Static stress changes associated with normal faulting earthquakes in South Balkan area

ORIGINAL PAPER

Static stress changes associated with normal faulting earthquakesin South Balkan area

E. Papadimitriou Æ V. Karakostas Æ M. Tranos ÆB. Ranguelov Æ D. Gospodinov

Received: 7 April 2006 / Accepted: 20 October 2006� Springer-Verlag 2006

Abstract Activation of major faults in Bulgaria and

northern Greece presents significant seismic hazard

because of their proximity to populated centers. The

long recurrence intervals, of the order of several hun-

dred years as suggested by previous investigations,

imply that the twentieth century activation along the

southern boundary of the sub-Balkan graben system, is

probably associated with stress transfer among neigh-

bouring faults or fault segments. Fault interaction is

investigated through elastic stress transfer among

strong main shocks (M ‡ 6.0), and in three cases their

foreshocks, which ruptured distinct or adjacent normal

fault segments. We compute stress perturbations

caused by earthquake dislocations in a homogeneous

half-space. The stress change calculations were per-

formed for faults of strike, dip, and rake appropriate to

the strong events. We explore the interaction between

normal faults in the study area by resolving changes of

Coulomb failure function (DCFF) since 1904 and hence

the evolution of the stress field in the area during the

last 100 years. Coulomb stress changes were calculated

assuming that earthquakes can be modeled as static

dislocations in an elastic half-space, and taking into

account both the coseismic slip in strong earthquakes

and the slow tectonic stress buildup associated with

major fault segments. We evaluate if these stress

changes brought a given strong earthquake closer to, or

sent it farther from, failure. Our modeling results show

that the generation of each strong event enhanced the

Coulomb stress on along-strike neighbors and reduced

the stress on parallel normal faults. We extend the

stress calculations up to present and provide an

assessment for future seismic hazard by identifying

possible sites of impending strong earthquakes.

Keywords Coulomb stress � Foreshock—main shock

triggering � Greece and Bulgaria

Introduction

The Krupnik–Kresna, SW Bulgaria, earthquake of

April 1904, with a reported magnitude of Ms7.8

(Christoskov and Grigorova 1968), was one of the

strongest events in the eastern Mediterranean region,

causing extensive damage in a broad area that has been

added to the casualties caused by its stronger foreshock

(Ms7.2) that occurred just 23 min earlier. The next

‘‘twin’’ seismic excitation in Bulgaria took place in

1928 with a foreshock of M6.8 on 14 April, following by

E. Papadimitriou (&) � V. KarakostasGeophysics Department,Aristotle University of Thessaloniki,54124 Thessaloniki, Greecee-mail: [email protected]

V. Karakostase-mail: [email protected]

M. TranosGeology Department,Aristotle University of Thessaloniki,54124 Thessaloniki, Greecee-mail: [email protected]

B. Ranguelov � D. GospodinovGeophysical Institute, Bulgarian Academy of Sciences,Sofia, Bulgaria

B. Ranguelove-mail: [email protected]

D. Gospodinove-mail: [email protected]

123

Int J Earth Sci (Geol Rundsch)

DOI 10.1007/s00531-006-0139-x

its stronger mainshock (M7.0) just four days later (18

April). The 1931 Valandovo earthquake (M6.7), lo-

cated in the border area between northern Greece and

Former Yugoslavian Republic of Macedonia (FY-

ROM), was preceded by a strong foreshock (M6.0) one

day before, although there is no information on the

fault associated with this latter event. The Thessalo-

niki, northern Greece, earthquake of 20 June 1978

(M6.5) was preceded by a moderate (M5.8) event

originated in a neighbour fault segment on 23 May

1978. This repeated pair-event occurrences motivated

the investigation on possible foreshock—main shock

triggering.

Each earthquake alters the state of stress in its sur-

roundings, and it is natural to investigate the stress

changes associated with the first shock in each pair of

events in order to evaluate the potential for the con-

sequent event occurrence. Coulomb stress changes are

widely used in the literature to seek for fault interac-

tion between large magnitude earthquakes as well as to

model aftershock patterns and seismicity rate changes

over long-time windows (Harris 1998; Stein 1999

among others). Such studies exploring fault interaction

have been compiled for the broader Aegean area

during the last years for strike-slip events (Nalbant

et al. 1998; Papadimitriou and Sykes 2001; Papadimi-

triou 2002), normal faulting (Papadimitriou and

Karakostas 2003; Papadimitriou et al. 2005) as well as

between strike-slip and subduction earthquakes

(Messini et al. 2005). The results of the above studies

revealed that the vast majority of the earthquakes

whose triggering was inspected, were located inside

areas of positive static stress changes.

In the present study, we investigate the stress per-

turbations caused by the stronger earthquakes that

occurred in the area of Bulgaria and northern Greece

since 1904, when the first strong main shock of the

instrumental era occurred. This involves incorporation

of both tectonic loading and coseismic slip. Therefore,

the scope of this research is to define the geometry of

the faults associated with the occurrence of strong

earthquakes (M ‡ 6.0) and their kinematics, and then

by making use of this information to compute the static

stress changes for the whole study area. The computa-

tions reveal the stress interactions between the faults

and are important in determining where the next strong

events, outside the seismogenic volume, might occur.

The identification of currently being stress enhanced

sites contributes to the evaluation of future seismic

hazard in the study area. In the case of seismic excita-

tions involving multiple strong events in addition to the

main shock, like the earthquake pairs in our study area,

such investigation comprising determination of a strong

earthquake location and mechanism, together with the

knowledge of nearby major faults, may be useful in

post-event imminent hazard estimates.

Seismotectonic setting

North of the western continuation of the North Ana-

tolian Fault the Aegean extensional regime is present

in south Balkan region, namely central and southern

Bulgaria, northern Greece, FYROM and eastern

Albania (Fig. 1). The southern Balkan forms the

northern part of the Aegean extensional system al-

though deformation is not as great as in the southern

Aegean part (Burchfiel et al. 2000; Tranos et al. 2006).

To the east, the region of central Turkey lying between

the north and east Anatolian Faults is actively moving

westward relative to Eurasia as a coherent unit (Rei-

linger et al. 1997; McClusky et al. 2000). In western

Turkey this westward–southwestward moving region is

affected by extensional tectonics that is more prevalent

into the Aegean and mainland Greece due to the

north–south spreading of the Aegean microplate

(McKenzie 1972). This extension occurs as far north as

central Bulgaria. Recent investigations (Burchfiel et al.

2000; Nakov et al. 2001) suggest that the northern

boundary of the Aegean extensional regime passes

through north central Bulgaria, supporting McKenzie’s

interpretation based on limited seismological data that

a poorly defined crustal boundary passes through that

same region (McKenzie 1972).

