Geophysical Journal International
Geophys. J. Int. (2013) 194, 1–17
Advance Access publication 2013 April 23
doi: 10.1093/gji/ggt059
Jannis Makris,1 Joanna Papoulia2 and Tamara Yegorova3
1 GeoPro
GmbH, Hamburg, Germany. E-mail: [email protected]
of Oceanography, Hellenic Centre for Marine Research, Athens, Greece
3 Institute of Geophysics, National Academy of Sciences, Kiev, Ukraine
2 Institute
Accepted 2013 February 11. Received 2012 December 24; in original form 2011 March 15
SUMMARY
A 3-D density model of Greece was developed by gravity modelling constrained by 2-D
seismic profiles. Densities were defined from seismic velocities using the Nafe & Drake and
Birch empirical functions for the sediments, crust and upper mantle. Sediments in the North
Aegean are 6 km thick, and are deposited in transtensional basins developing by dextral strike
slip motion of the North Anatolian Fault. The Cyclades, central Aegean Sea, are free of
sediments. South of Crete, in the Libyan Sea, sediments are approximately 11 km thick. At
the western Hellenides sediments of up to 8 km thickness have been accumulated in basins
formed by crustal bending and southwestwards thrusting of the Hellenic napes. At a deeper
crustal level variations of crustal type and thickness cause density variations explaining large
part of the observed gravity field. The North Aegean domain is characterized by a 24-km-thick
continental crust, including sediments, whereas the western Cyclades, in central Aegean area,
have a slightly thickened crust of 26 km. Crustal thicknesses vary between 16 km in the deep
Ionian and Cretan Seas to 40 km in the western Hellenides. In western Crete crust is 30–32 km
thick, thinning eastwards to only 26 km. The deep Ionian basin, the Mediterranean Ridge, as
well as most of the Libyan Sea are underlain by oceanic crust. In western Turkey the crust
thickens from 30 km along the coast to 34 km to the interior. A third deeper level of density
variations occurs in the upper mantle. Subduction of the oceanic lithosphere below the Aegean
continental domain destabilizes the thermal field, uplifting the isotherms by convection and
conduction below the Aegean Sea. Consequently, volume expansion of the upper mantle and
lithological changes reduce its density and depress the gravity intensity. This low density–
velocity upper mantle extends from the Sporades islands in the North Aegean to the Cretan
Sea, occupying the space between the cold subducted Ionian oceanic lithosphere and the
Aegean continental Moho. Upper mantle densities vary from 3.24 g cm–3 in the Aegean area
to 3.29 g cm–3 below western Greece and the Ionian and Libyan Seas.
Key words: Gravity anomalies and Earth structure; Continental margins: convergent;
Dynamics: gravity and tectonics; Crustal structure; Europe.
1 I N T RO D U C T I O N
Crustal deformation in Greece is controlled by two main tectonic
processes: extension in the Aegean Sea and compression along the
western Hellenides. Transtensional basins are developing in the
northern, central and southern parts of the Aegean domain, due to
dextral wrench faulting associated with the North Anatolian Fault
(e.g. Armijo et al. 1999). The ongoing intense deformation along
the Aegean shear zones is expressed by intense seismic activity (e.g.
Makropoulos 1984; Papazachos & Papazachou 1997). The western
Hellenides are deforming by thrusting, normal and wrench fault
C
ing, and westwards movement of the Hellenic napes due to crustal
shortening. This is expressed by high seismicity and intense tectonic deformations of the upper crust and sediments (e.g. Aubouin
et al. 1976; Makris 1978; McKenzie 1978; LePichon & Angelier
1979; Jacobshagen 1986; Papoulia & Makris 2010).
To understand the regional crustal structure and the tectonic relation between the Aegean region and the western Hellenides, we
developed a 3-D density model of the crust and sediments and reconstructed the density distribution in the crust and upper mantle.
Density models for the crust of Greece, based on gravity modelling, were previously published by Makris (1977), Makris &
The Authors 2013. Published by Oxford University Press on behalf of The Royal Astronomical Society.
1
GJI Geodynamics and tectonics
A 3-D density model of Greece constrained by gravity
and seismic data
2
J. Makris, J. Papoulia and T. Yegorova
Figure 1. Topographic and bathymetric features of Greece and the surrounding Seas. Location of refraction seismic profiles is marked in red, blue, and black
lines. Profiles P1 to P12 were used to constrain the 3-D density model. White lines indicate locations of vertical density cross-sections of the 3-D model (shown
in Figs 11 and 12).
Stobbe (1984), Tsokas & Hansen (1997) and Tirel et al. (2004).
They were, however, either constrained by regionally sparse seismic data or were obtained by inversion of the gravity field using
potential field techniques of non-uniqueness. Thus these models
were rather crude or ambiguous. Recently obtained data on crustal
thickness from active and passive seismic experiments provide a
good background for reinterpreting the gravity field and developing a new 3-D density model of the crust and upper mantle of
Greece. These include information from receiver function analysis
of P- and S-wave time arrivals (Saunders et al. 1998; Li et al. 2004;
3-D density model of Greece
Sodouti et al. 2006) and 2-D velocity models along seismic profiles, shown in Fig. 1, from active seismic experiments onshore and
offshore mainland Greece (De Voogd et al. 1992; Bohnhoff et al.
