Geophys. J. Int. (2005) 162, 257–268
doi: 10.1111/j.1365-246X.2005.02640.x
Crustal velocity and Moho structure beneath the Gulf of Corinth,
Greece
Barry C. Zelt,1 Brian Taylor,1 Maria Sachpazi2 and Alfred Hirn3
1 School
of Ocean & Earth Science & Technology, Department of Geology and Geophysics, University of Hawaii, 1680 East-West Road,
Honolulu, HI 96822, USA. E-mail: [email protected]
2 Geodynamic Institute, National Observatory of Athens, Athens, Greece
3 Laboratoire de Sismologie Expérimentale, Département de Sismologie, Institut de Physique du Globe de Paris, Paris, France
SUMMARY
The Gulf of Corinth (GOC), Greece is a continental rift with high rates of seismicity and extensional strain. How this strain is accommodated in the crust and whether there are variations
in the mechanism along strike remain open questions, in part because of a lack of wide-angle
reflection/refraction studies that constrain crustal velocity structure. In 2001, an extensive multichannel seismic survey was conducted within the GOC, one component of which included
the wide-angle recording of sources from within the gulf at stations on land surrounding the
gulf. In this paper we use wide-angle data in two separate, but allied, studies to constrain crustal
velocities and depth to the Moho. A 2-D inversion of refraction and reflection traveltimes along
an axial profile through the GOC constrains the shallow crustal velocity structure, images the
Moho at 29 km depth in the east, dipping to 39 km in the west, and images the eastward
subducting African slab beneath the western GOC at a depth of 74 km. The 1-D average of the
2-D velocity model was used in a tomographic inversion of PmP reflection times to solve for
depth to the Moho throughout the Corinth region. This model shows generally thick, isostatically compensated crust (≥37 km) beneath the Hellenide Mountains, except immediately south
of the GOC, and a singular region of thin crust (<30 km) beneath the Perahora Peninsula at the
eastern end of the gulf. A comparison with Moho depths derived from gravity inversion shows
a general agreement with crust thickening from east to west, but a number of differences in
detail. The 3-D crustal thickness variations are more complex than those predicted by either
pure shear or simple shear models of continental extension and suggest significant pre-rift
structural variability.
Key words: crustal structure, Greece, Gulf of Corinth, Moho, tomography, velocity.
I N T RO D U C T I O N
The Gulf of Corinth (GOC; Fig. 1), which separates the Peloponnisos from mainland Greece, is an active continental rift on the
western edge of the Aegean, which itself is one of the most active
extensional continental regions in the world (Armijo et al. 1996).
Extension throughout the Aegean started in the late Oligocene
(Jolivet 2001); however, the Corinth rift is primarily a Quaternary
feature (Armijo et al. 1996). Earthquake focal mechanisms through
the GOC region show a pattern of east–west-trending normal faulting, consistent with geodetic measurements which show that extension is directed north–south across, and is mainly localized beneath,
the gulf (Billiris et al. 1991; Clarke et al. 1997, 1998; Davies et al.
1997; Briole et al. 2000). There is, however, considerable debate
concerning the geometry of crustal faulting that is accommodating the deformation, especially in the light of focal mechanisms
(Hatzfeld et al. 1996) and microseismicity studies (Rigo et al. 1996;
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2005 RAS
Rietbrock et al. 1996) which suggest that, in the western and central parts of the gulf, slip is occurring on low-angle (<30◦ ) normal
faults.
The crustal velocity structure, including variations in crustal
thickness, are important parameters that can help resolve issues
relating to the mode of extension. Such issues include whether the
crustal thinning occurs offset from, or beneath, surficial rifts (i.e.
simple versus pure shear) and the relationship between upper crust
faulting and lower crust deformation (e.g. Wernicke 1985; Ruppel
1995). To date, the only detailed map of Moho topography beneath
the Corinth rift comes from the inversion of gravity data (Tiberi et al.
2001). Furthermore, there are no published reports of crustal-scale
velocity variations in the immediate GOC region. Here, in addition
to using wide-angle reflection/refraction data to constrain a 2-D
crustal velocity model for an along-axis profile through the gulf,
we present a map of Moho topography derived from tomographic
inversion of wide-angle reflection traveltimes.
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GJI Tectonics and geodynamics
Accepted 2005 March 14. Received 2004 December 19; in original form 2004 May 5
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B. C. Zelt et al.
12
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Elevation (km)
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Figure 1. Topographic map of central Greece centred on the Gulf of Corinth (GOC). Dotted lines within the GOC denote locations of the ship track and shot
points. The thick east–west line within the gulf represents combined shot lines 8 and 23 used in 2-D modelling. White and black dots are land stations. Black
dots are stations used in 2-D modelling. Both black and white dots (except stations 36 and 48) were used in 3-D modelling. Large numbers are station numbers
referred to in the text. Small numbers are the number of PmP data used in the 3-D modelling. + symbols between stations 38 and 49 show approximate bounce
points for Ps phase (reflections from subducting African Plate). Triangles are centres of volcanic activity: s, Susaki solfataric field, m, Methana arc volcano. A,
Athens; C, Cephalonia; GOP, Gulf of Patras; NGOE, North Gulf of Evia; P, Patras; PP, Perahora Peninsula; SG, Saronic Gulf. The inset map shows region of
the larger map (rectangle) relative to the Greece/Aegean region.