The northern boundary of the Aegean extensional

system is defined by the sub-Balkan graben consisting

of nine east–west trending grabens located in central

Bulgaria that are complexly faulted with faults on

both flanks, but the main faults occur along the

northern side (Tzankov et al. 1996). Master graben-

forming faults are south dipping, low-angle normal

faults (dipping c. 30� or less) which have been cut by

steeper (60–80�) young to active normal faults at the

base of Stara Planina Mountains. Most of the active

and young faults, being normal, with a few of them

with strike-slip component, strike generally E-W,

WNW-ESE, NW-SE and rarely NE-SW, and they

bound basins of quaternary sedimentation (Nakov

et al. 2001). There is a continuous to discontinuous

line of faults that trend through central Bulgaria and

mostly lie along the south flank of the Stara Planina

Mountains. The faults bound the sub-Balkan graben

system for which Tzankov et al. (1996) suggested that

the faults dip gently south and have a long-term slip

rate of 1–2 mm/year. The faults continue both east to

the Black Sea, where they bound narrow and less


123

well-defined half grabens, and west, where they bound

the wide and well-developed Sofia graben. North of

this belt of faults there is little evidence for significant

faulting except for the Gorna Oriahovitsa fault zone

(Fig. 2) that exhibits earthquake activity but little

evidence for surface faulting. From field observations

the faults are mainly normal, with strikes indicating

general N-S extension within most of Bulgaria. Indi-

vidual faults have variations in strike that suggest

there may be strike-slip components on the NW- to

NE-striking faults. Most of the faults form linear

mountain fronts, indicating that most of the present

day topography of Bulgaria is the result of extensional

faulting.

The overall kinematic pattern shown by GPS results

confirm that the N-S to NNE-SSW extension present

within central and northwest Bulgaria marks the

northern boundary of the Aegean extensional region

(Kotzev et al. 2001) and that most of recent topogra-

phy of Bulgaria is of extensional origin, as suggested by

Zagorchev (1992), Tzankov et al. (1996), and Burchfiel

et al. (2000). Studies on earthquake focal mechanisms

Fig. 1 The broader southBalkan geographic regionwhere the study area isindicated by the rectangle.The dominant stress field isindicated by arrows. NAFNorth Anatolian Fault, NATNorth Aegean Trough

Fig. 2 Study area with themain fault systems (afterMatova et al. 1996; Tranoset al. 2003). Active faultsassociated with strong mainshocks occurring since thebeginning of the twentiethcentury are depicted by gray(this study) and the onesactivated in the past by whitecolor (Papazachos et al.2001). Fault plane solutionsare shown as lower-hemisphere equal areaprojections. The occurrencedate (year/month/date) ofeach event is annotated by thefocal spheres. SBGS sub-Balkan graben system


123

also indicate that active tectonism is dominated by

north–south extension, with rare indications of strike-

slip motion (Van Eck and Stoyanov 1996). The east–

west trending faults show only evidence for dip-slip

movement (Tranos et al. 2006) and there is little or no

field data to support strike-slip displacement on most

of them. Some faults of north–eastern and north–wes-

tern strike show evidence for strike-slip and others can

be interpreted to have a strike-slip component (Bur-

chfiel et al. 2000).

Figure 2 shows the major active normal fault zones

of the study area along with the fault segments that are

known to be associated with the occurrence of strong

earthquakes (Matova et al. 1996; Papazachos et al.

2001). Faults activated during the twentieth century,

the period that the present study covers, are shown in

gray, contradistinguishing them from the ones previ-

ously activated and shown in white. The main fault

zones (Matova et al. 1996; Tranos et al. 2003) are also

shown by thinner black lines. The fault plane solutions

of the respective main shocks are also shown by the

faults. The long-term slip rates on them were taken

equal to 1 mm/year following the above-mentioned

relevant results and more recent work of Kotzev et al.

(2006). Information on the fault location and geometry

is given in Table 1.

Stress evolutionary model

The changes in the stress field associated with coseis-

mic slips and long-term tectonic loading on the major

faults of the study area are calculated, following the

approach of Deng and Sykes (1997). The changes in

the Coulomb failure function (DCFF) depend on both

changes in shear stress, Ds, and normal stress, Dr[(modified from Scholz (1990)]:

DCFF ¼ Ds þ l0Dr ð1Þ

Here l¢ is the apparent coefficient of friction. Both Dsand Dr are calculated for the fault plane of the next

earthquake in the sequence of events, whose triggering

is inspected. The change in shear stress, Ds is positive

for increasing shear stress in the direction of relative

slip on the observing fault; Dr is positive for increasing

tensional normal stress. When compressional normal

stress on a fault plane decreases, the static friction

across the fault plane also decreases. Both positive Dsand Dr move a fault toward failure; negative Ds and Drmove it away from failure. A positive value of DCFF

for a particular fault denotes movement of that fault

toward failure (that is, the likelihood that it will rup-

ture in an earthquake is increased). The stress calcu-

lations are performed for an isotropic elastic half-space

(Erikson 1986; Okada 1992). The shear modulus and

Poisson’s ratio are fixed as 33 GPa and 0.25, respec-

tively. The selection of the value of the apparent

coefficient of friction, l¢ is based on previous results. A

value of l¢ equal to 0.4 was chosen and considered

sufficient throughout the calculations (King et al.

1994a; Nalbant et al. 1998).

Interseismic stress accumulation between large

events is modeled by ‘‘virtual negative displacements’’

along major faults in the entire region under study

using the best available information on long-term slip

Table 1 Major fault segments in the study area and long-term slip rates

Segmentnumber

Name Center Strike(deg)

Dip(deg)

Length(km)

Depth(km)

Faulttype

Slip Rate(mm/year)

Reference

Latitude(�N)

Longitude(�E)

1. Ierissos 40.53 23.97 93 53 50 0–15 N 1 12. Stivos W. 40.66 23.18 278 46 24 0–15 N 1 23. Stivos E. 40.66 23.35 265 40 10 0–15 N 1 24. Valandovo 41.30 22.50 270 53 38 0–15 N 1 15. Kocani 41.82 22.57 280 45 30 0–15 N 1 36. Krupnik W. 41.81 22.92 265 45 28 0–15 N 1 37. Krupnik E. 41.85 23.16 243 45 15 0–15 N 1 38. Kyustendil 42.33 22.73 256 53 60 0–15 N 1 19. Sofia 42.81 23.42 110 53 61 0–15 N 1 110. Chirpan 42.21 25.43 108 37 38 0–15 N 1 411. Popovitsa W. 42.28 24.70 300 67 21 0–15 N 1 412. Popovitsa E. 42.14 25.02 291 28 32 0–15 N 1 413. Gorna

Oriahovitsa43.13 25.75 280 37 22 0–15 N 1 5

1 Papazachos et al. (2001); 2 Tranos et al. (2003); 3 This study; 4 Bonchev and Bakalov (1928); 5 Alexiev and Georgiev (2002)


123

rates. These virtual dislocations are imposed on the

faults with sense of slip opposite to the observed slip.

