Perceptible earthquakes in the broad Aegean area

www.elsevier.com/locate/tecto

Tectonophysics 371 (2003) 175–186

Perceptible earthquakes in the broad Aegean area

G.Ch. Koravosa, I.G. Mainb, T.M. Tsapanosa,*, R.M.W. Mussonc

aGeophysical Laboratory, School of Geology, Aristotle University of Thessaloniki, 54006 Thessaloniki, GreecebDepartment of Geology and Geophysics, University of Edinburgh, West Mains Road, Edinburgh EH9 3JW, UK

cBritish Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 3LA, UK

Received 22 October 2002; accepted 16 May 2003

Abstract

A probabilistic estimate of seismic hazard can be obtained from the spatial distribution, of earthquake sources, their

frequency–magnitude distribution and the rate of attenuation of strong ground motion with distance. We calculate the

earthquake perceptibility, i.e. the annual probability that a particular level of ground shaking will be generated by earthquakes of

particular magnitude, by weighting frequency–magnitude data with the predicted felt area for a given level of ground shaking at

a particular magnitude. This provides an earthquake selection criterion that can be used in the anti-seismic design of non-critical

structures. We calculate the perceptibility, at a particular value of isoseismal intensity, peak ground acceleration and velocity, as

a function of source magnitude and frequency for the broad Aegean area using local attenuation laws. We use frequency–

magnitude distributions that were previously obtained by combining short-term catalogue data with tectonic moment rate data

for 14 tectonic zones in Greece with sufficient earthquake data, and where contemporary strain rates are available from satellite

data. Many of the zones show a ‘characteristic earthquake’ distribution with the most perceptible earthquake equal to the

maximum magnitude earthquake, but a relatively flat perceptibility between magnitudes 6 and 7. The maximum perceptible

magnitude is in the fastest-deforming region in the middle of the Aegean sea, and tends to be systematically low on the west in

comparison to the east of the Aegean sea. The tectonic data strongly constrain the long-term recurrence rates and lead to low

error estimates (F 0.2) in the most perceptible magnitudes.

D 2003 Elsevier B.V. All rights reserved.

Keywords: Perceptibility; Ground motion; Seismic hazard; Aegean area

1. Introduction

Different physical, numerical, and statistical mod-

els have been applied over the last decade in order to

define the seismotectonic environment and the future

behaviour of a region (e.g. Ito and Matsuzaki, 1990;

Cowie et al., 1993; Lomnitz-Adler, 1993; Rundle and

0040-1951/03/$ - see front matter D 2003 Elsevier B.V. All rights reserve

doi:10.1016/S0040-1951(03)00216-6

* Corresponding author.

E-mail address: [email protected] (T.M. Tsapanos).

Klein, 1993; Kagan and Jackson, 1994; Main, 1995).

These models have various proportions of determin-

istic and probabilistic elements, and are capable, in

principle of validation in particular regions once

sufficient data become available. All of these models

predict that the frequency–magnitude distribution at

low magnitudes takes the form of the Gutenberg–

Richter law

logN ¼ a� bm; ð1Þ

d.

G.Ch. Koravos et al. / Tectonophysics 371 (2003) 175–186176

where a and b are model parameters and N is the

cumulative number of earthquakes with magnitude

greater than or equal to m. This relation defines the

annual probability of occurrence of a given earth-

quake magnitude in a given region, and is a strong

primary constraint on seismic hazard. However, the

distribution at high magnitudes is less certain due to

the small number of data points for the rare, large

events.

A variety of different approaches have been adop-

ted in an attempt to overcome this problem of uncer-

tainty in recurrence rates for large events in the

calculation of seismic hazard (Burton et al., 1984;

Makropoulos and Burton, 1985; Goes and Ward,

1994; Lutz and Kiremidjian, 1995; Shimazaki et al.,

1999; Papazachos, 1999; Papaioannou and Papaza-

chos, 2000). In particular, geological and geodetic

data on deformation rates can be used to place strong

constraints on the long-term recurrence. For example,

Goes and Ward (1994) applied a seismicity model

based on the concept of fault segmentation and the

physics of static dislocation, that allows for stress

transfer between fault segments. The application of

their model allows to obtain recurrence statistics and

long-term probability estimates for shocks with

Mz 6.0 on the San Andreas Fault. In this approach,

constraints are provided by geological and seismolog-

ical observations of segment lengths, characteristic

magnitudes and long term recurrence rates. Interaction

between faults is also a theme in the stochastic

earthquake occurrence model (Lutz and Kiremidjian,

1995). This process can be used to estimate the hazard

due to large, spatially and temporally, dependent

earthquakes. A similar stochastic method known as

‘‘renewal process’’ is described by Shimazaki et al.

