Tectonophysics 734–735 (2018) 96–118
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Tectonophysics
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Tearing, segmentation, and backstepping of subduction in the Aegean: New
insights from seismicity
T
⁎
G.M. Bocchinia, , A. Brüstleb,f, D. Beckerc, T. Meierd, P.E. van Kekene, M. Ruscicd,
G.A. Papadopoulosa, M. Rischeb, W. Friederichb
a
Institute of Geodynamics, National Observatory of Athens, Lofos Nymfon 1, 11810 Athens, Greece
Institute of Geophysics, Ruhr University of Bochum, Bochum, Germany
c
Institute of Geophysics, University of Hamburg, Bundesstr. 55, 20146 Hamburg, Germany
d
Institute of Geosciences, Christian-Albrechts University Kiel, Otto-Hahn-Platz 1, 24118 Kiel, Germany
e
Department of Terrestrial Magnetism, Carnegie Institution for Science, 5241 Broad Branch Road, NW, Washington, DC 20015, United States
f
Institute of Geophysics, University of Stuttgart, Azenbergstraße 16, 70174, Stuttgart, Germany
b
A R T I C LE I N FO
A B S T R A C T
Keywords:
Hellenic Subduction Zone
Hellenic Arc
Nubian slab geometry
Slab tearing
Slab segmentation
Subduction backstepping
This study revisits subduction processes at the Hellenic Subduction Zone (HSZ) including tearing, segmentation,
and backstepping, by refining the geometry of the Nubian slab down to 150–180 km depth using well-located
hypocentres from global and local seismicity catalogues. At the western termination of the HSZ, the Kefalonia
Transform Fault marks the transition between oceanic and continental lithosphere subducting to the south and to
the north of it, respectively. A discontinuity is suggested to exist between the two slabs at shallow depths. The
Kefalonia Transform Fault is interpreted as an active Subduction-Transform-Edge-Propagator-fault formed as
consequence of faster trench retreat induced by the subduction of oceanic lithosphere to the south of it. A model
reconstructing the evolution of the subduction system in the area of Peloponnese since 34 Ma, involving the
backstepping of the subduction to the back-side of Adria, provides seismological evidence that supports the
single-slab model for the HSZ and suggests the correlation between the downdip limit of the seismicity to the
amount of subducted oceanic lithosphere. In the area of Rhodes, earthquake hypocentres indicate the presence of
a NW dipping subducting slab that rules out the presence of a NE-SW striking Subduction-Transform-EdgePropagator-fault in the Pliny-Strabo trenches region. Earthquake hypocentres also allow refining the slab tear
beneath southwestern Anatolia down to 150–180 km depth. Furthermore, the distribution of microseismicity
shows a first-order slab segmentation in the region between Crete and Karpathos, with a less steep and laterally
wider slab segment to the west and a steeper and narrower slab segment to the east. Thermal models indicate the
presence of a colder slab beneath the southeastern Aegean that leads to deepening of the intermediate-depth
seismicity. Slab segmentation affects the upper plate deformation that is stronger above the eastern slab segment
and the seismicity along the interplate seismogenic zone.
1. Introduction
The identification of the Hellenic Subduction Zone (HSZ) as a convergent plate boundary was proposed nearly 50 years ago when the first
seismological evidence supporting the northward subduction of the
African oceanic lithosphere beneath the Aegean continental lithosphere
was obtained from the analysis of fault plane solutions (Papazachos and
Delibasis, 1969), and distribution of deep hypocentres (Caputo et al.,
1970). The tectonic setting and the active deformation in the area were
first described in pioneering works of McKenzie (1970, 1972, 1978)
from the analysis of the seismicity and focal mechanisms in the
⁎
Corresponding author.
E-mail address: bocchini@noa.gr (G.M. Bocchini).
https://doi.org/10.1016/j.tecto.2018.04.002
Received 4 October 2017; Received in revised form 23 March 2018; Accepted 3 April 2018
Available online 07 April 2018
0040-1951/ © 2018 Elsevier B.V. All rights reserved.
Mediterranean region.
Since then numerous geophysical studies have been carried out in
the Aegean that have led to a better image of the subducting slab and its
relationship to seismicity. The Nubian slab, currently subducting beneath the Aegean region, shows a well-developed Wadati-Benioff zone
down to 150–180 km depth (Caputo et al., 1970; Papazachos and
Comninakis, 1971; Knapmeyer, 1999; Papazachos et al., 2000b). In
earlier works, the down-dip limit of the Wadati-Benioff zone seismicity
was considered to indicate the down-dip limit of the subducting slab
(e.g., Le Pichon and Angelier, 1979). Results from seismic body wave
tomography have imaged a subducting slab in the upper mantle (e.g.,
Tectonophysics 734–735 (2018) 96–118
G.M. Bocchini et al.
southeastern Aegean and is defined as significant deviation from a
laterally continuous, highly curved slab, to explain the increasing slab
dipping angle from west to east (Papazachos and Nolet, 1997; Meier
et al., 2007; Brüstle, 2012; Sodoudi et al., 2015). However, the properties and location of the transition between the segments are largely
unresolved.
We refine the geometry of the seismically active Nubian slab proposed by Ganas and Parsons (2009) by using the global International
Seismicity Centre (ISC) catalogue (ISC Bulletin, 2015) and local seismicity catalogues from temporary networks covering the southeastern
Aegean (EGELADOS; Brüstle, 2012), the central Aegean (CYCNET;
Bohnhoff et al., 2004, 2006; Brüstle, 2012) and eastern Crete (LIBNET;
Becker et al., 2010).
Distribution of earthquake hypocentres allows imaging with higher
resolution (< 10 km for temporary catalogues and < 15–20 km for
global catalogues) the geometry of the subducting slab than bodywaves and surface-waves tomography. However, it is limited in depth
by the down-dip limit of the Wadati-Benioff zone. In this study, we use
well-located hypocentres from global and temporary local seismicity
catalogues to revisit the slab complexity and propose a refined 3-D
model of the Nubian slab subducting beneath the Aegean region and
western Anatolia. After providing an overview on the tectonic setting of
the study area (Section 2) by describing the current kinematics (Section
2.1) and the tectonic evolution of the subducting system with a main
focus on western Greece (Subsection 2.2), we introduce the seismicity
datasets on which this study is based (Section 3). The results are presented in Section 4, where we present the refined geometry of the
seismically active subducting slab (based on specific analysis of the
seismicity distribution at the eastern and western termination of the
HSZ and in the southeastern Aegean) and thermal models for the two
identified slab segments of the HSZ. The refined geometry of the seismically active slab is related to properties of the seismicity, upper plate
deformation, and subduction related processes in the discussion
(Section 5).
Spakman et al., 1988; Piromallo and Morelli, 2003) that may penetrate
into the lower mantle down to 1400 km depth (e.g., Spakman et al.,
1993; Bijwaard et al., 1998; van der Meer et al., 2017), suggesting a
significant aseismic extension of the slab. Thus, the down-dip limit of
the Wadati-Benioff zone seismicity is not related to the maximum depth
reached by the subducting slab but to its properties (Meier et al., 2004a,
2007).
The geometry of the eastern and western termination of the HSZ is
still highly debated. At the western termination of the HSZ, the
Kefalonia Transform Fault is related to the transition from oceanic to
continental subduction to the south and north of it, respectively (e.g.,
Papanikolaou and Royden, 2007; Royden and Papanikolaou, 2011;
Legendre et al., 2012; Pearce et al., 2012). Different interpretations
have been proposed to describe the western termination of the HSZ,
including: (1) the presence of a vertical tear along the boundary between oceanic and continental lithosphere (Suckale et al., 2009), with
the formation of a Subduction-Transform-Edge-Propagator fault or
STEP-fault, namely the Kefalonia Transform Fault, above the edge between the two slabs (Govers and Wortel, 2005); (2) the presence of a
slab detachment (i.e., horizontal discontinuity) propagating from north
towards south and formed in response to subduction of continental lithosphere (Wortel and Spakman, 2000); and (3) a smooth transition
between the two slab segments, without the presence of a tear between
them at least at depths shallower than 100 km (Pearce et al., 2012;
Halpaap et al., 2018). Thus, the geometry of the slab in the broader area
of the Gulf of Corinth and the transition between the two slab segments
remain still poorly understood.
At the eastern termination, beneath southwestern Turkey, the presence of a WNW-ESE trending thrust fault striking parallel to the coast
was proposed to represent the easternmost termination of the HSZ
(Papazachos, 1996). Tomographic studies have suggested, with variable
resolution, the presence of a vertical tear separating the HSZ from the
Western Cyprus Subduction Zone (e.g., de Boorder et al., 1998;
Piromallo and Morelli, 2003; van Hinsbergen et al., 2010b; Biryol et al.,
2011; Legendre et al., 2012; Govers and Fichtner, 2016). The presence
of a vertical tear between the Aegean and Cyprus arcs is also supported
by observations on the surface, including: (1) the presence of uncontaminated highly alkaline volcanism; (2) high-temperature metamorphic domes with axes parallel to the direction of extension in the
Aegean, (3) and migration of granitoid intrusions towards southwest
observed in western Anatolia (Jolivet et al., 2015 and references
therein). Based on the interpretation of the Neotethyan paleogeography
and on the number of distinct subducted slabs different ages have been
proposed for the development of the tear: Miocene (e.g., van
Hinsbergen et al., 2010b; Jolivet et al., 2013; Pourteau et al., 2016);
Eocene-Oligocene (e.g., Bartol and Govers, 2014); and late Cretaceous
(e.g., Gürer et al., 2018). Although there are no doubts about the existence of the tear its geometry is not yet well-defined and needs to be
clarified to relate it with upper plate deformation and more specifically
with the presence of faults on the surface.
Since the pioneering paper on STEP-faults (Govers and Wortel,
2005), several studies have suggested the presence of a NE-SW striking
STEP-fault in the area of the Pliny and Strabo trenches resulting from
slab tearing between the HSZ and the Western Cyprus Subduction Zone
(e.g., Govers and Wortel, 2005; Dilek and Altunkaynak, 2009; van
Hinsbergen et al., 2010b; Biryol et al., 2011; Özbakır et al., 2013;
Govers and Fichtner, 2016). However, the presence of a NE-SW striking
STEP-fault at the eastern termination of the HSZ is in conflict with the
existence of a NW dipping subducting slab in the area of Rhodes as
indicated by seismicity (e.g., Brüstle, 2012) and by receiver function
studies (Sodoudi et al., 2006, 2015). Furthermore, Faccenna et al.
(2014) suggested horizontal tearing of the slab in the area of Rhodes.
Thus, there remain several questions regarding the presence of a NW
dipping slab in the area of Rhodes, and of a NE-SW striking STEP-fault
in the area of the Pliny and Strabo trenches.
A first-order segmentation of the slab has been proposed in the
2. Tectonic setting
In this section, we summarise the tectonic evolution of the HSZ as it
is related to the distribution of seismicity in the discussion.
2.1. Recent kinematics of the Hellenic Subduction Zone
The convergence rate between the Aegean region and the Nubian
plate (~35 mm/yr) exceeds the convergence rate between Africa and
Eurasia (~5–10 mm/yr) due to the rapid southwestward motion of the
Aegean region, with velocity increasing trenchward with respect to
Eurasia (McClusky et al., 2000; Reilinger et al., 2006, 2010; Fig. 1). The
present day geodetically-determined velocity field indicates that the
southeasternmost part of the Aegean region, namely the Rhodes block,
is moving towards southeast at rates up to about 10 mm/yr with respect
to the central and southern Aegean region and the Peloponnese
(Reilinger et al., 2010). The inferred GPS velocity field is driven by slab
rollback, and describes westward extrusion of the Anatolian region, in
combination with counterclockwise rotation and internal deformation
of the Aegean-Anatolian region (McClusky et al., 2000; Reilinger et al.,
2006, 2010).
Recent plate kinematics and the strong curvature of the plate
boundary imply an increasing obliquity of the plate convergence from
west to east in the HSZ. The plate convergence occurs perpendicular to
the arc to the west of Crete and its obliquity increases from about
20–30° south of central Crete to about 40–50° in the eastern forearc, in
the region of Rhodes, causing left-lateral motion at borders of forearc
slivers (Bohnhoff et al., 2005). Such forearc slivers are delimited by enechelon bathymetric troughs, namely the Ptolemy, Pliny, and Strabo
trenches (Fig. 1), which have been interpreted to have formed as result
of the ongoing rollback of the slab and the oblique slip between Crete
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Fig. 1. Main tectonic features of the Hellenic Subduction Zone (HSZ). The two white arrows show the relative motion of the Aegean region and the Nubian plate with
respect to stable Eurasia (McClusky et al., 2000). The yellow stars indicate the epicentral locations of the largest historical earthquakes in the HSZ: the Mw 8.3,
365 CE SW of Crete and the Mw 8.0, 1303 CE SE of Crete (e.g., Ambraseys, 2009; Papadopoulos, 2011), and of the largest tsunamigenic earthquake of the last century
(Ms = 7.4, 1956 Amorgos earthquake; Makropoulos et al., 1989). NAT, North Aegean Trough; KTF, Kefalonia Transform Fault; NHSZ, Northern Hellenic Subduction
Zone; HSZ, Hellenic Subduction Zone; WCSZ, Western Cyprus Subduction Zone; sed. arc, Sedimentary Arc (or outer Hellenic Arc). The dotted magenta line in the
Ionian Sea indicates the transition between oceanic and continental lithosphere in the Ionian Sea. Continuous lines with barbs indicate the active deformation fronts.
