the kordofan earthquakes, central sudan

Journal of African Earth Sciences, Vol. 6, No. 4, pp. 573-581, 1987 0731-7247/87 $3.00 + 0.00 Printed in Great Britain Pergamon Journals Ltd.

The Kordofan earthquakes, central Sudan

R. A. CLARK and S. E. BROWNE*

Department of Earth Sciences, The University, Leeds LS2 9JT, U.K.

(Received for publication 13 October 1986)

Abstract---Geophysical studies have revealed extensive rift systems throughout Central Africa, among them the White Nile Rift running SE-NW across central Sudan. Few earthquakes have been detected teleseismically in a r e a s of Sudan away from the East African or Red Sea Rift systems. Relocations presented here confirm three earthquakes to have occurred at shallow depth at the NW extreme of the White Nile Rift, around Jebel Dumbeir, Kordofan province. Ambiguity in a teleseismic P-wave first-motion solution for the 9 October 1966 0648GMT event is resolved by use of the P/pP/sP relative amplitude method. The fault plane is near-vertical, striking to 005--025 ° with a downthrow to the SE and a minor sinistral strike-slip component. This is consistent with ground deformation and location of an aftershock. Its downthrow direction appears to reflect subsidence of the White Nile Rift and central Sudan. Its strike is common with present-day activity in the Ethiopian Rift, suggesting that a near ESE-WSW stress regime persists across central and southern Sudan.

INTRODUCTION

IN CENTRAL and southern Sudan, large areas of Tertiary- Quaternary sediments northwest of the Ethiopian Plateau have been shown by geophysical methods to overlie an extensive system of elongate sedimentary basins of Cretaceous age or older (Fig. 1). Early interpretations of these features as fault-bounded basins (Mitwalli 1969, Khattab 1975, Strojexport 1977, Ali and Whitely 1981) have been supplemented by regional gravity and seismic surveys, including those undertaken for hydrocarbon exploration (Flege 1982) and by Leeds University (Browne and Fairhead 1983, Bermingham et al. 1983). The axial negative Bouguer anomaly associated with the sediments is now recognised as being superimposed on a broader positive anomaly ascribed to crustal thinning.

In east central Sudan (Fig. 1), Browne et al. (1985) propose a series of sedimentary basins, the White Nile Rift, trending SE from Bara to the White Nile thence south to Adar, where hydrocarbon exploration wells have proved at least 4545 m of sedimentary cover (Nicod 1982, 1983). Crustal extension of some 10 km NE-SW is inferred. Crustal thinning appears confined to the southern parts of the Rift. Browne et al. (1985) have further argued that the White Nile Rift, and also the Blue Nile and Southern Sudan Rifts, may be terminated by a major structural lineament, here termed the Nyala- Khartoum Lineament, an eastward continuation of the Foumban-Bake-Birao shear zone that traverses central Africa (Louis 1970).

The principal areas of seismicity in the Sudan are the extreme northeast and the far south, associated with the Red Sea and East African Rift Systems respectively, and are discussed by Fairhead and Stuart (1982).

Elsewhere, few instrumentally determined hypocentres are known, as supported by a seismograph at Nyala, west Sudan, which recorded no local earthquakes (other than some associated with the Jebel Marra vol-

* Present address: BP Petroleum Development Ltd, Britannic House, Moor Lane, London EC2Y 9BU, U.K.

canic province) during two years of operation from April 1980 (Browne eta l . 1985). Ambreseys and Adams (1986) review historic seismicity. They note that earthquakes have been felt at settlements along the Nile Valley (although it is not clear whether these were local or larger distant events), and have located two large earthquakes approximately (to +7 ° in latitude and longitude) using the early global network of Milne seismographs. Presumed surface wave records gave 15°N 35°E, Ms = 5.3, for the 25 Jan 1908 earthquake and 10°N 27°E, Ms = 5.5, for that of 30 May 1910 (Ms = surface wave magnitude). These lie in northeast Sudan, close to the Ethiopian border, and in the southwest within the South- ern Sudan Rift.

