Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece

Abstract

We deployed an on/offshore seismic network consisting of 36 three-component stand-alone digital stations in the Saronikos Gulf and surrounding area and observed the microseismic activity from 17 August to 9 December 2001. The geometry and location of the network were designed to observe the offshore activity to connect known active faults onshore with those offshore. An average of 15 events per day were recorded by a minimum of six stations. The focal parameters of 739 events could be defined. Within the network area we located 306 events with an rms travel-time residual of about 0.2 sec. The epicenters in the Saronikos Gulf are aligned mainly along an east–west trending fault with a length of nearly 30 km. This fault is located north of Aegina and could produce earthquakes as large as M s 6.5. It is the offshore continuation of a seismic zone located in east Corinthia. A second active fault was identified east of Aegina, trending northeast–southwest. It is probably linked to the northeast–southwest striking zone of the Salamis-Fili seismic region, activated by the September 1999 event of Athens. Outside the network we mapped an active area at the Parnis mountain, still active after the M s 5.9 event, and a second one further northeast at the western coast of Evia. This zone is close to Chalkis and the recently active Psachna area. The network also mapped intense activity at the northeastern coast of Evia, close to the city of Kymi and the area of Skyros, which was seriously damaged by an M s 5.8 earthquake in 2001. The onshore/offshore local network identified several regions of very high seismic activity not detectable by regional networks, and demonstrated once more the advantage of deploying local seismic networks for mapping active deformation in very short periods of time.

Full text

920
Bulletin of the Seismological Society of America, Vol. 94, No. 3, pp. 920–929, June 2004
Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece
by J. Makris, J. Papoulia, and G. Drakatos
Abstract We deployed an on/offshore seismic network consisting of 36 three-
component stand-alone digital stations in the Saronikos Gulf and surrounding area
and observed the microseismic activity from 17 August to 9 December 2001. The
geometry and location of the network were designed to observe the offshore activity
to connect known active faults onshore with those offshore. An average of 15 events
per day were recorded by a minimum of six stations. The focal parameters of 739
events could be defined. Within the network area we located 306 events with an rms
travel-time residual of about 0.2 sec. The epicenters in the Saronikos Gulf are aligned
mainly along an east–west trending fault with a length of nearly 30 km. This fault
is located north of Aegina and could produce earthquakes as large as M
s
6.5. It is
the offshore continuation of a seismic zone located in east Corinthia. A second active
fault was identified east of Aegina, trending northeast–southwest. It is probably
linked to the northeast–southwest striking zone of the Salamis-Fili seismic region,
activated by the September 1999 event of Athens. Outside the network we mapped
an active area at the Parnis mountain, still active after the M
s
5.9 event, and a second
one further northeast at the western coast of Evia. This zone is close to Chalkis and
the recently active Psachna area. The network also mapped intense activity at the
northeastern coast of Evia, close to the city of Kymi and the area of Skyros, which
was seriously damaged by an M
s
5.8 earthquake in 2001. The onshore/offshore local
network identified several regions of very high seismic activity not detectable by
regional networks, and demonstrated once more the advantage of deploying local
seismic networks for mapping active deformation in very short periods of time.
Introduction
The region of Attica and, in particular, that of the city
of Athens have long been considered aseismic (Sieberg,
1932). This was mainly attributed to the fact that the his-
torical center of the City, built on schist of high rigidity, is
not vulnerable to earth tremors. However, since Athens has
been constantly expanding, densely urbanized areas haveen-
croached into seismogenic regions and Athens has become
vulnerable to earthquakes. Today, according to estimates of
Papadopoulos and Arvanitides (1996) the seismic risk of
Athens is among the highest in Greece. This was verified
dramatically by the M
s
5.9 earthquake of 7 September 1999,
which was the costliest natural disaster of Greece in recent
history.
