2; Fig.
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Fig. 15. Possible plate-tectonic configurations for the recent situation in the SE-Carpathians in which earthquake hypocentres (black dots) are located beneath the accretionary wedge,
not beneath the Miocene suture.
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
Fig. 16. Gravity models for three of the plate-tectonic configurations of Fig. 15. (a, b) Subduction beneath the Miocene suture (Fig. 15a) as well as subduction beneath the foredeep
(Fig. 15b) results in a lateral shift of the gravity anomaly minimum between measured and modelled data for both Bouguer and free-air anomaly. (c) Subduction beneath the Miocene
suture followed by delamination (Fig. 15d) gives a better fit between measured and modelled gravity anomaly data concerning the position of the minimum. Nevertheless, the
modelled anomalies are larger than those measured, indicating that the basin and the Moho are not as deep as assumed in the model. However, adjustments have been made in order to
improve the fit (Fig. 13d).
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
15b) would fit better for the earthquake locations, but
now the accretionary wedge lies completely on the
upper plate. This would be reasonable if large backthrusts (towards the west) could be found, especially
at the western end of the accretionary wedge. However, the majority of the nappes have a movement
direction towards the east and no indications are found
that the whole accretionary wedge was (back)thrusted
onto the Tisia – Dacia block. Thus, the accretionary
wedge needs to be located on the lower plate. Model 3
combines model 1 and 2 by assuming two subduction
zones (Fig. 15c), but this means that twice the amount
of convergence is needed. Furthermore, two different
accretionary wedges as well as two different foredeeps
should develop. This is not the case. In model 4,
subduction took place at the same locality as in model
1 (Fig. 15d). However, continental collision was
followed by delamination along a horizontal tear
either at Moho depth or somewhere deeper in the
lithospheric mantle (Girbacea and Frisch, 1998).
Thus, the subducted slab rolled back towards the east
until it reached its present location about 100 km from
the Miocene suture zone.
The different models can be tested with gravity
modelling. We excluded model 3 from this test due to
its unrealistic, double amount of convergence. For the
other three models, we assumed that the tectonic
evolution was the same until continental collision
(as described in Sections 4.4 and 4.5), but the postcollisional evolution and the location of the slab are
different. In model 1, with subduction beneath the
Miocene suture, misfit between modelled and measured Bouguer anomaly data occurs in the foredeep
and in the Transylvanian basin (Fig. 16a). Model 2
shows a better fit for the load basin (which is in this
case not equivalent to the Transylvanian basin), but
the foredeep and the accretionary wedge of the model
are now located too far east, so that they produce a
misfit in this region (Fig. 16b). Model 4 gives the
best-fitting results with respect to the Bouguer anomaly data (Fig. 16c) and it also best explains the vertical
movements in the SE-Carpathians in the last few
million years. Fission track data indicate that the
southern part of the Eastern Carpathian accretionary
wedge started to uplift at about 5 Ma (Sanders et al.,
1999). At the same time, the southern part of the
foredeep basin (Focsani) subsided at its maximum rate
(Artyushkov et al., 1996; Matenco et al., 2003). These
77
opposite vertical movements in close proximity are
best explained by delamination of the lithospheric root
beneath the accretionary wedge (causing uplift) and
by a heavy lithospheric root still hanging beneath the
foredeep (causing subsidence). The fact that delamination only occurred a few million years ago is also
responsible for the still low heat flow above the
delaminated region. According to Bird (1979), heat
from asthenospheric material at lower-crustal levels
takes several million years to reach the surface (maximum surface heat flow after ca. 9 Ma for a 30-kmthick crust) and produces a heat flow anomaly of only
15– 25 mW m 2.