The magnitude is incremented according to the long-

term slip rate of the fault. This virtual negative slip is

equivalent to constant positive slip extending from the

bottom of the seismogenic layer to infinite depth.

Hence, tectonically induced stress builds up in the

vicinity of faults during the time intervals between

earthquakes. All computed interseismic stress accu-

mulation is associated with the deformation caused by

the time-dependent virtual displacement on major

faults extending from the free surface up to the depth

at which earthquakes and brittle behavior cease

(~15 km).

Fault length and average displacement are two

parameters necessary for the model application. The

distribution of slip is actually non-uniform along a

fault, but we are interested in its average value, as

well as for a fault length expressing the main rupture.

For this reason, we used scaling laws suggested by

Papazachos et al. (2004) for continental dip slip faults,

in order to calculate both these values, as a function of

the magnitude, M, of the particular earthquake:

log L ¼ 0:50M � 1:86 ð2Þ

log u ¼ 0:72M � 2:82 ð3Þ

that give the fault length, L (in km), and average

displacement, u (in cm), as a function of the mainshock

moment magnitude, Mw. These values were checked

and found in agreement with Wells and Coppersmith

(1994). Information on the focal parameters of the

events taken into account in the stress calculation

model is given in Table 2.

Time-dependent stress field changes

Stress changes, i.e., values of DCFF, since 1904 are

computed at a depth of 8 km. This depth was chosen to

be several kilometers above the locking depth (15 km)

in the model. This is in agreement with King et al.

(1994b) who found that seismic slip peaks at mid-

depths in the seismogenic zone. The depths of the

larger (M ‡ 6.0) earthquakes that occurred in the

broader Aegean region, for which reliable determina-

tion of the focal parameters exists based either on

waveform inversion or recordings of local seismic

networks, range from 8 to 13 km (Papazachos et al.

1998). From studies of aftershock sequences for which

reliable determination of the aftershocks focal param-

eters also exists, it is evident that the majority of their

foci are located in a seismogenic layer extending from

a depth of 3 to 15 km, some reaching a depth of 20 km.

Considering all of the above information, the depth of

the seismogenic layer in our calculations is taken to be

in the range of 3–15 km for all of the events we mod-

eled. The rupture models are approximated by rect-

angular surfaces with two edges parallel to the Earth’s

surface. Although the rupture is more complicated and

slip varies along its segments, it is believed that the

above approximation is sufficient to identify areas of

stress changes, when they are computed for distances

far from the causative fault. In Fig. 3 pure green indi-

cates no significant change in CFF blue regions denote

negative changes in CFF and inferred decreased like-

lihood of fault rupture, and are called stress shadows

(Harris and Simpson 1993, 1996). Yellow–red regions

represent positive DCFF and increased likelihood of

fault rupture, and are called stress bright zones. Sha-

dow zones and bright zones are specific to strike, dip,

and rake of the fault that experiences the DCFF.

1904 Kresna earthquakes The earthquake of 4 April

1904 is one of the largest shallow twentieth century

events on land in the Aegean back arc region, which

caused extensive damage and the geological effects

included landslides, rock falls, liquefaction of the

ground, changes in water and stream flow, and the

surface faulting (Ambraseys 2001). Magnitude esti-

mates assigned an Ms as large as 7.8 for the main shock

and 7.2 for the foreshock, respectively (Christoskov

and Grigorova 1968), which considered exaggerated as

many earthquakes of that era. Pacheco and Sykes

(1992) give a seismic moment of 0.44·1027 dyn cm,

which fully agrees with Dineva et al. (2002) who sug-

gest Mw6.8 and 7.0, for the foreshock and mainshock,

respectively. Ambraseys (2001) suggests an Ms7.2 by

instrumental data reappraisal and 7.1 from reassess-

ment of the intensity distribution (Mo = 3.1·1026 and

8.1·1026 dyn cm, respectively). Papazachos and Papa-

zachou (2002) give a magnitude of 7.0 for the fore-

shock and of 7.3 for the mainshock, on the basis of

macroseismic observations. The main shock was pre-

ceded by 6 months foreshock sequence and followed

by aftershock activity that lasted until the beginning of

1907 (Dineva et al. 1998). After 1907, the seismic

activity in this area decreased [with only six moderate

events of M = 5.0–5.5 (Sokerova et al. 1989)].

Since neither the surface break nor the source faults

of the 1904 events were identified, there are contro-

versies on the geometry and kinematics of the related

rupture zone. In particular, Meyer et al. (2002) based

on morphotectonic indexes suggested that the 20 km


123

long NE–SW striking Krupnik fault should be associ-

ated with this seismic activity, among the three most

conspicuous ones in the same region, since Bansko–

Predela and Kocani faults are far from the epicentral

area. In an investigation on the static stress changes

calculations by Ganas et al. (2005) the faults that are

suggested as possible ones to accommodate the cosei-

smic slip is the eastern part of the Krupnik fault for the

first event, while the main shock is attributed to the

western part of the Bansko–Predela fault that was

evidently triggered by stress transfer.

An effort is made in the present paper to reexamine

the available information and use recent tectonic

analysis results in order to define the causative faults of

the two strong events. According to macroseismic

descriptions of Ambraseys (2001) and isoseismals of

Shebalin (1974) and Papazachos et al. (1997) the area

that has been more affected by both the foreshock and

mainshock is extended from Kocani that is located in

FYROM, up to Krupnik that is located in SW Bul-

garia. This rupture zone had a total length of almost

70 km. According to Ambraseys (2001), the region

mostly affected by the foreshock was in the moun-

tainous area in FYROM, where there is some evidence

that the first shock caused widespread, and in places

serious damage with casualties. The main shock added

to the damage, which now extended over an area of

about 40 km, which straddles the modern FYROM–

Bulgaria borders. The combined epicentral area gen-

erated by the foreshock and the main shock extends in

an almost E-W direction for about 65 km, from

Krupnik to near Kocani in FYROM, being consistent

with a 7.2 magnitude earthquake and a seismogenic

depth of 10 km.