Prior to this event and during the instrumental period of
Greek seismology, no significant earthquake activity had
been reported. The strongest shocks (M
s
6.0) to affect Athens
were located in the Thiva-Oropos zone at a distance of more
than 50 km from the historic center of the city (Papazachos
and Papazachou, 1997). Events of rather small magnitude
recorded from time to time in the vicinity of Athens, like
those of 3 April (M
s
4.2) and 4 December (M
s
3.5) 1965 and
a cluster of events ranging from 3.8 to 4.5 in November 1997
[see the National Observatory of Athens (NOA) bulletins],
did not attract public interest. It was therefore assumed that
a detailed examination of the microseismicity of the Athens
area was unnecessary.
The 1999 event, however, clearly demonstrated that the
low rate of seismicity alone is not a safe criterion to assess
the seismic potential of a region. Only a detailed tectonic
analysis and mapping of active faults and soil conditions
contribute toward a reliable seismic risk assessment. Had
such studies been undertaken before 1999, they certainly
would have revealed the potential risk of the active tectonic
structures that ruptured during the 7 September 1999 seismic
crisis.
Public awareness and the need for a thorough under-
standing of the regional seismic potential after the 7 Septem-
ber 1999 event motivated the Earthquake Planning and
Protection Organization of Greece to support an onshore-
offshore microseismicity survey in the region of the Saron-
ikos Gulf, whose purpose was to identify active faults on-
and offshore and to define their seismic potential. A local

Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece 921
seismic network was deployed consisting of eight ocean bot-
tom seismographs (OBS) and 28 land stations (Fig. 1). This
network operated for 45 days.
Both marine and land stations were equipped with
SEDIS seismic recorders of GeoPro, Hamburg (Makris and
Moeller, 1990; GeoPro technical reports). These instruments
are capable of recording six seismic channels on hard disc
with 10-GB capacity. Onshore, an integrated Global Posi-
tioning System (GPS) receiver provides coordinates and tim-
ing at each location. Marine stations are housed in a 17-inch
glass sphere and record seismic signals directly on the sea
floor. A gimbal-mounted three-component geophone system
and a deep sea hydrophone are used for this purpose. The
instruments are autonomous and can be acoustically ad-
dressed by a surface-mounted system. Timing is provided
by an internal quartz clock. By measuring the quartz tem-
perature continuously and GPS time at the beginning and end
of the survey, time-drift corrections are obtained for each
instrument.
Regional Tectonics
The region of Attica as presented in the 1:50,000 scale
geological maps of the Institute of Geology and Mineral Ex-
ploration of Greece (IGME) (1989), belongs to an autoch-
thonous metamorphic basement of Paleozoic-Mesozoic age,
consisting of marbles and schists (see also Jacobshagen,
1986). Remnants of a “tectonic cover” have been identified
in several localities. A major tectonic boundary, dominating
the area, separates the mountains of Parnis to the northwest
from those of Pendeli and Immitos to the southeast. These
two are characterized by metamorphic rocks, mainly marbles
and schists belonging to the Cyclades Massif, which dip to
the northwest, below the nonmetamorphic limestones of
northwestern Attica. Normal faults along the Parnis moun-
tain are considered to be the eastward extension of the Cor-
inthiakos Gulf fault system and are of roughly east–west
orientation. During historic times and the more recent period
of instrumental earthquake observations in Greece, consid-
erable earthquake activity has been observed along this zone,
whereas the southeastern part of Attica appears to be nearly
aseismic (Galanopoulos, 1960, 1961; Makropoulos et al.,
1989; Papazachos and Papazachou, 1997).
The Saronikos Gulf is a neotectonic basin, divided by a
very shallow north–south-oriented platform into a western
and an eastern part. The islands of Methana, Angistri, Ae-
gina, and Salamis (see Fig. 1) are situated on this platform.
The western part of the Gulf consists of two basins. The
Epidaurus Basin is of west-northwest–east-southeast orien-
tation and more than 400 m deep, whereas the Megara Basin
is of east–west orientation with relatively shallow depth,less
than 250 m. The marginal extensional faults of the Epidaurus
Basin strike parallel to the bathymetry and have a 350-m
throw, forming a symmetric graben. The Megara Basin is
formed by east–west to east-northeast–west-southwest mar-
ginal faults. The Eastern Saronikos gulf is dominated by
northwest–southeast faults of relatively small throws and al-
ternating shallow basins and plateaus (Papanikolaou et al.,
1988).