SEISMIC EXPERIMENT
In 2001 July the University of Hawaii, National Observatory of
Athens, and Institut de Physique du Globe de Paris conducted a
seismic survey of the Gulf of Corinth, Greece. A dense grid of
about 50 multichannel seismic (MCS) lines was recorded within the
gulf (Fig. 1) using the R/V Maurice Ewing (Sachpazi et al. 2002;
Goodliffe et al. 2003; Taylor et al. 2003; Weiss et al. 2003; Pi Alperin
et al. 2004; Zelt et al. 2004). The seismic source was a 138 l, 20-gun,
tuned air gun array. Shot spacing was 50 m with an average time
between shots of 21 s. The water depth beneath the shots ranged
from 40 to 870 m. Bathymetry along each line was measured using
Hydrosweep DS2 multibeam sonar.
In addition to the MCS component of the experiment, a network of
40 temporary stations was deployed across southern Greece (Fig. 1)
specifically to record the marine shots. These stations were deployed
and operated under a contract to Landtech Enterprises. All of the
stations that provided usable data comprised an Earth Data PR2400
portable field recorder and a GeoSIG VE-53DH three-component
seismometer. The maximum shot–receiver offset was 180 km. In
addition, we also use data recorded at two stations of the Corinth
Rift Laboratory Project (Lyon-Caen et al. 2001). In general we found
that the seismic signals were strongly attenuated, suggesting that,
ideally, for an experiment of this scale in this region a larger seismic
source would be very beneficial.
PREVIOUS VELOCITY MODELS
A N D E S T I M AT E S O F M O H O D E P T H
There have been few published reports of crustal velocity and Moho
structure in the vicinity of the GOC. Clément (2000) and Clément
et al. (2004) used wide-angle reflections recorded at two stations
from a shot profile within the gulf to derive Moho depth estimates
of 40 km beneath the western end of the gulf (near our station 49) and
32 km along the northern shore at the eastern end of the gulf. About
100 km south of the GOC, Makris (1978) interpreted refraction data
from a profile crossing the Peloponnisos. He found that the velocity
at the top of the crystalline basement was 6 km s−1 , the velocity
at the base of the crust was 6.5–7.0 km s−1 , the Moho increases in
depth from 26 km in the east to 46 km in the west before decreasing
rapidly near the western shore of the Peloponnisos and the upper
mantle velocity was ∼7.7 km s−1 . He also noted the presence of
late arrivals interpreted as reflections from a subcrustal reflector at
65 km depth beneath the western Peloponnisos. At about the same
latitude, Clément et al. (2000) looked at data recorded along an
onshore–offshore profile and noticed a reflective layer, interpreted
as the lower crust, which terminated at the upper mantle at an inferred
depth of ∼20 km. At the latitude of the GOC, but further west, Hirn
et al. (1996) studied the crustal structure of the Ionian Islands region
using marine MCS reflection profiles combined with recordings at
land stations. Their MCS data imaged a 12 km deep reflector from
outboard of Cephalonia to the island’s southeastern tip, where it dips
strongly eastwards. This was interpreted as the top of the subducting
African slab which, at this location, is still overlain by upper Aegean
Plate crust. From wide-angle reflection data recorded at stations of
both sides of the Gulf of Patras, they inferred a depth to Moho of
24 km under the western Gulf of Patras at 21.25◦ , SSE of our station
36.
There have been a number of earthquake tomography studies that
include the Corinth region (Papazachos et al. 1995; Drakotas et al.
1997; Papazachos & Nolet 1997); however, cell sizes are typically
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2005 RAS, GJI, 162, 257–268
Gulf of Corinth Moho structure
∼50–100 km horizontally and 10 km vertically, and thus both the
horizontal and vertical resolution of these studies is poor. A tomographic image of the subducting African slab (Papazachos & Nolet
1997) shows a depth of 60–65 km in the western Peloponnisos, with
a shallow eastward dip of ∼ 10◦ , increasing rapidly to 35◦ in the
central Peloponnisos.
Tiberi et al. (2001) inverted gravity data to create a map of Moho
depth in the vicinity of the GOC. They found northwest–southeasttrending features superimposed on a general crustal thickening from
∼25 km in the east to ∼40 km in the west. They interpret an offset
between relatively thin crust near the northern margin of the gulf
with the main axis of the gulf as support for the hypothesis that
extension has been accommodated by a low-angle normal fault at
12 km depth, as suggested by Rietbrock et al. (1996) and Rigo et al.
(1996), and a ductile lower crust.
2 - D T R AV E L T I M E I N V E R S I O N A L O N G
A N E A S T – W E S T P RO F I L E
For the 2-D study, we use shots from two east–west MCS profiles,
8 and 23 which, combined, form the longest (85 km) continuous
profile through the Gulf (Fig. 1). We use data recorded at eight stations: 17–20 to the east of the gulf, and 36, 38, 48 and 49 to the
west of our shots. The latter two stations are Corinth Rift Laboratory stations TEM and TRI. This east–west transect of the gulf is
vastly superior to other potential 2-D profiles that could be created
from our shot/receiver distribution: it is the longest possible transect
(250 km), the overall data quality is the best and it has the most inline shots. For these reasons, we concentrate on this single profile.