Structural elements in the map have been taken from Jolivet et al. (2013). (For interpretation of the references to colour in this figure legend, the reader is referred to
the web version of this article.)
Fault in the late Miocene-early Pliocene (7–5 Ma) with most of the
offset occurring after about 5–4 Ma. Differential slab rollback and
trench retreat due to the subduction of more buoyant continental and
less buoyant oceanic lithosphere have been invoked to explain the development and propagation of the Kefalonia Transform Fault (e.g.,
Royden and Papanikolaou, 2011). van Hinsbergen and Schmid (2012)
argue that the present offset of the deformation fronts to the north and
to the south of the Kefalonia Transform Fault (~100–150 km) has not
been entirely accommodated by right-lateral displacement along the
fault (Royden and Papanikolaou, 2011), but it has rather formed by a
combination of the accretion of the pre-Apulian Unit (~110 km wide)
in the area of the Ionian Islands and about 40 km of displacement along
the fault since about 5–4 Ma. Additional 20 km of displacement to the
Kefalonia Transform Fault zone have been accommodated by the Thesprotiko-Aliakmon fault zone between about 7 and 3.5 Ma (van
Hinsbergen and Schmid, 2012).
The transition between the HSZ and the Northern Hellenic
Subduction Zone is characterised by rapid extension in the Gulf of
Corinth, an E-W striking active continental rift, separating the
Peloponnese from northern Greece (e.g., Armijo et al., 1996). The Gulf
of Corinth developed about 3.5 Ma (e.g., van Hinsbergen and Schmid,
2012) and it is currently deforming due to activity on E-W to NW-SE
oriented normal faults (e.g., Bell et al., 2009). Geodetic data show extension rates up to 10–15 mm/yr in the western part of the gulf (e.g.,
Briole et al., 2000; Avallone et al., 2004). Different mechanisms have
been proposed to explain the initiation of the rifting. These include: (1)
extension associated to the rollback of the Nubian slab (e.g., Jolivet
et al., 2010); (2) gravitational collapse of overthickened crust beneath it
(Le Pourhiet et al., 2003); and (3) the southwestward propagation of
the North Anatolian Fault into the Aegean (Armijo et al., 1996, 1999).
and Rhodes (ten Veen and Kleinspehn, 2002; Meier et al., 2007). The
term trench or Hellenic trench often used to refer to such bathymetric
features is not optimal because it does not represent the superficial
expression of an oceanic trench in the usual sense as initially thought
(e.g., Jongsma, 1977; Le Pichon and Angelier, 1979).
Recent marine magnetic data (e.g., Granot, 2016) and tomography
studies (e.g., Boschi et al., 2009; Legendre et al., 2012) have suggested
the presence of very old oceanic lithosphere in the Eastern Mediterranean between Sicily and west of Cyprus (270–230 Ma, Müller et al.,
2008; 340 Ma in the Herodotus basin, Granot, 2016; 220–230 Ma in the
Ionian basin, Speranza et al., 2012). This oceanic lithosphere, namely
the Eastern Mediterranean Oceanic Lithosphere (e.g., Legendre et al.,
2012), is subducting beneath the HSZ and the Western Cyprus Subduction Zone (WCSZ in Fig. 1) where it is forming the Nubian part of
the slab.
The subduction zone present to the north of the Kefalonia Transform
Fault has been referred to as “Northern Hellenic subduction” (Royden
and Papanikolaou, 2011) to discriminate it from the subduction zone to
the south of the Kefalonia Transform Fault. Following this suggestion,
from here onward we refer to the subduction zone to the north of the
Kefalonia Transform Fault as the Northern Hellenic Subduction Zone,
and to the subduction zone to the south of the Kefalonia Transform
Fault as the Hellenic Subduction Zone or HSZ (NHSZ and HSZ in Fig. 1).
The slower trenchward motion of the overriding plate to the north
of the Kefalonia Transform Fault compared to the Aegean region, induces lower convergence rates (5–10 mm/yr) at the Northern Hellenic
Subduction Zone compared to the HSZ (e.g., Cocard et al., 1999;
Hollenstein et al., 2008). Numerical modeling results (Royden and
Papanikolaou, 2011) and geological observations (van Hinsbergen
et al., 2006) have suggested a development of the Kefalonia Transform
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G.M. Bocchini et al.
the multiple-slab model (e.g., Bonneau, 1982; Shanov et al., 1992;
Ricou et al., 1998); and (2) the single/continuous-slab model (singleslab model from now onward; e.g., Faccenna et al., 2003; Meier et al.,
2004a; van Hinsbergen et al., 2005; Jolivet and Brun, 2010). The singleslab model that nowadays is widely accepted (e.g., Faccenna et al.,
2003; Meier et al., 2004a, 2004b; van Hinsbergen et al., 2005; Jolivet
and Brun, 2010), is built on the presence of a continuous N dipping
anomaly, extending down to about 1400 km depth, which has been
interpreted as the subducting slab (e.g., Spakman et al., 1993; Bijwaard
et al., 1998; van der Meer et al., 2017). In this model, the continuation
of the subduction and the formation of a single slab are caused by the
backstepping of the subduction from the front-side to the back-side of
the colliding microcontinents (e.g., Faccenna et al., 2003; Meier et al.,
2004a, 2004b; van Hinsbergen et al., 2005; Jolivet and Brun, 2010).
The opening of the Gulf of Corinth together with the Amvrakikos
graben added additional 20 km of displacement to the Kefalonia
Transform Fault zone leading to a total displacement of about 80 km
(van Hinsbergen and Schmid, 2012).
2.2. Tectonic evolution of the HSZ
Herein we briefly summarise the evolution of the subduction history
of the HSZ with a main focus on the area of western Greece. In the
Aegean region, Africa-Europe convergence involved the subduction of
three different oceanic basins namely the Vardar (Jurassic–Cretaceuos),
the Pindos (Eocene), and the present-day subducting eastern
Mediterranean oceanic basins, with interposed continental domains
(e.g., Dercourt et al., 1986; Robertson et al., 1996; Ricou et al., 1998).
After the closure of the Vardar Ocean (late Cretaceous), nappe stacking
indicates the continuation of the subduction despite the presence of
microcontinents entering the trench (e.g., Faccenna et al., 2003; Meier
et al., 2004a; van Hinsbergen et al., 2005; Jolivet and Brun, 2010). Such
nappes, nowadays piled-up in the upper plate, represent upper crustal
portions delaminated from the underlying continental or oceanic lower
crust and mantle lithosphere sufficiently dense to subduct (e.g.,
Faccenna et al., 2003; Meier et al., 2004a; van Hinsbergen et al., 2005;
Jolivet and Brun, 2010).
Gaina et al. (2013) introduced the term “Greater Adria” to include
all platform and basin units that existed between the Eastern Mediterranean Ocean, the Alpine Tethys Ocean and the Vardar-Sava-IzmirAnkara Ocean. In western Greece, from east to west, Greater Adria or
Adria consists of: (1) the Pelagonian Unit; (2) the Pindos Unit; (3) the
Gavroro-Tripolitza Unit and Phyllite-Quarzite Unit; (4) the Ionian Unit
and Plattenkalk Unit; and (5) the pre-Apulian Unit (or Paxos Unit) (van
Hinsbergen et al., 2005 and references therein). The accretion of the
continental units belonging to the Pelagonian zone started in the late
Cretaceous following the closure of the Vardar Ocean and continued
until the late Paleocene-early Eocene (e.g., van Hinsbergen et al., 2005
and references therein). It was followed by the subduction of the Pindos
basin which started in the early Eocene (e.g., van Hinsbergen et al.,
2005 and references therein) and continued until its closure in the late
Eocene (34–37 Ma, Stampfli and Borel, 2004). From here onward we
assume an age of 34 Ma for the closure of the Pindos basin. In the
Oligocene the Pindos Unit was underthrusted by the Tripolitza Unit that
was simultaneously underthursted by the Ionian Unit (van Hinsbergen
et al., 2005). Between about 15 Ma and 4 Ma folding, thrusting, and
accretion of the pre-Apulian zone occurred in the Ionian Islands (van
Hinsbergen et al., 2006 and therein references). The pre-Apulian zone,
interpreted as the slope of the Apulian platform, did not extent much
further south than Zakynthos (Finetti, 1982; Underhill, 1989; van
Hinsbergen and Schmid, 2012) where the Eastern Mediterranean
Oceanic Lithosphere was subducting since at least 13 Ma (e.g., Finetti,
1982; Underhill, 1989).
The outcropping of the Pindos Unit, the Gavroro-Tripolitza Unit and
Phyllite Quarzite, and of the Ionian Unit and Plattenkalk in western and
eastern Greece (e.g., van Hinsbergen et al., 2005; van Hinsbergen and
Schmid, 2012) suggest a similar tectonic evolution along the entire HSZ
at least until the emplacement of these units.
The complete accretion of the Ionian zone to the upper plate was
followed by a diachronous onset of subduction of the eastern
Mediterranean oceanic lithosphere. The onset of oceanic subduction
started earlier to the east than to the west: at about 35 Ma beneath
southwestern Turkey, at about 20 Ma beneath south of Crete and about
4 Ma in the area of the Ionian Islands (van Hinsbergen and Schmid,
2012 and references therein). Moreover to the north of the current
position of the Kefalonia Transform Fault, oceanic lithosphere subduction has never initiated, and the westernmost part of Adria exposed
in northern Greece (i.e., Apulian platform) is still subducting.
To explain the repeated subduction of oceanic basins, two endmember models have been proposed for the eastern Mediterranean: (1)
3. Data
In this section we describe the seismicity catalogues on which this
paper is built, namely the International Seismological Centre (ISC)
global catalogue (Section 3.1), and the EGELADOS, CYCNET and
LIBNET temporary local catalogues (Section 3.2). We discuss the location accuracy, the magnitudes of completeness, the recording periods,
and discuss the distribution of the seismicity in the different datasets.
3.1. The ISC catalogue
The HSZ is a very seismically active structure in the Mediterranean
region. Historical catalogues covering > 2000 years (e.g., Papazachos
and Papazachou, 1997; Papazachos et al., 2000a, Papazachos et al.,
2010; Guidoboni and Comastri, 2005; Ambraseys, 2009) report numerous large magnitude earthquakes including the tsunamigenic
M ~ 8.3, 365 CE and M ~ 8.0 1303 CE earthquakes that occurred to the
southwest and to the southeast of Crete, respectively (Fig. 1).
The global catalogue of the ISC Bulletin (2015) shown in Fig. 2
provides an explanatory overview of the current seismicity in the HSZ.
The catalogue contains instrumental seismicity since 1964 with a
magnitude of completeness (Mc) of about 4.0 for the Aegean area,
roughly estimated from a visual inspection of the magnitude-frequency
distribution of the earthquakes. When comparing hypocentres from the
ISC catalogue with those of temporary local seismic networks in the
seismogenic zone south of Crete, Meier et al. (2004b) found that ISC
locations were on average systematically shifted northeastward and
downward by about 15 km. This is likely an indication for a bias in the
ISC locations due to high seismic velocities in the slab.
Strong shallow seismicity occurs in the outer part of the outer
Hellenic Arc (or Sedimentary Arc) along an arcuate belt, parallel to the
plate boundary, and extending up to 50–100 km from the coast line
towards the Mediterranean Sea (Fig. 2). The hypocentres within this
belt are located along and above a landward dipping interface identified as the interplate seismogenic zone. Interplate seismicity can be
identified from the distribution of the ISC catalogue hypocentres (see
Fig. 4 of the Supplementary material) if a bias towards larger depths of
the ISC hypocentres is taken in account (Meier et al., 2004b). The interplate seismogenic zone as well as the seismicity along the outer
Hellenic Arc have been investigated by a number of temporary seismic
networks (e.g., Hatzfeld et al., 1989, 1993; Delibasis et al., 1999;
Sachpazi et al., 2000; Meier et al., 2004b; Becker et al., 2010). The
spatial distribution of the shallow seismicity is observed to be similar to
that of other subduction zones with high activity along the plate interface. Interplate seismicity recorded by temporary local networks has
been observed to occur between about 10 to 15 km depth in the region
of the Ionian Islands (Sachpazi et al., 2000) and between about 20 to
40 km depth south of Crete (Meier et al., 2004b) where the down-dip
limit of the seismogenic zone is roughly located along the southern
coastline of the Island (Meier et al., 2004b; Becker et al., 2010). In
addition, strong microseismic activity within the overriding continental
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G.M. Bocchini et al.
Fig. 2. International Seismological Centre (ISC)
seismicity catalogue for the period 1964–2014 (ISC
Bulletin, 2015). Only earthquakes with vertical and
horizontal errors ≤10 km are plotted in the map and
cross-sections. Earthquakes are colour coded according to depth and are sized according to magnitude. The seismicity occurring 50 km on each side of
the profiles (continuous black lines in the map) is
projected into the cross-sections. The dotted magenta
line in the seismicity cross-sections indicates the top
of the subducting slab that for cross-sections 4 and 5
has been improved by using earthquake hypocentres
from temporary local networks. (For interpretation
of the references to colour in this figure legend, the
reader is referred to the web version of this article.)
in Fig. 2 show the main features of the Wadati-Benioff zone of the
subducting Nubian slab that include: (1) a steeper slab in the east
compared to the west resulting in a longer active part of the slab in the
western part (cross-sections 2, 3, and 4 in Fig. 2) compared to the
eastern part (cross-section 5 in Fig. 2); and (2) stronger intermediatedepth seismic activity in the easternmost part (cross-section 5 in Fig. 2)
compared to the west of it (cross-sections 2, 3, and 4 in Fig. 2).
crust occurs along the Ptolemy, Pliny, and Strabo trenches from the
surface down to the plate interface (Delibasis et al., 1999; Meier et al.,
2004b; Becker et al., 2010). Beneath the forearc and more precisely
beneath the outer Hellenic Arc (e.g., below Crete), microseismicity in
the upper plate is confined to the upper about 20 km (e.g., Delibasis
et al., 1999; Meier et al., 2004b).
