TECTONOPHYSICS
Tectonophysics 272 (1997) 269-290
ELSEVIER
Lateral variations in lithosphere strength in the Romanian Carpathians:
constraints on basin evolution
Anco Lankreijer a,*, Victor Mocanu b, Sierd Cloetingh a
a Institute of Earth Sciences, Vrije Universiteit, Amsterdam, Netherlands
b Department of Geophysics, Bucharest University, Bucharest, Romania
Received 1 August 1995; accepted 6 June 1996
Abstract
Lithospheric strength profiles constructed through extrapolation of rock mechanics data and constrained by crustal
geophysics data, demonstrate the existence of important spatial variations of thermomechanical properties in different
domains of the Romanian Carpathians, the Pannonian basin and the Transylvanian depression. These models show
important lateral variations in lithospheric rigidity along the Southern and Eastern Carpathian foreland, compatible with
inferences from flexural modelling studies. The modelling predicts an absence of mantle strength and the presence
of weak lithosphere in the Pannonian and Transylvanian basins. A NW-SE-trending rheological transect is presented,
connecting the Moesian platform, the Vrancea region of the Carpathians, the Transylvanian basin and the Apuseni
Mountains with the Romanian part of the Pannonian basin. This transect demonstrates the existence of a close correlation
of the inferred spatial variations in mechanical properties of the Romanian lithosphere with observed patterns in
large-scale gravity anomalies and with tectonic units. A notable feature of the theological models is the presence of an
intra-Carpathian weak zone. The combination of a strong foreland lithosphere with a weak lithosphere in the inner part
of the Carpathian/Transylvanian/Pannonian system sheds light on a number of key features observed in the evolution
of sedimentary basins in these areas. Remarkable differences in rheology predictions between flexural models and
depth-dependent theology models indicate that the crustal lithosphere is detached from the mantle part of the lithosphere in
the Carpathian foreland.
Keywords: rheology; lithosphere; basin evolution; Pannonian basin; Moesian platform; Carpathians
1. Introduction
Rheology controls the response of the lithosphere
to basin formation processes (e.g., Vilotte et al.,
1993). Over the last few decades important advances
have been made in the quantification of the thermomechanical properties of the continental lithosphere
* Corresponding author. Fax: +31 20 646-2457.
through the construction of rheological models, constrained by laboratory rock mechanics data (e.g.,
Carter and Tsenn, 1987). The predictions of these
models can be verified by comparison with independent strength measures, for example elastic plate
analogues to model the lithospheric response to surface loads, such as foreland basins (McNutt et al.,
1988). We compare EET estimates derived from
foreland basin modelling with those derived from integration of the thickness of the mechanically strong
0040-1951/97/$17.00 © 1997 Elsevier Science B.V. All rights reserved.
PII S 0 0 4 0 - 1 9 5 1 ( 9 6 ) 0 0 2 6 2 - 4
270
.4. Lankl~'ilcr
¢'t a/. /7i'clom~phwic~ 272 ~1997) 269 290
parts of strength envelopes (Burov and Diamcnt.
1995) and estimates }'or the thickness of the mechanically strong part of the crust based on thermal
properties.
Until recently, flexural modelling studies of the
Carpathian foreland were very rare (e.g., Royden ct
al., 1983a,b), and reliable gravity data were not accessible to constrain foreland flexure studies in these
areas. The recent release of these data, together with
the large body of work carried out through deep
seismic sounding on the structure and composition
of the crust and subcrustal lithosphere in Romania (RS.ileanu et al., 1994: Mocanu and RS.dulescu.
1994: Mocanu et al., 1995), enables the extension
of the approaches sofar explored only in western
and central Europe to the Romanian Carpathians and
surrounding areas.
In this paper we present strength proliles calculated for different segments of the Romanian
Carpathians and Transylvanian/Pannonian basins.
We concentrate on spatial variations in lithospheric
strength in these areas to examine the possible existence of a rheological control on basin evolution in
this region.
The Carpathian-Pannonian area has for long been
a key natural laboratory to validate quantitative
basin formation models (e.g., Stegena et al., 1975:
Horvfith et al., 1975: Sclater et al., 1980: Royden
and D6venyi, 1988; Van Balen et al., 1995: Lankrei.jer et al., 1995: Zoetemeijer et al., 1996: Matenco
et al., 1997). This is largely due to the availability
of an extensive, high-quality database, m particular
concerning the Pannonian basin. The presented theological profiles of the Romanian lithosphere provide
constraints on these models.
2. Tectonic f r a m e w o r k and constraints on crustal
and lithospheric structure
Detailed accounts of the tectonic development
of the modelling area are found in the papers by
Burchfiel (1980) for the Eastern Alpine-Carpathian
system, Royden et al. (1983b), Horvfith (1984, 1993)
and S~ndulescu (1988, 1994) for the CarpathianPannonian region.
The Carpathian mountain chain is bounded by
the European platform in the north and northeast
and by the Moesian, Moldavian and Scythian plat-
forms in the south and east. The Carpathian arc is
formed in a region dominated by collision between
the European, Moesian and Moldavian continental
margins and microplates associated with the colliding Apulian platform. Tomek (1988) suggests that in
late Mesozoic times, the arcuate region now formed
by the Carpathians was similar to the Black Sea.
Subduction and collision closed the small ocean and
formed the Carpathians, comparable to the recent
situation in the Banda Arc (Hamilton, 1979).
Deformation of the Carpathian mountain chain
started in the early Badenian (Burdigalian) and migrated in time from northwest to southeast along the
arc (15-0 Ma) (Sfindulescu, 1975).
Synchronous with compression in the Carpathian
arc, the intra-Carpathian region was subject to
an extensional tectonic regime (e.g., Rumpler and
Horvfith, 1988: Taft et al., 1992; Royden, 1993),
which is expressed by extremely thin lithosphere and
associated high heat flow values in the Pannonian
basin (D6ven3)i and Horv~th, 1988).
Extensional collapse of overthickened and, therefore, unstable crust (Taft et al., 1992: Horvfith, 1993)
and roll-back of the subducting European slab (Royden and Karner, 1984; Royden, 1993) in combination
with lateral escape (Ratschbacher et al., 1991 ) are the
driving mechanisms for formation of the extensional
basin in the Pannonian realm.