Our ability to constrain 2-D structure is limited by two factors:
(1) the acquisition geometry, which is less than ideal because there
is no overlap of shots and receivers and (2) the quality of data, which
is limited by the size of the seismic source and apparent strong attenuation in this region. We confidently identify three seismic phases:
refractions through the upper crust (Pg), reflections from the Moho
(PmP) and a later phase (Ps) which, based on its arrival time, am-
259
plitude and moveout characteristics, corresponds to reflections from
a deep, sub-Moho reflector. The maximum offset observed for the
Pg phase varies from 43–72 km; Pg is not observed at station 36
and is only tentatively observed at station 17. Pn (refractions through
the uppermost mantle) is not observed. PmP is generally prominent except at station 48, where it is very weak, and at station 36,
where its identification is indefinite (Fig. 2). Ps is a prominent, late
phase recorded only at station 36 (Fig. 2). Pick uncertainties between
50 and 250 ms were calculated using an automated scheme based
on the signal-to-noise ratio in a short (250 ms) time window around
the pick (e.g. Zelt & Forsyth 1994). Picks were binned such that
the average spacing between picks was 250 m. Table 1 gives the
statistics of the data.
We model the 2-D data using the traveltime inversion method
of Zelt & Smith (1992) in which a layered velocity model is
parametrized by boundary (depth) nodes and velocity nodes. We use
a layer-stripping approach, modelling successively deeper structures
while fixing shallower structures (e.g. Zelt et al. 1993). Our exiguous data and less than optimal recording geometry lead us to seek
the simplest model, i.e. with the fewest parameters, that fits the data.
To this end, for example, except where necessary, we do not allow
lateral velocity variations.
Ray coverage and a comparison between observed and calculated
traveltimes are shown in Fig. 3. The preferred final model is shown
in Fig. 4. Estimates of velocity and reflector depth uncertainty are
given in Table 2. The model origin corresponds to station 36 at
21.11176◦ E, 38.48812◦ N. The model comprises seven layers. Layer
1 is water. Layer 2 comprises syn-rift sediments. Our data do not
constrain velocities in this layer; we assigned a constant velocity of
2.5 km s−1 , which is the average velocity for the sediments as found
in a shallow tomography study of the gulf (Zelt et al. 2004). The
base of this layer is taken from MCS interpretations of lines 8 and
23 (Weiss 2004). Its thickness varies from 0 to 2.3 km, and it causes
significant delays and adds structure to the observed phases. It is
essential that this layer is accurately represented to avoid errors in
deeper parts of the model.
Figure 2. Receiver gather for station 36 and shot line 8+23. Data are bandpass filtered between 3 and 8 Hz and reduced at 8 km s−1 . The horizontal axis is
shot–receiver offset. The strong, late phase between 10.6 and 13.0 s is Ps, interpreted as a reflection from the subducting African slab. The weaker phase around
7 s is probably PmP. The inset nomogram shows the location of the receiver (black dot) and shot line (black line) in relation to other stations and shots.
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2005 RAS, GJI, 162, 257–268
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Table 1. Traveltime data for 2-D modelling and the goodness-of-fit statistics for the final model. Table
columns: Station, station number (Fig. 1); X , model distance of receiver (distance from station 36); Pg,
PmP and Ps, number/average uncertainty (ms) of Pg, PmP and Ps picks, respectively; Total, total number
of picks/average uncertainty (ms) for each station; χ 2 , normalized chi-squared misfit for each station;
dt rms , rms traveltime residual for each station. Last three rows: Total, number of picks/average uncertainty
for each phase; χ 2 , normalized chi-squared misfit for each phase; dt rms , rms traveltime residual for each
phase.
Station
36
38
49
48
20
19
18
17
Total
χ2
dt rms (ms)
36
38
X (km)
Pg
PmP
Ps
Total
χ2
dt rms (ms)
0.0
44.8
85.0
92.4
189.0
215.6
228.3
252.7
—
20/148
65/228
87/197
196/142
72/156
68/190
9/200
48/224
90/217
63/205
22/234
93/201
162/208
106/215
106/206
145/218
—
—
—
—
—
—
—
193/219
110/205
128/217
109/205
289/161
234/192
174/206
115/206
0.7
1.0
1.0
1.5
1.0
0.6
0.6
0.8
142
175
178
214
135
101
123
154
517/172
1.1
158
690/211
0.7
139
145/218
0.7
146
1352/197
0.9
147
0.9
147
49
48
20
19
18
17
Moho
a
W
E
36
36
38
17
18
36
Ps
38
49
49
48
20
19
20
49
38
48
49
49
18
17
PmP
17
19
18
19
48
Pg
b
20
Figure 3. (a) Rays traced through the 2-D model. Every third ray is plotted. Three phases, coded by colour and identified in panel (b), were used in the
modelling: Pg, PmP and Ps. Numbered triangles are station locations. No vertical exaggeration. (b) Observed data (short, coloured vertical lines with height
equalling twice the pick uncertainty) and calculated traveltimes (black dots). Numbers indicate the stations at which the data were recorded.
Layers 3 and 4, which extend to a maximum depth of ∼6.5 km,
are constrained by Pg. Beneath the gulf, velocities in layer 3, which
represents the basement to the current rifting phase, are 5.3 and
5.9 km s−1 at the top and bottom, respectively. Layer 4 is included
to affect a decrease in velocity gradient. The velocity at the base
of this layer is 6.0 km s−1 beneath the gulf and further east. In the
east, beneath stations 18 and 17, we have reduced the velocities
in layer 3 to 5.0–5.5 km s−1 to affect a traveltime delay to data
from these stations. The source of the causative velocity reduction,
however, cannot be constrained; it could, for example, be equally
likely to arise from the presence of a thin, surficial low-velocity
region. Similarly, faster velocities are required in the west to satisfy
the early Pg arrival times at station 38. To this end we increase
the shallow crustal velocities west of station 49 to 6.1–6.2 km s−1 .