Recent investigation on the SW Bulgaria terrain by

Tranos et al. (2006), indicated the existence of a major

E-W striking Kocani–Krupnik rupture zone dipping to

the north and having a length of more than 50 km, to

transect the region joining the Kocani and Krupnik

faults. This rupture zone with edges defined by the two

bends formed in the Kocani and Gradevo area is

considered as potentially associated with the 1904

earthquakes, since satisfies the above mentioned con-

ditions such as the length, the strike and the area of the

reported damage. The distribution of weak earth-

quakes in the area to the east of Krupnik fault shows a

seismogenic structure with a WSW-ENE strike,

revealing that this seismogenic structure has to be

prolonged at least 20 km to west–southwest on the

territory of FYROM.

Fault plane solutions for the 1904 events are not

available, and for this reason their rupture models

were derived from considering smaller magnitude

earthquakes focal mechanisms (Alexiev and Geor-

giev 2002), fault strike from structural analysis

mentioned above and considering a dip of 45�, which

is typical value for crustal normal faults. Additional

argument supporting the definition of the Kocani-

Krupnik rupture zone as related with the 1904

seismic sequence is that its almost E-W strike is

optimally oriented to the contemporary least princi-

pal stress axis (tension) governing the study area, as

recently defined by geodetic information (Kotzev

et al. 2001) and fault slip data analysis (Tranos et al.

2006).

Figure 3a shows the coseismic stress changes asso-

ciated with the 1904 foreshock for which a displace-

ment of 1.0 m and a fault length of 30 km were

Table 2 Rupture models for the earthquakes included in the stress calculation model

Date Time Latitude (u� N) Longitude (k� E) Depth (km) L (km) u (m) Mw Mechanism

Strike Dip Rake

1904, April 4 10:02:34 41.85 23.0 3–15 30 1.01 7.2 280 45 –82 16.8 2

1904, April 4 10:25:55 41.80 23.0 3–15 42 1.66 7.8 265 45 –93 17.0 27.0 37.2 4

1913, June 14 09:33:13 43.10 25.70 3–15 22 0.61 7.0 280 37 –84 56.4 2

1928, April 14 09:30:01 42.15 25.28 3–15 38 1.20 6.8 95 38 –90 61928, April 18 16:47:19 42.10 25.00 3–15 47 1.66 7.0 300 62 –65 61931, March 8 01:50:28 41.38 22.49 3–15 30 1.00 6.7 270 53 –93 41932, September 26 19:20:42 40.45 23.86 3–15 50 1.66 7.0 93 53 –93 41978, May 23 23:34:11 40.698 23.295 3–15 12 0.25 5.8 265 40 –83 71978, June 20 20:23:21 40.729 23.254 3–15 24 0.72 6.5 278 46 –70 7

1 Christoskov and Grigorova (1968); 2 Dineva et al. (1998); 3 Pacheco and Sykes (1992); 4 Papazachos and Papazachou (2002); 5Paskaleva et al. (1986); 6 Karakostas et al. (2006); 7 Soufleris and Stewart (1981)


123

computed. Rupture is taken to extend throughout the

seismogenic layer, i.e., from 3 to 15 km (Table 2). The

foreshock has enhanced the positive stress changes on

the Krupnik fault where the main shock occurred just

23 min later. The fault associated with the occurrence

of the main shock is approximated with two planar

orthogonal surfaces, following the descriptions of the

surface expressions and macroseismic descriptions,

giving a total fault length of 42 km (Table 2). Given its

faulting length, the seismogenic thickness, and the

seismic moment that corresponds to its magnitude, a

slip of 1.66 m was calculated. The old original

descriptions report that the fault vertical dislocations

reached up to 3–4 m that can be considered either as

primary coseismic displacements (Shanov et al. 1999)

or descriptions of landslides fissures extended to a total

length of 50–60 km (Ranguelov et al. 2001). Figure 3b

shows the state of DCFF just after the main shock

occurrence, where a broad bright zone appears cover-

ing the central part of Bulgaria and partially the fault

associated with the 1928 earthquakes.

1913, Gorna Oriahovitsa earthquake Fig. 3c shows

the accumulated static stress changes due to the

coseismic slip of the 1904 events and 9 years (1904–

1913) tectonic loading on the major faults of the study

area. This earthquake occurred at the northern part of

Bulgaria where E-W trending normal faults dip to the

north (Alexiev and Georgiev 2002), in an area of

positive DCFF due solely to the tectonic loading, be-

cause the 1904 events were too far to affect its loca-

tion. According to Matova et al. (1996) this area

belongs to the Fore-Balkan oblique (dextral-normal)

zone that dips 50–80�N. The main shock magnitude

has been redetermined as equal to 6.4 by Dineva

et al. (2002). Published fault plane solution is not

available for this event, for which a representative

normal faulting was considered (strike = 280�,

dip = 37� and rake = –84�, Papazachos and Papazac-

hou 2002) and fault length and coseismic slip, equal to

22 km and 0.60 m, respectively, were estimated from

relations (2) and (3).

1928 Chirpan and Plovdiv earthquakes In April

1928, the Upper Thracian Depression in southern

Bulgaria was struck by two destructive earthquakes

that occurred only 5 days apart associated with

faulting of total length equal to 54 km and maximal

height difference of 3.5 m, for which the performed

geodetic measurements show the entirely normal

faulting character (Yankov 1945). Information on

magnitudes and macroseismic intensities is given by

Christoskov (2000). During the generation of the

mainshock (M7.0), the ground rupture was extended

more to the west and turned to the northwest

(Papazachos and Papazachou 2002). Numerous after-

shocks were recorded during the following month, 14

of them with M ‡ 5.0 (van Eck and Stoyanov 1996).

The largest of those was an M5.7 earthquake on 25

April near the village of Gulubovo, ~50 km east of

the first shock (Vanneste et al. 2006). Isoseismals of

the three largest events of this seismic excitation are

constructed by Shebalin (1974) and Papazachos et al.

(1997). Karakostas et al. (2006) used information on

surface faulting (Bonchev and Bakalov 1928) and

distribution on surface deformation (Yankov 1945)

caused by the two events for the determination of the

slip distribution and confirmation of their faults

geometry.

Figure 3d shows the accumulated static stress chan-

ges just before the occurrence of the 1928 foreshock,

which is associated with the Chirpan fault located inside

a bright zone. Figure 3d differs from Fig. 3c in that 15

more years of tectonic loading is added in the stress

field calculation along with the coseismic slip of the

1913 earthquake. During the 14 April earthquake two

main parallel surface breaks formed, trending 100–

110�E with a distance between them equal to 15 km.