2001; Makris et al. 2001, 2003; Makris & Broenner 2001; Makris
& Papoulia 2009a,b, 2011; Makris 2010; Papoulia & Makris 2010).
The present paper on the 3-D density model for the crust and upper
mantle is based on a combination of 2-D and 3-D gravity modelling,
constrained by 2-D velocity models and empirical relations between
P-wave velocities Vp and densities ρ. It has been developed using
the following procedures:
(i) The geometry of the crust and sedimentary basins was constrained by 2-D velocity models.
(ii) Densities for gravity modelling were derived by converting
velocities of the seismic models into densities using the empirical
relations of Birch (1960, 1961) and Nafe & Drake (1963).
(iii) The Talwani et al. (1959) method for computing the gravity
effect of a 2-D density model is computed and compared to the
observed field. Uncertainties in estimating density from velocity,
can easily vary within 5–10 per cent, which limits the 2-D density
modelling results within this range.
(iv) From the 2-D density cross-sections a starting 3-D density
model was constructed and tested by 3-D gravity calculations. The
3-D model is described by a set of prisms of constant (5 × 6 ×
0.25 km3 ) dimensions.
(v) The gravity effect of the prisms is calculated using the Nagy
(1966) formula. For a given position of the model the effects of
all prisms are added. The validity of the model is provided by
comparing the calculated and observed gravity. The 3-D model is
modified until the best fit between the calculated and observed fields
is achieved.
In this paper we will present a 3-D density model for Greece, for
the area between 18–30◦ E and 34–42◦ N, following the procedure
described above.
2 G R AV I T Y S U RV E Y S A N D
REDUCTIONS
For evaluating the gravity data and compiling them in gravity maps
of Free Air (g ) and Bouguer anomalies (g ) it is essential to
have sufficient coverage of stations that permit to define the gravity
anomalies with the required resolution. In the following, we present
the gravity data used for developing the 3-D density model of Greece
that were obtained from terrestrial and marine surveys.
2.1 The terrestrial data
From 1971 to 1973 a regional gravity survey covered all Greek
mainland with a grid of 4 × 5 km2 . Details of these studies were
published by Makris et al. (1973) and Makris (1977). This survey
was organized by the University of Hamburg. The participants of
this survey were the Institute of Geophysics (IfG) of the University of Hamburg, the Institute of Geology and Mineral Resources
of Greece (IGME), the Institute of Geodesy of the Technical University of Athens, and the Institute of Physical Astronomy of the
University of Thessaloniki. Topographic data and absolute gravity
values of the first order gravity network of Greece were provided
by the Geographic Department of the Hellenic Army (GYS) that
also supported the field operations and helped to obtain necessary
permits. From 1975 to 1995 most sedimentary basins were covered by dense gravity surveys of one gravity station per square
3
kilometre. These studies were associated with geothermal and oil
exploration programmes, and were performed by IGME, the Public
Petroleum Corporation of Greece (PPC) and IfG. Finally, during
1999–2000, IfG, in cooperation with the University of Bochum, established 2000 gravity stations on Crete (Casten & Snoupek 2006;
Makris & Yegorova 2006). All terrestrial studies are connected to
the first order gravity network of Greece, which is linked to the
European calibration line. The gravity link between Frankfurt International Airport and Athens Hellenicon International Airport, East
Air Terminal, station A, gave an absolute gravity value of 980.058,
28 ± 0.08 mGal for station A.
2.2 The marine data
During 1965–1972 the Osservatorio Geofisico Sperimentale of Trieste (OGS) performed a regional gravity survey of the Western
Mediterranean Sea and the area of the Ionian, Libyan and Aegean
Seas. They used a Graf-Askania Sea gravimeter type GSS-2, and
Loran-C navigation (Allan & Morelli 1971; Finetti & Morelli 1973;
Morelli et al. 1975a,b,c). The Aegean and Libyan Seas were revisited by the R/V SONNE and R/V AEGAEO in the years 1982–1984
in a cooperation of IFG with HCMR. A linear type marine gravimeter Bodensee KSS-30/31 and GPS supported navigation were used.
The Aegean Sea was covered by a network of profiles, spaced at
1.5–3.0 km in the Cretan Sea to 3.0–5.0 km in its central and northern parts. For the region south of Crete, due to the poor coverage
of ship born data, we used a 2 min grid, satellite derived Free Air
gravity anomalies (Sandwell & Smith 1997). The eastern part of the
Mediterranean Sea was surveyed by the University of Cambridge
(1975) using a similar Graf-Askania gravimeter with that of OGS
and with very similar data accuracy and survey procedures.
The distribution of the regional gravity data onshore and offshore the Mediterranean countries and Greece are presented in
Fig. 2. These data, including a regional gravity map of Turkey, were
provided by UNESCO, International Oceanographic Commission,
and were used in compiling complete Bouguer gravity maps of
1:1 000 000 scale (Makris & Morelli 1994; Makris et al. 1998). It is
obvious that the regional gravity information from onshore Turkey
can only be used to estimate regional features of crustal thickness.