Our idealized delamination model results in modelled gravity anomaly curves which fit the measured
curves concerning the position of minima and maxima, but they fail to explain the magnitude of gravity
anomalies (the measured anomalies are smaller than
the modelled ones; Fig. 13c). For a better fit, we
adapted the model geometry to reality by reducing the
depth of the basin and of the Moho (40 km as seen
from the refraction seismic data of Hauser et al., 2001)
and by smoothing the topography (instead of the
vertical ‘‘wall’’ in Fig. 13c). As these processes
(erosion, basin subsidence) are still ongoing, isostasy
might not necessarily be achieved so that we made the
adaption without mass balancing. The resulting model
fits the measured data much better, at least at a larger
scale (Fig. 13d). Since our intention was to study
large-scale lithospheric processes, we are satisfied
with this result. Smaller misfits are caused by
smaller-scale structures that are not the topic of this
study.
5. Discussion
5.1. Modelling assumptions
5.1.1. Time
In our models, time is not included because timedependent deformation processes are not considered.
Isostatic adjustments are assumed to happen instantaneously. Thus, our results have no time scale. Nevertheless, sequences for sedimentation and erosion can
be determined. However, it should be kept in mind
that the sedimentation or erosion processes might be
interrupted by the next event before they reached their
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
final state. In the case of a sediment basin, which
should, for example, reach a maximum depth of 4 km
due to slab pull, subsidence might be interrupted after
2 km because slab break-off eliminated the driving
process for the subsidence. This example makes it
clear that our results give maximum values for the
case that enough time was available to complete each
distinct process.
5.1.2. Coupling between slab and overlying
lithosphere
In our models, we assume a strong coupling
between subducted slab and overlying lithosphere so
that the gravitational effect of the mass excess in the
lithospheric root is transferred to the upper plate and
thus causes subsidence. We make this assumption of
strong coupling even for slabs which descend deeper
than the lithospheric mantle (e.g. Fig. 2c). In this case,
slab and overlying lithosphere are separated in the
deeper regions by asthenospheric material (mantle
wedge) which behaves ductilely at geological deformation rates (ca. 10 15 s 1). This means that the
gravitational effect of the slab might be (partly) transformed into internal deformation of the mantle wedge
instead of causing downward movements of the
overlying crust. Thus, our calculations only give
maximum values for possible vertical movements at
the surface.
5.1.3. Model geometry
We used 2D-models with infinite extent in the
direction perpendicular to the profiles. This approach
is useful to get general results for any subduction or
collision zone. For our example from the SE-Carpathians, this approach gives only a first approximation
because seismic tomography shows that the presentday slab has only a small lateral extent (Martin et al.,
2001). This small lateral extent has no influence on
the vertical movements along the profile because
isostatic disequilibrium was compensated locally.
However, the calculated gravity field would be different, not only because of the reduced mass excess at
depth (smaller slab), but also because the basins
evolving above the slab would be laterally limited.
This would result in a reduced mass deficit. Thus, for
modelling the present-day gravity field in the SECarpathians, a 2.5D model with limited lateral extent
of the slab and the basins would be more appropriate.
On the other hand, it has to be taken into account that
the Carpathian slab was not always as confined as it is
today. During subduction, it had a larger lateral extent
which was successively reduced during collision (slab
break-off) (Sperner et al., 2001). This could be an
explanation for the fact that the Transylvanian basin,
which evolved as a load basin on the upper plate
during subduction, is broader than the Focsani basin
(foredeep basin on the lower plate; Fig. 9) that mainly
developed after continental collision when the slab
was already reduced in its lateral extent.
5.2. General results
Our modelling results give maximum values for
vertical surface movements related to oceanic subduction, continental collision, slab steepening, breakoff, and delamination. Due to the fact that we used
purely kinematic models, no time-dependence is included and consequently we cannot predict any rates
for subsidence or uplift. Nevertheless, our models
reveal characteristic sedimentary, structural and erosional patterns that depend on the tectonic evolution
of the study area. These patterns can be used to
identify subcrustal processes like oceanic subduction,
slab steepening or break-off. We used this method to
study the Tertiary tectonic evolution of the SE-Carpathians and found results that are in agreement with
other data (e.g. uplift data from fission-track analysis). For a more sophisticated test of the general
results, the method should be applied to other subduction or collision zones (e.g. Aegean region, Dinarides, Apennines).