Both were reported to have a throw of 0.3–0.4 m, down

to the south and down to the north, respectively

(Bonchev and Bakalov 1928). The continuous south-

dipping rupture had a 38 km length and an average

displacement of 40 cm, reaching a maximum of 50 cm

in the middle (DIPOZE 1931). Vanneste et al. (2006)

suggest that only the north break is the direct expres-

sion at the surface of the fault that generated the 14

April earthquake, which is a long regional fault striking

E-W and dipping to the south. Dimitrov et al. (2006)

failed in determining the fault plane solution from the

collected available data of P first arrivals while Jackson

and McKenzie (1988) suggested a focal mechanism

from surface ruptures (strike = 105�, dip = 45�,

rake = –90�), being in excellent agreement with a re-

cently determined smaller magnitude one (09/09/1991,

strike = 108�, dip = 37�, rake = –99�) by Alexiev and

Georgiev (2002). In our study, we adopted the solution

suggested by Karakostas et al. (2006) since it better

satisfies the reported surface deformation (strike = 95�,

dip = 45�, rake = –90�).

The foreshock occurrence has enhanced the stress

field at the location of the mainshock, as it is re-

vealed by comparing Fig. 3d and e, evidencing the

main shock triggering. The 18 April earthquake

generated a 53-km long system of discontinuous

breaks, trending 120–160�E, with throws up to 1.5 m,

and in one place even up to 3.5 m, down to the

north. Fault plane solutions published for this event

from Glavcheva (1984), VanEck and Stoyanov


123

Fig. 3 Coulomb stress calculations for normal faults at a depthof 8.0 km. The stress pattern is calculated for the faulting type ofthe next large event in the sample. Changes are denoted by thecolor scale at bottom (in bars). The faults associated with theoccurrence of the events included in our evolutionary model aredepicted by black color. a Coseismic Coulomb stress changesassociated with the 1904 foreshock. b The stress field is invertedaccording to the fault plane solution of the 1904 main shock.c Stress evolution until just before the 1913 event. d State of

stress just before the occurrence of 1928 foreshock. e State ofDCFF before the 1928 mainshock. f Stress evolution just beforethe 1931 earthquake. g State of stress before the 1932 Ierissosevent. h Stress pattern just before the 23 May 1978 event. i TheDCFF just before the 1978 mainshock. j Coulomb stressevolution until 2005 for normal faulting. k Same as j, exceptl¢ = 0.2. l Same as j, except l¢ = 0.6


123

(1996), and Dimitrov and Ruegg (1994) indicated a

WNW-ESE oriented normal fault with a significant

dextral component. The data used for the solution

were only from remote stations, since there were no

data from proximate stations, and most probably, this

fact resulted in the solution of normal/strike-slip

character (Alexiev and Georgiev 1996). The more

recent determination, based on newly collected data

by Dimitrov et al. (2006), gives a northeast-dipping

fault in good agreement with the main surface rup-

Fig. 3 Continued.


123

ture (strike = 300�, dip = 67�, rake = –124�). This

solution involves a dextral strike-slip motion that

disagrees with the N-S extension prevailing in

the area and a NW-SE fault strike. A solution

determined for a smaller magnitude earthquake

(strike = 291�, dip = 28�, rake = –90�) by Alexiev and

Georgiev (2002) is in good agreement with the latter

solution as far as the fault strike concerns but

keeping the normal character of the faulting. Kara-

kostas et al. (2006) found that a fault with a mean

strike = 300�, dip = 62�, rake = –65� is the most

suitable to reproduce the observed deformation field

when comparing modeled displacements with those

obtained from geodetic measurements by Yankov

(1945). These fault parameters are in agreement with

both surface ruptures and the orientation of the axis

of maximum extension in the area (Kotzev et al.

2006) resulting to a left-lateral strike slip component

added to the normal faulting. In order to assign a

more detailed geometry we considered two segments

for the source approximation differing in strike.

1931 Valandovo earthquake The largest foreshock

(M6.0) occurred on 7 March, one day before the

mainshock (Papazachos and Papazachou 2002). Dine-

va et al. (1998) refer on both strong events and on the

aftershock sequence lasting until 1932, noting the lack

of historical information on reactivation of the area in

the past. Since there is no other source of information

for double event and faulting descriptions, we consid-

ered only the mainshock in our evolutionary model,

keeping in mind this ‘‘twin’’ occurrence of strong

earthquakes in our study area. It means that, like in the

previous cases a strong but smaller magnitude earth-

quake triggers the neighbor fault segment where the

stronger mainshock occurred.

Figure 3f shows the evolved stress field before the

occurrence of this event, computed according to its

fault plane solution (strike = 270�, dip = 53�, rake =

Fig. 4 Accumulated static stress changes associated with thetectonic loading on the major faults and the coseismic slip of theearthquakes taken into account in the stress evolutionary model,resolved onto the rupture plane of the next strong event.Contour lines are accompanied with corresponding values ofstress changes in bars. Rectangles denote the rupture areas,

considered as rectangular surfaces with two edges parallel to theEarth’s surface, for a The 1904 Kresna main shock. b The 1913Gorna Oriahovitsa event. c The 1928 Chirpan foreshock. d The1928 Plovdiv mainshock. e The 1931 Valandovo earthquake.f The Hierissos earthquake. g The 1978 Thessaloniki foreshock.h The 1978 Thessaloniki main shock


123

–93�). A very small bright zone started to be created at

the western part of its associated fault, meaning that

the 27 years (1904–1931) tectonic loading was not en-

ough to obliterate the stress shadow created by the

1904 coseismic slips, revealing that these two parallel

normal fault zones were inactive for much longer time

to allow strain accumulation and lack of stress shad-

ows.

1932 Ierissos earthquake Fig. 3g shows the accumu-

lated stress changes just before this event, calculated

for its fault plane solution (strike = 93�, dip = 53�,

rake = –93�). It is associated with the WNW-ESE

Stratoni fault in the east part of the Chalkidiki penin-

sula (Pavlides and Tranos 1991), which is located inside

an area of positive DCFF, due solely to the tectonic

loading since this fault is located far from the past

ruptures and it is not affected by their coseismic slip.

Its coseismic slip resulted in the creation of stress-en-

hanced areas at the location of the 1978 earthquakes

(Fig. 3h).

1978 Thessaloniki earthquakes Triggering of the

foreshock in the main shock occurrence has been evi-

denced by Tranos et al. (2003), their approach being

incorporated in the present study. The 23 May 1978

foreshock (M5.8), with a fault length of 11 km in full

agreement with the total length of the ground ruptures

associated with this event aligned along the Geraka-

rou–Stivos fault and the aftershock distribution, cre-

ated bright zones located along the main fault zone

both to the west and east of its rupture. The main shock

(M = 6.5) occurred in the region where positive chan-

ges in DCFF had the largest values, showing that the

foreshock enhanced stress and brought the adjacent

fault segment closer to failure (Fig. 3i).