Sediments and lateral variations of density in the mantle cannot be
resolved with sufficient accuracy by these gravity data.
The accuracy of the gravity surveys depends strongly on the
time they were carried out. Accuracy of ±0.1 mGal was derived
for onshore Greece surveys. This is due to the method used to define altitudes. These were obtained by combining micro altimetry
and absolute altitude values from TP—triangulation and levelling
points. The marine data collected by the Askania gravimeter, due
to cross coupling effects have ±3 mGal accuracy, while those collected by the Bodensee linear gravimeter and GPS navigation, are
characterised by a better accuracy (<1 mGal). This aspect of the accuracy of the anomalous gravity field should be kept in mind when
we deal with the accuracy of modelling and the best fit between
observed and computed gravity values.
Data were reduced to Free Air (g ) and complete Bouguer
anomalies (g ) using the reduction techniques described by Jung
(1961), Makris (1971) and Makris et al. (1998). g and g anomalies were computed by the formulas:
g = g − γ0 + δgF ,
and
g = g + δgr + δgB ,
4
J. Makris, J. Papoulia and T. Yegorova
Figure 2. Distribution of gravity stations in Greece and other Mediterranean countries used to compile Bouguer gravity maps.
where γ 0 is the theoretical gravity according to the International
Formula of 1967, g is the measured gravity adjusted to the firstorder gravity net connected to Athens Hellenicon Airport, East Air
Terminal Station A: g = 980.05828 ± 0.08 mGal, δgF is the freeair reduction: 0.3086 (hs − ho ) mGal with: hs = altitude of the
gravity station in m, ho = reduction level = 0 m, δgB is the Bouguer
reduction; the Bouguer masses are reduced spherically to Hayford
Zone O2 (0–166.7 km) with uniform density of 2.67 × 103 kg m–3
according the formula of Cassinis et al. (1937), δgT = topographic
reduction, computed in a system of geographic coordinates with the
formulas of Jung (1961) and Nagy (1966). Constant density ρ =
2.67 × 103 kg m–3 was also used for the topography according to
the international standards. Water density for the marine areas is
ρ = 1.05 × 103 kg m–3 .
The Bouguer anomaly map presented in Fig. 3a, is interpolated
with 10 mGal isoline interval and shows strong lateral variations.
The g values of +300 mGal in the Ionian Sea changes gradually northwards to negative gravity values of −120 mGal at the
western Hellenides. A steep horizontal gradient of approximately
4 mGal km–1 was observed in the transition from the Ionian Sea to
the islands of Cephalonia and Zakynthos off western Greece. The
g values along the northern part of western Hellenides strikes
NW–SE, following the trend of the main Mountain ranges of Pindos (see Figs 3a and b). To the south of the line defined by the
Sperchios (SB in Fig. 3b) and Amvrakikos (AG) Gulfs, the field
trend changes to nearly N–S orientation towards the Patraikos Gulf
(PG; see Fig. 3b). This orientation change is probably caused by
the two dextral faults of Cephalonia (CF) and Andravida (AF) that
displace this part of the Hellenides westwards (see also Papoulia
& Makris 2010). At the Peloponnese coast close to the Andravida
fault, the field reorients itself to a NW–SE direction until the Laconikos Gulf (LG), southeast Peloponnese (see Fig. 3b), and by
slightly eastwards bending is terminated at the western coast of
Crete.
At the Cretan Sea, north and northeast of the island of Crete, the
trend of the g values follows the E–W direction of the bathymetrytopography and reaches a maximum value of +160 mGal (see
Fig. 3a). In this part of the Aegean domain the Moho is encountered
between 16 and 18 km. South and southwest of Crete Moho reaches
26 km depth, while below western Crete it is encountered at more
than 30 km depths (Makris 1977; Bohnhoff et al. 2001). The crust
consists of the continental crust of Crete and the subducted Ionian
oceanic crust, which is still attached to it.
The 80-mGal isoline at the Cyclades is NW–SE oriented, parallel to the western Hellenides. In the northern Aegean area a NE–
SW-oriented gravity high of 100 mGal follows the North Aegean
Trough. At the Chalkidiki Peninsula and the Thermaikos Gulf the
g anomalies have NW–SE orientation and range from +60 to
0 mGal. Western Turkey is characterized by negative g values
that gradually decrease eastwards, reaching −75 mGal values at the
eastern margin of the area, at 30◦ E. The pattern of the anomalous
gravity field follows the surface topography of continental Turkey
with a gradual increase of negative values eastwards. In summary,
the Bouguer anomalies in Fig. 3a distinguish the main structural
domains and elements of the Hellenides and therefore express
the mass distribution within the sediments, crust and upper mantle and provide information on the tectonic processes that formed
them.