This paper describes relocated epicentres for three recent events (Table 1) in the Umm Ruwaba/Jebel Dumbeir region of Kordofan Province, central Sudan, and a short-period P wave focal mechanism solution for the largest, to examine them in relation to the newly- recognised tectonic environment (Browne et al. 1985). Two of the three are in the sequence described by Qureshi and Sadig (1967), near J. Dumbeir, at 0648GMT (mb 5.1) and 1028GMT (mb 4.1) on 9 October 1966. Damage reports imply Mercalli intensities of IV- VI at most locations. A maximum of VIII was estimated on the basis of water ingress to dry wells. From series of en echelon tension gashes found in clayey superficial deposits and tentative isolines of intensity, sinistral strike-slip motion striking N020°E was inferred. Smaller earthquakes were occurring until April 1967 at least. The ISC epicentre of the third earthquake, at 0731 GMT on 17 March 1974, mb 4.6, falls some 100 km north of the J. Dumbeir region.

Sykes (1970) included the 1966 earthquakes in relocations of Indian Ocean earthquakes, but presented the revised epicentres only in map form. Focal depths were 'less than 70 km'. Shudofsky (1985) included the 0648GMT 9 October 1966 event in a catalogue of East African earthquakes, but found the teleseismic Rayleigh and long-period P phases inadequate for use in a focal mechanism solution.

573

22"

(a)

° ,mmbm

I ""

•

4D

/

Jebe

l M

ar

~ B

...~

e~'*

" Bm

rm..'

ll

f,,Q

m

-

. /

~,¢.

¢"

• •

p°x

• to

wn

.m

~w

~te

rco

urs

e

• •

•

" r.

"lff

~r

~ ,

p.7

•

ghar

toum

# /

(b)

L~

-.d

~q

©

~n

km

| I

I I

I I

0 50

0

Fig.

1.

(a)

Gen

eral

loc

atio

n m

ap o

f th

e Su

dan

show

ing

volc

anic

and

int

rusi

ve c

entr

es (

soli

d ar

eas)

and

the

Cen

ozoi

c se

dim

ents

mar

king

the

Rif

t sy

stem

s (s

tipp

le).

(b)

IS

C a

nd r

evis

ed e

pice

ntre

s fo

r th

e th

ree

eart

hqua

kes

stud

ied

here

, in

re

lati

on to

mac

rose

ism

ic d

ata

and

geol

ogy.

Dat

es/t

imes

of t

he e

arth

quak

es a

re g

iven

adj

acen

t to

the

rel

ocat

ed e

pice

ntre

.

The Kordofan earthquakes, central Sudan 575

Table 1. Hypocentres and origin times for the three Jebel Dumbeir earthquakes, as published by the ISC and recomputed here using the JED method (see text)

JED JED ISC (aftershocks only) (synthetic master)

9 October 1966 (mainshock)

Magnitude 5.1 No. of stations 127 Latitude 12.6300N + 3 km Longitude 30.750°E + 4 km Depth 0 + 21 km Origin time 06-48-39 + 3.5 s Shift

9 October 1966 (aftershock)

Magnitude 4.1 No. of stations 33 Latitude 12.660°N +- 6 km Longitude 30.9100E + 8 km Depth 50 + 19 km Origin time 10-28-28 + 1.8 s Shift

17 March 1974 Magnitude 4.6 No. of stations 39 Latitude 13.320°N + 5 km Longitude 30.880"E + 7 km Depth 33 km* Origin time 07-31-25.8 + 0.4 s Shift

B

B

m

12.637°N + 5 km 30.834°E + 5 km 0km* 06-48-39.1 _+ 0.3s 9 km/085 °

12.645°I'4 + 15 km 12.750"N + 7 km 31.041°N + 24km 30.962°E + 8km 78 + 47 km 37 + 14 km 10-28-31.3 + 4.5 s 10-28-27.8 + 1.6 s 14 km/096 ° 12 km/031 °