Whether these higher velocities are present west of station 38 we
cannot say with our data; they are not inconsistent with our data.
For simplicity we extend the high velocities to the west end of our
model.
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2005 RAS, GJI, 162, 257–268
Gulf of Corinth Moho structure
261
Figure 4. 2-D velocity model with no vertical exaggeration. Numbered diamonds are stations used in the modelling. Numbers within the model, and inside
boxes above the model, are P-wave velocities at the top or bottom of the indicated layer. These velocities represent averages over regions that can be distinguished
by similarity of colour. Velocities in layer 2 and within the upper mantle (white region) are not directly constrained by the modelling; these layers have no
vertical velocity gradient. Heavy black lines denote portions of boundaries that are sampled by reflections. Broken lines on the west side of the model represent
shallowing of the Moho to 24 km beneath station 36 and the approximately 30◦ dip of the African Plate required to reach a depth of 12 km beneath Cephalonia;
both based on Hirn et al. (1996). The thick grey line shows the position of the Moho (M) obtained from the 3-D inversion of PmP data.
Table 2. Estimated uncertainties for velocities and depths that were solved
for in the 2-D inversion. The value of velocity or depth at a velocity or
boundary node was perturbed through a range of values (both negative and
positive). The range which provides a statistically equivalent fit (using an
F-test) is an estimate of parameter uncertainty. Estimated uncertainties for
Moho and slab depth assume that the overlaying velocity structure is accurate.
Parameter/region of model
Velocity in layer 3
Velocity in layer 4 west of X = 85 km
Velocity in layer 4 east of X = 85 km
Depth to Moho near X = 40 km
Depth to Moho between X = 60 and 185 km
Depth to Moho near X = 220 km
Depth to slab between X = 50 and 75 km
Estimated
uncertainty
± 0.1 km s−1
± 0.15 km s−1
± 0.1 km s−1
± 1.1 km
± 0.7 km
± 2.0 km
± 1.0 km
The remainder of the crust is represented by layer 5. We have
no refractions through this layer to constrain velocity. Velocities at
the top of this layer were set to the same values as at the base of
layer 4 (6.0–6.2 km s−1 ). PmP times were inverted for depth to
Moho using various fixed and constant values for the velocity at
the base of the crust. Although the PmP data were consistent with
velocities between 6.5 and 7.0 km s−1 , we consider the optimal
value to be 6.7 km s−1 because, as detailed below, this value also
provided the best overall fit to our inversion of PmP for 3-D Moho
topography. It is possible to fit the PmP times at all stations without resorting to lateral velocity variations in layer 5; however, the
converse is not true. We could not fit the PmP data by assuming a
flat (or flattish) Moho and allowing reasonable lateral velocity variations in layer 5. PmP times from all stations except 36 constrain
a Moho which dips to the west from 28.5 km beneath station 19 to
38 km beneath station 49. Using a faster or slower velocity at the
base of the crust does not affect this trend: depths would vary by
about ±0.5 km and ±1.0 km beneath stations 49 and 19, respectively, for variations in the average velocity at the base of layer 5 of
±0.2 km s−1 . Our identification of PmP recorded at station 36 requires that the Moho shallows significantly in the west beneath station 38 to 32 km.
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2005 RAS, GJI, 162, 257–268
The Ps phase recorded at station 36 constrains a deep, subcrustal
reflector at 74 km depth between stations 38 and 49. This depth
assumes that the average upper mantle velocity, which is unconstrained in our model, is 7.9 km s−1 . The reflector depth would be
71 or 76 km for upper mantle velocities of 7.7 and 8.1 km s−1 , respectively. As shown in Fig. 4, the reflector, where sampled by rays,
is flat; however, an eastward dip of no more than ∼5◦ is allowed by
the data.
3-D TOMOGRAPHIC INVERSION FOR
MOHO DEPTH
We confidently identify PmP arrivals at 28 stations. PmP, where
present, is generally not difficult to recognize because intracrustal
reflectivity is typically very low and the only coherent, relatively
strong phase is probably PmP. The prominent PmP and lack of
intracrustal reflectivity in the vicinity of the GOC have also been
noted by Clément et al. (2004). To reduce the size of the 3-D volume
we do not use PmP data from station 36 in the inversion. Including
data from this station has a predictable effect: the Moho would be
imaged at ∼32 km approximately beneath station 38.
PmP is observed between offsets of 50–140 km. Arrival times
between different stations vary quite significantly; e.g. for offsets in
the range of 90–100 km there is a difference of 3 s in PmP arrival
times across the network. Exemplary recordings of PmP at stations
on the west and east side of the array are shown in Fig. 5. Note the
large difference in arrival times for station 38 in the west (Fig. 5a;
8–8.6 s) compared with station 21 in the east (Fig. 5b; 5.5–6.2 s),
suggestive of significant Moho topography (and/or lateral velocity
variation).
To provide the best areal coverage of PmP reflection points, for
each station, where possible, PmP was picked on a number of shot
lines. On average, shots from five lines were picked for PmP reflections at each station. Picks were made on undecimated data, but the
data were binned prior to inversion to give an effective shot spacing
of 1.5 km. After initial inversion tests, data extremely inconsistent
with nearby PmP picks, possibly due to misidentification of the PmP
phase on noisy shot lines, were culled, leaving a total of 1534 PmP
arrival times with an average uncertainty of 198 ms. Fig. 1 shows
the number of PmP data for each station.
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B. C. Zelt et al.