Seismicity cross-sections from the ISC catalogue (ISC Bulletin, 2015)
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manually picked for the region east of 26°E (Brüstle, 2012). The final
event catalogue was obtained with NLLoc (Lomax et al., 2000) using a
1-D velocity model (i.e., 1-D minimum velocity model) and the corresponding station corrections obtained by applying the program VELEST
(Kissling et al., 1994). This has led to an accurate earthquake catalogue
with an average semi-major axis of the 68%-confidence ellipsoid
≤10 km for about 3000 earthquakes (Brüstle, 2012). The entire EGELADOS catalogue for the southeastern Aegean consists of about 5400
earthquakes (Brüstle, 2012).
The CYCNET network (September 2002–September 2005) consisted
of 22 seismic stations deployed in various network configurations in the
central part of the Hellenic Volcanic Arc (Bohnhoff et al., 2004, 2006).
The network was extended by selected permanent stations of the
GEOFON network, predominantly on Crete, to obtain a better azimuthal coverage of the southern part of the Hellenic Volcanic Arc. To
make the earthquake locations from the CYCNET network consistent
with those from the EGELADOS network the entire dataset was relocated with the same 1-D velocity model used for the EGELADOS
network (Brüstle, 2012). The entire CYCNET catalogue consists of about
6800 events out of which about 4000 have an average semi-major axis
of the 68%-confidence ≤10 km (Brüstle, 2012). Brüstle (2012) estimated a magnitude of completeness (Mc) of about 2 for the EGELADOS
and CYCNET catalogues.
The LIBNET network, deployed to the south of eastern Crete, consisted of five separated observation phases carried out between July
2003 and June 2004 during which up to eight OBSs, in various configurations, were jointly operating with five temporary short-period
stations on Crete (i.e., Messara network) and several permanent broadband stations on Crete and surrounding islands (Becker, 2007; Becker
et al., 2010). The LIBNET catalogue consists of > 2600 earthquakes and
has a Mc of about 1.8–2.1 in the region of Crete and of the Ptolemy
trench, increasing to 2.5 to the south and southeast of Crete in the region of the Strabo trench (Becker, 2007; Becker et al., 2010). The
seismicity recorded during the three temporary local seismic experiments is shown in Fig. 3 (see Fig. 1 of the Supplementary material for
the configurations of the temporary local seismic networks).
Seismicity cross-sections obtained from temporary local seismicity
catalogues provide a sharper and clearer image of the subducting slab
(Fig. 3) compared to the ISC catalogue (Fig. 2). The distribution of the
Wadati-Benioff zone seismicity, which shows a steeper dipping angle in
the area of Karpathos-Rhodes (cross-section 3 in Fig. 3) than in the area
of Crete (cross-section 1 in Fig. 3), confirms the observations from the
ISC catalogue. Shallow seismic activity (depth ≤ 50 km) is more
abundant to the south of eastern Crete (cross-sections 1 in Fig. 3)
compared to the forearc region to the southeast of Karpathos and
Rhodes (cross-sections 3 in Fig. 3). In the region between eastern Crete
and Karpathos two separated alignments of intermediate-depth seismicity are observed (cross-section 2 in Fig. 3). Additional analysis and
further discussion of this seismological feature along with the proposed
interpretation is given in Section 4.3.
Fig. 3. Seismicity recorded by the EGELADOS (circles), CYCNET (triangles) and
LIBNET (diamonds) temporary networks. Only well-located earthquakes (errors ≤ 15 km) are plotted. Seismicity occurring within 20 km towards each side
of the profiles is plotted into the cross-sections.
3.2. Seismicity recorded by temporary networks
Temporary local networks provide on the one hand more accurate
(errors ~ 5–10 km) hypocentral solutions than those from regional and
global seismicity catalogues (e.g., the ISC catalogue), allowing for detailed studies of seismogenic structures, but on the other hand they do
not describe the long term variability of the seismic activity.
We use the EGELADOS temporary local earthquake catalogue
(Brüstle, 2012) to refine the geometry of the slab in the southeastern
Aegean. To enrich the dataset of temporary local seismicity, we also use
observations of the CYCNET (Bohnhoff et al., 2004, 2006; Brüstle,
2012) and LIBNET (Becker et al., 2010) temporary local seismic experiments (Fig. 3). The EGELADOS experiment lasted for about
1.5 years (October 2005–March 2007) during which a total of 75 land
stations and 23 Ocean Bottom Seismometers (OBS), uniformely distributed in the Aegean region, were operating (Friederich and Meier,
2008). The earthquakes recorded by the EGELADOS network and used
in this study were automatically detected (Küperkoch et al., 2011) and
4. Results
It this section we clarify the geometry of the subducting slab at the
western (Section 4.1) and at the eastern terminations (Section 4.2) of
the HSZ, as well as beneath the southeastern Aegean (Section 4.3) by
analysing the spatial distribution of the seismicity from global and
temporary local catalogues. We further present a refined geometry of
the seismically active subducting slab (Section 4.4). Finally, we calculate the thermal structure of the two identified slab segments of the HSZ
to investigate pressure (P) and temperature (T) conditions at the
earthquake hypocentres.
4.1. Western termination of the Hellenic Subduction Zone
Herein we use seismicity from the ISC catalogue to clarify the slab
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shear-waves attenuation study (Konstantinou and Melis, 2008), a wideangle experiment (Zelt et al., 2005), and a local earthquake tomography
study (Halpaap et al., 2018).
The Kefalonia Transform Fault is located above the edge of the
seismically active slab of the HSZ (Fig. 4). Therefore, it has been suggested that the Kefalonia Transform Fault represents a STEP-fault
(Govers and Wortel, 2005). It is however not extending to the north of
the Gulf of Corinth where a sharp drop in seismic activity, indicates the
edge of the seismically active slab at depth below about 120 km.
The opening of the Gulf of Corinth and of the Amvrakikos graben,
decoupling the upper plate continental units in northern Greece from
those in the Peloponnese in an area of lateral heterogeneity, may have
caused the cessation of the dextral strike-slip deformation along the
Thesprotiko-Aliakmon fault zone above the edge of the subducting
oceanic lithosphere. In fact, according to the reconstruction of van
Hinsbergen and Schmid (2012), the Gulf of Corinth and the Amvrakikos
graben have accommodated the same amount of deformation (20 km in
the last 3.5 Ma) as the Thesprotiko-Aliakmon fault zone (20 km from 7
to 3.5 Ma). At about 4 Ma the onset of oceanic lithosphere subduction in
the area of the Ionian Islands (van Hinsbergen et al., 2006), and the
associated increased slab rollback, may have caused the development of
a STEP-fault, namely the Kefalonia Transform Fault, to the west of the
Gulf of Corinth. The Thesprotiko-Aliakmon fault zone could represent a
former STEP-fault, which developed on top of the edge of the seismically active slab before the opening of the Gulf of Corinth (Fig. 4d).
geometry in the broader area of the Kefalonia Transform Fault, at the
western termination of the HSZ. The downgoing plate at the western
termination of the HSZ is characterised by the transition from the
Eastern Mediterranean Oceanic Lithosphere to the Adriatic continental
lithosphere (e.g., Papanikolaou and Royden, 2007; Royden and
Papanikolaou, 2011; Legendre et al., 2012; Pearce et al., 2012).
A study of scattered wave signals along two linear arrays has imaged
a weaker about 20 km thick negative anomaly down to 70–100 km (top
and Moho of the slab, respectively) beneath northern Greece, and a
sharper about 8 km thick negative anomaly down to about 80 km (top
of the slab) beneath the Peloponnese (Pearce et al., 2012). The two
layers have been interpreted as continental (thicker and weaker) and
oceanic (thinner and sharper) crust, respectively. Differences in the lithospheric thickness and average shear wave velocities between the
slabs subducting beneath northern Greece and the Peloponnese have
also been highlighted by seismic tomography (e.g., Legendre et al.,
2012).
Seismicity cross-sections indicate the extension of the seismically
active slab to the north of the Gulf of Corinth (cross-sections 1, and 7 in
Fig. 2). An abrupt lateral termination of the intermediate-depth seismicity is very likely associated with the transition between oceanic and
continental subduction (see also Halpaap et al., 2018). Very few intermediate-depth earthquakes are observed in the slab subducting beneath Northern Greece (Fig. 4). The boundary between the seismically
active and seismically almost inactive slabs is observed to be located
slightly to the south of the Thesprotiko-Aliakmon fault zone (Fig. 4d).
Intermediate-depth seismicity beneath the Gulf of Corinth indicates that
a horizontal tear in the slab (Wortel and Spakman, 2000) is unlikely at
depths shallower that 120 km. Instead, the Gulf of Corinth is found
above a sharp boundary between a Wadati-Benioff zone extending
down to about 180 km to the south of the Gulf (cross-section 2 in Fig. 2;
Fig. 4) and a Wadati-Benioff zone extending down to about 120 km
depth to the north of the Gulf (cross-section 1 in Fig. 2).
The continuity of the seismically active slab beneath the Gulf of
Corinth is consistent with other independent observations including a
4.2. Eastern termination of the Hellenic Subduction Zone
Also the eastern termination of the Hellenic Subduction Zone is of
considerable complexity. Herein we use well-located earthquake hypocentres from global and temporary local seismicity catalogues to
clarify the geometry of the slab segments subducting beneath the
southeastern Aegean and western Anatolia as well as to investigate the
existence of a slab tear between them.
Seismicity cross-sections from temporary local (cross-section 3 in
Fig. 4. ISC seismicity and slab isodepths at the western termination of
the HSZ. Only hypocentres with estimated vertical and horizontal errors
≤10 km are plotted. The seismicity is
plotted at 20 km intervals as indicated
at the left top of each map. Continuous
lines with triangles indicate the active
deformation fronts. HSZ, Hellenic
Subduction Zone; NHSZ, Northern
Hellenic Subduction Zone.
Location of the fault zones shown into
panel d are taken from van Hinsbergen
and Schmid (2012).
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continuously distributed seismicity between 50 and 100 km depth
(cross-section 3 in Fig. 5). Beneath southwestern Anatolia the direction
of the dipping angle of the Wadati-Benioff zone, well-defined down to
150 km depth, changes towards NE (cross-section 6 in Fig. 2). An
aseismic area at depths larger than about 20–40 km is observed between the slab segments subducting beneath the HSZ and the Western
Cyprus Subduction Zone, respectively (cross-section 8 in Fig. 2). We
interpret this aseismic zone located beneath southwestern Turkey as a
tear in the slab. According to tomographic images (e.g., Piromallo and
Morelli, 2003; van Hinsbergen et al., 2010b; Biryol et al., 2011;
Legendre et al., 2012; Govers and Fichtner, 2016) and mangnetic
anomalies (Granot, 2016), south of the tear old oceanic lithosphere is
present in the eastern Mediterranean. Thus, the same oceanic lithosphere is subducting to the east of Karpathos, at the HSZ, and at the
Western Cyprus Subduction Zone. For geometrical reasons a tear is
required between the two slab segments that are subducting in different
directions, namely towards NW and NE, respectively. The existence of a
tear between the HSZ and the Western Cyprus Subduction Zone is
consistent with a rather fuzzy slow velocity anomaly imaged by tomography studies in the area (e.g., de Boorder et al., 1998; Piromallo
and Morelli, 2003; Biryol et al., 2011; Legendre et al., 2012; Govers and
Fichtner, 2016). Seismicity allows for determining the geometry of the
slab segments and the tear with a higher resolution than tomography
down to about 180 km depth. The tear is about 250 km wide at 100 km
depth (cross-section 8 in Fig. 2). Below about 180 km depth, the tear
can only be resolved by tomographic imaging. A sharp ending of the
intermediate-depth seismicity (depth > 50 km) east of Rhodes indicates the easternmost termination of the HSZ (Fig. 5). This termination is trending almost parallel to the southwestern Anatolian coast at
about 50–80 km depth and it is observed to shift progressively towards
west with increasing hypocentral depths (Fig. 5).
4.3. Seismological evidence for slab segmentation in the southeastern
Aegean
Well-located hypocentres from temporary network catalogues
(EGELADOS, and CYCNET) are analysed to investigate the geometry of
the slab in the southeastern Aegean where several studies observed a
sudden increase of the slab dipping angle (Papazachos and Nolet, 1997;
Meier et al., 2007; Brüstle, 2012; Sodoudi et al., 2015).