Several volcanoes and volcanic outcrops of Plio-
Quaternary age are also situated in the Gulf and are part of
the western active Aegean volcanic arc. The westernmost
group of volcanic outcrops are those of the Loutraki-Sousaki
area. They are locally more than 70 m thick (Freyberg,
1973). The peninsula of Methana consists mainly of a large
volcanic dome and volcanic rocks of Pliocene and Quater-
nary age which are widely exposed (Gaitanakis and Dietrich,
1995). The most recent eruption is dated around 250 B.C.
and was desrcibed by Strabo (Stothers and Rampino, 1983).
To the northwest they appear to truncate all but the upper-
most 100 m of Quaternary sediments (Papanikolaou et al.,
1988). Volcanic rocks are also found to the south of the
island of Poros (Fytikas et al., 1986). Clearly, volcanic ac-
tivity in the Gulf is extensive, recent, and very close to Ath-
ens. It is poorly understood and its hazard has never been
assessed.
From displacements at faults and the distribution of sed-
iments it is shown that the Saronikos Gulf is more active in
the west than in the east. This has also been verified by the
microseismic activity, as will be presented subsequently.
The most seismically active area in the vicinity of Ath-
ens is the Corinthiakos Gulf, which exhibits deformation
typical of the extension that dominates central Greece and
the Peloponnese (Jackson et al., 1982; Collier et al., 1992;
Armijo et al., 1996), and an annual extension rate on the
order of 1 to 7 mm. About 20% of this extension is linked
to east–west-striking normal faults of eastern Corinthia (Pa-
pazachos and Kiratzi, 1996). The continuation of this fault
system to the east was also mapped by high-resolution,
single-channel reflection seismic profiles (Papanikolaou
et al., 1988) offshore in the Saronikos Gulf. A significant
part of the observed seismicity coincides with this system of
faults, as will be discussed subsequently.
Microseismicity
Within 45 days of field observations we recorded 739
microearthquakes above a threshold magnitude of M
L
0.3
(Fig. 2), which were observed by at least six stations.
Magnitudes were defined by the coda length of the seis-
mograms, calibrated by using earthquakes recorded also by
the Seismograph Network of the NOA. For the magnitude
estimates, we derived a mean function from four stations
having similar site response and applied this relationship to
the full temporary network (Fig. 3).
The locations of the hypocenters were obtained using
HYPOINVERSE software (Klein, 1989). HYPOINVERSE
allows the application of spatially varying local velocity
models for the hypocenter location. The velocity model used

922 J. Makris, J. Papoulia, and G. Drakatos
Figure 1. Location of the on and offshore seismic network used to record mi-
croearthquake activity in the area of Saronikos gulf, Greece. Stations are represented
by triangles.
was determined by a controlled source seismic experiment
between Kythnos and Eratini, crossing the Saronikos Gulf
from southeast to northwest (Makris et al., 2002). However,
to avoid local geological complexities, we adopted a mean
velocity model, as presented in Table 1.
Within the network most travel-time residuals after lo-
cation have an rms 0.2 sec. This corresponds to an epicentral
accuracy of the order of Ⳳ2 km, depending on the local
velocity model used, whereas the hypocentral depths are af-
fected by an error twice this value. Outside the network,
errors are significantly higher (Fig. 4).
The frequency-magnitude distribution of the recorded
events is shown in Figure 5. Within the network the thresh-
old magnitude of completeness is M
L
1.7. Outside the net-
work this increases to M
L
2.3.
To demonstrate the advantage of using local networks
over regional ones in assessing the seismic activity in this
area, we present in Figure 6 a comparison between historic
events (Fig. 6a), instrumentally recorded events published
by the NOA regional network (Fig. 6b), and events recorded
by this temporary local network (Fig. 6c). The advantage of
using a local network of densely spaced stations is obvious.
Even a short period of observations is enough to resolve
active structures that a regional network would require tens
of years to identify due to the different threshold of com-
pleteness.