Accumulated stress changes The evolved stress field

due to the long term tectonic loading on the major

faults in our study area and the coseismic slips of all

strong events that occurred in Bulgaria and northern

Greece since 1904 up to date and described above, is

shown in Fig. 3j. This figure evidences that the large

part of the study area is located inside stress shadows.

Since the long-term slip rates are assigned low values,

as geodetic information and infrequent occurrence of

strong events manifest, it is derived that in the near

future it is not expected that the stress shadows will be

eroded in many places.

Testing the values of l¢ A lower apparent coefficient

of friction (l¢ = 0.2) is used in Fig. 3k. The stress pat-

tern corresponding to the lower l¢ (= 0.2) is very sim-

ilar to that depicted in Fig. 3j, which was calculated

with a higher l¢–value (= 0.4). Fig. 3l shows the stress

pattern calculated for a higher apparent coefficient of

friction (l¢ = 0.6). The resulting pattern remains also

unaffected, without allowing the discrimination of the

most proper value of l¢, in accordance with previous

investigators as mentioned above. Therefore, the value

of the apparent coefficient of friction selected and used

is suitable for the calculations performed in the present

study.

In order to examine the effect of the evolutionary

model on the incoming ruptures, the cumulative static

stress changes were calculated onto the fault plane of

each event, just before its occurrence, and shown in

Fig. 4. Although precise hypocenter location or details

about rupture initiation are not available in our case,

and in this case it is not feasible to correlate them with

the position of the maximum positive DCFF, it is evi-

dent in all cases except one (the case of 1931 Val-

andovo earthquake, Fig. 4e) that the assumed rupture

surfaces are inside stress enhanced areas. Even in the

exception mentioned above, the stress shadow started

to be reduced in a small area at the western part of the

assumed rupture. We may with caution make the

speculation here, that the strong foreshock of this

seismic sequence probably took place onto this area

and that the stress changes resulted after its occurrence

brought closer to failure the 1931 main shock.

Stress transfer by foreshocks to the future rupture

zone was exhaustively investigated in the case of Loma

Prieta earthquake (Perfettini et al. 1999). Although

two moderate events that occurred only 11 km from

the main shock hypocenter did not bring the future

rupture zone closer to Coulomb failure, they reduced

the normal stress onto the fault plane at the site where

the greatest slip subsequently occurred in the Loma

Prieta earthquake. It is then suggested that these

foreshocks are more likely to have influenced the slip

distribution on the main shock fault. In the present

study, details on the slip distribution are not known

and therefore such correlation is not feasible. Never-

theless, the foreshocks of the 1904, 1928 and 1978

probably brought closer to failure their counterpart

main shock, as it is evidenced by the resulted positive

Coulomb stress changes (‡1 bar in all cases) on the

future rupture (Fig. 4a, d, h, respectively).

Discussion and conclusions

It is now widely accepted that there is a causal relation

between the static Coulomb stress perturbations

resulting from earthquake slip and the triggering of

subsequent events. Studies on earthquake triggering up

to now primarily concern main shock–main shock and

main shock–aftershock interaction through the calcu-

lations of Coulomb stress changes. The present study


123

evidenced the triggering in foreshock–main shock case.

The short-term influence of stress perturbations caused

by the strong foreshocks in three cases (1904, 1928 and

1978 seismic excitations) is obvious. Investigation was

performed aiming in understanding the interaction of

stresses and earthquakes in the study area by resolving

the stress tensors for preferred rupture type following

Deng and Sykes (1997). This is important in the for-

ward problem, where a ‘‘predictive’’ Coulomb stress

map might look quite different depending on assump-

tions about the orientation of the active structures.

Additionally, these maps are more reliable since they

comprise the tectonic loading, which in timescales of

decades play a major role in earthquake occurrence.

The stress field was calculated since 1904 for our

study area and shown in chronological order. The re-

sults shed light to the important topic of temporal

clustering and spatial migration of large events, sug-

gesting that stress interactions in this fault network had

an important role in determining which faults would

fail in the sequence of events. In our study area, al-

though the known active faults exhibit long recurrence

intervals, a burst in strong earthquake occurrence took

place in particular during 1904–1932. Previous investi-

gations for the explanation of triggered earthquakes

concern among others the rate change and spatial

distribution of large aftershocks that occurred outside

the classical aftershock zones (Parsons 2002). From

calculations of shear stress changes on subsequent

earthquake rupture planes it was derived that about

61% of earthquakes that occurred near the M ‡ 7.0

shocks are associated with calculated shear stress in-

creases, while ~39%, are associated with shear stress

decreases, the cutoff defined at |Ds| ‡ 0.001 MPa. The

triggered earthquakes obey an Omori law rate decay

that lasts between ~7–11 years after the main shock. In

our case, the locations of the next seismic excitation is

relatively far from the previous one and therefore

similar calculations showed negligible shear stress

changes on their rupture planes due to the previous

dislocations. Nevertheless, the Coulomb stress instead

of the shear stress changes, when incorporate tectonic

loading, give plausible explanation for the consequent

occurrences.

The most important issue addressed, the ‘‘strong

event pairs’’ that seems to be a quite characteristic

behavior in the study area, is perfectly explained by

stress transfer between adjacent fault segments. Laying

along strike these fault segments, the foreshock

occurrence triggered the occurrence of the consecutive

main shock. Immediately after the first strong event,

calculations of DCFF might have been useful in sug-

gesting the potential locations of their strong counter-

parts. This would have required a very quick

determination of the first event geometry and mecha-

nisms in the sequence of events; a goal now made

possible by modern seismograph networks both

regionally and worldwide. Also needed would be a

detailed knowledge of nearby faults large enough to

generate significant earthquakes.

Acknowledgments The stress tensors were calculated using theDIS3D code of S. Dunbar, which later improved by Erikson(1986) and the expressions of G. Converse. The GMT system(Wessel and Smith 1998) was used to plot the figures. The paperhas been greatly benefited by the revision of Martin Brady andan anonymous reviewer and the editorial assistance of Prof.Wolf-Christian Dullo. This study was supported by the bilateralresearch project between Greece and Bulgaria EPAN-M.4.3.6.1and BG-9/05. Geophysics Department contribution 681.