Figure 5. (a) Receiver gather for station 38 and shot line 30. Data are bandpass filtered between 3 and 8 Hz and reduced at 8 km s−1 . The horizontal axis is the
azimuth measured clockwise from receiver to shot. Offsets range from 68 km in the north to 72 km in the south. Note the strong, clear Moho reflection (PmP)
at about 8 s. (b) Receiver gather for station 21 and shot line 8+23. Same plotting parameters as (a) except the horizontal axis is shot–receiver offset. Note that
PmP arrives much earlier (at about 6 s) compared with (a), suggesting a significantly thinner crust in the east. The inset nomogram in both panels shows the
location of the receiver (black dot) and shot line (black line) in relation to other stations and shots.
To solve for depth to Moho we use the tomography method of
Zelt & Barton (1998), extended to include reflections (Zelt et al.
1999). This method, as applied to the inversion of reflection data
to constrain reflector depth, penalizes the roughness and size of the
reflector with respect to a starting model. The amount of regularization is automatically chosen so that the appropriate data misfit is
achieved on the final iteration. Traveltimes are calculated by finite
difference (Hole & Zelt 1995) on a uniform grid with a node spacing of 1.5 km. The inversion is calculated on a grid with a lateral
node spacing of 3 km and vertical node spacing of 1.5 km. The
reflector is free to assume any depth at each grid node location.
The forward and inverse grid sizes were chosen to be small enough
to allow accurate traveltime calculations and to adequately repre-
sent lateral Moho depth and velocity variations while maximizing
computational efficiency.
In general, PmP traveltime residuals could be mapped into Moho
depth variations and/or crustal velocity variations. Using only PmP
data, it is not advisable to solve simultaneously for depth and velocity
because the data are not capable of determining how the residuals
should be partitioned. Thus, we start by using a fixed 1-D velocity
model, based on an average of the 2-D velocity model for the east–
west profile, and solve only for Moho depth. We stripped off the
water and sedimentary layers and moved the shots to the base of
the rift-onset sedimentary sequence determined by Weiss (2004).
Traveltime through the water and sediments was removed from the
observed data. This traveltime was estimated by ray tracing through
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2005 RAS, GJI, 162, 257–268
Gulf of Corinth Moho structure
263
Table 3. 1-D velocity model
used in 3-D inversion for
Moho depth. Velocities are
linearly interpolated between
each specified point. 0 km
depth corresponds to sea level.
Depth
(km)
Velocity
(km s−1 )
−3.0
3.0
4.5
6.0
33.0
51.0
5.3
5.9
6.0
6.1
6.7
6.7
a simple water–sediment–crust model using the true shot–receiver
offset to give the proper take-off angle from the shot and the correct
ray path from the shot to the base of the sediments. The preferred 1-D
velocity model (Table 3) increases from 5.3 km s−1 at the top of the
model (3 km above sea level) to 6.7 km s−1 at 33 km depth, with an
average velocity to 33 km depth of 6.25 km s−1 . This model produced
the best fit to the PmP data. 1-D models with average velocities
within ±0.1 km s−1 of the preferred model resulted in Moho models
which statistically fit the data equally well and which look similar,
except depths are shifted by about ±1.5 km, respectively.
The final Moho model, shown in Fig. 6, has a normalized χ 2
misfit of 1.7 and a root mean square (rms) traveltime residual of
dt rms = 250 ms. The crust is thickest in the west beneath the Hellenides, with a Moho depth up to 40 km, and relatively thin in the
east, with a zone of very thin crust (25 km) beneath the Perahora
Peninsula. A comparison of Moho depths along the 2-D profile is
shown in Fig. 4 (grey line): depths are similar in the west, but the
Moho of the 3-D model is 1–2 km deeper east of X = 120 km. The
differences are not significant given the combined uncertainties of
the two methods. Because a fixed, 1-D velocity model was used in
the inversion, the topographic variations in this model are probably
maximal, assuming that the 1-D model is an accurate representation of the average velocity structure. Allowing traveltime residuals
to map into both velocity and depth would produce a model with
less topography. With a fixed 1-D velocity model we find the data
are underfitted (χ 2 > 1), which suggests that the crustal velocity
structure is not laterally homogeneous.
Model assessment
To test if, with ‘reasonable’ lateral velocity variations, this Moho
model could fit the data to within the assigned uncertainties, we
tried inverting the PmP data for crustal velocity variations while
keeping the Moho of Fig. 6 fixed. The forward and inverse grid sizes
were the same as described in the previous section; these sizes are
small enough to adequately represent lateral velocity variations that
our coarsely spaced data are capable of resolving. The data cannot
discriminate the depth range over which lateral velocity variations
occur. Thus, we allowed velocities throughout the entire crust to vary
by up to ±1 km s−1 relative to the 1-D model in Table 3. The resultant
model (Fig. 7) produced a misfit of χ 2 = 1 (dt rms = 189 ms) with
reasonable velocity variations. Shallow velocity variations near the
stations (Fig. 7a) effectively represent station corrections. Relative
to the 1-D starting model, velocities are typically faster down to
about 15 km, then slower at greater depths. We stress, however, that
the velocity model is non-unique and thus caution against attaching
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Figure 6. Map of depth to the Moho from inversion of PmP traveltime
data using a fixed 1-D velocity model (Table 3). The contour interval is 1
km. Thick contour lines are drawn at depths of 30 and 40 km. Triangles
are stations used in the inversion. Black dots are the locations of reflection
points for the five non-linear iterations of the tomographic inversion. Only
regions of the Moho within 5 km of a reflection point are unmasked. The
misfit for this model is χ 2 = 1.7. The small region labelled ‘32 km’ on the
left-hand side of the model indicates depth to the Moho beneath station 38
based on 2-D modelling.
too much significance to the velocity variations. Most importantly,
this model demonstrates that the Moho reflector in Fig. 6 together
with this reasonable, smooth velocity model, fits the data to within
the estimated uncertainties.