(1999) and applied to the recurrence of large earth-

quakes in Japan.

Estimates of seismic hazard depend not only on

the recurrence rates of events of different magnitudes,

but also on the attenuation of seismic energy with

distance (due to a combination of elastic geometric

spreading and anelastic absorption). Burton et al.

(1984) combined earthquake magnitudes obtained

through Gumbel’s third asymptotic distribution of

extreme values with regional intensity attenuation

laws to determine the relative ‘perceptibility’ of

earthquakes of different magnitude at a given level

of ground shaking. The earthquake perceptibility

P(x |m) is defined as the conditional probability that

a site perceives ground shaking at least of level x

(where x may be intensity, peak ground acceleration,

velocity, etc.) given the annual probability of occur-

rence of an earthquake of magnitude m (Burton,

1978). Hence the most perceptible earthquake is the

event that is most likely to occur and be perceived or

felt at a given level of ground motion at a site or in

an area.

The concept of the ‘‘most perceptible earthquake’’

essentially involves a trade-off between small earth-

quakes, which are common but not widely felt, and

large earthquakes, which are uncommon but felt

widely. For any site, there will be a magnitude that

represents a maximum, determined by the variables

governing frequency and impact.

The concept is analogous to that of ‘‘design

earthquake’’ in hazard analysis. In a very recent

study Tsapanos (2003) applied the ‘‘design earth-

quake’’ concept and estimated the ‘‘design PGA’’

values for a seismic hazard scenario in the main

cities of the island of Crete. A given level of ground

motion hazard is associated with a specific probabil-

ity (the typical output of a hazard study for design

purposes), so it is reasonable to ask which earth-

quake parameters are most likely to produce this

ground motion. Is the hazard mostly from small local

earthquakes or large distant ones? The answer will be

different for different sites. This problem of disag-

gregation of hazard can be addressed analytically

(McGuire, 1995) or through simulation (Musson,

1999).

Makropoulos and Burton (1985) used the notion of

perceptibility to assess seismic hazard in Greece using

attenuation laws for ground acceleration. Most recent-

ly Papaioannou and Papazachos (2000) estimated

time-dependent hazard based on the whole earthquake

catalogue (not just the extreme values). They calcu-

lated the conditional probability of occurrence for

macroseismic intensity for 144 Greek cities, based

on a method that assumes fluctuations in b to be much

lower than those in a in Eq. (1) above (Papazachos,

1999). This is based on the hypothesis that b values

depend on material properties and the seismotectonic

setting and therefore should vary relatively smoothly

in space.

There have been many studies of seismic hazard in

Greece, due both to the availability of different types

G.Ch. Koravos et al. / Tectonophysics 371 (2003) 175–186 177

of data and practical concerns for the population. On a

global scale, Greece is ranked sixth in terms of

seismic hazard (Tsapanos and Burton, 1991). The

country has often experienced strong and catastrophic

earthquakes, with significant casualties and damage to

buildings infrastructure, e.g. the recent Athens earth-

quake (m = 5.9, 7 September 1999). The geographical

distribution of seismic hazard, based on such seismic

sources, has been determined by Papazachos et al.

(1993), and more recently quantified as a function of

various design parameters. These include: the maxi-

mum expected macroseismic intensity (Shebalin et al.,

1976; Papaioannou, 1984), the peak ground acceler-

ation or velocity (Algermissen et al., 1976; Makro-

poulos and Burton, 1985), and the duration of strong

ground motion (Margaris et al., 1990; Papazachos et

al., 1992).

In this paper we calculate the perceptibility of

ground motion (intensity, peak ground acceleration

and velocity) for the Aegean area. We use frequency–

magnitude relations from a recently-published data set

of historical and instrumentally recorded earthquakes,

constrained by recent satellite data on deformation

rates in Greece and the surrounding areas (Holt et al.,

2000; Koravos et al., 2003). We allow the distribution

to deviate from Eq. (1) at high magnitudes, and

calculate the annual probability of ground shaking

at a particular level as a function of earthquake

magnitude, based on appropriate ground attenuation

laws for intensity acceleration and velocity. The

results follow from a previous paper (Koravos et al.,

2003) where we used a new catalogue for earthquake

recurrence in Greece and the surrounding area to

determine the best-fitting frequency–magnitude rela-

tions some of the 16 tectonic zones defined by a

geodetic study of crustal deformation rates (Holt et

al., 2000). These authors considered two zones to

have insufficient data for the comparison of seismic

and tectonic information rate (zones 15 and 16), and

Koravos et al. (2003) found a further two (zones 1

and 2) insufficient for analysis of maximum magni-

tude. The deformation rates were used to constrain

extrapolations from the frequency–magnitude rela-

tion, and to estimate the maximum credible magni-

tude, a crucial parameter of interest to earthquake

design engineers. Here we combine these frequency–

magnitude distributions with local attenuation laws to

predict the distribution of perceptibility of ground

motion as a function of magnitude for the same

tectonic zones.