The Late Cretaceous to recent tectonic history
of the Transylvanian depression is characterized by
multiple compressional events, separated by minor
extensional events (Ciulavu and Bertotti, 1994). The
Transylvanian depression is regarded as the vertical
superposition of Neogene mostly extensional basins
associated with extended crust, on top of inverted,
Upper Cretaceous-Palaeogene basins (Taft et al.,
1993: Proust and Hosu, 1995). The most recent thermotectonic event resetting the thermotectonic age in
the Transylvanian crust was a phase of pre-SenonJan folding (80-100 Ma) (Demetrescu and Veliciu,
1991). The Apuseni Mountains, bounding the Transylvanian depression in the west, can be regarded as
a Cretaceous ophiolitic sutural belt associated with
Tethyan microplate collision (SS.ndulescu, 1984).
The Carpathian foredeep, located in front of the
Eastern and Southern Carpathians, overlies a basement consisting of Moesian and Moldavian platforms. The sedimentary fill of the foreland basins
A. Lankreijer et al. / Tectonophysics 272 (1997) 269-290
reaches a maximum thickness of around 20 km.
The width of the foreland area changes from 25
km in the north to approximately 100 km in the
bend zone, i.e. the transition between the Eastern
and Southern Carpathian forelands. The foredeep
developed during the later stages of Cretaceous to
Tertiary convergence, contemporaneous with the formation of the Eastern Alps and the Carpathians.
The Outer Carpathians include deep marine flysch
deposits as part of the accretionary complex separating Europe from the colliding terranes (Tomek,
1988). Exploration drilling has shown that rocks
of the European foreland extend southward beneath
the molasse foredeep of the Western and Eastern
Carpathians (Birkenmajer, 1986). Deep seismic profiles show that the platform strata are located beneath
12 to 15 km of Outer Carpathian flysch (Tomek et
ah, 1987, 1989; R~ileanu et al., 1994; Mocanu and
RSdulescu, 1994; Mocanu et al., 1995).
Recent gravity models for the Carpathian foreland
(Szafi~in et al., 1995) suggest that the subducted slab
has been detached and assimilated in the Western
and Eastern Carpathians, whereas in the Southern
Carpathians the crustal slab is still present.
The Moldavian platform represents a Proterozoic
heterogeneous craton, separated from the Moesian
platform by the North Dobrugea orogen of Baikalian
age. The Moesian platform consolidated during the
Cadomian orogeny, and consists of a complex of
reworked Sveko-Karelian and younger schist units
covered with an almost complete section of Cambrian to recent sediments. The thermal age is defined by epi-Hercynian metamorphosis and folding
(1.0-1.6 Ga)(Demetrescu and Veliciu, 1991). Recent
palaeostress analyses document a present-day N-S
compressional stress regime (Bergerat, 1995; Hippolyte and S~ndulescu, 1995), the so-called 'Wallachian' phase, indicating continuing convergence.
The Vrancea area is the locus of many deep
(50-220 km) earthquakes, indicating that brittle conditions prevail at those depths. These earthquakes are
possibly related to the termination of the subduction
process (Fuchs et al., 1979).
The thermal structure of the lithosphere gives an
indirect indication on the thermotectonic age (Sclater
et al., 1980), and on the rheology of the lithosphere.
Heat flow in the Romanian territory ranges from 30
mW m -2 for the lowest values in North Dobrogea,
271
to over 100 mW m -2 in some parts of the Pannonian
basin area. The mean value for North Dobrogea
is 60 mW m -2. The Neogene volcanics zone in the
Eastern Carpathians shows also high heat flow values
(100 mW m-Z). The Transylvanian depression is a
relative cool area with heat flow values below 60
mW m -2. The Moesian foreland shows a transition
in heat flow from <50 mW m -2 in the eastern part
to > 80 mW m -2 in the western part. The Moldavian
platform is characterized by heat flow values around
50 mW m -2 (Veliciu et al., 1985; Demetrescu and
Veliciu, 1991). This already gives an indication on
the lateral variations in lithospheric strength in the
study area; the hot Pannonian basin will be weak,
whereas the foreland areas, which are relatively cool,
will show a strong rheology.
3. Construction of lithospheric strength profiles
Laboratory rock mechanics studies demonstrate
a dependence of rock-strength on temperature and
pressure (e.g., Goetze and Evans, 1979; Ranalli,
1995). Within the mechanically strong part of the
lithosphere, it is possible to define an upper region
where the strength of the lithosphere is defined by
criteria for brittle failure (Byerlee's Law). Creep processes become dominant at temperatures exceeding
roughly half the melting temperature of rock (Carter
and Tsenn, 1987). Therefore, the strength in the
lower part of the lithosphere and the lower parts of
the crust is governed by the temperature distribution.
Dislocation climb and dislocation glide (Dorn Creep)
are the main flow mechanisms occurring in the lower
part of the lithosphere (Goetze and Evans, 1979).
Extrapolation of flow laws and laboratory failure
criteria (i.e., Byerlee, 1978; Brace and Kohlstedt,
1980), adopting estimates for tectonic strain-rates
and thermal gradients at different depths provides
a first-order description for the strength distribution within the lithosphere. For each depth interval
strengths for both brittle and ductile deformation are
calculated, with the lesser of these representing the
limiting strength of the lithosphere at that particular depth level (e.g., Beckman, 1994; Ranalli, 1995;
Burov and Diament, 1995).
Critical input data for the modelling are crustal
composition and the thermal structure of the lithosphere. Although, the construction of strength pro-
,4. Lankreijer et al./Tectonophys cs 272 (1997) 269-290
272
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A. Lankreijer et al./ Tectonophysics 272 (1997) 269-290
files invokes a number of intrinsic uncertainties, resuits of many recent studies support the extrapolation
of micro-physical models from a laboratory scale to
a lithosphere scale (i.e., Ranalli and Murphy, 1987;
Cloetingh and Banda, 1992; Beekman, 1994; Burov
and Diament, 1995). We adopted a four-layer rheological model for the Romanian lithosphere, consisting of a sedimentary layer (where present), a granite
layer representing the upper crust, a diabase layer
representing lower crust and olivine representing the
mantle, for dry rocks. Wet rheology is represented
by (wet) granite, diorite and wet dunite, respectively
(Table 1). Fig. lc illustrates the difference in calculated strength for dry and wet rheology for location
G 1. The strength envelope calculated for a wet rheology shows a detachment at the base of the crust,
whereas dry rheology predicts a coupled behaviour
for crust and lower lithosphere. The EET calculated
from the wet strength envelope is remarkably thinner (14 km) than the EET estimated based on dry
rheology (25 km).