3-D density model of Greece
5
Figure 3. (a) Complete Bouguer gravity map of Greece and surrounding Seas. Reference level is 0 m and mass reductions, topographic and Bouguer, were
computed on a spherical Earth to Hayford zone O2 (0–166.7 km). A uniform density of 2.67 × 103 kg m–3 was used for the mass reductions. Contour interval
is 10 mGal. For further explanations see text. Rectangle indicates a detailed map shown in (b). (b) Enlarged map of the Bouguer field presented in (a) (area
coordinates: 36.30-42.00N, 18.30-23.00E). Gravity trend changes along the Pindos Mountains towards Peloponnese, due to the southwest displacement of the
Hellenides caused by the Cephalonia (CF) and Andravida (AF) dextral strike slip faults.
6
J. Makris, J. Papoulia and T. Yegorova
3 S T RU C T U R E O F T H E C RU S T F R O M
S E I S M I C R E F R A C T I O N D ATA
Velocity cross-sections of the crust and sediments have been derived from seismic data collected in two different periods. The
older data were acquired during 1971–1974 (Makris & Vees 1977;
Makris 1978). More data were collected in 1980s in the North
Aegean Trough (Ginzburg et al. 1987) and offshore Corfu (Makris
& Thiessen 1984); they are shown in Fig. 1 by blue lines. These data
were acquired by explosive sources and provided the regional geometry of the Moho. They were also used to constrain the thickness
of sediments in the North Aegean and the Ionian Seas. Accuracy
of crustal thickness and thickness of the sediments obtained by
two-point ray tracing techniques is ±4–5 per cent.
A new systematic study of the sediments, crust and upper mantle by wide aperture reflection/refraction seismic observations was
initiated in the mid-1990s (red lines in Fig. 1). As energy source,
tuned airgun arrays were used with volumes varying from 44 to
120 l. Results were presented in several publications mentioned in
the introduction. Two examples of these velocity cross-sections are
shown in Fig. 4, delineating the structure of sediments and crust in
the Ionian, and Cretan-Libyan Seas. From the Zakynthos profile, it
is obvious that the oceanic crust of the Ionian Sea (African crust)
is subducted below the western Hellenides. The continental Moho
below mainland Greece is approximately 35 km deep. Nearly, 8 km
of sediments mainly composed by metamorphic limestones were
mapped (Makris & Papoulia 2009a, 2011). The NE–SW oriented
profile between Libya and Crete shows a thickened continental crust
below western Crete, with a Moho at 30 km. The subducted oceanic
crust follows the continental Moho morphology. Sediments thicken
from south Crete to the Mediterranean Ridge, where they reach their
maximum thickness of approximately 11 km (Makris & Broenner
2001). Accuracy of crust and sediments thickness is 2–3 per cent.
This accuracy improvement was achieved by using dense nets of
recording stations. Thus the Common Station and Shot Gathers better constrained velocity and thickness of the models than the older
data.
Finally, a crustal velocity model on a seismic line across the
Mediterranean Ridge (De Voogd et al. 1992; black line in Fig. 1),
and a density model constrained by these data (Truffert et al. 1993)
were also used to develop the initial 3-D density model.
In addition to the refraction seismic profiles, estimates of crustal
and lithospheric thickness from Sodouti et al. (2006) for 65 locations
in the Aegean region, and three crustal values from Saunders et al.
(1998) for western Turkey obtained from P- and S-wave receiver
function analysis, were considered and compared with our results.
4 3-D DENSITY MODEL CONSTRAINED
B Y S E I S M I C D ATA A N D T H E B O U G U E R
G R AV I T Y F I E L D
Gravity modelling of the 3-D density distribution was performed using the GRAVMAG software developed for computing 3-D gravity
and magnetic fields (Tchernychev & Makris 1996). This program
is based on the Talwani et al. (1959) algorithm for calculating the
gravity effect of a prism. GRAVMAG is a tool, which allows computing potential fields from a model represented by a 3-D grid of
identical rectangular prisms (see also Makris & Yegorova 2006).
For our model of Greece we used the prism of 5 × 6 × 0.25 km3
size. The cells (prisms) are divided into classes of densities defining
density domains constrained by seismic and geological information.
This allows generalizing the model, avoiding assigning density val-
ues at each cell individually. It permits a significant simplification in
computing and editing the model, by calculating the gravity effect
of a given prism as a function of distance for unity density. The
gravity effect of a body at a given point at the surface is obtained
by computing the distance of each individual prism to the given
surface point. Using the gravity effect of each prism multiplied by
the body density, the total gravity effect for a given surface position
is then derived by summation:
δgi = f (δg pr /ri ),
g B = ρi
δgi ,
where δg p is the gravity effect of the elementary prism, ri are its
Cartesian coordinates and gB is the gravity effect of the entire
body.
GRAVMAG uses relative densities for the gravity calculations,
which implies the definition of a reference density for the model.
This density is defined as the average density of an ‘average’ column
of the model. It includes all densities from sea water to the upper
mantle. The average density used for the 3-D computations is 3.0 ×
103 kg m–3 . Side effects were also considered by calculating the
gravity effect of each layer by extending it beyond the limits of
the model. A simple correction for an infinite slab was applied
for that. This correction smoothes the intensity of the gravity field
at the edges of the model, which otherwise would be terminated
abruptly causing gravity picks of large amplitudes. Computations
were accomplished on a flat Earth, by transforming the Geographic
into Cartesian coordinates using the GMT software (Wessel & Smith
1998).