Within the network and its immediate vicinity we lo-
cated 306 events (Fig. 7). Most of them are within the upper
15 km of the crust, whereas only a few were found at depths
exceeding 30 km. The distribution of the foci is illustrated
in Figure 8. The southwest–northeast oriented profile (Fig.
8) extends from eastern Peloponnese (point A) crossing the
Saronikos Gulf to central Evia (point B). Hypocenters con-
tained in a 100-km-wide zone, parallel to line AB, were pro-
jected in the surface defined by AB and the zaxis extending
to 40 km depth. Their distribution shows some distinctzones
of foci concentration that can be associated with known
faults. At 90 km (see Fig. 8) the hypocenters are located at
the “Aegina fault” as presented in Figures 2 and 7. The linear
distribution of the foci along the fault extends to a depth of
more than 15 km and marks a deep fracture of the thinned
Saronikos crust. The next concentration, 20 km further

Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece 923
Figure 2. Microseismic activity recorded by the Saronikos seismic network in a 45-
day period. The map shows the epicentral positions of 739 events, each having been
identified by a minimum of six stations. Depth of the hypocenters is indicated by the
color scale.
northeast coincides with the faults east of the island of
Salamis, which were reactivated by the “Athens Event” of
7 September 1999. They do not seem to have a deep exten-
sion and probably originate at the thrust front between the
Cyclades Massif and west Attica. In the Parnis area, 30 km
further northeast, we encounter the Fili fault zone. The seis-
micity shows that all faults around the Parnis epicenter of
7 September 1999 are still active and that the activityextends
through the complete crust.
At 45-km distance to the northeast we cross an active
seismic zone that is associated with the Chalkis area (see
Fig. 2). The Psachna region nearby is presently very active

924 J. Makris, J. Papoulia, and G. Drakatos
Figure 3. Calibration functions used to define lo-
cal magnitudes by the coda-duration method. The av-
eraged values were derived from functions defined at
four different location of comparable recording con-
ditions and therefore influenced by similar site effects.
Table 1
Local Velocity Model Used in the Location of Events
(Makris et al., 2002)
Velocity Vp (km/sec) Depth (km)
4.7 0.0
5.7 3.0
6.8 7.0
8.1 17.0
and has hosted several M
s
4.5 events. Another 35 km to the
northeast we encounter the Kymi fault zone with intense
crustal seismicity. This system of faults separates the north-
ern part of the island of Evia from the Skyros basin and
truncates the entire crust. Large-magnitude events have been
reported from this zone (Makropoulos and Burton, 1984;
Papazachos and Papazachou, 1997). Finally, around the
Skyros area, 65 km to the northeast off the coast, we en-
counter a seismically active region with a diffuse distribution
of foci. This area is notorious for its intense seismicity and
has hosted several large earthquakes (see, e.g., Makropoulos,
1978; Makropoulos et al., 1989). The large distance of the
Saronikos network from the island of Skyros and the limited
period of monitoring does not permit us to associate the
recorded microseismicity with the existing faults. A special
study with a local network is needed because the position
and size of the active faults around Skyros are not yet clearly
mapped.
In general, below Peloponnese and particularly the Sa-
ronikos gulf, the foci are located much deeper than below
Attica, Evia, and the Aegean domain (see Fig. 8). This phe-
nomenon points to the existence of two regional stress fields
that dominate the deformation of the Hellenides and the Ae-
gean region. Deeper seismicity, at subcrustal levels, as ex-
pressed below the Saronikos Gulf, is driven by the subduc-
tion of the lonian oceanic plate below Greece. Shallow
seismicity, on the other hand, concentrated at crustal levels,
is associated with extensional stresses due to deformation of
the upper crust. The oceanic lithosphere is subducted to the
northeast, whereas the continental crust and thinned litho-
sphere are uplifted and stretched westward. The crustal de-
formation is very complex involving rotation of blocks,
normal faulting, and crustal fragmentation. Significant west-
ward motion of the Hellenic napes creates a complex topog-
raphy that complicates even more the tectonic phenomena,
defining a third level of deformation at shallow depth, as
recognized by GPS observations (Kahle et al., 1995). This
complex deformation has been extensively discussed in the
geological and geophysical literature (see, e.g., Le Pichon
and Angelier, 1979; McKenzie, 1978; Rotstein, 1985) and
is also suggested on Figure 8.