13.392°N + 19km 13.660°N + 10kin 30.986°E + 26 km 31.286°E + 10 km 22 + 49 km 0 km* 07-31-25.5 + 6.1 s 07-31-25.1 + 0.7 s 14 km/055 ° 58 kin/050 °

Latitude and longitude are in degrees, depth in kilometres. Asterisks indicate constrained parameters. Following normal practice at the ISC and elsewhere, depths were fixed at zero if they became negative during iteration. Shifts of JED solutions are quoted as a distance and bearing relative to the original ISC location. Uncertainty bounds are 95% confidence limits for the JED results or standard errors for the ISC solutions, and are given in kilometres for depth, latitude and longitude, and in seconds for the origin time. Magnitude is body-wave magnitude as quoted by the ISC.

EPICENTRE LOCATIONS

The ISC hypocentres of the three White Nile Rift earthquakes are given in Table 1. Figure 1 shows them in their geological setting, together with the macroseismic results of Qureshi and Sadig (1967). The ISC solution for the 1966 main shock lies just a few kilometres west of the ground rupture, both being to the west of J. Dumbeir itself.

These have been refined by application of the Joint Epicentre Determination method (JED) of Douglas (1967), whereby locations are made relative to a 'master event ' whose latitude, longitude, origin time and depth are fixed. Some of these parameters may be fixed for the remaining events also. The unrestrained hypocentre parameters, and also travel-time residuals at the re- ceivers, are solved for in a least-squares sense. This performs the location independently of systematic errors induced by receiver correction terms.

First arrival times were taken from ISC Bulletins, excluding any with residuals greater than 5 s as probable mis-picks, and all P-diffracted readings due to their uncertain travel-time curve. Azimuthal distribution is uneven- -no more than 11 stations lie to the south and west - - though there is a wide range of epicentral distances in the remaining coverage.

Revised absolute locations for all three events of

Table I were found by taking as the fixed master event a set of synthetic Jeffreys-Bullen travel times, from the ISC location of the 0648GMT 9 October 1966 event to all stations observing any of the three. Station corrections, after Lilwall and Douglas (1970) and Poupinet (1979), were applied to remove receiver correction errors.

The solutions show all three events falling on a roughly common trend (Fig. 1, Table 1). The main shock of 9 October 1966 now lies at about 12.64°N 30.83°E, closer to the observed ground rupture and more centrally within the isoseismal lines. Its aftershock lies along the trend of the surface breaks, at 12.75°N 30.96°E, as does the 1974 earthquake, this having been shifted some 60 km from the ISC location to 13.66°N 31.29°E. Almost all of the reporting stations lie in the azimuth of the relocation vectors, i.e. to the north and east: of the few to the southwest, several (e.g. KIC/LIC, BUL) have the marked travel-time residuals typically associated with cratonic environments.

An estimate of the depth of the 0648GMT 9 October 1966 main shock has been obtained by repeating the JED procedure with this as the 'master ' event, but with a range of constrained depths, from zero to 40 km in steps of 5 km. When the depth is closest to its actual value, the solution precision should be greatest and the sum of squared residuals a minimum (cf. Mitchell and Landisman 1969, solving for seismograph calibration

576 R . A . CLARK and S. E. BROWNE

f u ~ ' - - V K A = ~ : : - = : - : : ~_ z '

. . . . CLL_ _ _ . . . . " . . . . . .

M O X

STU

UME

J E R

+

Q U E

L I C

L o w e r - h e m i s p h e r e e q u a l - a r e a p r o j e c t i o n

T a k e o f f angles assume source layer v e l o c i t y 6 . 2 k m l e e c

Record length Is 25a e x c e p t V K A , C L L , A L E ,

$ C H , a n d 8 H I where minute marks are Included

B U L K O D

Fig. 2. Short-period vertical seismograms of the 0648GMT 9 October 1966 earthquake used in focal mechanism determination. Compressionai ground motion is recorded as upward movement at all stations except VKA. The

non-WWSSN instruments have a generally higher gain at frequencies above 1 Hz than the WWSSN seismometers.

constants). By this criterion, the optimum depth for the main shock is 5-10 km. The solution degrades slightly for zero depth and markedly for 15 km and deeper.