An extensive geophysical data base is available
for major parts of Romania. Recent studies on lithospheric and crustal thicknesses (Enescu, 1987; R~dulescu, 1988; Horvfith, 1993; R~ileanu et al., 1994),
heat flow (Demetrescu, 1982; Demetrescu and Veliciu, 1991), the geothermal regime (Veliciu et al.,
1985; Demetrescu and Andreescu, 1994) and crustal
composition (Cornea et al., 1981; Demetrescu and
Veliciu, 1991; R5ileanu et al., 1994) are used here as
constraints in the construction of rheological profiles
in different tectonic domains of Romania.
Geothermal gradients along the profile are based
on interpretations by Sollogub et al. (1993) and
Demetrescu et al. (1981), and were calculated for the
other locations, using heat flow (after Negut, 1984;
273
Veliciu et al., 1985) and isotherms (Sollogub et al.,
1993; Demetrescu and Andreescu, 1994). The lithosphere-asthenosphere boundary (Tm) is assumed to
coincide with the 1300°C isotherm which is different
than the assumption used for this by Sollogub et al.
(1993) where the lithosphere-asthenosphere boundary roughly matches the 1200°C isotherm (Fig. 2).
Fig. ld shows that these differences in Tm do not
change the geotherm drastically and therefore do not
have a significant effect on strength. The distribution
of temperature in the crust and upper part of the
mantle lithosphere have a more pronounced effect on
strength predictions. Another liability of the adopted
interpretation of isotherms (Sollogub et al., 1993) in
the foreland area and the Vrancea zone is unable to
explain the occurrence of deep earthquakes at depths
where temperatures exceed 800°C.
Observations on strain rates (Carter and Tsenn,
1987; Van den Beukel, 1990) indicate a range of
10-17 < k < 10-12 S-1. A strain rate of 1 x 10-15 s -1
is adopted, in our modelling, since it is commonly
observed for extensional and compressional settings
(Carter and Tsenn, 1987). Fig. la,b shows variations
in strength profiles with varying strain rates resulting in different EET estimates. Higher strain rates
yield higher predicted strengths, and result in a mechanical coupling of upper and lower crust in the
strength envelope for location G1 (Fig. lb). Strainrates are typically assessed within the accuracy of
one order of magnitude. Such uncertainties in estimation change the strength by no more than 10%.
Differences between dry and wet rheologies are incorporated in the powerlaw creep laws (Table 1),
yielding lower strengths for wet rheologies. Fig. 5d
illustrates the effect of dry and wet rheology on the
strength predictions.
Fig. 1. (a) A four-layer rheological model consisting of a sedimentary layer, a granite layer representing the upper crust, a diabase layer
representing lower crust and olivine representing the mantle, for dry rocks. Wet theology is represented by (wet) granite, diorite and wet
dunite, respectively (for failure criteria for these rocks see Table 1). Strength envelope for location Trl (Transylvanian depression) along
the profile, adopting different strain rates, varying from k = 10-13 s -1 to 10-17 s -1 . Decoupling between crust and mantle disappears
with lower strain rates. (b) Strength envelope for location GI (Baraolt Mts.) along the profile for different strain rates, varying from
= 10 -13 s -1 to 10 -17 s -1 . Midcrustal decoupling disappears when strain rates decrease below 10-16 s -1 . (c) Strength envelope for
location G1, showing the difference between dry and wet rheology. The 'wet' strength envelope shows a detachment at the base of the
crust, whereas the 'dry' rheology predicts a coupled behaviour for crust and lower lithosphere. The calculated differences in EET are
listed in Table 2. (d) Strength envelope for location G1, showing the difference in predicted strength between a 1200°C and 1300°C
definition of the base of the lithosphere. The different geotherrns are shown. As can be seen from these strength envelopes, the difference
in strength is negligible.
/~
R - 1
1 +fl(R-
R-1
1)
strike slip faulting
for thrust faulting
for normal faulting
Powerlaw creep: ocreep = (k/Ap)V"exp[ Ep/nRT]
c~ =
R
R-I
Brittle failure: O-brittle : ctpgz(l -- ).)
I
2700
50
0.25
1.9
140
7.94 × 10 16
granite
wet
s-
m s ?
J mol 1 K i
W m 2
°C
Units
R = I(l +.L2) j / ~ - AI 2
2700
50
0.25
3.3
186
3.16 × 10 -26
Density
Young's modulus
Poisson's ratio
Powerlaw exponent
Powerlaw activation energy
Pre-exponential constant (powerlaw)
Failure~Creep functions
granite
dry
Petrology
,o
E
v
n
Ep
Ap
Upper crust
I,ithology
7),~
f,
i'
L
q.,
9.81
8.314
30-100
1300
0.6
10- J5
~0.35
Acceleration of gravity
Universal gas const.
Surface heatflux
Temperature base lithosphere
Static friction coeff,
Strain rate hydrostatic pore fluid
Factor (Pw/P)
g
R
Value
Definition
2900
70
0.25
3.05
276
6.31 x 10 2o
diabase
dry
Lower crust
Table 1
Material properties used for rheology models; after Carter and Tsenn (1987) and Goetze and Evans (1979)
2900
90
0.25
2.4
212
1.26 x 10 16
diorite
wet
3300
70
0.25
4.5
535
7.94 x 10 IS
dunite
dry
Mantle
3300
70
0.25
3.6
498
3.98 × 10 -25
dunite
wet
kg m ~
GPa
kJ tool
P a N s-J
Units
5t x )
%
4~
275
A. Lankreijer et al./ Tectonophysics 272 (1997) 269-290
Pannonian
Basin
ApuseniMountains
Trensylvanian Depression Eastern Carpathians
Fore-deep
Northern
Dobrogea
WNW
540
P1
500
450
A1
400
350
TR3°_°
1
250
H1 2OOGIDet01so
loo
Detl
Det2
Det3
VR1
v15O
ESE
D2°
Fig. 2. Lithosphericcross-section through the North Dobrogea orogen, Carpathian foreland, Vrancea area, Transylvaniandepression,
Apuseni Mountains and the Pannonian basin, showing lithospheric structure (after Demetrescu and Veliciu, 1991). Triangles mark
locations of calculatedstrengthprofiles(Fig. 1).