In Table 1 we present the average densities used for the 3-D
gravity computations. These are 2.4 × 103 kg m–3 for the sediments,
2.6 × 103 kg m–3 for compacted lithologies and meta-sediments
(i.e. lower limestones mapped onshore western Greece and Crete),
2.59 × 103 kg m–3 for the metamorphic rocks of the Cyclades to
2.90 × 103 kg m–3 for the oceanic crystalline crust, and 3.24–3.29 ×
103 kg m–3 for the upper mantle.
These data are plotted in Fig. 5 that presents the relationship between density and P-wave velocity. It is obvious that the density of
igneous rocks depends on the mean atomic weight of the mineralogical components of the rocks. It is therefore possible that density
values can be up to 5 per cent different from the mean density for
a given velocity. We have adopted a mean value as indicated by
the colour bars in Fig. 5. The length of the bars shows the range
of variability of the velocity–density dependence. It is obvious that
in the sediments density variations can exceed 5 per cent that was
estimated for the igneous rocks.
The gravity effect of this initial model was computed and compared with the observed field (Fig. 3a). Local differences between
calculated and observed values ranged initially from 50 to 100 mGal.
In order to improve the model, we sliced it in a series of 19 crosssections: 14 EW-oriented and spaced approximately at 30 on latitude and 5 NS-oriented, spaced at 2◦ on longitude. Density and
structure of each section was modified by directly comparing it
to the next closest seismic line and 2-D gravity calculations were
performed. Thus, the differences between computed and observed
values were minimized.
The optimized 2-D density models were further used to improve
the 3-D density model. This was achieved by modifying the density
perpendicular to the computed 2-D lines for a number of prisms,
and recalculating their gravity effect. This procedure was repeated
until the best fit between computed and observed gravity anomalies
was obtained. During this process it became obvious that the crust
and sediments alone cannot explain the regional gravity trend. Mass
3-D density model of Greece
7
Figure 4. Examples of two seismic lines from the Ionian Sea to western Greece (Makris & Papoulia 2009a) (upper part) and from the Libyan margin to Crete
(Makris & Broenner 2001) (lower part). These type of crustal velocity models obtained by wide angle seismic experiments are used to constrain 2-D density
models. For further explanations see text.
deficiency had to be introduced in the mantle below the Aegean Sea
by decreasing its density by 0.05 × 103 kg m–3 compared to that of
the surrounding area. In this way we satisfied the long wave length
of the Bouguer gravity field. This has also been reported by Makris
(1977), who showed that the regional trend of the Bouguer gravity
field over the Aegean region was 50 mGal higher than that observed
if only crustal effects are considered. The calculated gravity anomalies are presented in Fig. 6, and the residuals between calculated and
8
J. Makris, J. Papoulia and T. Yegorova
Layer
Vp (km s–1 )
ρ × 103 (kg m–3 )
Sediments
Sediments
Compacted sediments (including the layer of lower limestones)
>4–4.5
4.6–6.0
2.4
2.6
Consolidated crust
Metamorphic rocks of Cyclades
Metamorphic rocks of Crete
Continental crust (Aegean Sea, mainland Greece, western Turkey)
Oceanic crust (Ionian Sea, Mediterranean Ridge, Libyan Sea)
4.8–6.9
6.0–6.3
6.0–6.7
6.7–7.0
2.59
2.72
2.82
2.90
Upper mantle
Table 1. Average densities of the model layers used for the 3-D gravity modelling. Colours correspond to
those used in the density models.
Upper mantle 1
Upper mantle 2 (Ionian Sea, western Greece)
Upper mantle 3 (Aegean Sea)
7.9–8.0
8.0
7.7–7.8
3.27
3.29
3.24
we calculated for the various regions of the model the gravity effect
by allowing density changes of 5 per cent. In Table 2 we present
the gravity effects for the Ionian Sea, the Aegean Sea and the central Aegean area. As can be seen in the table an increase of 5 per
cent in density causes gravity effects in the order of 22–33 mGal.
Such differences between observed and calculated gravity are too
large and would cause unacceptable discrepancies in the geometry of the model as constrained by the seismic data. We discussed
that the observed gravity in the worst case of older marine data is
within ±3 mGal accurate. The calculated gravity effects resulting
from density variations of 5 per cent are far beyond the accuracy
of the observed gravity data or the crustal constrains posed by the
seismic models derived from active seismic experiments.
5 THE 3-D DENSITY MODEL
OF GREECE
Figure 5. Empirical relation between P-wave velocities and densities for
sediments, igneous and metamorphic rocks. Sediment values were taken
from Nafe & Drake (1963), while crustal and upper mantle rocks are from
Birch (1960, 1961). Numbers in brackets indicate mean atomic weights
of the various lithologies. Colour bars mark the velocity range for various
densities (see also Table 1). Further explanation in the text.
observed anomalies are presented in Fig. 7. The residual map is an
objective way to present the quality of the synthetically calculated
model since it presents the absolute discrepancies between observed
and calculated field. The regional features of observed (Fig. 3a) and
calculated (Fig. 6) fields are in very good agreement. Short wave
length anomalies were not always modelled, since they are only partially constrained by seismic observations. Spacing between seismic
lines is too large for constraining local gravity effects.