Seismicity cross-sections from temporary network catalogues (i.e.,
EGELADOS, and CYCNET) show a shallower and less seismically active,
N-NNE dipping slab in the area of Crete (cross-section 1 in Fig. 5) and a
steeper and more seismically active, NW dipping slab in the area of
Karpathos-Rhodes (cross-section 3 in Fig. 5). Similarly to Brüstle (2012)
we observe a vertical offset of about 30–40 km between the two
aforementioned cross-sections (top of Fig. 6). This vertical offset in the
seismicity is not observed at depths shallower than 60–70 km which
probably indicates that the discontinuity in the slab does not propagate
towards shallower depths (top of Fig. 6). The different dipping angles
observed in the seismicity are in agreement with a regional body-wave
tomography study which has imaged a less steep (25° dip) and wider
slab segment in the western part, and a steeper (35° dip) and narrower
slab segment in the eastern part (Papazachos and Nolet, 1997).
A double seismic zone is present in the region in between the less
steep Wadati-Benioff zone in the area of Crete and the steeper WadatiBenioff zone in the area of Karpathos-Rhodes (cross-section 2 in Fig. 5).
The double seismic zone is only found in the transition zone between
the two segments at about 60–140 km depth and does not extend to the
west or to the east (cross-sections 1 and 3 in Fig. 5). On the basis of
these observations, together with the fact that a double Wadati-Benioff
zone has never been recognised before in the HSZ, we conclude that the
described double seismic zone is related to the segmentation of the slab.
In fact, a seismicity cross-section parallel to the strike of the slab in the
area of Rhodes, indicates the underthrusting of the eastern steeper and
deeper segment beneath the shallower and less steep western segment
Fig. 5. Seismicity along the subducting Nubian slab (depth ≥ 50 km) recorded
at the EGELADOS and CYCNET local temporary networks. Only well-located
earthquakes (errors ≤ 15 km) are shown. The dotted black lines in the crosssections indicate the top of the slab while the dotted black lines in the map
represent the top of the slab isodepths (the depth is indicated in the yellow
boxes); the red line indicates the boundary between the western and the eastern
segment (the dotted portion of the red line indicates the possible southward
prolongation of the boundary between the two segments for which no expression can be found in the observed seismicity); the dotted blue line indicates the
eastern termination of the HSZ. Hypocentres located 25 km towards each side of
the profiles are projected into the cross-sections. Focal mechanisms (Global
CMT, 2016) of three large recent earthquakes which ruptured in the region are
shown in the map: (a) the 01-22-2002 (Mw 6.1) with centroid depth of 90 km;
(b) the 04-01-2011 (Mw 6.1) with centroid depth of 63 km; and (c) the 03-272015 (Mw 5.1) with centroid depth of 86 km. Stars in the map view (highlighted by an arrow) and in the cross-sections indicate the location of two
events, one in the upper and one in the lower seismic zone, for which waveforms are compared at the station ANAF (red triangle) in Fig. 8. (For interpretation of the references to colour in this figure legend, the reader is referred
to the web version of this article.)
Fig. 5) and global (cross-section 5 in Fig. 2) seismicity catalogues show
clear evidence for a NW dipping Wadati-Benioff zone down to about
180 km in the area of Rhodes. There is no indication for a horizontal
tear or slab break-off (Faccenna et al., 2014) because of the
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mechanisms is small, their solutions are consistent with the hypothesis
that the steeper eastern segment is underthrusting the western segment
resulting in a compressional regime in the lower seismic zone where the
two slab segments collide.
Three moderate magnitude intermediate-depth earthquakes have
recently occurred at the depth range of the identified double seismic
zone (~60–140 km), namely: (a) the Mw 6.1, 22-01-2002 earthquake;
(b) the Mw 6.1, 01-04-2011 earthquake; and (c) the Mw 5.1, 03-272015 earthquake (Fig. 5) (Global CMT, 2016). A centroid depth of
90 km, 63 km, and 86 km has been obtained for these three earthquakes, respectively (Global CMT, 2016). Revisited hypocentral solutions from the National Observatory of Athens (NOA, http://bbnet.gein.
noa.gr/) report depths of 104 km, 63 km, and 67 km, respectively.
These three earthquakes occurred close to each other (epicentral distance < 20 km) in the area where the top of the slab is expected at
about 65 km. Therefore, the 2011 and 2015 earthquakes, which exhibit
an almost pure strike-slip failure, could be associated to the western
segment or they could have occurred at the transition between the two
segments. The 2002 earthquake, which exhibits a strong thrust component and it is deeper, could be associated to the underthrusted part of
the eastern segment beneath the western segment. In fact, the difference in depth, observed especially between the largest of the three
analysed earthquakes, namely the 2002 and the 2011 earthquakes,
cannot be justified by only taking the location errors into account.
Fig. 6. Seismological evidence for a vertical discontinuity in the slab. The location on the map of the cross-sections reported in this figure is shown in Fig. 5.
At the top the offset in the intermediate-depth seismicity is shown by overlapping profiles 1 and 3 at the Sedimentary Arc (Fig. 1). At the bottom is shown
a seismicity cross-section parallel to the slab isodepth in the area of Rhodes
which reveals a discontinuity of the slab and an overthrusting of the western
segment on top of the eastern segment. Earthquake hypocentres from the
EGELADOS (circles) and CYCNET (triangles) catalogues with errors smaller
than 20 km are shown in the cross-section. The dotted line indicates the location of the top of the slab. W into the cross-sections indicates the width in km of
the projected seismicity towards each side of the surface traces of the profiles.
4.4. Revisited geometry of the seismically active subducting slab
Observations of seismicity are well suited to define the 3-D geometry of the HSZ down to a depth of about 150 -180 km. Based on the
results from the previous three sections we propose a refined model of
the seismically active slab subducting beneath the Aegean region and
southwestern Anatolia. The resolution is of about 5–10 km in the
southeastern Aegean due to the use of temporary local seismicity catalogues, and somewhat lower, of about 10–20 km, in the western part of
the HSZ and beneath southwestern Anatolia due to the use of the ISC
global seismicity catalogue.
The revisited geometry of the active slab has been obtained by refining the slab model of Ganas and Parsons (2009) in the broader area
of the western termination (Fig. 4) and in the southeastern Aegean
(Fig. 6), where the depth of the top of the slab has been carefully detected. Ganas and Parsons (2009) constrained the slab geometry using
earthquake hypocentres (Papazachos et al., 2000b; Meier et al., 2004b),
receiver function data (Li et al., 2003), and results from seismic profiling (Bohnhoff et al., 2001).
Intermediate-depth seismicity in cold subduction zones tends to
occur in the subducting oceanic crust (Abers, 1992; Kirby, 1995; Abers
et al., 2014). Therefore, we measure the mean depth of the WadatiBenioff zone seismicity and substract 5 km to estimate the depth of the
plate interface.
At shallow depths (slab isodepths ≤ 20 km) from the west of the
Peloponnese to the south of western Crete we do not change the model
by Ganas and Parsons (2009). To the south of eastern Crete we used the
LIBNET catalogue (Becker et al., 2010) to refine the 20 and 40 km slab
isodepths. The 40 km slab isodepth has been constrained with the results of Meier et al. (2004b) to the south of western Crete and with
those of Sachpazi et al. (2016a) beneath the Peloponnese. To the
southeast of Karpathos-Rhodes, we draw the slab isodepths shallower
than 65 km in continuity with those to the south of Crete and parallel to
the deeper slab isodepths in the area, obtaining a geometry, which is
consistent with the one proposed by Papazachos et al. (2000b). Because
the Northern Hellenic subduction is aseismic we rely completely on the
scattered waves images there (Pearce et al., 2012).
The model is presented in Fig. 9. In contrast to previous models
(e.g., Gudmundsson and Sambridge, 1998; Knapmeyer, 1999;
Papazachos et al., 2000b; Ganas and Parsons, 2009), the oceanic slab of
the HSZ is found to extend to the north of the Gulf of Corinth. The
(bottom of Fig. 6). Thus, we associate the upper seismic zone of the
identified double seismic zone to the top of the slab in the western
segment, while the lower seismic zone to the top of the slab in the
eastern segment. This is in agreement with the stress regime deduced
from focal mechanisms of intermediate-depths intraplate earthquakes
which indicate along-strike compression and active shortening acting
on the Nubian slab (Benetatos et al., 2004; Bohnhoff et al., 2005; Shaw
and Jackson, 2010).
To check if the observed double seismic zone was an artifact derived, for example, from the use of two different seismicity catalogues
(i.e., EGELADOS and CYCNET networks), or due to earthquake mislocations, the seismic events belonging to it have been relocated with
HypoDD, an algorithm which improves the relative location between
seismic events (Waldhauser, 2001). Both differential travel-times from
high-precision cross-correlation (P and S) and phase arrival times (P
and S) listed in the earthquake catalogues have been used for the relocation. Relocated hypocentres did not significantly shift from their
initial position (shift < 5 km, Fig. 7) and their association to the upper
or lower seismic zone does not depend on the catalogue from which the
events were taken (Fig. 6).
Furthermore, the waveforms of the earthquakes belonging to the
upper or to the lower seismic zone and having similar epicentral coordinates show clear differences in the S-P travel times (Fig. 8 and in
Fig. 2 of the Supplementary material). These differences must be caused
by different depths of the events because the events show similar epicentral locations (Fig. 5) and the epicentral distance of the deeper
event, with respect to the seismic station at which the waveforms are
compared (ANAF, Fig. 5), is actually smaller than that of the shallower
event. This supports the hypothesis of a double seismic zone and an
offset of the slab segments of about 30–40 km.
Focal mechanisms of small magnitude earthquakes (Ml ≤ 3.8), recorded by the EGELADOS and CYCNET networks and having at least ten
first-motion readings (Friederich et al., 2014), show an extensional
regime for four earthquakes in the upper seismic zone and a compressional regime for the only available focal mechanism in the lower
seismic zone (Fig. 8). Although the number of available source
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G.M. Bocchini et al.
Fig. 7. Zoom in the observed double seismic zone. Gray coloured circles represent original locations, coloured circles the relocated hypocentres using HypoDD.
Available focal mechanisms (Friederich et al., 2014) have been associated to the events in the double seismic zone. W into the cross-sections indicates the width in km
of the projected seismicity towards each side of the surface traces of the pofiles. (For interpretation of the references to colour in this figure legend, the reader is
referred to the web version of this article.)
by Granot (2016) we suggest a similar age of about 340 Ma for the slabs
subducting beneath the eastern Aegean and southwestern Anatolia
(Fig. 9). The age of the slab subducting between the Kefalonia Transform Fault and east of Crete (western slab segment of the HSZ, Fig. 9) is
constrained by magnetic anomaly data from Speranza et al. (2012)
which inferred an age of 220–230 Ma for the oceanic crust in the Ionian
basin.
transition towards the Northern Hellenic Subduction Zone remains
elusive because of the lack of seismicity. In Fig. 9 this is indicated by a
question mark. We will further discuss this in Section 5.2 of the paper.
The plate interface is also poorly constrained in the Gulf of Patras to the
east of the Kefalonia Transform Fault due to the low seismic activity.
However a local seismic tomography indicates the presence of a subducting slab in the area (Halpaap et al., 2018). Here we draw the
northernmost edge of the oceanic slab by connecting the northeast tip
of the Kefalonia fault with the 65 km slab isodepth.
A first-order segmentation is revealed for the area of Karpathos
(Section 4.3) separating the slab subducting beneath the HSZ into a
western and eastern segment. The boundary between the two segments
is quite clear at depths larger than 60 km where it is marked by a
vertical offset in the seismicity distribution, but cannot be observed at
shallower depths as indicated by a question mark in the map (Fig. 9). A
received fuction study has suggested smaller scale segmentation within
the western slab segment of the HSZ beneath the Peloponnese (Sachpazi
et al., 2016a). We do not consider this in our revised slab model as it is
not resolved by the ISC catalogue.
Finally, the geometry of the eastern segment is well-constrained by
microseismicity recorded by temporary networks that indicates a welldefined NW dipping Wadati-Benioff zone. The eastern slab segment of
the HSZ is separated from the slab segment subducting beneath
southwestern Anatolia by a vertical tear in the Nubian slab that is
opening towards larger depth (Fig. 9). Following marine magnetic data
4.5. Thermal structure for the slab segments of the Hellenic Subduction Zone
In this subsection we characterise the thermal structure of the two
slab segments (Section 4.3) and the pressure and temperature (P-T)
conditions of the earthquake hypocentres along them.
The model set-up and numerical solution closely follows Syracuse
et al. (2010) and van Keken et al. (2011). A concise description of the
modeling is provided in the appendix of Wei et al. (2017). In brief, we
use a combined kinematic-dynamic model where the slab is modeled
with constant speed and velocity vectors that are parallel to the slab
surface. The overriding plate is assumed to have zero velocity to 50 km
depth. The mantle wedge below this depth is modeled as an incompressible, infinite-Prandtl-number fluid with a viscosity that is
based on the combined diffusion and dislocation creep of dry olivine.
The thermal structure of the subduction zone is modeled by solving the
governing equations of conservation of mass, momentum, and thermal
energy as described in Syracuse et al. (2010) and Wei et al. (2017). We
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errors are small due to the use of this high resolution mesh and the
resulting uncertainties in the temperature in the slab are < 5 °C.
We estimate from the tectonic reconstruction of Jolivet et al. (2013)
subduction rates of 19 mm/yr and 23 mm/yr, over the last 20 Ma, for
the eastern and western segments of the HSZ, respectively. This accounts for the decreasing relative velocities at the plate interface from
west to east due to the increasing obliquity.