Based on the above presented geophysical data
we have constructed a rheological profile along a
NW-SE-trending cross-section connecting the Moesian platform, the Carpathian foreland belt and the
Transylvanian and Pannonian basins (location given
in Fig. 3). In addition we have calculated strength
profiles for a number of positions in different segments of the Carpathian fold belt and the Transylvanian and Pannonian basins in order to examine
along-strike variations within these tectonic units
(Fig. 3).
Estimates for the effective elastic thickness (EET)
of the lithosphere are obtained using three different
techniques. Foreland basin modelling yields values
for EET directly coupled to flexural rigidity (i.e.,
Zoetemeijer et al., 1990). Based on the observation
that crustal rocks lose their strength at temperatures
exceeding 300-600°C, the depth of, for instance,
the 400°C isotherm can be used as another rough
indicator of the EET (Watts et al., 1982; Cloetingh and Banda, 1992). Integration of the thickness
of the mechanically strong parts of the strength
envelope, obtained through rheological modelling,
yields a third estimate for EET (Burov and Diament,
1995).
4. Strength variations along a NW-SE profile
through Romania
In the following section, the strength distribution calculated along the profile Crisineu Cris-Galati
(Fig. 2) is described from southeast to northwest.
Fig. 4 shows the calculated strength envelopes for
twelve different positions along this profile. In the
cool, eastern part of the profile, including Northern Dobrogea (D2), the Eastern Carpathian foredeep
(V1) and the Vrancea Mountains (VR1), we predict the strongest lithosphere. The EET distribution
(Fig. 5c) shows corresponding high values (50 km)
for this area. The adopted geothermal gradient (after
Sollogub et al., 1993) is based on heat flow measurements, which can be strongly influenced by the
blanketing effect of the extremely thick sedimentary
sequence covering this particular area.
The fault zone, at location 25 km (between D2
and V1) (Fig. 2), offsetting the base of the crust,
corresponds to a minimum in the strength envelope
(Fig. 4) at a depth of 35-45 km. Our models are
not in agreement with the traditional interpretation
of the continuation of this fault zone from 20 km to
70 km of depth, since the interpreted fault crosscuts
276
A. ixmkreijer et al./Tectonophysics 272 (1997) 269-290
122
0
'2%RAINE
26f
'28
100km
HUNGARY
MOLDA VIA
47
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.-.,,-..,
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.
.
.
"Bucharest
.
M~SIANPLATFORM
BULGARIA
•
Locations strength envelopes
Profile P1-D2
,<Y>~
~ . Neogene Magmatic Arcs
Fig. 3. Tectonic map of Romania, showing the main tectonic units, and location of the lithospheric cross-section (Fig. 2). Dark shading
marks orogenic areas (Dobrogea orogen, Carpathians and Apuseni Mountains), light shading denotes Neogene basins (Pannonian basin,
Transylvanian depression, Eastern and Southern Carpathian foreland basins. Platform areas (Moldavian and Moesian platform) are
represented by intermediate shading. Locations of strength profiles are represented by dots (o). N.D.O. = North Dobrogea orogen, P1 =
Crisineu Cris (Pannonian basin), A1 = Apuseni Mountains, Trl = Medias (Transylvanian basin), H1 = Rupea. V1 = Ramnicu Sarat, D2
= Dobrogea.
the predicted strongest parts of the lower lithosphere
(50-60 km depth) and the strongest part of the crust
(<30 km).
The Eastern Carpathians are characterized by a
thickening of the lower and upper crust in combination with higher geothermal gradients (Fig. 2.).
The strength calculations predict that the lower lithosphere looses its strength progressively from east to
west along the profile in this area (VR1-G1) (Fig. 4).
The calculated strength envelopes for sites DET1,
DET0 and G1 predict a very weak crustal part of
the lithosphere, in combination with a very limited
amount of strength predicted for the mantle part of
the lithosphere. This strength minimum (Fig. 5a) can
also be observed in EET estimates based on the depth
to the 400°C isotherm (Fig. 5c). The thinning of these
EET values is not so pronounced as the decrease
in EET values obtained through integration of the
strong layers in the strength envelope, since the
strength minimum is not only caused by relatively
high temperatures, but also by the thickness of the
crust. Crustal rocks at a depth of 40-60 km where,
in this case, temperatures reach values between 400
and 700°C, have almost no strength, whereas lower
,
0 ItBm~X~
,
~
I
~
Carpathians
L
.
m n ~
VR1
DET3 Focsani
Vrancea
DET2 Vrancea
n, v L , , ,
Transylvdepressi
anian on
0,
,
,
Foreland orogen
North
Dobrogea
RamniSarat
cV1
u 0.D0brooea
D2.
Carpathians
For discussion see text.
"--3
-..O
qb
bO
O~
5
ilQ''tii!i'Bar
i''i
° ~aOltG1~c)}
~ ~~~]~tVrancea
DETO''! !
Fig. 4. Strength profiles for selected locations along the profile (Fig. 2), including geothermal gradient model used for strength calculation based on the temperature
distribution as shown in Fig. 2. The Pannonian basin and Transylvania are characterized by a relatively weak rheology, whereas the foreland area shows a stronger rheology.
u
Pannoni
basi
n an Apuseni
mountains
~"'/I ,,0t=~.= o~..=
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=~All~ 11~'
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10 TR1 !!''t'i~
=
Lithospheric strength along NW - SE profile
A. Lankre(jer el al./Tectonophysics 272 (1997) 269-290
278
18f Pannonian
basin
Apuseni Mts.
Transylvaman
depression
16 ~
14f"
12~~
10 F
8~
6
4
Foredeep
/'~
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• o--DRY-compression
--- ~
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500
450
400
0
350
300
250
Distance (kin)
200
150
100
~-100
50
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~ - ~ :~ ~
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Depth to 400 °C EET
DDREETcompr.
DDR EET tens.
40 ~
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200
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J.......
350
i ~ . . . . . . . .