A characteristic of most of the studied processes is
wave-like vertical movements at the surface. In some
cases, like oceanic subduction, this wave simply
produces an expanding basin (only subsidence), while
other cases, like slab steepening, result in more
complicated movement patterns (subsidence followed
by uplift). In addition to different vertical movement
directions (up or down), the wavelength and amplitude of the wave might also vary with time. Slab
steepening, for example, results in a lithospheric root
limited to a smaller area, but with a larger thickness.
The result is a smaller, but deeper basin at the surface.
In contrast, in the case of delamination, the area
remains constant, but the slab length increases so that
the basin size remains constant while the depth
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
increases with ongoing delamination. In both cases
(slab steepening and delamination), the basin axis
migrates parallel to the horizontal component of slab
movement (rollback) so that depocentre migration
might be an indication of processes acting in the
mantle (if other causes for this migration can be
excluded).
In contrast to the sedimentary record, which might
reflect the (long-term) tectonic evolution of an area,
gravity data give only information about the presentday situation. Nevertheless, they are useful for testing
different tectonic evolution scenarios by comparing
the resulting (final-stage) gravity data with measured
values. In general, the gravity data (especially the
free-air anomaly) reflect vertical surface movements
because basins and uplifted regions have a more
pronounced influence on gravity than deep-seated
structures like a subducted slab. The Vrancea slab,
as revealed by seismic tomography, has a modelled
positive gravity effect of about 20 mGal (Hackney et
al., 2002), while the measured gravity anomaly (Bouguer and free-air) reaches negative values of about
120 mGal. These negative anomalies originate
mainly from the Focsani basin.
5.3. SE-Carpathians
5.3.1. Type of subduction
For a long time, it has been debated whether
subduction beneath the Carpathians was oceanic or
continental (e.g. Girbacea and Frisch, 1998; Girbacea
and Frisch, 1999; Linzer, 1996a,b; Pana and Erdmer,
1996; Pana and Morris, 1999). Doubts about oceanic
subduction are mainly based on the fact that no
ophiolithes are preserved from the Miocene Carpathian subduction. However, slab retreat produced a
low-stress subduction zone in the Carpathians so that
the entire oceanic lithosphere disappeared into the
mantle; only the sedimentary cover was (partly)
scraped off and thrust onto the European foreland.
Indications for slab retreat come from the occurrence
of back-arc extension in the Pannonian basin, while at
the same time compression took place along the
Carpathian arc (thrusting in the accretionary wedge).
This simultaneous occurrence of compression and
extension is typical for slab retreat and low-stress
subduction zones (e.g. Doglioni, 1992; Royden,
1993).
79
Our modelling results show that continental subduction would result in an uplift of 2.8 km in the
region where the crust was doubled due to underthrusting of the upper plate by the subducting plate
(Fig. 6f; Section 3.2). However, topography of the SECarpathians is mainly in the range of 0.5 –1.5 km
above sea level. Highest values occur in the Outer
Carpathians, i.e. in the accretionary wedge, and not in
the basement of the upper plate as should be the case
if continental subduction had taken place. Furthermore, we would expect a large BA anomaly with long
wavelength and high amplitude (Fig. 6f). Only the
latter is valid for the SE-Carpathians with its strong
but narrow BA anomaly (Fig. 12b). In addition, new
refraction seismic data show that the Moho in the SECarpathians reaches a maximum depth of ca. 40 km
(Hauser et al., 2001, 2002), thus indicating that only
minor crustal thickening took place. Similar results
were achieved by Lillie et al. (1994) when they
compared the situation in the Eastern Alps and the
Western Carpathians. They concluded from their
modelling, in comparison with measured gravity and
topography data, that the continental collision in the
Carpathians was a ‘‘soft’’ one with only 50 km of
underthrusting. In summary, we conclude that the
evolution of the SE-Carpathians can be best explained
by oceanic subduction followed by ‘‘soft’’ continental
collision with only minor crustal thickening.