With just three earthquakes, discussion of trends is rather conjectural but it is tempting to note that all three revised locations lie roughly on the trend of the Nyala- Khartoum lineament. There is also now internal consist- ency of epicentres for the 1966 events, macroseismic data for the main shock, and the location of the aftershock relative to it. This suggests any systematic bias in the JED procedure due to azimuthal variation of source correction is small (as also concluded by Shudofsky 1985). The new results are thus preferred to the ISC locations.

FOCAL MECHANISMS USING ONSET POLARITY

Twenty-one short-period vertical component P-wave seismograms (15 WWSSN and six others; Fig. 2) were collected for the 0648GMT 9 October 1966 event. No records having adequate signal-to-noise ratios were obtained for its aftershock or the 17 March 1974 earthquake. The majority of the stations are in Europe and Asia, and are thus concentrated in the north and east quadrants of the focal sphere: some, in North America, are almost antipodal. Three African stations lie to the south and west. At most, the energy is seen to be in the first 5 s or so of the signal. There is also sometimes a

prominent arrival at about 15 s after P (e.g. at SHL). Several close stations, such as JER (epicentral distance 19.49°), exhibit triplication. Signal-to-noise ratios at a few stations (e.g. SDB, SHL) are good but as low as 1.5-2.0 elsewhere.

The 'first-motion' technique of focal mechanism solution (Byerley 1926, Stauder 1962) projects the polarity of the first arrival at teleseismic stations onto the focal sphere, the source layer P velocity being used to fix take-off angles; 6.2 km s -1 was assumed here. The suite of polarities reported through the ISC failed to yield a coherent solution. After comparison of waveforms at all stations, 12 polarities were picked with confidence. These few do constrain the style of faulting. Consistent solutions (Fig. 3a) fall into two groups, which accommo- date sinistral strike-slip motion (Qureshi and Sadig 1967) on a sub-vertical fault plane only if it strikes more east-west (060-080 ° ) than the ground deformation. The second domain of solutions shows normal faulting, striking to 010-020 ° (the trend observed in ground rupture), with a minor strike-slip component of ambiguous sense.

FOCAL MECHANISM USING THE RELATIVE AMPLITUDE METHOD

The first-motion technique cannot be applied univer- sally. A body-wave magnitude in excess of 5.5 is typically required for detection at a suite of stations suitably

N

N

(A)

x Q

UE

x N

DI

x 5H

L

x C

HG

co

mp

res

sio

ne

i x AA

E

BUL

(B)

dil

ata

tlo

na

l

x L

IC

x

x $C

H

?Y

llL

8°,

(a)

x S

HI

x D

UE

× N

DI

x $H

L

x CIS~i P

O0

y KOD

x AA

E

co

mp

res

sio

na

l

(b

0 0 ('3

~-t

r~

ero

up

t

N

|9-

str

lke

-sn

p o

n ¢.

TO

dep

. di

p to

NW

/NE

ero

vp

2 N

|O

dego

NN

E-8

8W

no

rmal

fsu

lt

wit

h m

inor

str

ike-

slip

co

mpo

nent

low

er

he

mis

ph

ere

e

qu

al-

are

a

pro

jec

tio

ns

co

mp

res

sio

na

l fi

eld

s

ha

de

d

on

in

se

t

cro

ss

es

-

sta

tio

n

loc

ati

on

s

sq

ua

res

-

nu

ll

ve

cto

rs

Fig.

3.

(A)

Ster

eogr

aphi

c re

pres

enta

tion

s of

the

foc

al m

echa

nism

s co

mpa

tibl

e w

ith

the

firs

t-m

otio

n da

ta o

nly.

The

y w

ere

dete

rmin

ed u

sing

the

FA

LT

pro

gram

(se

e te

xt)

wit

h a

10 °

grid

on

dip,

str

ike,

and

slip

ang

le a

nd a

ll am

plit

udes

unb

ound

ed.

The

tw

o gr

oups

of s

olut

ions

are

sum

mar

ized

in

cart

oon

form

bel

ow.