The subducting plate has been modeled as an oceanic lithosphere
with a crustal thickness of 7 km and an age of 220 Ma along the profile
across eastern Crete (Profile 1, Fig. 10) and 340 Ma along the profile
across Karpathos (Profile 3, Fig. 10) according to magnetic anomaly
results of Speranza et al. (2012) and Granot (2016), respectively. The
position the top of the slab along two transects crossing the western and
the eastern slab segments of the HSZ is constrained by using well-located hypocentres from temporary local seismic networks (cross-sections 1 and 3 in Fig. 5). The thermal structure at the trench is derived
from the GDH1 plate model (Stein and Stein, 1992). We note that the
thermal structure of the slab and the heat flow predicted from the
GDH1 plate model does not show significant variations for oceanic
plates older than 100 Ma by construction (Stein and Stein, 1992). The
incoming plate is modeled with an average 6 km thick sediment layer
on top (e.g., Kopf et al., 2003) that gradually thins to 1 km at 15 km
depth and remains constant towards larger depths.
The overriding plate is modeled as an extended and thinned continental crust. We assume an upper crust of 15 km (with radiogenic
heating of 1.3 × 10−6 W/m3) and 15 km of middle crust (with reduced
radiogenic heating of 0.27 × 10−6 W/m3). The overriding plate structure evolves by the half-space cooling from the top boundary condition
(T = 0). The model is evolved for 20 Myr, which is sufficient for the
slab to reach a steady-state thermal structure. To match the model with
the heat flow of the overriding plate away from the volcanic arc
(Jongsma, 1974; Fytikas and Kolios, 1979), we assume an initial condition for the overriding plate which is that of an oceanic lithosphere
Fig. 8. Waveforms of two earthquakes belonging to the upper (top 3 traces, red
star in Fig. 5, ML 2.1, depth 78 km) and lower (bottom 3 traces, yellow star in
Fig. 5, ML 2.3, depth 125 km) double seismic zone between Crete and Karpathos. The waveforms were recorded at station ANAF (Fig. 5) and a 3rd order
Butterworth filter between 2 and 25 Hz is applied. Traces are aligned according
to the respective P-onset of the event at the station. (For interpretation of the
references to colour in this figure legend, the reader is referred to the web
version of this article.)
use a finite element method for the spatial discretization of the equations. The finite element mesh has a resolution of 1 km in the regions
with the strongest thermal gradients and we use coarsening up to 10 km
resolution in the near-isothermal regions. The numerical discretization
Fig. 9. Refined geometry of the Nubian
slab (modified after Ganas and Parsons,
2009) inferred from global (ISC Bulletin,
2015) and temporary local (EGELADOS,
CYCNET, and LYBNET) seismicity catalogues. Dotted black lines indicate the
top of the slab isodepths. We assume
that intermediate-depth seismicity occurs within the subducting oceanic crust
(Abers, 1992; Kirby, 1995; Abers et al.,
2014) to revisit the geometry of the top
of the slab at depths larger than 50 km.
Dotted red lines indicate main discontinuities observed in the distribution
of the seismicity. The continuous red
line indicates the boundary between the
western and eastern slab segments subducting at the Hellenic Subduction Zone
(HSZ). KTF, Kefalonia Transform Fault;
HSZ, Hellenic Subduction Zone; NHSZ,
Northern Hellenic Subduction Zone;
WCSZ, Western Cyprus Subduction
Zone. (For interpretation of the references to colour in this figure legend, the
reader is referred to the web version of
this article.)
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Fig. 10. Pressure-temperature conditions of earthquakes in Fig. 3. Profiles 1 and 3 refer to the western and eastern slab segments of the HSZ, respectively (for the
location of the profiles refer to map in Fig. 3 or 5). The profiles to the left show the 2-D thermal structure calculated for the two slab segments and the distribution of
the seismicity along them, the top of the oceanic crust is indicated with a thick continuous black line. The plots to the right illustrate modeled pressure-temperature
conditions of earthquakes along the two cross-sections and the slab surface (red line) and slab Moho (blue line) pressure-temperature paths. (For interpretation of the
references to colour in this figure legend, the reader is referred to the web version of this article.)
the two segments although a discussion of the trends is more robust
(Wei et al., 2017). The pressure conditions are logically less affected by
uncertainties and show quite significant differences between the two
segments. The pressure conditions of intermediate-depth earthquakes in
the eastern slab segment (bottom-right, Fig. 10) are higher (4–6 GPa)
than those in the western slab segment of the HSZ (3–4 GPa; upperright, Fig. 10). The higher pressures found for the intermediate-depth
earthquakes in the eastern slab segment compared to the western slab
segment of the HSZ, together with the penetration of the seismicity
deeper in the slab, indicate the presence of an older slab in the area in
agreement with the results of Granot (2016). This suggests a similar
trend to that observed for intermediate-depth seismicity in the Tonga
arc (Wei et al., 2017).
with variable age that evolves within the 20 Myr model time. We ignore
the volcanic arc region as we assume that the highly scattered heat flow
near the volcanoes is due to advective magmatic input to the crust that
is not modeled here.
The slab couples to the mantle wedge at 80 km depth (following
Wada and Wang, 2009) which causes a focused cornerflow that brings
hot mantle in contact with the slab at depths greater than the decoupling point. The resulting models for the two profiles are shown in the
left column of Fig. 10. The cold anomaly clearly defines the subducting
slab. Sharp thermal gradients characterise the transition of the thermal
structure of the slab below the decoupling point with a rapid increase in
slab surface temperature at a pressure of about 2.5 GPa (red curve in
right column of Fig. 10). The thermal models indicate a slightly colder
eastern slab segment (profile 3, Fig. 10) compared to the western slab
segment (profile 1, Fig. 10) of the HSZ. The relatively short time of
20 Myr of cooling for the upper plate is sufficient to model main
properties of the surface heat flow. This is consistent with the time over
which, according to the single-slab model for the Aegean, the upper
plate has been cooling (20 Ma onset of oceanic subduction south of
Crete, van Hinsbergen and Schmid, 2012). The depth of the 1300 °C
isotherm (i.e., thermal lithosphere-asthenosphere boundary) in the
backarc area (60–80 km, Fig. 10) is roughly consistent with the lithosphere-asthenosphere boundary depth (60 km) inferred by seismological studies (Endrun et al., 2011; Sodoudi et al., 2015) supporting the
idea of newly formed mantle lithosphere beneath the backarc (Sodoudi
et al., 2015), and the single-slab model (Faccenna et al., 2003; Meier
et al., 2004a; van Hinsbergen et al., 2005; Jolivet and Brun, 2010).
In the absence of independent constraints we base the geometry of
the slab on the earthquake locations. Thus, the few earthquakes that are
located at the top of the slab very likely occurred inside it and may be
due to mislocation of the top of the slab (Fig. 10). Intermediate-depth
seismicity along the western slab segment of the HSZ seems to be
concentrated close to the top of the slab (profile 1, Fig. 10), possibly
within the crust, while along the eastern slab segment of the HSZ the
seismicity penetrates deeper in the slab and many events appear to
occur in the lithospheric mantle (profile 3, Fig. 10).
We computed the pressure-temperature conditions of the earthquakes occurring along the two slab segments using the thermal models
and report these on the right side of Fig. 10. Due to the strong thermal
gradients in the slab, relatively small errors in location, can lead to
relatively large errors in temperature. Therefore we cannot derive
strong conclusions about the temperature in a given earthquake along
5. Discussion
In this section we interpret the main results of this study and discuss
their implications. We relate the refined geometry of the seismically
active slab and its segmentation to the along-strike properties of the
seismicity at shallow and intermediate-depths (Section 5.1) as well as to
upper plate deformation (Section 5.2). We compare the evolution of the
slabs beneath northern Greece and the Peloponnese and relate it to the
presence of continental and oceanic subduction, respectively (Section
5.3). Furthermore, we discuss the presence of STEP-faults in the
southeastern Aegean (Section 5.4). Finally, we discuss the relation between the slab geometry and asthenospheric flow (Section 5.5).
5.1. Slab segmentation and properties of seismicity
The refined geometry of the HSZ proposed in Section 4.4 shows a
first-order slab segmentation between Crete and Karpathos with the
presence of a steeper dipping slab below Karpathos-Rhodes and a less
steep dipping slab below Crete and the Peloponnese. This poses the
question whether this geometry is also reflected in the seismic energy
release and seismic coupling.
To answer this question we investigate the occurrence of interplate
as well as intermediate-depth events on different time and magnitude
scales. While the instrumental ISC catalogue, when restricted to welllocated events, can give information over a longer time span (~50 year)
with reasonable depth control, temporary networks have the ability to
highlight seismically active areas due to their low magnitude of completeness and better location accuracies. On the other hand, historic
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Fig. 11. (a) Absolute moment density of interplate (diamonds) and intermediate-depth earthquakes (crosses) for cross-sections shown in Fig. 2. (b) Moment density
(diamonds) and event density (crosses) ratios between interplate (depth < 60 km) and intermediate-depth (d > 60 km) earthquakes for cross-sections shown in
Fig. 2.
magnitude (Md) or with unknown magnitude type have been excluded
from the calculations.
The seismic moment density of the interplate events shows a general
decrease from west to east with the strongest moment density release in
the intermediate-depth seismicity that is observed for profile 5 at the
eastern slab segment of the HSZ (Fig. 11a). The ratio between the
moment density of intermediate-depth and shallow events shows a
general increase from west to east (cross-sections 2 to 6, diamonds in
Fig. 11b). A similar trend is also observed in the ratios between the
number of intermediate-depth and interplate events (crosses in
Fig. 11b). Moreover, seismic moment density of shallow events is
higher than the seismic moment density of intermediate-depth events in
the cross-sections across the western slab segment of the HSZ (2 to 4),
while in Sections 5 (eastern slab segment of the HSZ) and 6 (Western
Cyprus Subduction Zone) it is the opposite (Fig. 11a, b). This means
that in the eastern slab segment of the HSZ, and in the slab subducting
beneath southwestern Anatolia, intermediate-depth seismicity is more
pronounced compared to the shallower seismicity in the interplate
seismogenic zone. These results are of course sensitive to the specific
definition of interplate as well as intermediate-depth events and the
location and width of the chosen profiles. However, the trend of eastward decreasing moment release densities of interplate seismicity and a
corresponding increase of the ratio of the moment density of intermediate-depth to interplate seismicity is also observed for other profile
widths (see Fig. 6 of the Supplementary material).
Lower seismic energy release at the plate interface towards the east,
as suggested by Fig. 11a, is also in agreement with observations of the
temporary local networks south of Crete and southeast of KarpathosRhodes. Over the duration of the EGELADOS seismic experiment only
few seismic events were recorded at plate interface depths (upper crosssection, Fig. 12). In contrast, the interplate seismogenic zone south of
eastern Crete, located in the region of the western segment of the HSZ,
has shown strong seismic activity during the LIBNET seismic experiment (lower cross-section, Fig. 12). Although it should be noted that the
station coverage of the LIBNET network reached further offshore
compared to the EGELADOS network (Fig. 12 and Fig. 1 of the Supplementary material), resulting in a larger lower completeness
threshold for interplate seismicity in the more distant offshore regions
catalogues, due to their long duration, might span the entire seismic
cycle of the study region but generally lack reliable event locations.
The seismic moment release of interplate as well as intermediatedepth events present in the ISC catalogue (reviewed ISC Bulletin, 2015)
along profiles 2–6 of Fig. 2 are depicted in Fig. 11a. We only consider
ISC catalogue events with magnitudes larger than the magnitude of
completeness (Mw4). We omit profile 1 here because it is not fully located in a region with oceanic subduction. To identify the respective
events, the topography of the downgoing slab as given in Fig. 9 is used.
The known offset between the ISC catalogue and local seismic networks
(Meier et al., 2004b) used to obtain the slab topography in this paper is
accounted for by shifting the slab geometry 15 km downwards. Events
occurring within a perpendicular distance of 100 km towards both sides
of the profiles and within a depth interval of 10 km above and 20 km
below the interpolated 3-D slab geometry are identified as interplate
events when their ISC depths are between 10 and 60 km (see Fig. 4 of
the Supplementary material). We choose a 10 km interval in the upward
and 20 km interval in the downward direction to account for location
errors in the ISC catalogue and the fact that the interplate seismicity
spans a finite depth interval extending downwards from the top of the
plate interface. Events deeper than 60 km occurring at an epicentral
position where the top of the slab is deeper than 60 km are considered
as intermediate-depth events (see Fig. 4 of the Supplementary material). Calculated seismic moment densities and ratios, shown in Fig. 11
(and Fig. 6 of the Supplementary material), are normalised with the
width of the seismicity considered across each profile and with the
approximate length of the interplate and intermediate-depth seismicity.