The shallow Moho (25 km) beneath the Perahora Peninsula on
the isthmus separating the Corinth and Saronic gulfs is, topographically, the most significant feature of the Moho model. To investigate the necessity for this feature we performed a test in which the
Moho depth model of Fig. 6 was clipped to a minimum depth of
30 km. This interface was held fixed and PmP times were inverted
to solve for velocity structure. The resultant model provided a χ 2
misfit of 1, suggesting that the very shallow depths can be traded off
against velocity. A region of significantly faster velocities centred
around stations 21 and 22 is required relative to the model shown in
Fig. 7. Velocities between depths of 3 and 15 km, however, are 6.7–
6.9 km s−1 , perhaps somewhat faster than might reasonably be expected, suggesting that while some trade-off between depth and velocity is permitted, a shallow Moho depth (<30 km) at this location
is required by the data.
A comparison of our Moho model with the model of Tiberi et al.
(2001), derived by inversion of gravity residuals, reveals major differences. For example, in the region where our tomography results
constrain the Moho, the model of Tiberi et al. (2001) is, on average, 5.5 km shallower. Using the 1-D velocity model of Table 3, the
Tiberi et al. (2001) model produces a misfit of χ 2 = 41 (dt rms =
1180 ms). This is significantly worse than our model which provides a misfit of χ 2 = 1.7 (dt rms = 250 ms) using the same velocity
model. To test if, with ‘reasonable’ lateral velocity variations, the
model of Tiberi et al. (2001) could fit the data to within the assigned
uncertainties, we tried inverting the PmP data for crustal velocity variations while keeping the Tiberi et al. (2001) Moho fixed.
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Figure 7. Velocity model derived by inverting PmP traveltimes while keeping the Moho model of Fig. 6 fixed. Horizontal slices at six depths show the
difference between final model and the 1-D model of Table 3. The contour interval is 0.05 km s−1 . Thick contours are ±0.2 km s−1 . The velocity value indicated
on each slice is the velocity of the starting 1-D model (Table 3) at that depth. Note that the velocity perturbations are generally smooth, and the resultant
velocities are reasonable at all depths. This velocity model, combined with the Moho depth model of Fig. 6, produces a misfit of χ 2 = 1.0.
The final velocity model, which gives a χ 2 misfit of 1.6 (dt rms =
234 ms), is quite unrealistic, with large regions of both unreasonably
slow (<5 km s−1 at 7.5 km depth) and fast (7.3 km s−1 at 18 km
depth) velocities in the centre of the gulf. We reran this test after
increasing the average depth of the Tiberi et al. (2001) model by
5.5 km (giving a mean depth of 34.5 km—the same as the preferred
Moho model in Fig. 6) at all points to see if a static shift would
make their model consistent with the PmP data. The resulting velocity model gave a misfit of χ 2 = 1.3 (dt rms = 220 ms), significantly
worse than our preferred Moho model (χ 2 = 1, dt rms = 189 ms;
Fig. 6) together with reasonable crustal velocity variations (Fig. 7).
Furthermore the resultant velocity model has unreasonable velocity
values, e.g. >7 km s−1 beginning at 12 km depth beneath the GOC.
D I S C U S S I O N O F R E S U LT S
Our ability to constrain velocity structure is limited by the data, from
which we have identified only three phases. Because of this, we have
looked for the simplest model, i.e. with the least amount of structure,
which fits the data. We were able to fit the data for the 2-D profile with
a very simple model (Fig. 4). The only significant velocity variations
are in the shallow crust (<6 km depth); the faster velocities in the
west and slower velocities in the east are required primarily to delay
or advance Pg arrival times relative to stations close to the gulf. The
PmP data could be fitted well without requiring significant lateral
velocity variation. The Moho dips to the west from about 28.5 km
beneath station 19 to 38 km beneath station 49. Further west, beneath
station 38, PmP data from station 36 images the Moho at 32 km. The
location of this very rapid change in crustal thickness corresponds to
the western edge of the NNW-trending Hellenides mountain chain,
and suggests an abrupt termination of a relatively thick crustal root.
Another onshore–offshore seismic experiment (Hirn et al. 1996)
found Moho depth to be 24 km just south of our station 36 in the
Gulf of Patras, as shown in Fig. 4. Makris (1978) documented an
even more abrupt and more substantial thinning of the crust further
south on the Peloponnisos at approximately the same location relative to the Hellenides. The Moho map of Tiberi et al. (2001) also
suggests significantly thinner crust to the west in the vicinity of our
station 38.
We interpret the late and strong Ps phase recorded at station 36
as reflections from the subducting African Plate. We image the reflector between stations 38 and 49 (Fig. 1) at a depth of 74 km,
without significant dip, but we cannot ascertain whether the reflections originate from the top or bottom (Moho) of the subducting
crust. Papazachos & Nolet (1997) present a 3-D velocity model for
the Hellenic subduction zone region derived from inversion of local
earthquake events and show an east–west slice that is roughly parallel and about 25 km south of our profile. Their P-velocity tomogram
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Gulf of Corinth Moho structure
images the top of the subducting crust beneath the Peloponnisos at
∼60 km, dipping very slightly (3◦ ) to the east, and with a velocity anomaly thickness of 10–20 km. The depth of our Ps reflector
would be consistent with a reflection from the base of the subducting
crust as imaged by Papazachos & Nolet (1997). Otherwise, if the
Ps reflector originates from the top of the slab, this would suggest a
very significant decrease in depth (∼10 km) of the slab from north
to south over a short distance (25 km). The slab as imaged by Papazachos & Nolet (1997) does not change its dip until it reaches the
eastern end of the GOC, where the dip sharply increases to ∼30◦ .