(B)

As

(A)

but

for

the

rela

tive

P: p

P:

sP d

ata

and

a 5 °

gr

id.

Gtl

"-

4 -.,

.,I

578 R. A. CLARK and S. E. BROWNE

Table 2. Data used in focal mechanism determination for 9 October 1966 mainshock

Epicentrai distance Azimuth P amplitude/ pP sP

Station (degrees) (from event) polarity amplitude amplitude

AAE 8.65 114 (Triplication. Aftershock P polarity +ve?) JER 19.49 011 (Triplication. P +ve) SHI 26.36 047 (Poor record. P +ve) IST 28.35 357 4-8, +ve 4-10 0-10 SDB 32.23 212 12-18, -ve 6-18 0-20 BUL 32.63 184 3-8, +ve 3-12 3-15 QUE 37.75 057 4-15, +ve 6-20 0-25 STU 40.23 338 1-5, ? 0-5 0-10 POO 41.86 076 0-2, ? 1-5 0-5 KOD 45.81 088 0-3, ? 1-5 0-5 NDI 46.02 062 2-4, +ve 3-7 0-5 KON 49.49 346 -re. Amplitudes unbounded. UME 51.65 354 2-8, +ve 0-5 0-5 SHL 58.75 068 4-12, +ve 10-20 5-12 CHG 65.64 075 1--4, +ve 2-5 0-5

LIC 35.83 263 5-17, ? 5-15 2.5-20 VKA 37.48 344 1-3, ? 1.5--4 1-5 MOX 41.01 341 1-3.5, +ve 2-5.5 0-7.5 CLL 41.18 343 0.5-2, ? 0-1.5 0-5 ALE 77.98 352 0-2, ? 1-3 0-4 SCH 84.08 325 1-3, ? 1-3 0-4

Non-WWSSN stations are grouped separately at the bottom of the table. Amplitudes are in arbitrary units consistent for a given station. Compressional onsets positive, dilatational negative. All pP and sP amplitudes are of uncertain polarity, indicated by ? for P phases.

distributed in azimuth and distance, and for confident recognition of polarities.

Instead, we employ here the relative P/pP/sP amplitude method of Pearce (1977, 1979, 1980), which is appropriate to small earthquakes, and utilises a greater amount of the source information present in a seismo- gram. For crustal earthquakes observed at teleseismic distances, the free-surface reflections pP and sP undergo no differential amplitude losses relative to P except those incurred during their path to and from the focal depth to the free surface, and their reflection from it. Their polarities and relative amplitudes therefore con- tain a great deal more information about the attitude of the double couple source than simple first motions. Even absent depth phases or limiting amplitudes set from coda or noise amplitudes become useful data. Observations from two or three teleseismic stations alone are often sufficient to derive a well-defined focal mechanism solution (e.g. Pearce 1980).

This dataset (Fig. 2) provides 21 stations' P : pP, P : sP and pP : sP amplitude ratios (Table 2). For three close-in stations, a polarity only was distinguishable; amplitude ratios were left unbounded. Eight stations gave measur- able amplitude ratios, but ambiguous polarity. No pP or sP polarities could be determined. 100% confidence limits took full account of noise levels at the appropriate frequency before the first arrival and possible inter- ference from the overshoot of preceding arrivals. Further difficulties with sP are noted below.

To map these data onto the focal sphere, accounting for source-surface amplitude losses and at correct take- off angles, various source region parameters have to be specified. Chosen values were: focal depth, 5 km or P-pP time c a 2 s; source layer velocity 6.2 km s- 1; surface layer velocity 4.0 km s-l; surface layer density 2300 kg

m -1. Predicted amplitude losses from anelastic attenuation on the source-surface path amounted to no more than a few percent for even the lowest realistic Q values. Using the program F A L T (Pearce 1979), the observed amplitude ratios are compared to those computed for all possible orientations of a Savage (1966) double couple source on a 5 ° grid of strike, dip and slip angle; a total of 93,312 possible orientations. Those predicting relative amplitudes within the 100% confidence limits of all data constitute the focal mechanism solution. Uncertainty bounds on the data are thus quantitatively reflected in the finite size of the solution region(s) in orientation space.