Correlation of Seismicity with Tectonic Elements
Within the Saronikos Network
To correlate the seismicity with the tectonic elements
we used two sources of tectonic information: one is the Seis-
motectonic Map of Greece, published by IGME (1989) for
the onshore areas, and the other is the submarine neotectonic
map of Saronikos (Papanikolaou et al., 1988). Although ex-
tensive literature exists on the regional deformation of the
Hellenides (e.g., Papazachos and Delibasis, 1969; Papaza-
chos and Comninakis, 1971; Taymaz et al., 1991; Stiros,
1993) it is the two sources mentioned here that best describe
the Saronikos area.
The first observation worth mentioning is that the east–
west-oriented fault between the islands of Aegina and Sa-
lamis is also marked by significant seismic activity of shal-
low events. The seismicity can be followed onshore into east
Corinthia, delineating an active fault of more than 30 km
length. Such a fault can produce events on the order of mag-
nitude M
s
6.5 (see also Kiratzi et al., 1985). The fact that
historic catalogs (e.g., Papazachos and Papazachou, 1997;
Ambraseys and Jackson, 1990) do not report events from
this offshore area (see Fig. 3) is due to bias of observations
based on land-locked locations only.
Significant activity was also identified at the eastern
coast of Aegina, along a northeast–southwest-trending lin-
eament, mapped by Papanikolaou et al. (1988). At the vol-
canic areas of Methana and Aegina, on the other hand, we
did not map shallow seismicity worth mentioning. Although
many foci locations below the Saronikos Gulf are subcrustal
and associated with the subducting oceanic slab below the
Hellenides (see also Leydecker, 1975; Makropoulos, 1978;
Makris and Roever, 1986), we could not associate the seis-
mic activity with magma movements or hydrothermal pro-
cesses in the volcanic zone. This is probably due to the in-
adequate density of the local network, which did not permit
the identification of events of so small magnitude and vol-
canic noise.

Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece 925
RMS Distribution Within the Network
0
20
40
60
80
100
120
140
160
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0
RMS (sec)
Number of Events
RMS Distribution Outside the
Network
0
10
20
30
40
50
60
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0
RMS (sec)
Number of Events
ERH Distribution Within the Network
0
20
40
60
80
100
120
140
0.0 1.0 2.0 3.0 4.0 5.0 6.0 7.0 8.0
ERH (Km)
Number of Events
ERH Distribution Outside the Network
0
20
40
60
80
100
120
0.0 1.0 2.0 3.0 4.0 5.0 6.0 7.0 8.0
ERH (Km)
Number of Events
ERZ Distribution Within the Network
0
20
40
60
80
100
120
140
0.0 1.0 2.0 3.0 4.0 5.0 6.0 7.0 8.0
ERZ (Km)
Number of Events
ERZ Distribution Outside the Network
0
20
40
60
80
100
120
0.0 1.0 2.0 3.0 4.0 5.0 6.0 7.0 8.0
ERZ (Km)
Number of Events
Figure 4. rms distribution within and outside the Saronikos Gulf seismic network.
Of interest is the area of Mount Parnis where the 1999,
M
s
5.9 “Athens Event”occurred. We observed a great con-
centration of microearthquakes close to the 1999 epicentral
location of the mainshock, indicating that the aftershock ac-
tivity is still ongoing and the stresses responsible for this
event have not yet been completely relaxed. Also the area
of Spata, in southeastern Attica is seismically active along a
fault of northeast–southwest orientation, extending offshore
in the southern Evoikos Gulf (see Figs. 2, 6c, and 7).