References

Alexiev G, Georgiev Tz (1996) Geodynamic problems of theKraishte–Sredna Gora morphostructural zone. Probl GeogrBulg Acad Sci 4:131–140

Alexiev G, Georgiev Tz (2002) Quaternary and recent geody-namic deformations of the territory of Bulgaria. ProblGeogr Bulg Acad Sci 1–2:30–48

Ambraseys NN (2001) The Kresna earthquake of 1904 inBulgaria. Annali Geof 44:95–117

Bonchev S, Bakalov P (1928) Les treblements de terre dans laBulgarie du sud les 14 et 18 avril 1928. Rev Vulg Geol Soc1(2):51–63

Burchfiel CB, Nakov R, Tzankov T, Royden LH (2000)Cenozoic extension in Bulgaria and northern Greece: thenorthern part of the Aegean extensional regime. Geol SocLondon Special publication(173):325–352

Christoskov L (2000) Energy and source parameters of thestrong Bulgarian earthquakes after 1900. Rep Geod, War-saw Univ technol 3(48):15–20

Christoskov L, Grigorova E (1968) Energetic and space–timecharacteristics of the destructive earthquakes in Bulgariaafter 1900. Bull Inst Geoph XII:9–107

Deng J, Sykes LR (1997) Evolution of the stress field in southernCalifornia and triggering of moderate-size earthquakes: a200 year perspective. J Geophys Res 102:9859–9886

Dimitrov D, Ruegg J-C (1994) The 1928 Bulgarian earthquakes:fault geometry from geodetic data and modeling. Firstinternational symposium on deformations in Turkey, Istan-bul, pp921–932

Dimitrov DS, Chabalier J-B, Ruegg J-C, Armijo R, Meyer B,Botev E (2006) The 1928 plovdiv sequence (Bulgaria): faultmodel constrained from geodetic data and surface breaks(submitted manuscript)

Dineva S, Sokerova D, Michailov D (1998) Seismicity of south–western Bulgaria and border regions. J Geodyn 26:309–325

Dineva S., Battlo J., Mihaylov D, van Eck T (2002) Sourceparameters of four strong earthquakes in Bulgaria andPortugal at the beginning of the twentieth century. J Seismol6:99–123

Direction for Support and reconstruction of the area damaged bythe 1928 earthquakes (DIPOZE) (1931) Report on theactivities undertaken from April 25, 1928 until November 1,1931 (in Bulgarian). State Press, Sofia, 421 pp


123

Erikson L (1986) User’s manual for DIS3D: a three-dimensionaldislocation program with applications to faulting in theearth. Masters Thesis, Stanford University, Stanford,167 pp

Ganas A, Shanov S, Drakatos G, Dobrev N, Sboras S, Tsimi Ch,Frangov G, Pavlides S (2005) Active fault segmentation insouthwest Bulgaria and Coulomb stress triggering of the1904 earthquake sequence. J Geodynamics 40:316–333

Glavcheva R (1984) Characteristic of the destructive earthquakeof April 18, 1928 (M = 7.0) in southern Bulgaria. Geophys JBulg Acad Sci XVI:38–44

Harris RA (1998) Introduction to special section: stress triggers,stress shadows, and implications for seismic hazard. JGeophys Res 103:24347–24358

Harris RA, Simpson RW (1993) In the shadow of 1857: anevaluation of the static stress changes generated by theM8 Ft. Tejon, California, earthquake. Eos Trans. AGU74(43) Fall Meet, Supplementary 427

Harris RA, Simpson RW (1996) In the shadow of 1857: Theeffect of the great Ft. Tejon earthquake on subsequentearthquakes in southern California. Geophys Res Letts23:229–232

Jackson JA, McKenzie DP (1988) The relationship betweenplate motions and seismic tensors, and the rate of activedeformation in the mediterranean and middle east. GeophysJ Int 93:45–73

Karakostas VG, Papadimitriou EE, Gospodinov D, Ranguelov B(2006) Slip distribution of the 1928 Chirpan and Plovdivmain shocks and earthquake triggering. Sixth internationalconference of SGEM, 119–127

King GCP, Stein RS, Lin J (1994a) Static stress changes and thetriggering of earthquakes. Bull Seism Soc Am 84:935–953

King G, Oppenheimer D, Amelung F (1994b) Block versuscontinuum deformation in the western US, earth planet. SciLett 128:55–64

Kotzev V, Nakov R, Burchfiel BC, King R, Reilinger R (2001)GPS study of active tectonics in Bulgaria: results from 1996to 1998. J Geodynamics 31:189–200

Kotzev V, Nakov R, Georgiev Tz, Burchfiel BC, King RW(2006) Crustal motion and strain accumulation in westernBulgaria. Tectonophysics 413:127–145

Matova M, Spiridonov H, Rangelov B, Petrov P (1996) Majoractive faults in Bulgaria. J Earthq Pred Res 5:436–439

McClusky S, Balassanian S, Barka A, Demir C, Georgiev I,Hamburger M, Hurst K, Kahle H, Kastens K, Kekelidze G,King R, Kotzev V, Lenk O, Mahmoud S, Mishin A,Nadariya M, Ouzounis A, Paradisis D, Peter Y, Prilepi M,Reilinger R, Sanli I, Seeger H, Tealeb A, Toksoz MN, VeisG (2000) GPS constraints on crustal movements anddeformations in the eastern mediterranean (1988–1997):implications for plate dynamics.J Geophys Res 105:5695–5719

McKenzie DP (1972) Active tectonics of the mediterraneanregion. Geophys J R Astron Soc 30:109–185

Messini AD, Papadimitriou EE, Karakostas VG, Baskoutas I(2005) Stress triggering in thrust and strike slip earthquakesalong the Hellenic Arc. Workshop on fracture dynamics:theory and applications to earthquakes. In: Honor ofProfessor Udias, Madrid, 26–28 September 2005 (abstract)

Meyer B, Armijo R, Dimitrov D (2002) Active faulting in SWBulgaria: possible surface rupture of the 1904 Strumaearthquakes. Geophys J Int 148:246–255

Okada Y (1992) Internal deformation due to shear and tensilefaults in a half-space. Bul Seism Soc Am 82:1018–1040

Nakov R, Burchfiel BC, Tzankov Tz, Royden LH (2001) Latemiocene to recent sedimentary basins in Bulgaria. Geol SocAm Map Chart Ser MCH088F

Nalbant SS, Hubert A, King GCP (1998) Stress couplingbetween earthquakes in northwestern Turkey and the northAegean Sea. J Geophys Res 103:24469–24486

Pacheco JF, Sykes LR (1992) Seismic moment catalog of largeshallow earthquakes, 1900 to 1989. Bull Seismol Soc Am82:1306–1349

Papadimitriou EE (2002) Mode of strong earthquake recurrencein central Ionian Islands (Greece). Possible triggering due toCoulomb stress changes generated by the occurrence ofprevious strong shocks. Bull Seismol Soc Am 92:3293–3308