2. Data used and method applied

Papazachos et al. (2000) compiled a comprehen-

sive catalogue of instrumental and historical earth-

quakes in Greece and the surrounding area for the

time span 550 BC–1999 AD. The magnitudes are

quantified in terms of the seismic moment magnitude

scale which we should denote m. The completeness of

the catalogue and the corresponding errors in magni-

tude, depth and epicenter estimates are given by

Papazachos et al. (2000). The published catalogue is

available at (http://www.jjlahr.com/iaspei/europe/

greece/the). In a later study (Koravos et al., 2003), it

was shown that some of the large events included in

Papazachos et al. (2000) catalogue have systematical-

ly higher values. Because they were calibrated using

the events at the beginning of the 20th century (1900–

1912 or so) which were not corrected for the effect of

using undamped narrow-band seismic instruments

(Abe and Noguchi, 1983). For example, we corrected

the event of 1904 (41.80N 23.00E) that occurred in

zone 14 from ms = 7.7 to ms = 7.1, as well as the event

of 1905 (040.26N 24.33E) generated in zone 10 from

ms = 7.5 to ms = 6.8, based on the instrumental deter-

mination of Abe and Noguchi (1983) and Ambraseys

(2001). Since these events are amongst the largest

recorded events in the catalogue, they also have a

disproportionate effect on the calibration of the his-

torical catalogue for the largest large events. This

systematic effect is on average of the order of � 0.4

(Koravos et al., 2003) and this was then applied to the

historical data. We also added one additional year of

data (for the year 2000) from the preliminary annual

bulletins of the seismological station of the Aristotle

University of Thessaloniki. The data were divided

geographically into the 14 tectonic zones considered

to have adequate data by Holt et al. (2000), and used

to estimate the range of complete reporting for differ-

ent magnitude thresholds, the frequency–magnitude

relation and maximum magnitudes, constrained by the

tectonic deformation rates (Koravos et al., 2003). Fig.

1 depicts the examined zones and shows the spatial

distribution of the epicenters in the study area for the

time period 550 BC–2000 AD. For consistency the

http:\\www.jjlahr.com\iaspei\europe\greece\the

Fig. 1. Map of the area of study, divided into 16 numbered regions after Holt et al. (2000), showing the spatial distribution of the epicentres of

shallow seismicity (hV 50 km) in the Aegean and the surroundings for 550 BC–2000 AD. The data are from Papazachos et al. (2000).


numbers of the tectonic zones correspond to those in

Holt et al. (2000).

Here we calculate the perceptibility based on the

felt area at a given level of seismic intensity I

(modified Mercalli scale), ground motion acceleration

a in cm/s2 and velocity v in cm/s. The attenuation

formula for I as a function of hypocentral distance R

for the broad Aegean area is:

I ¼ 1:063þ 1:5222Ms � 1:1021lnR� 0:0043R ð2Þ

where Ms is the surface magnitude (Musson, 2000).

The corresponding attenuation law of Eq. (2) in terms

of epicentral distance D in km (for D>h) is:

I ¼ 1:063þ 1:5222Ms � 1:1021ln� ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi

D2 þ h2p �

� 0:0043� ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi

D2 þ h2p �

ð3Þ

where h is the depth in km.

The attenuation relationship for horizontal peak

ground acceleration, ag (in cm/s2) and velocity, vg(in cm/s), in the broad Aegean is:

lnag ¼ 4:37þ 1:02Ms � 1:65lnðD þ 15Þ þ 0:31S

þ 0:66P* ð4Þ

lnvg ¼ �0:18þ 1:29Ms � 1:62lnðD þ 10Þ þ 0:22S

þ 0:73P* ð5Þ

(Theodulidis and Papazachos, 1992), where Ms is the

surface magnitude, D is the epicentral distance in km,

the site effect S takes the value of zero at ‘‘alluvium’’

sites and one at ‘‘rock’’ sites. The parameter P*

accounts for the scatter in the data about the best

fitting line. P* = 0 for 50 percentile values and = 1 for

84 percentile values. The attenuation law of Eqs. (3),

(4) and (5), is shown in Fig. 2A, B and C, respectively

for Ms = 6.0, h = 10 km, S= 0.5 and P* = 0.