Outside the network we recorded high seismicity in
three areas (see Fig. 2). One is the northeastern part of the
Corinth Basin, toward Thiva. This area is well known to
seismologists for its intense activity and, in particular, the
destructive earthquake sequence of 1981 that severely dam-
aged this region. The second area is the northeastern part of
central Evia (Fig. 2), north of Kymi as mentioned previ-
ously. This coastal part of the island and its immediate in-
terior are seismically very active. We could not define the
faults precisely, since the network is located at a distance of
80 to 100 km. The lateral velocity variations in the crust and
sediments are intense and cause significant systematic errors
in locating the hypocenters. Assuming a lateral homoge-

926 J. Makris, J. Papoulia, and G. Drakatos
Cumulative Magnitude Distribution
Within the Network
1
10
100
1000
0.0
0.4
0.8
1.2
1.6
2.0
2.4
2.8
3.2
3.6
Magnitude (ML)
logN(M)
Cumulative Magnitude Distribution
Outside the Network
1
10
100
1000
0.0
0.4
0.8
1.2
1.6
2.0
2.4
2.8
3.2
3.6
Magnitude (ML)
logN(M)
Limit Magnitude of
Com
p
leteness: ML 1.7
Limit Magnitude of
Com
p
leteness: ML 2.3
Figure 5. Frequency-magnitude distribution
within and outside the Saronikos Gulf seismic net-
work.
neous velocity model and reducing the three-dimensional
structure to one dimension only, location uncertainties in-
crease and errors may easily become systematic.
The third region we want to refer to is that to the east
and northeast of the island of Skyros (see Fig. 2). Here the
high seismicity recorded is associated with the aftershock
sequence of the M
s
5.8 “Skyros Earthquake”of 2001. A
precise definition of the associated active faults is also not
possible for the reasons stated previously.
Conclusions
The Saronikos local seismic network, despite its limited
operation of 45 days, produced important results permitting
several conclusions on the existence and location of active
faults previously unknown or inaccurately mapped. The
shallow seismicity within the network is, to a great extent,
associated with the deformation and opening of the Corin-
thiakos Gulf, which is one of the most rapidly deforming
areas of Greece. This is particularly true for the east–west
fault extending from Corinthia across the Saronikos Gulf,
between the islands of Salamis and Aegina. The deep, sub-
crustal seismicity, extending to a minimum of 100 km depth,
is obviously linked to the interaction of a subducted oceanic
Ionian Sea lithosphere with the mantle below Corinthia, the
Saronikos Gulf, and western Attica. The subducting litho-
sphere is limited to the north by the volcanic area of Li-
chades, in the northern Evoikos Gulf. Deep seismic events
reported by the Greek Seismological Services and our ob-
servations indicate that this is the northernmost limit of the
subducting slab. All the remaining seismicity is crustal. It is
controlled to the east, in the central and north Aegean Sea,
by transtension and normal faulting along the Corinthia Rift
due to rearrangement of crustal blocks escaping the com-
pression along the western Hellenides.
Use of a local velocity model obtained by a controlled-
source seismic experiment (Makris et al., 2002) significantly
increased the location accuracy. The seismicity defined by
observations of the seismological network of Greece delin-
eates active faults by the concentration of large-magnitude
events observed over long periods of time. Local networks
can significantly improve the picture by detecting and better
locating small-magnitude events over short periods of time.
This permits rapid assessment of the seismic hazard of an
area. The necessity for an accurate three-dimensional veloc-
ity model for the territory of Greece cannot be stressed
enough. It is mandatory if we wish to more accurately define
the active faulting and provide the engineering community
with the reliable seismic hazard parameters needed.
Acknowledgments
We are grateful to two anonymous reviewers for their useful com-
ments, which contributed to the improvement of our article. We acknowl-
edge the assistance provided by GeoPro-Hamburg for the field operations
and the instruments used in the network. We thank Captain C. Handras and
the crew of the R/V AEGAIO for their support during the marine opera-
tions. We specially thank Mr. D. Becker and Mr. S. Stoyanov from GeoPro-
Hamburg for the data processing and preparation of figures. Our field
operations were financed by the Earthquake Planning and Protection
Organization of Greece.
References
Ambraseys, N. N., and J. A. Jackson (1990). Seismicity and associated
strain of central Greece between 1890 and 1988, Geophys. J. Int. 101,
663–708.
Armijo, R., B. Meyer, G. King, A. Rigo, and D. Papanastasiou (1996).
Quaternary evolution of the Corinth rift and its implications for the
late Cainozoic evolution of the Aegean, Geophys. J. Int. 126, 1,
11–53.