5.3.2. Correlation of sedimentation and tectonics
We interpret the sedimentation in the Transylvanian basin (at least the Tertiary part of the basin) to be
related to the subduction process beneath the Carpathian arc. Sedimentation in the Transylvanian basin
shows a hiatus from the end of the Cretaceous to the
beginning of the Eocene. This Palaeocene hiatus is
probably related to a change in the large-scale tectonic
situation. Sediments older than the hiatus correlate
with the late Cretaceous subduction which took place
beneath the Alpine – Carpathian arc. Then, at the end
of the Cretaceous, the buoyant (thinned) continental
margin of the European platform started to enter the
subduction zone in the Eastern Alpine area (Stampfli
et al., 2002). In the Carpathian region, oceanic lithosphere was still available in an embayment in the
southern margin of the European continent (Sperner et
al., 2001). However, subduction of this oceanic lithosphere was blocked by the continental collision in
80
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
the Alps which is responsible for a slower northward
movement of the Adriatic microplate and a nearly
fixed position of the slab. In the late Eocene, slab
break-off occurred in the Alps (Wortel and Spakman,
1992) and most probably also removed the Carpathian
slab. Thus, the Carpathian slab was released and a
new phase of subduction started. We interpret the
sediments younger than the Palaeocene hiatus (i.e.
Eocene and younger) to be correlated with this younger phase of subduction.
Tertiary subduction started with a flat dip (ca.
21j) which can be calculated from the position of
the subduction-related volcanism relative to the plate
boundary (Fig. 11). At this time, thick sediments
were mainly deposited in the western part of the
Transylvanian basin (above the thickest part of the
lithospheric root; Fig. 13a). During the Middle
Miocene, the depocentre migrated SE-ward, which
we interpret as a consequence of slab steepening
with SE-ward migration of the lithospheric root (Fig.
13b). Slab steepening is proven by its present-day
vertical situation which is visible from intermediatedepth seismicity and seismic tomography (Martin et
al., 2001; Sperner et al., 2001). Removal of the
lithospheric root from the western parts of the
Transylvanian basin led to uplift and erosion (since
12 Ma) of the Apuseni mountains and neighbouring
parts of the Transylvanian basin (Sanders et al.,
1999). The NE –SW orientation of the present-day
slab (Martin et al., 2001) and the distribution of
Sarmatian sediments (13.0 – 11.5 Ma) in the Transylvanian basin (Sandulescu et al., 1978) are parallel to
each other (Fig. 17). This indicates that slab orientation was already NE – SW during Sarmatian subduction retreat, as proposed in the geodynamic
model of Sperner and the CRC 461 team (in press).
Uplift and erosion of the southern part of the
Eastern Carpathians started at about 5 Ma (Sanders
Fig. 17. Sketch map of the SE-Carpathians showing the distribution of Miocene sediments in the Transylvanian basin (Sandulescu et al., 1978)
and (preliminary) results from seismic tomography at 110 – 150 km depth (Martin et al., 2001). Note the parallelism (NE – SW) between the
high-velocity body and the axis of Miocene sediments in the basin.