Fig. 2. Ground motion attenuation with distance for (A) modified

Mercalli intensity, (B) peak horizontal ground acceleration and (C)

peak horizontal ground velocity, for m= 6.0 and h= 10 km.


The surface-wave magnitude Ms is related to the

moment magnitude m by

m ¼ 0:56Ms þ 2:66 : 4:2VMsV6:0

m ¼ Ms : 6:0VMsV8:0ð6Þ

(Papazachos et al., 1997). The felt area of an event of

magnitude m producing an intensity of at least I at

distance r from the epicentre is:

AðIÞ ¼ p½rðIÞ�2 ð7Þ

The annual perceptibility of ground motion at a level I

(a or v) is then

PðI=mÞ ¼ FðmÞAðIÞ=Amax ð8Þ

where F is the incremental annual frequency of

occurrence of a magnitude in the range mF dm/2,and Amax is the total area of study.

3. Results

The perceptibility for the 14 zones is shown in Fig.

3 for I=VIII and a typical focal depth h = 10 km. The

solid lines represent the functional form for Eq. (8),

given the frequency distributions of Koravos et al.

(2003) and the attenuation laws described above,

plotted against the data. The finite focal depth implies

that there is a minimum threshold magnitude at the

surface that can produce ground shaking at a given

intensity (Main, 1995). In Fig. 3, this threshold

magnitude for perceptibility of I=VIII is above mag-

nitude 6.4. For the first two zones (zones 1 and 2) of

the examined area, the absence of data above the

threshold magnitude of 6.4 for Greece (Koravos et al.,

2003) made it impossible to estimate the maximum

credible magnitude. These regions are not considered

further. Above the threshold magnitude, most of the

remaining areas show a steady increase in percepti-

bility to a relatively flat local maximum or inflexion

point, and then a peak at the maximum magnitude.

For these plots, the most perceptible magnitude is

equal to the maximum magnitude. This is due to the

fact that the best fitting frequency–magnitude distri-

bution for these zones has a characteristic peak at the

largest event sizes (Koravos et al., 2003), consistent

Fig. 3. The perceptibility P=F(m)A(I)/Amax for the different tectonic zones identified by Holt et al. (2000), predicted from the frequency–

magnitude data and curve fits of Koravos et al. (2003), using I =VIII and h= 10 km.


Fig. 4. Perceptibility of ground motion at 0.2 g peak horizontal ground acceleration, assuming S= 0.5 for the site conditions and P*= 0 in Eq. (4).


Fig. 5. Perceptibility of ground motion at 10 cm/s peak horizontal velocity, assuming S= 0.5 for the site conditions and P*= 0 in Eq. (5).



with a ‘supercritical’ distribution (Main, 1995). Zone

9 in contrast shows a distinct maximum at m = 6.9

and zone 10 at m = 7.4, although zone 10 has a wide

range of magnitudes (m = 7.0–7.7) where the percep-

tibility remains constant. Main (1995) showed that

the perceptibility remains relatively constant for all

magnitudes above the threshold for the ‘critical’

(Gutenberg–Richter) distribution. This is consistent

with the results for zone 10.

The perceptibility for ground-motion acceleration

at a level of 0.2 g, assuming an average site condition

S = 0.5 and P* = 0 is shown in Fig. 4. In this case the

threshold magnitude of perceptibility is m = 5.6. Most

of the curves show features that are qualitatively

similar to the corresponding curves for intensity VIII

in Fig. 3. However, the peak probability occurs

systematically at lower magnitudes for acceleration

0.2 g in Fig. 4. For example, in zone 9 the most

Fig. 6. Spatial distribution of the examined zones and the perceptible magn

is also given in parenthesis for both magnitudes.

perceptible magnitude mp is in the range 6.1(F 0.1)

for acceleration cf. 6.9(F 0.1) for intensity. These low

error values in mp reflect the strong constraint of

deformation rates on the results. In zone 10 the broad

maximum in perceptibility has a lower and narrow

range for acceleration (m = 6.1–6.4), when compared

to the range of intensity (m = 7.0–7.7).

The perceptibility for ground-motion velocity at a

level of 10 cm/s for S = 0.5 and P* = 0 is shown in Fig.