300
25(1
±J
200
: L ~ ~
100
~_/ ~ ~ ~ ~ 0
50
0
Fig. 5. (a) Total integrated strength (surfaces of strength envelopes) for the different positions along the profile (Fig. 2). Note the
differences between values for total integrated strength for dry and wet rheologies and between compressional and tensional strength.
These variations are caused by the dissimilarities in brittle failure criteria (tension-compression) and by different creep laws between wet
and dry rocks in ductile deformation. Table 1 lists used parameters. (b) Bouguer anomalies along the profile (after gravity map Fig. 9b).
Note negative scale. A positive correlation can be made with the total integrated strength (Fig. 5a). Strong lithosphere is capable of
supporting large gravity anomalies. (c) Diagram showing relation between different EET estimates: (1) defined by the depth to the 400 °
isotherm; (2) integrated thicknesses of mechanically strong layers for a depth-dependent rheology, for dry and wet rheology, respectively.
For discussion see text.
lithospheric rocks (olivine, dunite) would have added
considerable strength at this depth interval under the
same conditions.
The fault zone indicated on location 170 km in
the profile (Fig. 2) coincides with a strength minimum in the strength envelope DET! (below 35 km).
This fault zone can be interpreted as an expression
of the mechanical decoupling in the weak zone, between the foreland area (DET2 and eastward) and
the Transylvanian lithosphere (H 1 and westward).
The Transylvanian depression (H1-A1) shows a
strong lithosphere, with relatively large strengths
predicted in the upper part of the lower lithosphere.
The strength profile calculated for a site in the
Apuseni Mountains shows a substantial reduction
in strength in comparison to the Transylvanian de-
pression, due to a slight thickening of the crust of the
Apuseni Mountains.
In the Pannonian part of the profile very high
geothermal ~adients prevail, corresponding to the
presence of thin lithosphere yielding extremely low
strength values.
Fig. 5 illustrates the relation between total integrated lithospheric strength, gravity and EET (based
on the 400°(2 isotherm). The Bouguer maximum
(Fig. 5b) between km 50 and 150 in the foreland
area corresponds to a flexural couple in the downbending foreland due to loading of the lithosphere.
Fig. 5c shows the EET variation along the profile
based on the depth to the 400°C isotherm, coinciding with the mechanically strong part of the crust.
A correlation appears to exist between this EET
A. Lankreijer et al./Tectonophysics 272 (1997) 269-290
and the total integrated strength. EET values derived
by integration of the mechanically strong parts of the
strength envelope (see below) show the same relation
(Fig. 5).
Table 2
Thickness (km) of the mechanical strong layers (hi, h2 and h3
for upper, lower crust and subcrustal part of the lithosphere,
resp.) in the calculated strength envelopes
Location
4.1. Comparison with EET estimates
Values for EET indicate the magnitude of the
elastic rigidity of the lithosphere, expressed as the
thickness of a fully elastic layer floating on a viscous
substratum. The EET strongly affects the behaviour
of the lithosphere when loaded with, for example, a
sedimentary basin. Different methods for estimating
the EET, each have their specific problems. EET
estimates obtained from foreland flexure represent
a palaeo-situation and do not reveal changes perpendicular to the foreland basin strike. Integrating
the thicknesses of mechanically strong layers in the
strength envelope has the disadvantage that smallscale inhomogeneities in rheology inherited, for example, from older faults are not taken into account
potentially changing the overall strength. In addition assumptions on strain rate and estimates on the
thickness of the mechanically strong layers, have
potentially a large impact on the estimated EET.
Estimates of EET based on the depth to the 400°C
isotherm are somewhat empirical.
We adopt a pressure-scaled minimum yield
strength (10 MPa/km) (Cloetingh and Burov, 1996)
to define the thickness of the mechanically strong
layers (hi, h2, h3) for the upper and lower crust and
the subcrustal part of the lithosphere, respectively
(Table 2). It is important to note that the thickness of
h i is difficult to constrain, since very little is known
about the state of stress in the lithosphere. Both
horizontal, in-plane stresses and flexural stresses related to vertical loads may increase stress levels and
therefore decrease hi by 30-50% (E.B. Burov, pers.
commun., 1995). Therefore Eq. 1 gives the upper
estimate of EET. The presented strength envelopes
suggest that in a number of sites the crust is mechanically decoupled from the upper mantle (Figs. 3-8).
The strength envelopes for Buzau, Pitesti, Titu, H1
and V1 do not show such a decoupling for a dry
rheology, yielding a very high thickness for the mechanically strong part of the lithosphere.
Following the method of Burov and Diament
(1995), it is possible to estimate the EET of the
279
Beba
Buzau
Coman
Crai
Falti
Pitest
Streh
Tere
TgMur
Titu
Urzi
Valea
A1
D2
DET0
DET1
DET2
DET3
G1
H1
PI
TR1
V1
VR 1
Dry
Wet
EET (kin)
hi
h2
h3
hi
h2
h3
Dry
Wet
12
63
36
27
32
84
33
12
20
71
33
20
20
36
27
28
43
45
25
80
19
29
79
42
1
0
0
0
0
0
0
1
1
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
20
15
8
0
22
0
0
0
29
0
15
20
0
0
19
22
0
0
2
19
0
22
8
31
31
22
19
72
27
8
17
28
27
16
11
29
16
23
36
38
14
33
14
24
25
24
0
0
0
0
4
0
0
0
0
0
0
0
4
0
0
0
0
0
0
33
0
0
11
7
0
18
15
15
2
0
20
0
0
25
18
0
5
8
3
0
9
12
2
0
0
13
27
11
12
63
38
29
32
84
36
12
20
71
39
20
23
38
27
28
44
47
25
80
19
32
79
43
8
33
32
24
19
72
30
8
17
34
29
16
12
29
16
23
36
38
14
42
14
25
33
25
Effective elastic thickness (EET) in km is calculated for the
combined effect of the multilayer, using Eq. 1. Locations are
given in Fig. 1.
lithosphere based on the multi-layer elastic plate
model. The effective elastic thickness of (EET) the
plate consisting of n detached (non-welded) layers,
with a thickness (h) is:
EET(n)=
,-.., Ah~
(1)
i=1
Comparison between EET estimates derived from
foreland basin modelling (Matenco et al., 1997)
and EET estimates derived from integrating the
thicknesses of the mechanically strong layers in
the strength envelope (Burov and Diament, 1995),
demonstrates that although the estimates obtained by
these two independent methods may differ in magnitude, the overall pattern (EET decreases towards the
west in the Moesian platform) is similar (Fig. 6.).