The seismic moment release is estimated by using the moment
magnitude (Mw) equation of Kanamori (1977). We converted bodywaves (mb) and surface-waves magnitude (Ms) to Mw using the
equations by Scordilis (2006). Ml was converted to Mw following
Papazachos et al. (1997), who found that for earthquakes in Greece, Ml
is half a unit smaller than the Mw in the low-magnitude range
(4.5 ≤ Ml ≤ 6.0). However, Scordilis (2006) highlighted the difficulties
to determine the conversion between Ml and Mw due to different amplifications given by different station types used to calculate Ml. We
noticed that assuming Mw = Ml or Mw = Ml + 0.5 does not affect the
final results significantly. Events with only an associated duration
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Fig. 12. Shallow seismicity (0–70 km) in the southeastern Aegean. Only hypocentres of well-located earthquakes (err ≤ 20 km) from the LIBNET (diamonds),
EGELADOS (circles) and CYCNET (triangles) temporary local catalogues are plotted in the maps and cross-sections. Hypocentres located 20 km towards each side of
the profiles are projected into the cross-sections. Dotted black lines in the map to the left represent the top of the slab isodepths while the dotted black line in the
profiles to the right represent the top of the slab; the thick red line indicates the location of the first-order discontinuity in the Nubian slab. The earthquakes are
coloured according to depth and sized according to magnitude. Black symbols indicate the temporary local network stations. (For interpretation of the references to
colour in this figure legend, the reader is referred to the web version of this article.)
this earthquake is still debated, e.g., Shaw et al., 2008); the M 6.1,
1664 CE; the M 6.6, 1805 CE; M 6.9, 1810 CE; the M 6.8, 1886 CE; M
6.5, 1910 CE; M 6.1, 1913 CE; M 7, 1952 CE; M 6.5, 1972 CE; M 6,
1973 CE. In addition, there is the 2008 CE, Mw 6.8 Methoni interplate
earthquake (e.g., Sachpazi et al., 2016b) which is not reported in the
Papadopoulos (2011) catalogue since it did not cause damage on the
Island of Crete. Instead, only one earthquake, namely the M ~ 8,
1303 CE, is indicated as a possible candidate for a megathrust earthquake along the eastern segment of the HSZ (Papadopoulos, 2011;
Papadopoulos et al., 2012). According to historical sources the 1303 CE
earthquake occurred to the southeast of Crete (Guidoboni and Comastri,
1997) close to the eastern termination of the western segment. However, its hypocentral location is not accurately resolved. Thus, it cannot
be determined whether the 1303 CE earthquake ruptured above the
south of Crete when compared to Karpathos. However, the observed
differences cannot be fully explained by lateral variations in the completeness threshold. For example, in the region directly south of Karpathos, which is inside the EGELADOS network, hardly any interplate
earthquake is detected, while towards the west, between Karpathos and
Crete, the EGELADOS network is able to identify events at depths
comparable to those also obtained for events located with the LIBNET
network (cluster of green circles at about 27°E/34.75°N in Fig. 12).
Finally, also historic earthquake catalogues (Fig. 13) can be considered to constrain seismic energy release at the plate interface and at
intermediate-depths. According to the historical catalogue of
Papadopoulos (2011), several large earthquakes potentially occurred
along the main plate interface above the western slab segment of the
HSZ including the: M ~ 8.3, 365 CE (although the interplate nature of
Fig. 13. Historical seismicity in the area of Crete and Rhodes
(Papadopoulos, 2011; Papadopoulos et al., 2012). Earthquakes are sized according to magnitude and coloured according to depth: interplate (occurred along the plate interface), shallow (overriding plate), intraplate (subducting
plate), unknown origin. The slab isodepths (black dotted
lines) and the observed location of the first-order discontinuity in the Nubian slab (red line) are shown in the
map. Labels indicate earthquakes with estimated Mw ≥ 7.
(For interpretation of the references to colour in this figure
legend, the reader is referred to the web version of this article.)
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seismicity from the ISC catalogue (ISC Bulletin, 2015) while the geometry of the top of the slab beneath northern Greece is constrained by
using receiver function results beneath a dense 2-D array (Pearce et al.,
2012). Furthermore, at the present stage we define the position of the
deformation front, of the backstop, of the Pindos thrust (Fig. 1) and of
the main tectonic units according to tectonic maps available for the
area (van Hinsbergen et al., 2005; van Hinsbergen and Schmid, 2012;
Jolivet et al., 2013). The longitude bar given at the bottom of Fig. 14 is
consistent with the current position of the tectonic elements (e.g.,
backstop, deformation front, thrust zones, etc.), but we note that this
might be not the case at earlier temporal stages in case of absolute
movements of the entire region.
According to kinematic reconstructions by van Hinsbergen and
Schmid (2012) about 260–280 km of continental lithosphere belonging
to Greater Adria have been consumed between about 34 and 15 Ma
along northern Greece, in proximity of the Greek-Albanian border, and
about 440–460 km along the southwestern Peloponnese. To constrain
the nature of the slab subducting at the HSZ, the length of the oceanic
lithosphere subducted at the current active margin needs to be estimated. In the following we discuss at first estimates of the length of
subducted oceanic lithosphere from the geological record and kinematic reconstructions. Later, we compare these estimates with the
down-dip limit of the seismically active slab.
At about 15 Ma, oceanic lithosphere belonging to the Eastern
Mediterranean Ocean started to subduct beneath the southwestern
Peloponnese (BB′, Fig. 14b), while the accretion of continental units
belonging to Greater Adria, was still ongoing in northern Greece. The
initiation of the subduction of the Eastern Mediterranean Ocean at
about 15 Ma beneath the southwestern Peloponnese is supported by the
acceleration of the clockwise rotation (~30°) of western Greece and
Peloponnese relative to the Moesian platform since about 17–15 Ma
(van Hinsbergen and Schmid, 2012), as indicated by paleomagnetic
rotations (Kissel and Laj, 1988), and consequent acceleration of extensional processes in the backarc at about 15 Ma (Jolivet and Brun,
2010 and references therein). The onset of oceanic subduction represents the easiest way to explain the estimated increasing extensional
rates (Faccenna et al., 2014). In available tectonic reconstructions (e.g.,
Jolivet et al., 2013) oceanic lithosphere is not yet subducting to the
southwest of the Peloponnese at 23 Ma. Moreover, according to
Broadley et al. (2006), the Ionian zone in northwestern Greece was still
part of Africa until about 15 Ma. Thus, geological data are consistent
with an age of about 15 Ma or a bit earlier for the onset of oceanic
lithosphere subduction to the southwest of the Peloponnese. Numerical
modeling results have predicted about 300 km of subducted oceanic
lithosphere since its initiation in middle Miocene (Royden and
Papanikolaou, 2011). From the tectonic reconstruction of Jolivet et al.
(2013), we estimate about 360–380 km of oceanic lithosphere subducted in the last 15 Ma. We draw our sketches (Fig. 14) by using the
amount of subducted oceanic lithosphere obtained from the tectonic
reconstruction of Jolivet et al. (2013).
western or the eastern slab segment of the HSZ. We note that according
to the first scenario all megathrust events occurred above the western
slab segment and the current lower interplate seismicity above the
eastern slab segment would also be observable in the historic catalogue
for the last ~2000 yrs. Instead, if the 1303 CE earthquake ruptured
above the eastern slab segment of the HSZ, the low interplate seismicity
observed from instrumental catalogues may reflect temporal changes of
seismic coupling along the plate interface above the eastern slab segment of the HSZ. Temporal changes of seismic coupling have been
suggested by Becker and Meier (2010) for the plate interface southwest
of Crete. As stated earlier, the precise location of historic events is often
not well-constrained and thus has to be critically assessed. The results
from the analysis of the known historic seismicity, however, are in
agreement with results from instrumental global and temporary local
catalogues.
It is interesting to note that GPS measurements show a currently
faster retreat of the trench along the eastern segment of the HSZ compared to the western segment (Reilinger et al., 2010). Thus, the low
seismicity along the interplate seismogenic zone above the eastern slab
segment may either indicate that the relative plate convergence is largely accommodated by aseismic slip (i.e., low seismic coupling) or that
the plate interface is locked (i.e., high seismic coupling). Combining
these seismological observations with the segmentation of the subducting slab between eastern Crete and Karpathos proposed in this
paper (Section 4.3), we suggest lower coupling of the plate interface
above the eastern slab segment compared to the plate interface above
the western slab segment of the HSZ. This lower coupling is also evident
for the interplate seismogenic zone of the Western Cyprus Subduction
Zone (Fig. 11a). We attribute the lower interplate seismicity and the
lower coupling above the eastern slab segment of the HSZ to enhanced
slab rollback and faster trench retreat, caused by the steeper subduction
of denser and older oceanic lithosphere (e.g., Granot, 2016). In fact,
strong slab-pull forces associated to denser subducting lithosphere induce lower coupling along the interplate seismogenic zone, whereas
weaker slab-pull forces associated to lighter subducting lithosphere
induce higher coupling along the interplate seismogenic zone (Scholz
and Campos, 1995, 2012). The higher moment release at intermediatedepth for the eastern slab segment of the HSZ compared to the western
slab segment of the HSZ (Fig. 11a) may be related to an increasing slab
age from west to east. Older slabs may release higher seismic moment
in the intermediate-depth range due to stronger hydration.
5.2. Oceanic vs continental subduction in western Greece
To clarify the origin of the slab at the western termination of the
HSZ we relate its geometry to available kinematic reconstructions for
the region (van Hinsbergen et al., 2005; Royden and Papanikolaou,
2011; van Hinsbergen and Schmid, 2012; Jolivet et al., 2013). Specifically, we reconstruct the tectonic evolution of the subduction along
two profiles that are orthogonal to the deformation front and cross
northern Greece (AA′ to the north of the Kefalonia Transform Fault;
Fig. 14a) and the Peloponnese (BB′ to the south of the Kefalonia
Transform Fault; Fig. 14b), respectively. The two profiles are located
close and parallel to the two cross-sections analysed by Pearce et al.
(2012) to retrieve information about the Moho depth of the upper and
downgoing plates. According to tectonic and geodynamic reconstructions (e.g., van Hinsbergen et al., 2005; Royden and Papanikolaou,
2011; van Hinsbergen and Schmid, 2012), we consider three significant
temporal stages for the evolution of the area, namely: (1) the closure of
the Pindos basin (~34 Ma); (2) the onset of oceanic subduction which
occurred only along the southern profile beneath the Peloponnese
(~15 Ma); and (3) the present situation.
We start by drawing the current geometry of the subduction zones
along the two cross-sections (Fig. 14a and b lower panels). The geometry of the top of the slab beneath the Peloponnese is constrained by
using well-located (vertical and horizontal errors smaller than 10 km)
5.2.1. Present (lower panels in Fig. 14a and b)
The cross-section sketching the present situation in the area of
Peloponnese shows the presence of oceanic lithosphere down to about
200 km (BB′, Present, Fig. 14b) in accordance with the above estimate
from tectonic reconstruction. Note that this estimate corresponds
roughly to the down-dip limits of the seismically active slab which
extends down to 180–200 km depth (cross-section 2, Fig. 2). It is widely
accepted that intermediate-depth seismicity is an indication for hydrated oceanic lithosphere (e.g., Hacker et al., 2003; van Keken et al.,
2011). Phase transitions at depth between about 50 km and 300 km are
invoked to explain the occurrence of seismicity at these depths (e.g.,
Hacker et al., 2003; van Keken et al., 2011). Thus, alternatively we may
estimate the length of the oceanic part of the slab by measuring the
length of the slab from the backstop to the down-dip limit of the seismicity. This length is about 370 km beneath southwestern Peloponnese
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Fig. 14. Sketch showing the different tectonic evolution of the western HSZ to the
north (AA′) and to the south (BB′) of the
Kefalonia Transform Fault. The evolution of
the subduction zones is shown at three
temporal steps: (1) the closure of the Pindos
basin which has occurred about 34 Ma; (2)
the onset of subduction of the Eastern
Mediterranean Oceanic Lithosphere which
has occurred about 15 Ma and only in the
area of Peloponnese; (3) and the present
day. In the sketch showing the present situation, the Pindos thrust and the limits
between tectonic units (simplified in this
study) have been taken from van
Hinsbergen et al. (2005), van Hinsbergen
and Schmid (2012), and Jolivet et al.
(2013); the geometry of the Apulian slab
subducting beneath northern Greece is inferred from a scattered-waves study (Pearce
et al., 2012).
microcontinents has been delaminated from the accreted upper crustal
parts and has been subducted (Faccenna et al., 2003; Meier et al.,
2004a; van Hinsbergen et al., 2005; Jolivet and Brun, 2010).
In contrast to the Peloponnese (BB′, Present, Fig. 14b), the slab
subducting beneath northern Greece (AA′, Present, Fig. 14a) is composed at least down to about 200 km depth of continental lithosphere
which has been subducting in the area since about 34 Ma (closure of the
Pindos basin), consuming about 400 km of Greater Adria, including the
Tripolitza zone, the Ionian zone, the pre-Apulian zone, and the Apulian
platform which is still subducting (van Hinsbergen and Schmid, 2012;
AA′, Present, Fig. 14a). At depths larger than about 220 km the slab
might be composed of mantle lithosphere and lower crustal portions
belonging to the Pindos basin (AA′, Present, Fig. 14a). According to a
scattered waves study, continental mantle lithosphere and crystalline
continental lower crust are subducting beneath northern Greece (Pearce
et al., 2012; AA′, Present, Fig. 14a). The slab in this area is seismically
inactive (only three earthquakes with magnitudes above 3 are reported
by the ISC catalogue at the Greek-Albanian border; Fig. 4c) in contrast
to the slab in the HSZ (cross-sections 7, Fig. 2). This observation supports the conclusion that the seismically active part of the slab in the
HSZ is of oceanic nature.