In contrast, we observe Ps reflections only at station 36, yet if the
slab were indeed (nearly) horizontal beneath most of the GOC we
would expect to see similar phases on all of the western stations and
possibly stations 20 and 19. The lack of any evidence of a Ps phase
at these stations suggests that, at the latitude of our profile, the dip
of the slab increases quite sharply in the vicinity of station 49 near
the western part of the gulf as shown on Fig. 4. Again, if this is the
case, the slab geometry changes dramatically over the short distance
between our profile and the parallel profile of Papazachos & Nolet
(1997).
Subduction of the African Plate begins at the southern/southwestern edge of the Mediterranean Ridge (Kastens 1991)
(Fig. 1 inset). The dip of the slab is shallow until at least the Ionian
island of Cephalonia where Hirn et al. (1996) image the top of the
slab at a depth of 12 km. To reach a depth of 74 km where we image
it between stations 38 and 49 implies an average dip of the slab of
at least 30◦ west of Cephalonia, as shown in Fig. 4.
The inversion of PmP reflection times recorded at 28 stations
reveals that in the vicinity of the GOC Moho depth generally increases from east to west. Thick crust (>35 km) occurs beneath the
Hellenides, both north and south of the gulf, where elevation exceeds 2000 m; however, the crust is also relatively thick beneath the
western end of the gulf. Thus, while the regions of high elevation
appear to be isostatically compensated to some extent, regions of
low elevation do not conversely correlate with thin crust, suggesting that the Moho has not responded in an obvious manner to the
recent (Quaternary) phase of rifting. Just south of the gulf, however,
between X = 100–120 km, we do see a region of relative thinning,
with the Moho at ∼35 km. The thinning south of the main Corinth
depression suggests that a significant component of strain may be
accommodated on faults south of the main rift. MCS reflection data
from the western GOC, at about this longitude, show that the importance of faulting within the gulf diminishes here (Weiss 2004),
which may be expected if extension transfers to, or is partitioned
with, faults south of the gulf.
The thinnest crust occurs beneath the Perahora Peninsula, which
is a feature on the isthmus separating the Corinth and Saronic
gulfs. The central part of this region is dominated by the Gerania
Mountains, which rise to over 1 km elevation. Moho depth here decreases to 25 km; however, as noted previously, tests showed that the
thin crust imaged here may in part be a response to higher crustal
velocities in this region. Instead of the relatively slow-velocity crust
associated with the rift it is conceivable that velocities in this region, which lies at the northwestern end of the Hellenic Arc, are
more typical of an arc setting, where mid-and lower-crustal rocks
are, on average, about 0.3–0.4 km s−1 faster than typical rift settings (Holbrook et al. 1992). The nearest arc-related volcano is
50 km to the southeast on the island of Methana (Fig. 1); however, there is evidence of volcanic activity on the isthmus at Susaki,
which is a solfataric field (Simkin & Siebert 1994). Another potential source for relatively high velocities are ophiolitic rocks, which
crop out in this region. High velocities may trade-off somewhat
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265
with depth, but it is unlikely that the crustal thickness here is greater
than 30 km, meaning that, regardless of velocity variations, this
region does indeed have the thinnest crust. The thin crust and topographic high may be a sign that the lithosphere is strong here
and resistant to stretching and subsidence, or that the thin crust
may in fact be a product of extension. For now this question is
unresolved.
The region of the Moho constrained by PmP reflection points lies
entirely within the map of Moho depth produced by Tiberi et al.
(2001) from inversion of Bouguer gravity anomalies. They assume
that the residual gravity anomaly, after removing the effect of the
subducting slab, is due entirely to Moho topography. This is somewhat analogous to our assumption that all variations in PmP times
are due to Moho topography; however, because we include a good
1-D velocity model in our inversion, the depths we obtain are meaningful in terms of their absolute values. The gravity inversion maps
gravity anomalies to an interface that is tied to an assumed reference Moho depth, chosen as 30 km by Tiberi et al. (2001). We
found that, on average, the Moho depths of Tiberi et al. (2001) are
5.5 km shallower than the depths derived by traveltime inversion,
suggesting that a more appropriate value for their reference depth
would have been ∼35 km. They do point out that their results are
stable for different reference depths, and thus we will compare our
results with theirs after applying a 5.5 km shift. Within the region
constrained by reflection points (non-grey region of Fig. 6) there
is a general agreement about a regional increase in Moho depth
from east to west; however, the correlation between high elevation in the Hellenides with thick crust is stronger in our model.
Both models show relatively thick crust beneath the western end of
the GOC. The relatively very shallow Moho beneath the Perahora
Peninsula is not present on the gravity Moho; instead the gravity
Moho contains a very shallow feature about 40 km further east and
other shallow features both slightly north and south of the central
GOC. Tiberi et al. (2001) suggest that a zone of relatively thin crust
along the northeastern margin of the GOC represents an offset between extension-related crustal thinning and the main axis of the rift.