Comparison of integrated EET estimates obtained
_
i
~ C
.~
~
~ooo
~ h ~ h (MPa)
oCraiova
3CO0
CO
\
,
I,
,
U
IULIJ
mer~th (MPe)
500 l(~
tem~alufe (C)
0
Titu
slre~g~ (MPa)
Pitesti
(C)
IULU
strength(MPO)
U
Urziceni
IULIIj
Moesian tN
Platform
terAoe~e (C)
lera~e
',,,
Buzau
(C]
!
"4
i i
!
tefl'~,~e (C)
3000 0 5(]8106I)
I
0i
fempefo'~e
\ i !
4o0o
0
]eoo
$ffeng~h(MPa)
20oo
~o
,
L Ii
',
0 500 } #
i00!
75
50
ternper~eCC)
0 ~010~
Foreland
Basin
oRamnicu Sarat Vl
Fig. 6. Strength envelopes for the Southern Carpathians and comparison with EET estimates obtained through integration of the thicknesses of the mechanically strong
layers and EET estimates inferred from flexural studies of foreland bending (Matenco et al., 1997). Note the predicted strong lithosphere in the foreland area (Strehaia,
Pitesti, Buzau and Ramnicu Sarat) and the associated low geothermal gradients. Due to flexure and subduction relatively cold material is present at depth, depressing the
brittle-ductile transition in both the crustal and the mantle part of the lithosphere, thus yielding high strengths. Locations are indicated in Fig. 3.
:
Flexural EET (km)
Depth dependent EET Dry (km)
Depth dependent EET Wet (km)
Strehaia
&
-,q
txa
5
2~
281
A. Lankreijer et al. I Tectonophysics 272 (1997) 269-290
Eastern Carpathians
~ ~oFalticeni
I~ ~
N ¸
8O
Focsani VR1
3000
=tn~gth ~
25 ~
i
i L
0 ~O 1110
l~OelS~(O)
K~ Depthdependent EET Dry (km)
[ ] Depthdependent EET Wet (kin)
•
4~
Foreland
Scythian
Platform
~O
-,~
; ~ ,o~
,,,,,,~<~
Ramnicu Sarat V1
oComanesti
ca
I
i
Dobrogea D2
,0,
,o
N|
20
40t
4o
i:
0O
",4
CD
q0|
96
fool
11o ~
IN q080
J
I
0
i~o
~
Ill
~oc ~ ~
lllol
120
3000 ] 500 1000
~
0
I000
~engiil(MPa)
~(C)
~,
2O00
~
3O00
~
0 ~o 1~
ii£ I ~
100
~
~
!
lz 10~0~ n g0t h ( M ~ )1000 a~00 ~00 0 ~o I ~
ie'rtoaa~ (C)
Fig. 7. Strengthenvelopes for the EasternCarpathiansand comparisonwith EET estimatesobtainedthroughintegrationof the thicknesses
of the mechanically strong layers and EET estimates inferredfrom flexural studies of foreland bending. As in Fig. 6. the foreland area
shows the highest values for the yield strength.Locationsindicatedin Fig. 3.
by the latter method with EET estimates based on
the depth to the 400°C isotherm (Watts et al., 1982;
Cloetingh and Banda, 1992) demonstrates a difference in magnitudes with strong similarities in spatial
patterns of EET (Fig. 5).
5. Spatial variations of lithospheric strength in
tectonic units
Below, strength profiles are discussed for all tectonic units examined in this paper (Fig. 3) followed
by a comparison with EET estimates derived from
flexural modelling of the Carpathian foreland.
5.1. Moesian platform and Southern Carpathians
The strength envelope for Craiova (Fig. 6.), located on the boundary of the Moesian plate and
the Southern Carpathian foreland shows intermediate strengths. In Dobrogea the thickened crust leads
to lower strength values. The Southern Carpathian
foreland (Fig. 6) shows a high predicted crustal and
lithospheric strength (Strehaia, Pitesti and Titu), due
to a relatively low geothermal gradient.
In the Moesian platform and the Southern
Carpathian foreland (Fig. 6) we see a decrease in
strength towards the west, and an increase in strength
towards the north. High strength values are predicted
282
A. Lankre(ier et al./Tectonophysics 272 (1997) 269-290
Valea
,o
0
,
4o
,
'~J
,o ~.~,
~o~e~
Chisineu Cris P1
io:
'W
I
........
'~
~
'~
IX5
r,o~
I geothe'm
Ic~o
o
iooo
, ,
,,ooo
J1
)tr~m
lo~
o
I0o~
200o
3OOO
O
,i
mm~C)
Beba Veche
Transylvanian
Basin
Medias TR1
h)
',
i
r
2~
5
!
taro
o
I
1030
I(
I
Teremia Mare
@
k)
i
.__
D~
no
Depth dependent EET Dry (km)
Depth dependent EET Wet (km)
~ ~T
]m
qC~O
0
1030
21~00
I
J I
Pannonian
Basin
Fig. 8. Strength profiles calculated for sites in the Transylvanian and Pannonian basins. The high geothermal gradient in the Pannonia.
basin is associated with very low strength values; subcrustal strength is practically absent in this area. Locations indicated in Fig. 3.
A. Lankreijeret aL /Tectonophysics 272 (1997) 269-290
for the foreland area. Increasing sediment thicknesses in the foreland basin, towards the west also indicate a decreasing flexural rigidity towards the west.
5.2. Scythian platform, Eastern Carpathians and the
intra-Carpathian weak zone
The Eastern Carpathian foreland (Fig. 7.) (locations Urziceni, Buzau and V1 on the profile, and
Falticeni (Moldavia)) exhibit large strengths in relatively thick crust. The strength envelope calculated for Comanesti, located in the outermost Eastern Carpathians, like Vrancea (VR1 on the profile), points to an extremely strong crust and lithosphere. The Vrancea zone with its unusual deep
earthquakes, could be the continuation to depths
of around 200 km of this strong zone. Predicted
strength values decrease towards the north in the
Eastern Carpathian foreland, corresponding to a decreasing flexural rigidity as deduced from sediment
thicknesses in the foreland.