Collier, R. E., M. R. Leeder, P. J. Rowe, and T. C. Atkinson (1992). U-
series dating and tectonic uplift of upper Pleistocene marine sediments
from the Corinth and Megara basins, Greece, Tectonics 11, 1159–
1167.
Freyberg, B. V. (1973). Geologie des Isthmus von Korinth, Erlanger Geol.
Abh. 95, 183 p.
Fytikas, M., F. Innocenti, N. Kolios, P. Manetti, and R. Mazzuoli (1986).

Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece 927
Figure 6. A comparison among historical
(Papadopoulos et al., 2000: 450 B.C. to 1889,
light-gray circles; Makropoulos et al., 1989
and Papazachos and Papazachou, 1997: 1914–
1964, dark-gray circles and NOA bulletins
1965–1999: black circles), (a), instrumental
readings from NOA (b), and presently recorded
seismicity (c) in the Saronikos Gulf and adja-
cent areas.

928 J. Makris, J. Papoulia, and G. Drakatos
Figure 7. Location of events recorded within the Saronikos network. A minimum
of six records per event has been used to locate each: earthquake. Events with an rms
value exceeding 0.5 sec have been discarded. Active faults identified by our observa-
tions are marked in red. A discussion of the map is presented in the text.
The Plio-Quaternary volcanism of Saronikos area (western part of the
Aegean volcanic arc), Ann. Geol. Pays Hell. 33, 23–45.
Gaitanakis, P., and A. D. Dietrich (1995). Geological map of Methana
peninsula, scale 1:25 000, Swiss Federal Institute of Technology,
Zurich.
Galanopoulos, A. G. (1960). A catalogue of shocks with Io ⱖVI or Mⱖ
5 for the years 1801–1958, Seismological Laboratory, Athens Uni-
versity, Athens, Greece, 1–119.
Galanopoulos, A. G. (1961). A catalogue of shocks with Io ⱖVI for the
years prior to 1800, Seismological Laboratory, Athens University,
Athens, Greece, 1–19.
Institute of Geology and Mineral Exploration (1989). Seismotectonic map
of Greece, scale 1:500,000.
Jackson, J., J. Gagnepain, G. Houseman, G. King, P. Papadimitriou, C.
Soufleris, and J. Virieux (1982). Seismicity, normal faulting, and the
geomorphological development of the Gulf of Corinth (Greece): the
Corinth earthquakes of February and March 1981, Earth Planet. Sci.
Lett. 57, 377–397.
Jacobshagen, V. (1986). Geologie von Griechenland, Beitr. Regionalen
Geologie der Erde, 363 p.
Kahle, H. G., M. V. Muller, A. Geiger, G. Danuser, S. Meuller, G. Veis,
H. Billiris, and D. Paradisis (1995). The strain field in NW Greece
and the Ionian islands: Results inferred from GPS measurements, Tec-
tonophysics 249, 41–52.
Kiratzi, A. A., G. F. Karakaisis, E. E. Papadimitriou, and B. C. Papazachos
(1985). Seismic source parameter relations for earthquakes in Greece,
Pageoph 123, 27–41.
Klein, F. (1989). User’s Guide to HYPOINVERSE, a program for VAX
computers to solve for earthquake locations and magnitudes, U.S.
Geol. Surv. Open-File Rep. 89-314.
Le Pichon, X., and J. Angelier (1979). The Hellenic arc and trench system:
a key to the neotectonic evolution of the eastern Mediterranean area,
Tectonophysics 60, 1–42.
Leydecker, G. (1975). Seismizitatsstudien im Bereich des Peloponnes auf-
grund von Prazisionsherdbestimmungen, Ph.D. Dissertation, Univer-
sity of Frankfurt, Frankfurt, Germany.
Makris, J., and L. Moeller (1990). An ocean bottom seismograph for gen-
eral use. Technical requirements and applications, in Proceedings of
Symposium “Europe and the Sea,” Hamburg, J. Hoefeld, A. Mitzlaff,
and S. Polomsky (Editors), sponsored by the EC.