Before accepting solutions, all source parameters have been tested in a systematic fashion although Pearce (1984) demonstrates that, for shallow earthquakes at teleseismic distances, relative P/pP/sP amplitudes and their consequences for focal mechanism are only weakly dependent on knowledge of take-off angles and source layer velocity/density structure. 'Correct ' values were deemed to have been found when the greatest compatibility with a double couple solution was manifest as the largest number of solutions. Varying the severity of attenuation effects made little difference to the solutions except when quite unrealistic losses of 20-30% (or aver- age source-surface P-wave Q of less than 10) were invoked, which gave no solutions at all. Source and surface layer velocities, and surface layer density, were varied over the ranges 6.0-6.5, 2.0-5.0 km s -1 and 2000-2700 kg m -1 respectively. A similar or smaller number of fully compatible solutions were usually obtained: the extreme values gave none. Most impor- tantly, the general character of the solutions, in terms of style of faulting, did not change.

Correct identification of depth phases is pivotal in


obtaining a meaningful result from this technique. Focal depths from 5 to 40 km were tested, by picking the signal amplitudes in the appropriate time windows along each record. Forty kilometres corresponds to the deepest plausible crustal earthquake, the shallowest possible after-shock depth (Table 1), and the prominent phase some 15-17 s after P (e.g. SHL; Fig. 2) being pP. Most solutions were found for a depth of 5 km. At greater depths, none at all were obtained; one or two stations at least had to be disregarded to yield any.

The largest number of partially compatible solutions was given by a focal depth of 40 km. In this case, the coda energy before pP (i.e. at least 10-12 s after P) would have to be ascribed to near-source mode conversions (S-to-P at source Moho) and reverberations, and/or a complex source signature. Conversely, for the shallowest depth, an explanation is required for the later arrivals construed as pP for 40 km depth. Their delay relative to P is not incompatible with multiple reflections between Moho and free surface. To predict the approxi- mate amplitudes of these various phases, the modelling method of Hudson (1969a,b) and Douglas et al. (1972) was used with a simple horizontally-layered source crust and the focal mechanism solutions of Fig. 3a. This yielded amplitudes of up to 40% or so of those observed. Less difficulty was experienced in producing the multiple reflections than the S-to-P conversion. Particularly large predicted horizontal component amplitudes suggest that only a small amount of dip would be required for greater similarity. A further possibility, for either the 5 or 40 km case, is that the earthquake has a complex source signature and/or comprises several discrete events; a large number of disturbances were felt locally (Qureshi and Sadig 1967). Overall, the shallow (ca 5 km) depth is accepted because of: (1) the appearance of surface deformation (arguing against any great depth for such a relatively small earthquake); (2) the reasonable compatibility of synthetic and observed phases; and (3) the preference for that depth in the FALT focal mechanism solution and both the ISC and JED hypocentre solutions.

Using all P, pP and sP data, just 25 fully compatible solutions, or 0.006% of orientation space, were obtained. Inclusion of sP in fact contributed nothing to the solution because of its large confidence limits, required due to its greater inherent uncertainty from the variable effects of near-surface layers (e.g. Pooley et al. 1983) and, for some crustal models, the near-coincident travel time of sP and a Moho S-to-P conversion.

Compatible solutions are illustrated in Fig. 3b. The relative amplitude data reject the pure strike-slip solutions in Fig. 3a, and resolve ambiguity in the sense of the small strike-slip component of the normal faults. Acceptable nodal planes are (a) near-vertical striking to 005-025 °, and (b) sub-horizontal with dips of less than 10 ° to the NNE. In favour of nodal plane (b), it might be noted that a basement fault of the appropriate strike (though unknown dip) has been mapped (Fig. 1). Alter- natively, nodal plane (a), the revised epicentre (Table 1), and the trend of ground displacement (Fig. 1)

are mutually consistent. The senses of strike-slip motion agree, although normal faulting dominates the earthquake mechanism. No such compatibility can be claimed for nodal plane (b). It is by no means necessary that the displacement of superficial deposits mimic exactly that of the earthquake, due to their lower competence and refraction of the stress field near the free surface (e.g. Yielding et al. 1981), especially where the epicentre lies in a varied surface geology and topographic relief. Note that the surface breaks occur along the margin of the Jebel. The near-vertical nodal plane is thus preferred as the fault plane.