Lack of deep seismic profiles makes it difficult
to track the observed weak zone (DET2-G1) along
the Carpathians. The remarkable intra-Carpathian
strength minimum zone, predicted by the strength
modelling, indicates a zone of decoupling between
a strong foreland area and the Transylvanian depression. It also corresponds to the area where extreme
recent differential movements occur (Popescu and
Dr~goescu, 1986). Furthermore, the location of this
weak zone coincides with a deep crustal fault as
shown by Horvtith (1993) also indicating a decoupiing of the foreland area along this zone. Stress
indicator data by Hippolyte and S~_qdulescu (1995)
documented a decoupling between Neogene thrust
movements in the inner margin and the foreland area
of the Carpathians.
283
displays very low predicted amounts of strength, as
can be observed from inspection of the strength envelopes for Beba Veche, Valea lui Mihai, Teremia
Mare (Fig. 8) and Crisineu Cris (Figs. 4 and 8). The
strength minimum predicted in the Romanian part
of the Pannonian basin and the Transylvanian depression corresponds to thinned lithosphere in these
areas and continues into the Hungarian part of the
Pannonian basin.
5.4. Implications for deep structures
Reflections observed at mid-crustal levels have
been interpreted as large-scale detachment horizons
(R~leanu et al., 1994) at the locations: Cralova, Titu,
Urziceni, Falticeni, Tg. Mures and Teremia Mare.
Our strength models do not support the presence
of mid-crustal detachments for these locations. In
Urziceni and Titu, the depth of these mid-crustal reflectors corresponds with the strongest section of the
crust. In Tg. Mures, Falticeni and Cralova, a slight
strength minimum, when adopting a wet rheology,
is visible (Figs. 6-8), but more likely detachment
occurs in the lower crust where the strength is considerably lower. In Teremia, the strength is primarily
located in the uppermost 10 km of the crust. As a resuit, at this location crustal detachments are likely to
occur at depths deeper than 10 km. These mid-crustal
reflections are probably inherited. However, to generate a weak zone at this level, important changes
in the geotherm or the crustal structure must have
occurred. It is more likely that detachments occur at
or just above the Moho, and that Moho reflectivity
expresses the strength minimum observed at these
depths. The observed general increase of reflectivity
in the lower crust (R,~iileanuet al., 1994) is probably
due to subhorizontal ductile shear zones in these low
strength regions.
5.3. Inner Carpathian basins
5.5. Lithosphere strength map for Romania
The Transylvanian depression (Tg. Mures and
Medias H1 on the profile) show a relatively thin
lithosphere compared to the surrounding area. The
large difference between the strength calculated for
Tg. Mures (Fig. 8) and Medias (Figs. 4 and 8), separated only 50 km, is largely due to the steep gradient
in the lithospheric thickness (Horv~ith, 1993) in this
basin. The Romanian part of the Pannonian basin
Fig. 9a shows a contour map of the total integrated
strength for the Romanian territory calculated by integration of the yield strength over the thickness of the
lithosphere. Interpretation of the pattern, as visualized
by the contourlines for 1, 3, 7 and 15 x 1013 N]m,
follows crustal and lithospheric thickness (Horv~ith,
1993) and heat flow patterns (Veliciu et al., 1985).
(a)
Fig. 9. (a) Contour map of total integrated compressional strength of the lithosphere. The Pannonian basin and the Transylvanian basin are characterized by very weak
lithosphere. The foreland area shows a high strength. See text for discussion.
8.22
magnitude of total integrated compressional
+Strehaia strenght ('10 13 N/m), location
15*1013Nkn
7*1013Ntm
3*1013N/rn
1*1013N/m
Total integrated
compressional strength
2001ml
3GOkm
400km
Fig. 9 (continued). (b) Scheme of Bouguer gravity field (according to data from the Geological Insttute of Romania, completed after Rasca et al., 1995). Large negative
anomalies can be observed along the Southern Carpathian foreland, whereas the intra-Carpathian regions are characterized by much lower negative to neutral gravity
anomalies. Comparison with (a) shows the correlation of areas with a high strength with large Bouguer anomalies. Green colours indicate negative anomalies, red colours
relate to zero Bouguer anomalies.
lOOkm
20°E
I
tj
t~
~o
286
A. lxmkreijer et aL / Tectonophysics 272 (1997) 269-290
A strength minimum is predicted at the location
of the Transylvanian depression and especially the
Pannonian basin. The Carpathian foreland area is the
strongest of all regions analyzed, with values exceeding 15 x 1013 N/m. Platform areas show intermediate
strengths.
Large negative Bouguer anomalies exist along the
Southern Carpathian foredeep (Fig. 9b). A remarkable correlation exists between locations with high
strength and regions with a large negative Bouguer
anomaly ( - 1 3 5 mGal) (Fig. 9b), whereas areas with
low strength correspond with low gravity anomalies.
It is self-evident that the support of large gravity
anomalies requires a strong lithosphere and that very
weak areas are closer to isostatical equilibrium.
In Fig. 10, the EET and thermal age for selected
locations is plotted. A general increase in EET with
thermotectonic age can be observed. The predicted
EET values do not exceed the depth to the 600°C
isotherm as predicted for a cooling model of continental lithosphere (Cloetingh and Burov, 1996). The
upper limits of the boxes are defined by the EET calculated for a wet rheology, whereas the lower limits
indicate the EET for a dry rheology. The high values
of EET in locations V1, Titu and Pitesti are calculated for a coupled rheology, where crust and mantle
/
/
0
Mako
•
Tere
•
~
•
~
T
-
1
r
.
~
T •
,
,
~
,
l ~ -
10dc
Beba
Craiova
P1
~Craiova
3o
L
40 t
W
W
50
60
i
70 i
Titu
/
8°I
90
loo-,,
0
9°°°c
DDR-EET
Flex-EET
7oo°c \~oo°c
1000°C80(
L\i 1
200
400
600
800
1000
thermal age
1200
1400
1600
Fig. 10. Relation between EET and thermal age. EET data in Table 2. Thermal ages: Pannonian basin, 10--25 Ma (17.5 Ma);
Transylvanian depression, 80-100 Ma; North Dobrogea, 250-390; western Moesian platform, 400 Ma; Eastern Moesian platform, 700
Ma; Moldavian platform, 10(03-1600Ma. Some locations have not been plotted due to uncertainties of their thermal age. The isotherms
(after Cloetingh and Burov, 1996) are calculated for a cooling half-space model taking radiogenic heat production into account. The
upper limits of the boxes represent the EET calculated for a wet rheology, whereas the lower limits represent a dry rheology. Left and
right limits represent uncertainties in thermal age. FIex-EET represents EET values calculated for flexural models of the Carpathian
foreland (after Matenco et al., 1997).