In order to estimate errors that may occur due to the nonuniqueness in the empirical functions between velocity and density
The 3-D density model of Greece is presented in Figs 8, 9 and 10.
Fig. 8 shows the thickness of sediments. A major deposition centre
is located in the North Aegean Trough, along the Saros and Sporades basins, with 7.5–8 km thickness. These basins are NE–SW
oriented and were developed by dextral transtension in continuation
of the North Anatolian Fault. In the Chios basin further south, sediments are about 5 km thick. This basin of NW–SE orientation is not
influenced by active dextral transtension and is mainly controlled by
normal faulting (Makris & Papoulia 2009b). The largest sedimentary basins, mapped onshore western Greece, are NW–SE oriented
and exceed 10 km. They were deposited in a domain of crustal subsidence and were significantly deformed by westward sliding and
thrusting (see also Jacobshagen 1986). Further southwest, in the
Ionian Sea at water depth of nearly 4 km, sediments as thick as
6 km were mapped; they overlay the backstop of thin continental
crust that extends from the Hellenic Trench to the Mediterranean
Ridge. Southwest and southeast of Crete, along the Mediterranean
Ridge, sediments thickness exceeds 10 km. The two sedimentary
depots are separated by a basement uplift of the continental crust
extending for more than 100 km south of Crete (see also Makris &
Yegorova 2006).
The crust-mantle boundary (Moho; Fig. 9), shows very strong
lateral variations. We obtained Moho depths exceeding 40 km below
western Greece, while under the Cretan Sea, at the thinnest part of
the crust, Moho was encountered at16 km depth. The North Aegean
3-D density model of Greece
9
Figure 6. Calculated Bouguer gravity map from the 3-D density model constrained by results of wide angle active seismic experiments. The regional features
of the map are in very good agreement with the observed gravity field (Fig. 3a). Contour interval is 20 mGal.
domain is characterized by continental crust with a Moho at 24 km.
The Cyclades Massif separates the North Aegean domain from the
Cretan Sea. Moho at the western Cyclades is at 26 km thinning
westwards to only 18 km below Kythnos and the eastern Myrtoon
Sea (Makris & Papoulia 2009b).
The Western Hellenides along the Pindos Mountains, as stated
above, have the thickest crust in Greece (Moho depth exceeds
40 km). This zone extends southeastwards to central and south
Pelaoponnese, and is terminated at Central Crete. Moho below
western Crete is at 30–32 km, while eastwards it is encountered
at 26-km depth. To the southwest, in the deep Ionian Sea (at 4 km
water depth), the crust thins significantly and Moho is located at
16 km depth. Thickness of sediments ranges between 5 and 7 km.
Crust is of oceanic type, as revealed by active seismic observations
obtained by De Voogd et al. (1992) and Papoulia & Makris (2010).
In western Turkey the crust thickens and Moho is found at 30 km
depth along the coastal line and at 34 km inland. These values agree
with the results obtained by receiver function analysis of P and S
waves (Saunders et al. 1998) that defined crustal thickness at three
different locations.
Another presentation of the 3-D density model of Greece is given
in Fig. 10, showing the Moho surface (lower part), basement (middle part), and the topography-bathymetry (upper part). The model
shows strong correlation between the topographic-bathymetric features and the basement structure. The North Aegean depressions
at the Sporades and Saros basins (for location see Fig. 8) correlate
well with the westwards extension of the North Anatolian Fault. The
north Evia–Sperchios–Amvrakikos lineament separates the crust
of central and southern from that of northern Greece. It terminates
the rapidly south–southwestwards expanding Aegean domain (∼3.5
cm yr–1 ) relative to Nubia (Reilinger et al. 2010) from the crust of the
northern Hellenides, north of the Sperchios and Saros basins. The
most intense deformations of both basement and crust are observed
south of the Cephalonia fault and affect the complete Hellenic Arc
up to southwest of Crete (Papoulia & Makris 2010). The Rhodos
Abyssal plain, is part of the very thick sedimentary basin mapped
southeast of eastern Crete (see also Makris & Yegorova 2006), that
is terminated at southwestern Turkey (Fig. 10).
To better visualize details of the density distribution that cannot
be seen in the previous maps, we have extracted six cross-sections,
10
J. Makris, J. Papoulia and T. Yegorova
Figure 7. Residual gravity field obtained by subtracting the calculated gravity effect of the 3-D model (Fig. 6) from that observed (Fig. 3a). Contours are
in mGal. Residual amplitudes were mainly caused by discrepancies at the level of sediments since the seismic lines are sparsely distributed and the high
frequency part of the gravity field neglected. Discussion in text.