(cross-section 2, Fig. 2). We assume that there was no sedimentary
accretionary prism to the back-side of the Ionian Unit (or that its width
was not relevant) when the subduction of the Eastern Mediterranean
Ocean initiated (BB′, ~15, Fig. 14b). This is compatible with the present situation along the Northern Hellenic Subduction Zone where the
deformation front is located at the contact between the upper and lower
plate, leaving no space for the development of a sedimentary accretionary prism (AA′, Present, Fig. 14b). The seismologically estimated
oceanic slab length (~370 km) is in agreement with the oceanic slab
length (~360–380 km) estimated from tectonic reconstruction by
Jolivet et al. (2013). Also the seismological estimate is associated with
considerable uncertainties as the amount of hydration and the pressure
and temperature conditions in the slab may influence the present downdip limit of seismicity (e.g., Hacker et al., 2003). The agreement between both estimates is therefore remarkable and surprising. We
therefore suggest that the seismically active part of the slab in the HSZ
is indicating roughly the presence of oceanic lithosphere that has been
subducted at the current active margin.
Moreover, we observe that beneath the Peloponnese the top of the
subducting slab is located directly beneath the Moho of the upper plate.
That means the mantle lithosphere of the upper plate has been removed
during subduction (BB′, Present, Fig. 14b). In the sketch, we use a
standard incoming plate thickness of about 100 km for both Phanerozoic continental as well as old oceanic lithosphere because detailed
estimates are not available. The lithospheric thickness of the upper
plate at present is drawn in consistency with seismological observations, which indicate the lithosphere-asthenosphere boundary beneath
the central and northern Aegean at about 60 km depth (e.g., Endrun
et al., 2011; Sodoudi et al., 2015). The mantle lithosphere of the Aegean
upper plate has likely developed after the collision of Greater Adria
with Eurasia (Sodoudi et al., 2015). This is a consequence of the singleslab model according to which the original mantle lithosphere of the
5.2.2. 34 Ma (upper panels in Fig. 14a and b)
Cross-sections through the lithosphere for the two profiles are also
sketched for the time of the closure of the Pindos basin with the
Tripolitza zone entering the trench at about 34 Ma (e.g., Stampfli and
Borel, 2004). The sketch across northern Greece has been obtained by
restoring about 400 km of subducted continental lithosphere, measuring them from the Pindos thrust (AA′, 34 Ma, Fig. 14a). Across the
Peloponnese the sketch showing the tectonic setting at 34 Ma is obtained by restoring the subduction of about 440–460 km of continental
lithosphere, again measuring them from the position of the Pindos
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Özbakır et al., 2013; Govers and Fichtner, 2016).
By definition, a STEP-fault is the superficial manifestation (in the
upper plate) of a Subduction-Transform-Edge-Propagator or STEP in the
subducting slab (Govers and Wortel, 2005). The existence of a NE-SW
trending STEP-fault south of Rhodes has been suggested because of the
lack of clear evidence for the existence of an eastern segment of the HSZ
that is now provided by a detailed analysis of microseismicity in the
area. Thus, on the basis of seismological evidence, we rule out the
presence of a NE-SW trending STEP between the slab segments subdcuting beneath the HSZ and the Western Cyprus Subduction Zone and
therefore of a STEP-fault in the region of the Pliny-Strabo trenches. In
fact, the eastern termination of the slab subducting beneath the Aegean
region is found parallel to the NW-SE trending coastline of southwestern Anatolia (Fig. 9).
The geometry of the slab in the area of Rhodes (Fig. 9) allows for the
presence of a NW-SE trending STEP-fault along the southwestern Anatolian coast. The GPS-velocities support the existence of such a fault.
However, there are no indications for NW-SE trending seismically active STEP-fault along the coast line of southwestern Anatolia in the ISC
catalogue spanning the time 1964–2014, neither in the catalogues of
our local experiments (EGELADOS Oct. 2005–Mar. 2007). However, we
cannot exclude the possibility that it existed in the past or that it is
currently deforming aseismically.
Because of the absence of a NE-SW striking edge of the slab in the
area south of eastern Crete and Rhodes, sinistral strike-slip deformation
in the forearc is explained by increasing curvature and lengthening of
the forearc induced by slab rollback and oblique subduction (ten Veen
and Kleinspehn, 2002, 2003; Meier et al., 2007). Transtensional slip on
070° sinistral wrench zones both onshore and offshore between Rhodes
and Crete result from an interplay between the southwestwards
movements of the upper plate, arc-parallel stretching of the forearc (ten
Veen and Kleinspehn, 2002, 2003) and the rollback of the eastern slab
segment of the HSZ. The transtensional motion has started when the
inner forearc reached a certain threshold of obliquity (ten Veen and
Kleinspehn, 2002, 2003). In the Messara basin this mechanism was
active since about 3.4 Ma (ten Veen and Kleinspehn, 2003) while in the
Apolakkia basin on Rhodes Island, this deformational phase, which was
active both onshore and offshore, started at about 4–5 Ma (ten Veen and
Kleinspehn, 2002).
The existence of a subducting slab in the area of Rhodes is supported
by other independent observations. Oblique convergence between the
Nubian slab and the Aegean region, between eastern Crete and Rhodes,
has been invoked by Kopf et al. (2001) to explain the presence of mud
volcanoes, in the form of mud pies, in the Pliny-Strabo trenches and to
the southern margin of the Rhodes basin (Huguen et al., 2001, 2006),
and by ten Veen and Kleinspehn (2002) to explain the transpressional
deformation in the outer forearc, to the southeast of the Pliny trench.
Moreover, we interpret the GPS velocity field in the southeastern Aegean (calculated with respect to central Aegean and the Peloponnese),
showing vectors normal to the trench and velocities increasing trenchward (Reilinger et al., 2010), as consequence of the stronger rollback
of the eastern slab segment of the HSZ with respect to the western one.
Kontogianni et al. (2002) associated the Holocene uplift and subsidence
in Rhodes to the presence of off-shore (southeast of Rhodes) compressional structures parallel to the coast line. We suggest that a NW dipping slab in the area of Rhodes would provide and easy explanation for
the presence of such structures.
Two additional STEP-faults, both trending NE-SW, might be present
in the Western Cyprus Subduction Zone. One is above the western and
the other is above the eastern edges of the NE dipping slab subducting
in the area beneath southwestern Anatolia (Fig. 9). There is no evidence
for a seismically active STEP-fault above the western slab edge of the
Cyprus slab in the region of southwestern Antolia. A shallow seismically
active NE-SW trending zone is however found above the eastern slab
edge of the Western Cyprus Subduction Zone, in proximity of the
western coast of Cyprus (Fig. 2). This may be interpreted as a STEP-
thrust at the earlier stage (i.e., ~15 Ma), and about 360–380 km of
oceanic lithosphere (BB′, 34 Ma, Fig. 14b). To retro-deform the sketches
at earlier temporal stages we assume that the upper plate and therefore
the main thrust zones (i.e., Pindos thrust) remained fixed during continental collision (crustal units are accreted to the upper plate and the
delamination front migrates backward) while they migrate trenchward
during oceanic subduction due to slab-rollback and stretching of the
upper plate.
5.2.3. 15 Ma (panel in the middle of Fig. 14b)
At about 15 Ma, oceanic lithosphere belonging to the Eastern
Mediterranean Ocean started to subduct beneath the southwestern
Peloponnese, while the pre-Apulian lithosphere was subducting beneath northern Greece and the Ionian Islands (e.g., van Hinsbergen
et al., 2006). Thus, in the area of Peloponnese the subduction stepped
back to the back-side of Greater Adria. The cross-section sketching the
situation at about 15 Ma has been drawn by restoring about
360–380 km of subducted oceanic lithosphere and by shifting the upper
plate and the Pindos thrust by about 40 km towards northeast from
their current position (~15 Ma, BB′, Fig. 14b) according to van
Hinsbergen and Schmid (2012). At this stage the subducting slab segment beneath the Peloponnese, shown in the sketch, is entirely composed of continental lithosphere (BB′, ~15 Ma, Fig. 5b).
The reconstuction proposed in Fig. 14 clearly depicts the different
evolution of the subduction system to the north and to the south of the
Kefalonia Transform Fault. This is indicated by the presence of a seismically almost inactive continental slab subducting beneath northern
Greece (AA′, Present, Fig. 5a) and a seismically active oceanic slab
subducting beneath the southwestern Peloponnese (BB′, Present,
Fig. 5b). While the transition between oceanic and continental lithosphere on the surface is delimited by the Kefalonia Transform Fault, at
larger depths it is not well-defined and it is suggested, as in this study,
to be located at the northward termination of the intermediate-depth
seismicity (Halpaap et al., 2018). The abrupt termination of intermediate-depth seismicity towards the Northern Hellenic Subduction
Zone (Figs. 1, 4) may either be caused by a sharp transition in the
compositional properties of the subducting lithosphere, with the presence of a continuous slab (Pearce et al., 2012; Halpaap et al., 2018), or
by a trench-normal vertical tear between the oceanic and continental
slabs (Govers and Wortel, 2005; Suckale et al., 2009). The existence of
the Kefalonia Transform Fault points to localised deformation above the
slab edge at shallow depths. Furthermore, the presence of oceanic lithosphere beneath the continental crust in the Peloponnese and of
continental lithosphere beneath the continental crust in northern
Greece requires considerable relative displacement between the two
slab segments. In fact, according to the reconstruction proposed in
Fig. 14 the two slab segments along the shown profiles have been laterally displaced by about 200–250 km in the last 15 Ma. Although the
relative lateral displacement reduces towards the boundary between
the two slab segments, due to the more recent onset of oceanic subduction towards the Ionian Islands, it should still be larger than 100 km.
This favours the presence of a discontinuity, at shallow depth, between
the oceanic and continental slab. Additional tomographic and local
seismicity studies are needed to resolve unambiguously the structure in
the area.
5.3. STEP-faults in the southeastern Aegean
The existence of a NW dipping slab in the area of Rhodes, also
suggested earlier by studies of seismicity (Knapmeyer, 1999;
Papazachos et al., 2000b; Brüstle, 2012), and receiver function analysis
(Sodoudi et al., 2006, 2015) it is in conflict with the presence of a NESW trending STEP-fault in the region of the Pliny-Strabo trenches,
formed in response to the tearing between the HSZ and the Western
Cyprus Subduction Zone (e.g., Govers and Wortel, 2005; Dilek and
Altunkaynak, 2009; van Hinsbergen et al., 2010b; Biryol et al., 2011;
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Tectonophysics 734–735 (2018) 96–118
G.M. Bocchini et al.
rest of the Aegean region (Reilinger et al., 2010). Focal mechanisms
(Global CMT, 2016) of the larger magnitude earthquakes (Mw ≥ 5)
along this S-shaped seismicity zone exhibit an extensional regime
normal to the strike of this fault zone (Fig. 15). This supports the detachment of the Rhodes block from the central Aegean region.
Two dense clusters of seismic events are observed to the east of the
Island of Astypalea, and between the Islands of Amorgos and Santorini
(Bohnhoff et al., 2006), indicating that the deformation is not limited
along a sharp boundary but in a broader region (Fig. 15). A focal mechanism of an earthquake with Mw ≥ 5 in the region of the Astypalea
cluster (Global CMT, 2016) exhibits normal faulting on a fault plan
striking NE-SW, which is compatible with the focal mechanisms observed between Nisyros and Gökova (Fig. 15). The region between
Amorgos and Santorini is characterised by a series of grabens bordered
by NE-SW striking normal faults (Perissoratis and Papadopoulos, 1999).
The Ms = 7.4, July 9, 1956 Amorgos earthquake, the largest tsunamigenic earthquake recorded in the Aegean region in the last century,
occurred along one of these faults, namely the Amorgos fault
(Makropoulos et al., 1989; Perissoratis and Papadopoulos, 1999). The
most accepted focal mechanism solutions for the 1956 Amorgos
earthquake shows NE-SW striking normal faulting with the preferred
nodal plane being the SE dipping one (Okal et al., 2009; Brüstle et al.,
2014). The slip vector of the SE dipping fault calculated from the focal
mechanism of Brüstle et al. (2014) is consistent with the current upper
plate deformation pattern indicated by GPS. Thus, we suggest that the
1956 earthquake can be associated to the ongoing upper plate deformation observed in the southeastern Aegean, which is related to the
segmentation of the subducting slab between Crete and Karpathos and
to the enhanced slab rollback in the eastern slab segment of the HSZ.
Interestingly, the active Kos-Nisyros volcanic centre (e.g.,
Bachmann et al., 2010) is located along the north-northwestern part of
the seismologically delineated boundary between the Aegean region
and the Rhodes block. This boundary traces the deep basis to the
southeast of Kos and could be a near-surface reason for the final location of the sub-recent Kos-Nisyros volcanic centre (active from
~3–4 Ma, Bachmann et al., 2010).
Fig. 15. Deformation pattern observed in the southeastern Aegean region.
Black arrows show the GPS velocity vectors from Reilinger et al. (2010). Coloured focal mechanisms represent Harvard CMT for the period 1976–2015
with Mw ≥ 5 (Global CMT, 2016). The black and white focal mechanism is the
focal mechanism of the 1956 Amorgos earthquake (Brüstle et al., 2014), for
which is shown the slip vector calculated on the SE dipping nodal plane. In the
background shallow seismicity (0–50 km) recorded by the EGELADOS and
CYCNET temporary local networks is shown. Seismicity and focal mechanisms
are colour coded according to the depth and sized according to the magnitude.
The thick dotted black line indicates the S-shaped alignment of the upper plate
seismicity which is consistent with the boundary of the Rhodes block inferred
by GPS velocity vectors (Reilinger et al., 2010).