A. Lankreijer et aL /Tectonophysics 272 (1997) 269-290
are assumed to be welded together and behave as
one strong layer. From Fig. 10 it appears that the
lithosphere is probably detached at these locations
in the foreland area. So that the EET is closer to
the upper limits of the boxes for V1 and Titu, for a
detached wet rheology. For Pitetsti the thickness of
the detached EET is in the order of the crustal thickness (around 40 km). In the flexural foreland basin
stresses easily overcome the adopted 10 MPa/km,
leading to detachment. This causes the lithosphere to
be detached at these locations.
A good correlation between the calculated EET's
and the 400°C isotherm exists, if the EET values
in the foreland are assumed to be detached. These
temperatures correspond to crustal temperatures, indicating that crustal rocks control the rheology of the
Romanian territory.
6. Conclusions
The presented patterns of lithospheric strength
distribution in the Romanian lithosphere are largely
consistent with independent observations on flexural
rigidity. Although the absolute values of EET differ
depending on the method used, the spatial variations
in the strength pattern appear to be tectonically
significant.
Extrapolation of rock mechanics data constrained
by crustal geophysics and a comparison with EET
estimates points to an absence of subcrustal strength
in Pannonian and Transylvanian basins.
Depth-dependent rheology models, flexural studies and the analyses of the foreland architecture
indicate a strength maximum in the transition zone
of the foreland basins of the Southern Carpathians
and the Eastern Carpathians.
A remarkable weak zone is observed between the
Eastem Carpathians and the Transylvanian depression, indicating decoupling between these two areas.
Decoupling between crust and mantle lithosphere
is predicted for most localities, In HI, V1, Pitesti and
Titu, however, a mechanical coupling between crust
and mantle is predicted when adopting a pressurescaled yield strength of 10 MPa/km, yielding very
high values for EET.
The strength maximum calculated in the foreland area of the Carpathians (Fig. 9a) places important constraints on the evolution of foreland basins.
287
Downbending of the lower plate implies the introduction of relative cool material at greater depths,
thus increasing the strength. This mechanism puts
a limit to the rate of downbending. Bending rates
in excess of the thermal relaxation rate will lead
to increase of strengths in the plate automatically
blocking the movement by the increased flexural
rigidity. The vertical loads associated with the flexural basin, induce flexural stresses that reduce the
strength of a bending lithosphere severely (Cloetingh
and Burov, 1996; Bertotti et al., 1997).
Comparison of the rheology models and flexural
models suggests that the lithosphere in the foreland
basin is largely detached at the upper-lower crust
boundary and the lower crust-mantle boundary. The
increase in stress in the lithosphere due to flexure
facilitates detachment and must be taken into account when applying theological models to foreland
basins.
In general, EET values predicted by flexural models are considerably thinner than those predicted by
the integration of strong layers in the strength envelope. This is probably due to the increased stresslevels present in the flexed lithosphere.
Table 2 shows that in most locations the crust
and mantle lithosphere are mechanically decoupled
(hi < Moho depth) when adopting the pressurescaled yield strength of 10 MPa/km. However, in
Rupea (H1), Rarnnicu Sarat (V1), Buzan, Pitesti and
Titu the thickness of hi is larger than the crustal
thickness, indicating a coupling between crust and
mantle lithosphere when using the pressure-scaled
yield strength, thus predicting large EET values.
If we can extrapolate the calculated present-day
strength distribution to the geological past, we expect
local isostatic compensation for basins overlying a
very weak lithosphere, like the Pannonian basin.
The Transylvanian depression probably overlying a
cooler lithosphere, can partially be compensated in a
flexural manner.
The depth of necking can be defined as the level in
the lithosphere that does not move vertically during
extension in the absence of external forces (i.e.,
gravity, buoyancy forces). This level can either be
the strongest part of the lithosphere (Kooi et al.,
1992), the strongest crustal level (Van der Beek et al.,
1994) or an intermediate level in between two strong
layers (Spadini et al., 1995). The depth of necking
288
,4. l~mkr@er el al. / liwtonophysics 272 ~/ 997) 269-290
in the P a n n o n i a n b a s i n is p r o b a b l y s h a l l o w e r t h a n
that in the T r a n s y l v a n i a n d e p r e s s i o n , s i n c e the strong
layers are l o c a t e d c o n s i d e r a b l y h i g h e r in the crust.
F o r w a r d m o d e l l i n g s t u d i e s o f the a d j a c e n t M a k o
T r o u g h in the H u n g a r i a n part o f the P a n n o n i a n b a s i n
(Van B a l e n et al., 1995) i n d i c a t e low v a l u e s for E E T
( 4 - 8 kin) a n d d e p t h o f n e c k i n g (7.5 kin), w h i c h is
c o m p a t i b l e w i t h o u r findings.
Acknowledgements
V a l u a b l e d i s c u s s i o n s w i t h R. Z o e t e m e i j e r , G.
B e r t o t t i a n d R. v a n B a l e n are h i g h l y a p p r e c i a t e d . Rev i e w s b y G. R a n a l l i a n d a n a n o n y m o u s r e v i e w e r c o n t r i b u t e d to this p a p e r . T h i s r e s e a r c h was s p o n s o r e d b y
the I n t e g r a t e d B a s i n S t u d i e s ( I B S ) p r o j e c t , the Intern a t i o n a l L i t h o s p h e r e P r o g r a m m e ( I L P ) Task Force:
' O r i g i n o f S e d i m e n t a r y B a s i n s ' a n d the P e r i t e t h y s
p r o g r a m m e . V.M a c k n o w l e d g e d partial f u n d i n g by
S h e l l R o m a n i a . N e t h e r l a n d s R e s e a r c h S c h o o l o f Sedi m e n t a r y G e o l o g y p u b l i c a t i o n no. 9 5 0 7 1 2
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