three in NS and three in EW orientation (white lines in Fig. 1) from
the 3-D density model. The density cross-sections are presented in
Figs 11 and 12. The vertical columns in the models show the velocity
structure as defined by the seismic profiles (shown in Fig. 1). They
were used to constrain the density model. Starting from the east,
along line P1SN, sediments exceed 10 km thickness in the eastern
Libyan Sea, thinning rapidly under the Dodecanese and thickening
again to more than 5 km under western Turkey (see Fig. 11). The
crust south of Karpathos is thin continental and is truncated by the
Stravo sinistral strike-slip fault. It reaches 24 km under Karpathos,
and is strongly deformed due to crustal shortening. Under western
Turkey it gradually thickens to 30–32 km depth. The ‘soft’ Aegean
mantle mobilized due to the subduction of the Ionian Oceanic slab
below the Aegean domain has a density of 3.24 × 103 kg m–3 and
extends below the Dodecanese volcanic area and under part of
southwestern Turkey, Along line P2SN, sediment thickness ranges
from more than 10 km under the Libyan to 5 km under the North
Aegean Sea. The Cyclades at the central Aegean Sea are free of
recent sediments. The thinned continental crust under the Cretan
Sea thickens to about 28 km under the western Cyclades, thinning
again under the North Aegean Sea to about 23 km. It is strongly
deformed at the transition of the Aegean domain to the oceanic crust
of the Libyan Sea. North of the Sporades basin and its continuation
under mainland Greece the crust thickens rapidly and Moho is at
30 km. The Aegean ‘soft mantle’ terminates at the Sporades and
Saros basins.
Line P3SN extends from Cyrenaica (North Africa) to Albania.
The oceanic crust below the Ionian Sea and the Mediterranean
Ridge is terminated southwest of Zakynthos Island. To the north
of Zakynthos the continental crust thickens rapidly to nearly 40 km
3-D density model of Greece
11
Table 2. Errors in estimates of gravity effect of model layers.
Model layers
Upper and lower bounds (H)
of the layer (km)
H
(km)
Velocity range (Vp)
(km s−1 )
Vp
(km s−1 )
Density range (ρ)
(g cm−3 )
ρ
(g cm−3 )
Estimated gravity
(mGal)
4–11
11–17
7
6
3.8–4.2
6.8–6.9
0.4
0.1
2.38–2.42
2.90–2.94
0.04
0.4
11.7
10.4
Total: 22
3–15
15–21
12
6
4.0–4.4
6.8–6.9
0.4
0.1
2.38–2.42
2.90–2.94
0.04
0.04
20
10
Total: 30
0–2
2–26
2
24
5.9–6.0
6.5–6.6
0.1
0.1
2.57–2.60
2.81–2.84
0.03
0.03
3
30
Total: 33
Ionian Sea
Sediments
Crystalline crust (oceanic)
Mediterranean (Libyan) Sea
Sediments
Crystalline crust (oceanic)
Central Aegean Sea (Cyclades)
Cyclades rocks
Crystalline crust (continental)
Figure 8. Thickness of sediments (in km) derived from 3-D gravity modelling for Greece and surrounding area.
Moho depths under Albania. The mantle has a normal Vp velocity value of 8 km s–1 and a density of 3.29 × 103 kg m–3 , compared to the ‘soft mantle’ that was encountered below the Aegean
domain.
Lines P1WE to P3WE (Fig. 12) confirm the previous models. The
continental crust thickens significantly from 32 km under western
Crete to 40 km under the Pindos Mountains. The Aegean domain
‘soft mantle’ has its maximum thickness below the Cyclades and
12
J. Makris, J. Papoulia and T. Yegorova
Figure 9. Moho depth map (in km) derived from 3-D gravity modelling for Greece and surrounding area.
the Aegean volcanic belt, significantly thinning to the north, and is
terminated at the Sporades basin. As previously stated, the upper
mantle density is separated in two different domains: that of the
Ionian and Libyan Seas, and the western Hellenides with densities of 3.29 and 3.27 × 103 kg m–3 and that of the Aegean Sea of
low density (3.24 × 103 kg m–3 ) and velocity (soft mantle). This
is caused, as previously stated by subduction of the Ionian oceanic
lithosphere below the Aegean domain. This process is accompanied
by sediments transported at depth in the upper mantle. These sediments containing water trigger the mobilization and partial melting
of the asthenosphere causing the magmatic and volcanic activity at
crustal and subsurface levels. The Aegean volcanic belt developed
as a consequence of this process.
6 D I S C U S S I O N A N D C O N C LU S I O N S
The 3-D density model of Greece and surrounding regions, derived
by 3-D gravity modelling constrained by 2-D velocity models, reveals a very complex geometry of the mass distribution of the
Aegean Microplate and the western Hellenides as a consequence
of recent and on-going tectonic processes. The deep basins of the
north Aegean Sea and these of the Cretan Sea are associated with
the dextral transtension of the North Anatolian Fault and the tectonic processes of western Turkey (see Fig. 13). The fact that the
North Aegean basins are much deeper than those of the Cretan
Sea may be explained by the longer exposure of the former to the
deformation along the North Anatolian Fault. It is also possible
that the westward extension of the tectonic lineaments from western Turkey to the Aegean domain, are focused in the northern part
along a few strictly linear features, whereas in the south they are
dispersed over a wide zone. It could also be seen that the axial part
of the Sperchios–Amvrakikos basins relates to a distinct crustal lineament separating the southern from the northern part of Greece.