5. The threshold magnitude in this case is m = 5.5 and

the general pattern in this figure is similar to the

corresponding curves for acceleration in Fig. 4. In

three zones 4, 6 and 7 the perceptible magnitude

shows a difference of the order of F 0.1 in correlation

with perceptible magnitudes estimated from intensity

and acceleration.

Fig. 6 shows the spatial distribution of the studied

zones and the perceptible magnitudes for intensity,

itudes for intensity, acceleration and velocity. The standard deviation


acceleration and velocity with their uncertainties.

Inspection of this figure shows that perceptible mag-

nitudes are over 7.0 in almost all the regions with

sufficient data, indicating that the seismic hazard is

really dominated by the rare, large events in the

Aegean area.

4. Discussion

Papazachos and Kiratzi (1996) examined the seis-

mic deformation velocities for the area of Greece and

its surroundings. They found high seismic deforma-

tion velocities for the middle part of the Aegean sea

having in general north–south direction, high to

intermediate values for the eastern part and low

values for the western part of the Aegean area where

the direction changes to east–west. Crustal deforma-

tion strain rates were determined by Kahle et al.

(1998, 2000) based on the use of GPS measurements

in the eastern Mediterranean region. They found high

values for the area of the north Anatolian fault (170

nstrain year� 1) while lower (90–120 nstrain year� 1)

and lowest (less than 40 nstrain year� 1) values were

determined for the southeastern Aegean and central

and southwestern Aegean, respectively. Our results

are consistent with the more recent data of Holt et al.

(2000) deduced by the combination of SLR and earth-

quake mechanisms. The obtained results (P(x |m), mp)

confirm that local deformation rate exercises a first-

order control on the seismic hazard. The highest value

for mp is in the fastest-deforming part in the middle

of the Aegean sea and tends to be systematically low

on the west in contrast to the east of the Aegean sea

(Fig. 6). For example in zone 11 where the defor-

mation velocity is 24 mm/year according to Papaza-

chos and Kiratzi (1996) and 28 mm/year according

to Holt et al. (2000), mp magnitudes (deduced from

intensity, acceleration and velocity) are equal to 7.6.

The other one is zone 12 where the seismic defor-

mation velocity is 16 mm/year (Papazachos and

Kiratzi, 1996) and 17 mm/year (Holt et al., 2000),

mp magnitudes are equal to 7.4. The Aegean Sea

coincides from north to south with zones 14, 10, 11,

6, 7 and 3 of Holt et al. (2000). We observed that

generally the easternmost zones (12, 8 and 4,

corresponding to the western part of Turkey), have

larger mp values than the westernmost ones (13, 9, 5

and 1, i.e. mainland Greece). These results confirm a

very good correlation that exists between the seismic

deformation velocities, deformation strain rates and

mp.

5. Conclusion

The perceptibility of ground motion has been

estimated for seismic zones corresponding to a grid

used to determine tectonic deformation rates, for

intensity, acceleration and velocity. Such results

can be used to constrain seismic hazard based on

combining spatially variable frequency–magnitude

and satellite strain rate data with attenuation laws.

Most of the deformation zones examined here ex-

hibit a characteristic or supercritical frequency–mag-

nitude distribution, where the moment release and

the seismic hazard is dominated by largest magni-

tudes, and the most perceptible earthquake is equal

to the maximum magnitude. The tectonic data

strongly constrain the long-term recurrence, and lead

to rather low estimates of the error in perceptible

magnitudes, on the order of F 0.1 or F 0.2 magni-

tude units. Only two zones, zones 9 and 10 showed

critical behaviour with a most perceptible magnitude

in the middle of the perceptible range. The location

of this peak is lower for acceleration than intensity

in both of these zones. Three zones had insufficient

data for the analysis presented here, while in the rest

of the zones we observed a good correlation be-

tween the seismic deformation velocities and the

perceptible magnitudes. Zones of high seismic de-

formation rates are correlated with high values of

perceptible magnitudes (mp). The middle of the

Aegean has the highest deformation rate and the

highest value for the mp. Intermediate to high

deformation rates and mp values are found on the

eastern side of the Aegean area, while low defor-

mation rates are associated with relatively mp param-

eters on mainland Greece, in the western part of the

examined area.

Acknowledgements

We thank James Jackson and John Haines for

providing the coordinates of the tectonic zones, and


for access to the original principal average strain

rate data used to calculate the tectonic moment

release rate. The contribution of RMW Musson to

this paper is made with the permission of the

Executive Director of the British Geological Survey

(NERC).

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Perceptible earthquakes in the broad Aegean area

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