B. Sperner et al. / Tectonophysics 382 (2004) 51–84
et al., 1999), while at the same time subsidence
reached maximum rates in the southern part of the
Focsani basin (Artyushkov et al., 1996; Matenco et
al., 2003). We correlate these movements with a
further SE-ward movement and steepening of the
subducted slab towards its present-day vertical position (Fig. 13b). Its location beneath the foreland
indicates delamination along the Moho or at a deeper
level (Girbacea and Frisch, 1998) (Fig. 13c). The
vertically hanging slab produced a deep, symmetrical
basin in the foreland (Fig. 13c). This geometry is in
accordance with the shape of the Focsani basin,
which is characterized by symmetrical basin flanks
and still active normal faults along the basin margins
(Matenco et al., 2002). Uplift along its western flank
during the Quaternary (Matenco et al., 2002) indicates that SE-ward rollback (delamination) is still
active, thus moving the slab together with the
‘‘subsidence-uplift wave’’ further SE-ward. The symmetrical shape of the younger parts of the Focsani
basin is in contrast to ‘‘normal’’ foredeep basins
which show an asymmetric deepening of the basin
towards the accretionary wedge (i.e. towards the
plate boundary) due to flexural bending of the
foreland. The latter is caused by loading of the end
of the lower plate (slab pull) resulting in this
asymmetrical shape. Our interpretation is that the
state of (asymmetrical) flexural bending is, in the
Carpathians, already over because the formerly separated lithospheric blocks were welded during continental collision. They now react as one plate which
was loaded in the middle by the subducted lithosphere so that a symmetrical basin developed.
As discussed previously (end of Section 4.6.), we
adapted the Moho and the basin depth above the slab
to better fit the gravity anomaly data. Several reasons
are possible for the difference in Moho and basin
depths between the idealized and reality-adapted
model (Fig. 13c,d): (a) the slab might be shorter, (b)
the density contrast between slab and asthenosphere
might be smaller, (c) flexural rigidity might prevent a
larger subsidence, (d) the limited lateral extent of
structures like the Focsani basin and the subducted
slab, and (e) the slab has already started to decouple
from the overlying crust so that isostatic rebound led
to uplift of the region above the slab (Sperner and the
CRC 461 team, in press). The last possibility is a
realistic scenario that is supported by levelling data
81
that show uplift for this region (Joó, 1992; Popescu
and Lazarescu, 1988; Zugravescu et al., 1998).
6. Conclusions
Isostasy is one of the fundamental processes shaping the surface of the Earth. Mass deficit or mass
excess at the surface (e.g. ice sheets) or in the
subsurface (e.g. crustal roots of mountain belts) are
often responsible for vertical movements. Our modelling of oceanic subduction shows that the mantle
part of the lithosphere can trigger significant surface
movements on the upper (overriding) plate. After the
end of subduction and depending on the post-collisional evolution of the slab (steepening, delamination,
break-off), the lower plate might also be affected by
mantle-induced surface movements. The often wavelike vertical surface movements result in uplifted areas
and sedimentary basins with characteristic internal
structure (e.g. depocentre migration) which can be
used to reconstruct the tectonic evolution, especially
with respect to lithospheric processes active during
and after subduction.
For our example from the SE-Carpathians, we
propose a model that is consistent with the evolution
of various basins (Transylvanian basin—load basin,
Focsani basin—foredeep basin), with the present-day
position and dip angle of the slab as seen in seismicity
and seismic tomography, and with the gravity data.
The model assumes Tertiary subduction which started
with a shallow dipping slab (evolution of the Transylvanian basin on the overriding plate) followed by
slab steepening and (later) delamination, so that the
present-day position is subvertical and beneath the
foredeep (evolution of the Focsani basin on the
foreland). Our model gives a plausible and simple
explanation for the evolution of the Transylvanian
basin, a basin whose evolution was, until now, rather
enigmatic.
Acknowledgements
This work was done under the auspices of the
Collaborative Research Centre ‘‘Strong Earthquakes’’
(CRC 461) which is financed by the German Science
Foundation (DFG). Additional financial support from
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B. Sperner et al. / Tectonophysics 382 (2004) 51–84
the DFG was provided through a separate research
project (Sp 561/1 –2). B.S. highly appreciates the
hospitality experienced during her stay at the
Department of Geosciences at Oregon State University, Corvallis (USA). Thanks to all the researchers
who collected the data for the gravity map of Romania
during the past 30 years. Detailed reviews by Ron
Hackney and Hermann Zeyen are highly appreciated;
they significantly improved the manuscript.
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