The seismic activity over north Evia and the Sperchios valley is
reoriented along NW–SE trending faults compared to the NE–SW
trend observed in the North Aegean Sea (Papoulia et al. 2006). This
NW–SE trending system is connected at the Amvrakikos gulf with
the Cephalonia dextral strike slip fault. The Hellenides, south of
this line of major crustal change, are deforming much faster than
northern Greece as seen also by GPS observations (Kahle et al.
1998; Reilinger et al. 2010). Furthermore, the density model of the
Aegean domain, revealing the low-density upper mantle, is in good
agreement with the surface heat flow density values available for
the Aegean areas (Cermak et al. 1977; Cermak 1979), which are
two to three times higher than those observed in the Ionian Sea. The
isotherms below the Aegean domain are significantly elevated due
to conduction and convection processes (Makris & Stobbe 1984)
and relate to the low density asthenosphere that has intruded the
Aegean domain southwards of the North Aegean Trough. Crustal
structure, orientation and thickness of the sedimentary basins,
3-D density model of Greece
13
Figure 10. 3-D density model for Greece and surrounding area developed by 3-D gravity modelling (view from southeast). From top to bottom we present the
topography-bathymetry, base of sediments and depth to Moho.
derived by the 3-D modelling, are also confirmed by Tirel et al.
(2004). In particular, good agreement is obtained for the North
Aegean area and the central Cyclades. However, in the Cretan Sea
Tirel et al. (2004) show a crustal thickness of 22–23 km, which
contradicts our results showing that the Moho depth below central
Cretan Sea is not deeper than 16 km.
In our density models, the accuracy of which easily varies within
5–10 per cent, due to the ambiguous relation between velocities and
densities, we have not shown sediments overriding the down going
oceanic slab. The seismic data are not of sufficient resolution for
resolving sedimentary structures attached to the subducted oceanic
lithosphere. Also, the GPS observations suggest that the subducting
oceanic slab under most part of the Aegean domain is decoupled
from the overlying continental crust (Reilinger et al. 2010). This
implies that significant amount of sediments, including Miocene
evapourites, are involved in the subduction process. Consequently,
strain accumulation is very slow and earthquakes with very high
magnitudes, like that of 365 AD southwest of Crete, occur rarely.
Shaw & Jackson (2010) estimated that such events may occur at
a frequency of 1 event per 1000 yr, which obviously requires a
decoupled down-going oceanic slab from the overriding crust of
continental Hellenides.
Isostacy is another critical issue that can be addressed by the density model we have developed. Previous studies of Makris (1977,
2010), confirmed by the present models, have shown that isostacy
is strongly disturbed at crustal levels due to intense tectonic deformation. The eastern Ionian area and the western Hellenides are
controlled by significant mass deficiency due to the accumulation
of thick sediments in the down bended, shortened and thrusted continental crust. At deeper crustal levels and within the upper mantle,
below a depth of 30–40 km, isostatic balance is established except
for the collision front between the backstop and the Mediterranean
Ridge, and for the transition of the backstop to the thrusted and
folded Alpine sediments further east. The Aegean in contrast, being
an extensional area, tends to isostatic balance at more shallow levels. This is possible since the high temperature upper mantle permits
rapid readjustment of the stretched crustal units to the correct level
of subsidence, satisfying Airy’s model of isostacy.
14
J. Makris, J. Papoulia and T. Yegorova
Figure 11. Three NS-oriented density cross-sections obtained by 3-D gravity modelling (location of lines shown in Fig. 1). The observed and calculated
gravity are shown on the upper part of the drawings and the density models are presented in the lower part. Densities are in 103 kg m–3 for each geological unit,
derived from Fig. 5 and Table 1. Bars indicate location of 2-D seismic lines used to constrain the 3-D model and their crustal structure.
3-D density model of Greece
15
Figure 12. Three EW-oriented density cross-sections obtained by 3-D gravity modelling (location of lines shown in Fig. 1). Observed and calculated gravity
are shown at the upper part of the drawings and the density models are presented in the lower part. Densities are in 103 kg m–3 for each geological unit, derived
from Fig. 5 and Table 1. Bars indicate locations of 2-D seismic lines used to constrain the 3-D model and their crustal structure.
16
J. Makris, J. Papoulia and T. Yegorova
Figure 13. Thickness of sediments (in km) and main tectonic elements of Greece and surrounding area. Abbreviations are used for: NAFZ, North Anatolian
Fault; NAT, North Aegean Trough; SaB, Saros basin; SB, Sperchios basin; AG, Amvrakikos gulf; CF, Cephalonia fault; AF, Andravida fault; LG, Laconikos
gulf; NMTF, North Mani Transverse Fault; PA, Pre-Apulia zone; I, Ionian zone. Limits of the Ionian zone, Pre-Apulia zone and the Mediterranean Ridge are
indicated. Further discussion in text.
AC K N OW L E D G E M E N T S
Evaluation and modelling was supported by GeoPro Hamburg. We
acknowledge two anonymous reviewers for their constructive comments that helped to improve this paper.
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