Based on this they propose a model for extension which includes
a north-dipping upper crustal low-angle normal fault with ductile deformation in the lower crust. Our Moho model shows a
complicated pattern of depth variations that does not support this
simple interpretation. It is possible that pre-rift Moho topography complicates signals associated with extension-related crustal
thinning.
To explore further the possible relationships between crustal
thickness (Moho depth plus elevation, Fig. 8a) and elevation
(Fig. 8b), we cross-plot the two within the area of our tomographic
inversion (Fig. 8c). Assuming pointwise Airy isostasy, the ratio of
crustal thickness differences to positive elevations should equal the
crustal density divided by the density difference of mantle and crust.
A linear best fit to the observed data has a gradient of 4.16 and an
intercept of 32.8 km. Likewise, the ratio of crustal thickness differences to submarine depths should equal the crustal density contrast
with water divided by the density contrast of mantle versus crust.
A linear best fit to the observed data has a gradient of 2.03 and an
intercept of 34.2 km. Constraining the two relationships to have the
same crustal thickness at sea level (33 km at zero elevation), we plot
the combined best fit curve in Fig. 8(c).
Note that a subaerial gradient of 4.16 is equivalent to a crustal
density of 2.67 g cm−3 given a mantle density of 3.30 g cm−3 . The
former is the crustal density commonly used in Bouguer gravity
reductions. This value would predict a submarine gradient of 2.16 if
the bathymetry represented crustal depths. We know, however, that
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Figure 8. (a) Map of crustal thickness (i.e. Moho depth in Fig. 6 plus elevation). The contour interval is 1 km; heavy lines are the 30 and 40 km contours.
(b) Map of elevation. The contour interval is 0.5 km; the heavy line is the 1 km contour. (c) Plot of crustal thickness versus elevation for each point of the grid
constrained by seismic data (non-grey regions in other panels). The line represents a best estimate of crustal thickness versus elevation for regions both below
and above sea level. (d) Map of the difference between crustal thickness and predicted crustal thickness as predicted by the line in (c). In panels (a), (b) and (d)
triangles are stations used in the 3-D inversion to derive Moho depth.
there are up to 2.5 km of sediments beneath the GOC and that, if
sediments of average density 2 g cm−3 filled the rift basin, they would
produce a gradient of 1.16. Therefore, the observed intermediate
value (2.03) of the submarine gradient also appears reasonable.
The crustal thickness versus elevation data (Fig. 8c) show substantial variations from the simple gradients predicted assuming
Airy isostasy, indicative that the crust/lithosphere is able to support
loads (both positive and negative). Fig. 8(d) shows the spatial pattern of these variations. They are not random. We note the following
differences relative to Airy isostatic model predictions:
(1) The crust is differentially thickest beneath the western GOC
and thinnest beneath the Perahora Peninsula. The east–west gradient in differences along the GOC mimics the regional gradient in
topography and Moho depths outside the gulf (i.e. high/thick in the
Hellenides to low/thin towards the Aegean).
(2) As expected, short-wavelength topographic features are not
locally compensated. For example, the longitudinal river valleys at
X = 95 km and at the head of the Gulf of Itea, north of the GOC,
and at X = 110 km south of the GOC, interrupt the regional crustal
difference patterns with locally thicker anomalies.
(3) There are two well-resolved regions of crust thinner than predicted by local Airy isostasy. The first, north of the GOC, parallels
the North Gulf of Evia and is similar in location to that derived from
gravity modelling by Tiberi et al. (2001). The second, immediately
south of the GOC, follows the east–west normal fault trends there.
Both these regions also correlate with generally positive topography
(Fig. 8b), so the question arises as to the extent of undercompensation (or possibly lack of local compensation) represented by the
relative crustal thinning. Resolution of this issue is beyond the scope
of this study but will be investigated by joint inversion of the gravity
data and crustal structure in future work.
C O N C LU S I O N S
In this study we have presented a 2-D velocity model for an axial
profile through the Gulf of Corinth, and a regional map of Moho
depth derived from inversion of reflection traveltimes. Our results
contribute important information towards the understanding of the
deep structure in the GOC region. The 2-D analysis reveals a simple
velocity structure with a west-dipping Moho and a highly reflective
subducting African slab at ∼74 km depth beneath the western end of
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Gulf of Corinth Moho structure
the gulf. The 3-D analysis reveals thick crust beneath the Hellenides,
relatively thin crust immediately south of the GOC, and a singular
region of shallow crust centred on the Perahora Peninsula.
It is clear that the structural reality in the GOC (Fig. 8d) is more
complex than either uniform pure shear or asymmetric simple shear
models of crustal extension would predict. The region is not in local
Airy isostatic equilibrium. There are substantial inherited (pre-rift)
topography and crustal thickness variations that are superimposed
by a 3-D pattern of crustal thinning that does not simply correlate
with the location of upper crust brittle deformation. Notably, the
western GOC is little thinned, whereas both the areas immediately
northeast of and south of the GOC are.
AC K N OW L E D G M E N T S
Primary funding for this project was from the US NSF, with additional funding from Greece (NOA) and France (CNRS). We thank
the scientists and crew of the R/V Maurice Ewing cruise 0108. We
thank the Hellenic Coast Guard for facilitating our operations and
protecting the 6 km streamer with a chase boat. We thank Colin Zelt
for providing the reflection tomography code and for guidance on its
usage. We also thank Hélène Lyon-Caen for providing the recordings of our shots at the Corinth Rift Laboratory seismic stations.
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