5.5. Slab tearing, asthenospheric flow, and curvature of the HSZ
A characteristic feature of the HSZ is the asymmetric amphitheatrelike shape of the subducting slab as shown in the 3-D sketch in Fig. 16.
The slab has a stronger curvature to the east of Crete compared to the
western part of the subduction zone. Moreover, the along-strike length
of the slab is larger to the west of the point of maximum curvature
(located approximately south of Crete) compared to the east. The slab
dipping angle is larger in the southeastern Aegean. In the area of Karpathos, underthrusting of the narrower and steeper eastern slab segment (dark blue) beneath the wider and less steep western slab segment
(light blue) of the HSZ is observed (Fig. 16). The eastern slab segment of
the HSZ is separated from the slab in the Western Cyprus Subduction
Zone by a slab tear opening towards larger depths. The narrow gray
slab subducting to the north of the Kefalonia Transform Fault indicates
seismically almost inactive, subducting continental lithosphere.
The strong curvature and the asymmetry of the HSZ hint at a noncylindrical evolution of the subduction system that may be related: (1)
to the dynamics of the slab at depth (e.g., rollback, tearing, asthenospheric flow); (2) to the diachronous onset of oceanic subduction; (3) to
along-strike heterogeneities of the incoming slab (e.g., continental vs.
oceanic subduction). In the following we discuss the relation of the slab
shape to these processes.
The strong curvature is first of all related to slab rollback and fast
trench retreat. Fast slab retreat has been invoked as driving mechanisms
for the generation of strongly curved arcs in the Western Mediterranean
Sea (Faccenna et al., 2004). Numerical models have shown that narrow
slabs (length ≤ 1500 km) retreat fast and develop curved geometry,
concave towards the mantle wedge side, with trench retreat velocities
fault. This possible STEP-fault, similarly to the Kefalonia Transform
Fault, is located above the transition between subducting oceanic and
continental lithosphere.
5.4. Slab segmentation and upper plate deformation
It has been shown that sub-horizontal mantle flow induced by slab
rollback is able to produce tectonically significant shear stresses causing
upper plate deformation (Sternai et al., 2014). In the following we
discuss the relation between slab segmentation of the HSZ and the
deformation of the upper plate in the southeastern Aegean.
The upper plate deformation is analysed by combining: shallow
seismicity recorded by the temporary networks (EGELADOS and
CYCNET catalogues), focal mechanisms of larger earthquakes (Mw ≥ 5,
1976–2015), and geodetic data (Fig. 15). Shallow seismicity in the
overriding plate defines a peculiar S-shaped zone (Brüstle, 2012) which
stretches from the south of eastern Crete towards Nisyros, and from
Nisyros towards the Gulf of Gökova (southwestern Turkey), and seems
to delimit the borders of the Rhodes block indicated by GPS data
(Fig. 13, Reilinger et al., 2006, 2010). In fact, GPS velocity vectors
indicate southeastward drifting of the southeasternmost part of the
Aegean region, namely the Rhodes block, and its detachment from the
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G.M. Bocchini et al.
Fig. 16. 3-D sketch summarising the main findings of the study. Only the seismically active portions of slabs are shown with the exception of the slab subducting to
the north of the Kefalonia Transform Fault (gray) which is seismically almost inactive. The light blue slab represents the western slab segment of the HSZ. The dark
blue represents the eastern slab segment of the HSZ and the slab subducting at the Western Cyprus Subduction Zone which are separated by a vertical tear (slab
window). The white arrows show the relative motion between the slab segments identified in this study. Gray arrows indicate mantle flow. Zigzag patterns are used
to indicate that the slab does not terminate but extends in that specific direction. Vertical exaggeration by a factor of about 2. (For interpretation of the references to
colour in this figure legend, the reader is referred to the web version of this article.)
Turkey between 16 and 5 Ma (Kissel and Poisson, 1987; van Hinsbergen
et al., 2010a) may have contributed to the opening of the triangular
slab window as indicated in Figs. 9 and 16. The curvature and asymmetry of the HSZ have likely been further enhanced by the segmentation of the slab into a western and an eastern segment.
The model of asthenospheric flow through the tear and in the
backarc region as indicated in Fig. 16 is supported by observations of
seismic anisotropy. SKS splitting measurements show fast directions of
shear-wave propagation trending NE-SW or NNE-SSW from the Aegean
backarc to Anatolia, and NNW-SSE or NW-SE below northern Greece
and the Peloponnese (e.g., Jolivet et al., 2009; Evangelidis et al., 2011;
Paul et al., 2014). Beneath southwestern Turkey, Paul et al. (2014) have
imaged SKS anisotropy with fast orientations trending NW and have
related them to the toroidal flow through the tear. Flow of asthenospheric mantle of African provenance into the Aegean mantle wedge in
the backarc of the HSZ has been further supported by the prominent
geochemical signature of enriched subslab mantle component found in
the classical calcalkaline subduction-related magmas in the Kos-Nisyros
and Santorini volcanic centres (Klaver et al., 2016). Further, we suggest
that the high topography found above the slab tear (cross-section 8 in
Fig. 2) may be related to isostatic uplift associated to the presence of a
shallow asthenospheric mantle in the area of the tear.
maximum in the centre and decreasing towards the edges (Schellart
et al., 2007). In case of the HSZ, the slab retreated south-southwestward
with maximum velocities in the area of Crete and decreasing towards
the western (Kefalonia Transform Fault) and eastern (southwestern
Turkey) extremities (e.g., ten Veen and Kleinspehn, 2002).
The onset of the Eastern Mediterranean Oceanic Lithosphere subduction beneath the Aegean region has been suggested to be diachronous, becoming younger from east to west (van Hinsbergen and
Schmid, 2012 and references therein). It occurred at about 35 Ma to the
southwest of Anatolia, at about 20 Ma south of Crete, at about 15 Ma
southwest of the Peloponnese, and about 4–5 Ma in the area of the
Ionian Islands (van Hinsbergen and Schmid, 2012 and references
therein). This resulted in an east-west along-strike differential rollback
and slab retreat, and contributed to the asymmetry of the HSZ. The
clockwise rotation of western Greece which started about 25 Ma (van
Hinsbergen and Schmid, 2012) is related to this differential rollback.
The slab started to curve since 25–23 Ma, however it was after
15 Ma that the curvature of the slab has increased significantly (van
Hinsbergen and Schmid, 2012; Jolivet et al., 2013; Faccenna et al.,
2014). The acceleration of slab retreat and the increased curvature in
the east are indicated by upper plate extension (e.g., Faccenna et al.,
2003; Brun and Sokoutis, 2010; van Hinsbergen and Schmid, 2012),
and by clockwise rotation of western Greece and conterclockwise rotation of southwestern Anatolia (van Hinsbergen and Schmid, 2012).
The accelerated slab retreat and the development of a highly curved
slab are likely related to slab tearing beneath southwestern Anatolia
which became fully efficient at about 15 Ma (e.g., Brun and Sokoutis,
2010; Jolivet et al., 2015). Slab tearing allows for toroidal asthenospheric mantle flow around the edges of the slab and towards the north
through the slab tear. Toroidal asthenospheric flow is enhancing slab
rollback in proximity of the tear (e.g., Schellart et al., 2007). As a
consequence of the tearing, the subduction zone in the southeastern
Aegean and western Anatolia split into the Cyprus and Hellenic Subduction zones. The 25–30° counterclockwise rotation of southwestern
6. Conclusions
Distribution of seismicity is well suited to study the slab geometry
and active processes particularly at the top of subducting slabs. Using
well-located hypocentres from global and temporary local seismicity
catalogues, we propose a refined model of the seismically active Nubian
slab subducting beneath the Aegean region and southwestern Anatolia.
The model confirms the amphitheatrical shape of the slab as suggested
before. Other findings include: (1) the extension of the oceanic subducting slab to the north of the Gulf of Corinth (a horizontal tear is
unlikely to exist beneath the Gulf); (2) a discontinuity at shallow depths
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Tectonophysics 734–735 (2018) 96–118
G.M. Bocchini et al.
deformation of the overriding plate in the area of Rhodes, where the
Rhodes block is moving southeastward with respect to the central
Aegean and the Peloponnese.
between the seismically active oceanic slab and the almost seismically
inactive continental slab beneath northern Greece; (3) the presence of
the subducting slab directly below the continental crust of the Adriatic
Units beneath the Peloponnese supporting the single-slab model involving delamination and subduction of the mantle lithosphere and
eventually parts of the lower crust; (4) the presence of a NW dipping
subducting slab in the area of Rhodes which rules out the presence of a
NE-SW trending Subduction-Transform-Edge-Propagator between the
HSZ and the Western Cyprus Subduction Zone, and of a STEP-fault in
the region of the Pliny-Strabo trenches; (5) the existence of a vertical
tear in the slab below about 20–40 km between the eastern segment of
the HSZ and the slab subducting beneath southwestern Anatolia
opening towards larger depths; (6) a first-order segmentation of the
subducting slab in the region between Crete and Karpathos, with a less
steep and laterally wider slab segment to the west and a steeper and
laterally narrower slab segment to the east; and finally (7) an asymmetry in the slab curvature that increases towards the east.
A model for the evolution of the subduction is proposed for the area
of Peloponnese. The model shows the backstepping of the subduction
from the front-side of the accreted Adriatic Units (~34 Ma) to their
back-side (~15 Ma), with the onset of oceanic lithosphere subduction in
the Ionian Sea. To the contrary, in northern Greece ongoing thrusting of
upper continental units and subduction of the underlying lithosphere is
observed. According to the proposed tectonic evolution, a STEP-fault,
namely the Kefalonia Transform Fault, developed as consequence of
differential trench retreat triggered by the onset of oceanic subduction
beneath the Peloponnese. The opening of the Gulf of Corinth and of the
Amvrakikos fault zone (since 3.5 Ma), by decoupling the continental
units in northern Greece from those in the Peloponnese, caused the
cessation of the strike-slip activity along the Thesprotiko-Aliakmon
fault zone (active from 7 to 3.5 Ma) and its localisation along the
Kefalonia Transform Fault. The onset of oceanic subduction in the area
of the Ionian Islands at about 4 Ma likely allowed the reinitiation of
strike-slip deformation.
The diachronous onset of the Eastern Mediterranean Oceanic
Lithosphere subduction beneath the Aegean region, becoming younger
from east to west, caused along-strike differential rollback and slab
retreat and contributed to the current asymmetry of the HSZ. Slab
tearing beneath southwestern Anatolia accelerated slab rollback and
significantly increased the curvature of the HSZ since 15 Ma. Toroidal
flow around the slab edges further enhanced slab rollback in the
southeastern Aegean. The curvature and asymmetry of the HSZ has
likely been enhanced by the segmentation of the slab into a western and
an eastern segment.
Between Crete and Karpathos a double seismic zone reveals the
overthrusting of the western segment above the eastern segment of the
Nubian slab. This is caused by the increasing curvature of the slab towards larger depths which induces a compressional regime at intermediate-depth along the strike of the slab. Thermal modeling demonstrates that the seismicity penetrates deeper in the colder slab
subducting beneath the eastern segment of the HSZ, which is similar to
that observed for the Tonga subduction zone (Wei et al., 2017).
An increase in intermediate-depth seismicity towards the east, with
respect to the shallow seismicity, is observed. It likely reflects increased
slab hydration and decoupling in the interplate seismogenic zones of
the eastern slab segment of the HSZ and Western Cyprus Subduction
Zone compared to the western slab segment of the HSZ, possibly due to
fluid presence. This is very likely linked to the older age, and thus
colder temperature, of the oceanic lithosphere in the Mediterranean
south of western Turkey.
A change in the deformation of the upper plate in response to the
segmentation can be observed above the boundary between the western
and eastern slab segments of the HSZ. West of the boundary, in the
inner part of the Sedimentary Arc, the Aegean region is seismically not
very active while a stronger seismic activity in the overriding plate is
observed to the east of the boundary. GPS data suggest active
Acknowledgements
Gian Maria Bocchini, and Marija Ruscic have been funded by the
People Program (Marie Curie Actions) of the European Union 7th
Framework Programme FP7-PEOPLE-2013-ITN under REA grant
agreement no. 604713. Andrea Brüstle and Martina Rische have been
founded by the German Research Foundation (5483572) (DFG) within
the collaborative research centre 526 “Rheology of the Earth: From the
Upper Crust to the Subduction Zone”. We thank the GIPP (Geophysical
Instrument Pool of Potsdam at Geo-Forschungs-Zentrum, GFZ) for
providing portable seismic instruments for the CYCNET and the
EGELADOS temporary networks. We are grateful to Edi Kissling (ETH
Zurich) for his valuable support in the determination of the 1-D
minimum velocity model for the EGELADOS dataset and to Thomas
Pettke (University of Bern) for his feedback on the geochemistry of
volcanism in the Eastern Aegean. We thank Bernhard Stöckhert for
helpful discussions on the tectonic evolution of the HSZ. We thank
Douwe van Hinsbergen for his constructive review that helped to improve the manuscript substantially. We also thank an anonymous reviewer for the detailed comments on the manuscript. Figures have been
generated with the Generic Mapping Tools (Wessel and Smith, 1995)
and the background bathymetry using ETOPO1 and GMRT data.
Appendix A. Supplementary data
Supplementary data to this article can be found online at https://
doi.org/10.1016/j.tecto.2018.04.002.
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