Papadimitriou EE, Sykes LR (2001) Evolution of the stress fieldin the Northern Aegean Sea (Greece). Geophys J Int146:747–759

Papadimitriou EE, Karakostas VG (2003) Episodic occurrenceof strong (Mw ‡ 6.2) earthquakes in Thessalia area (centralGreece). Earth Planet Sci Lett 215:395–409

Papadimitriou EE, Sourlas G, Karakostas VG (2005) Seismicityvariations in southern Aegean, Greece, before and after thelarge (Mw7.7) 1956 Amorgos earthquake due to the evolvingstress. Pure Appl Geophys 162:783–804

Papazachos B, Papazachou C (2002) The earthquakes of Greece.Ziti Publications, Thessaloniki, pp 317

Papazachos BC, Papaioannou ChA, Papazachos CB, SavvaidisAS (1997) Atlas of Isoseismal maps for strong shallowearthquakes in Greece and the surrounding area (426BC–1995). Publ Geophys Lab, University of Thessalonoki, 4, 194pp

Papazachos BC, Papadimitriou EE, Kiratzi AA, Papazachos CB,Louvari EK (1998) Fault plane solutions in the Aegean Seaand the surrounding area and their tectonic implications.Boll Geof Teor Appl 39:199–218

Papazachos BC, Mountrakis DM, Papazachos CB, Tranos MD,Karakaisis GF, Savvaidis AS (2001) The faults that causedthe known strong earthquakes in Greece and surroundingareas during fifth century B. C. up to present. SecondConference Earthq Enging and Engin Seism 28–30 Septem-ber 2001, Thessaloniki, 1, 17–26

Papazachos BC, Scordilis EM, Panagiotopoulos DG, PapazachosCB, Karakaisis, GF (2004) Global relations between seismicfault parameters and moment magnitude of earthquakes.Tenth Congr Hellenic Geol Soc, Thessaloniki, 14–17, April2004, 539–540

Parsons T (2002) Global Omori law decay of triggered earth-quakes: Large aftershocks outside the classical aftershockzone. J Geophys Res 107. DOI 10.1029/2001JB000646

Paskaleva I, Ranguelov B, Simeonova S, Guergiev G (1986) Anapproach about the seismic hazards investigation in theGorna Oriahovitsa seismic zone for the design of reliablestructures. J Develop Urban Areas 6:34–43 (in Bulgarian)

Pavlides SB, Tranos MD (1991) Structural characteristics of twostrong earthquakes in the North Aegean: Ierissos (1932) andAgios Efstratios (1968). J Struct Geol 13:205–214

Perfettini H, Stein RS, Simpson R, Cocco M (1999) Stresstransfer by the 1988–1989 M = 5.3 and 5.4 Lake Elsmanforeshocks to the Loma Prieta fault: unclamping at the siteof peak mainshock slip. J Geophys Res 104:20169–20182

Ranguelov B, Rizhikova S, Toteva T (2001) Some data for themagnitude reevaluation of the strong earthquakes during1904 in Kresna–Kroupnik zone (SW Bulgaria). The Earth-quake (M7.8) Source Zone (Southwest Bulgaria), Professor:Drinov M Academic Publishing House, 266–269

Reilinger RE, McClusky SC, Oral MB, King RW, Toksoz MN,Barka AA, Kinik I, Lenk O, Sanli I (1997) Globalpositioning system measurements of present-day crustalmovements in the Arabia-Africa-Eurasia plate collisionzone. J Geophys Res 102:9983–9999


123

Scholz C (1990) The mechanics of earthquakes and faulting.Cambridge University, Cambridge 439pp

Shanov S, Boykova A, Stoev D (1999) Geophysical model of thecoseismic rupture from the April 4, 1904, Krupnik earth-quake, M = 7.8, SW Bulgaria. In: Second Balkan Geophys-ical Congress, Istanbul, Abstracts 232–233

Shebalin N. (editor), 1974) Catalogue of earthquakes, part III:atlas of isoseismal maps, UNDP–UNESCO survey Seismic-ity of the Balkan region. REM/70/172, Skopje

Sokerova D, Velichkova S, Dineva S (1989) Development of theKresna earthquake source in SW Bulgaria. Proceedings 4thInternational Symposium Seismicity and Seismic Risk.Bechyne, Czech., 93–100, September 4–9, 1989

Soufleris C, Stewart GS (1981) A source study of the Thessalo-niki (Northern Greece) 1978 earthquake sequence. GeophysJ R Astron Soc 67:343–358

Stein RS (1999) The role of stress transfer in earthquakeoccurrence. Nature 402:605–609

Tranos MD, Papadimitriou EE, Kilias AA (2003) Thessaloniki–Gerakarou Fault Zone (TGFZ): the western extension ofthe 1978 Thessaloniki earthquake (Northern Greece) andseismic hazard assessment. J Struct Geol 25:2109–2123

Tranos MD, Karakostas VG, Papadimitriou EE, Kachev V,Ranguelov B, Gospodinov D (2006) Major active faults ofsouthwest Bulgaria: implications of their geometry, kine-matics and the regional active stress regime. Spec Publ GeolSoc Lond 260:671–687

Tzankov TZ, Angelova D, Nakov R, Burchfiel BC, Royden LH(1996) The sub-Balkan graben system of central Bulgaria.Basin Res 8:125–142

Van Eck R, Stoyanov T (1996) Seismotectonics and seismichazard modeling for southern Bulgaria. Tectonophysics262:310–332

Vanneste K, Radulov A, DeMartini P, Nikolov G, Petermans T,Verbeeck K, Camelbeeck T, Pantosti D, Dimitrov D,Shanov S (2006) Paleoseismologic investigation of the faultrupture of the 14 April 1928 Chirpan earthquake (M6.8)Southern Bulgaria. J Geophys Res 111 B01303. DOI10.1029/2005JB003814

Wells DL, Coppersmith KJ (1994) New empirical relationshipsamong magnitude, rupture length, rupture width, rupturearea, and surface displacement. Bull Seism Soc Am 84:974–1002

Wessel P, Smith WHF (1998) New, improved version of thegeneric mapping tools released. EOS Trans. AGU, 79, 579

Yankov Y (1945) Level changes of the terrain caused by theearthquakes of April 14th and 18th 1928 in south Bulgaria.Ann Centr Mteor Inst 29–31:131–136

Zagorchev I (1992) Neotectonics of the central parts of theBalkan Peninsula: basic features and concepts. Geol Run-dsch 81:635–654


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Static stress changes associated with normal faulting earthquakes in South Balkan area

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