Tectonic Deformation and Microseismicity of the Saronikos Gulf, Greece 929
Figure 8. Seismicity along a northeast-southwest-oriented profile with all foci con-
tained in a 100-km-wide and 40-km-deep block, projected on the surface shown in the
figure. For further information, see text.
Makris, J., and P. Roever (1986). Struktur und heutige Dynamik der Lith-
osphaere in der Agais, in Geologie von Griechenland, V. Jakobshagen
(Editor), Gebruder Borntraeger, Berlin, 241–256.
Makris, J., J. Papoulia, G. Drakatos, and V. Karastathis (2002). Seismo-
logical investigations in the Saronikos gulf, Athens, XXVII General
Assembly of the European Seismological Society, Nice (Abstracts).
Makropoulos, K. C. (1978). The statistics of large earthquake magnitude
and an evaluation of Greek seismicity, Ph.D. Thesis, University of
Edinburg, 1–193.
Makropoulos, K. C., and P. W. Burton (1984). Greek tectonics and seis-
micity, Tectonophysics 106, 275–304.
Makropoulos, K. C., D. Drakopoulos, and J. Latousakis (1989). A revised
and extended earthquake catalogue for Greece since 1900, Geophys.
J. Int. 98, 391–394.
McKenzie, D. P. (1978). Active tectonics of the Alpine-Himalayan belt:
the Aegean Sea and surrounding regions, Geophys. J. R. Astr. Soc.
55, 217–254.
Papadopoulos, G. A., and A. Arvanitides (1996). Earthquake risk assess-
ment in Greece, in Earthquake Hazard and Risk, V. Schenk (Editor),
Kluwer, Dordrecht, The Netherlands, 221–229.
Papadopoulos, G. A., A. Vassilopoulou, and A. Plessa (2000). A Catalogue
of Historical Earthquakes for Central Greece: 480 B.C.–A.D. 1910,
Institute of Geodynamics, National Observatory of Athens, Publ.
no. 11.
Papanikolaou, D., V. Lykousis, G. Chronis, and P. Pavlakis (1988). A com-
parative study of neotectonic basins across the Hellenic arc: the Mes-
siniakos, Argolikos, Saronikos and Southern Evoikos Gulfs, Basin
Res. 1, 167–176.
Papazachos, B. C., and P. Comninakis (1971). Geophysical and tectonic
features of the Aegean arc, J. Geophys. Res. 76, 8517–8533.
Papazachos, B. C., and N. D. Delibasis (1969). Tectonic stress field and
seismic faulting in the area of Greece, Tectonophysics 7, 231–255.
Papazachos, B. C., and A. A. Kiratzi (1996). A detailed study of the active
crustal deformation in the Aegean and surrounding area, Tectono-
physics 253, 129–154.
Papazachos, B. C., and C. Papazachou (1997). The Earthquakes in Greece,
Ziti Ed., Thessaloniki, 304 pp.
Rotstein, Y. (1985). Tectonics of the Aegean block: rotation, side arc col-
lision on crustal extension, Tectonophysics 117, 117–137.
Sieberg, A. (1932). Erdbebengeographie, Verlag von Gustav Fischer, Ber-
lin, 102 pp.
Stiros, S. (1993). Kinematics and deformation of central and southwestern
Greece from historical triangulation data and implications for the ac-
tive tectonics of Aegean, Tectonophysics 220, 283–300.
Stothers, R. B., and M. R. Rampino (1983). Volcanic eruptions in the Med-
iterranean before A.D. 630 from written and archeological sources,
J. Geophys. Res. 88, 6357–6371.
Taymaz, T., J. Jackson, and D. McKenzie (1991). Active tectonics of the
north and central Aegean Sea, Geophys. J. Int. 106, 433–490.
GeoPro GmbH
20457 St. Annenufer 2
Hamburg, Germany
(J.M.)
Institute of Oceanography
Hellenic Centre for Marine Research
19013 Anavyssos
Attica, Greece
(J.P.)
Geodynamic Institute
National Observatory of Athens
11810 Thission
Athens, Greece
(G.D.)
Manuscript received 9 October 2002.