DISCUSSION

The J. Dumbeir earthquakes were either a response to localised stresses, or a localised response to regional stresses. By 'local' we mean an area encompassing the epicentres, for it seems improbable that three earthquakes isolated in the middle of a vast, almost entirely aseismic region could be unrelated. 'Regional' stresses in this context mean those apparent over many hundreds of kilometres.

There are two structural features in central Sudan to which the J. Dumbeir earthquakes could be related: the White Nile Rift and the Nyala-Khartoum Lineament (Fig. 1). Considering the supposed shear nature and NE-SW direction of the lineament, which contrast with the normal faulting and 005-025°E strike determined for the mainshock, it is unlikely that the earthquakes are directly related to movement along the lineament. Nevertheless, the J. Dumbeir earthquakes and those of 1908 and 1910 recognised by Ambreseys and Adams (1986) are SE of, and roughly collinear with, the Nyala- Khartoum lineament; possibly it still defines the north- westerly limit of continuing subsidence in the White Nile Rift and elsewhere.

The 1974 earthquake is thought to have occurred at shallow depth either within the consolidated sediments in the Bara basin (maximum depth estimated by Browne et al. (1985) from gravity modelling to be 1.5-2.0 km) or in the underlying basement. The present seismic data have been insufficient to determine a fault-plane solution and no ground deformation in the loose, surficial sands has been recorded which might point to the faulting mechanism. The 1966 earthquakes, however, took place in shallow basement, at the edge of the White Nile Rift, rupturing the clay surface. The derived fault-plane for the 0648GMT mainshock strikes almost perpendicu- lar to the rift axis (Fig. 4) and is more likely to be the result of adjustment to stresses set up by subsidence in the rift basin, rather than primary, extensional rifting.

North-northeast-south-southwest and NE-SW lin- eaments and faulting, presumably basement-controlled, have been noted in both the White Nile and Southern Sudan Rifts (R. F. Flege, Chevron Overseas Petroleum, pers. comm., Browne et al. 1985). The strike of contemporary faulting may, therefore, be strongly influenced by pre-disposing mechanical factors and may not indi-

580 R . A . CLARK and S. E. BROWNE

' ' ' g o . ' ' ' ' & -

1908 5" 0

1974

966 ! /

/

1 9 1 0 t

30" o ! ! I

I I

i - . _ ;

"4 Sudan "~

C . A . R . ~ ~ ""

5*

" ,....' .,,i,-7 . . . . --- . . .-. .-.- i/. .... . .... , ..,

Zalre ~ l ~ ganda \~

• I \ '.\

v / , ' ? " . ¢ : ~ i , ~ , , , ' . ° I I ~

Fig. 4. Focal mechanisms of the Kordofan earthquakes studied here, together with others in central East Africa compiled from Fairhead and Stuart (1982) and Shudofsky (1985) and the two new epicentres recognised by Ambreseys and Adams

(1986).

cate the causative stresses in a simple fashion. Browne (1984) has suggested that the approximately N-S fault trends interpreted from Bouguer anomalies may be reflecting the westward continuation of the crustal fabric which in East Africa Vail (1978) associated with the Mozambique Belt tectono-metamorphic event (Ken- nedy 1964).

It is interesting to view the earthquakes in central Sudan in the context of East African tectonics (Fig. 4). Let us assume that the pre-existing structure modifies but does not govern the fault plane strike. From the orientation of the 1966 fault-plane, it is unlikely the underlying stresses were associated with the SW-NE crustal extension in the Red Sea. The strike is, however, remarkably similar to the 019°E trend of faulting and rifting in Ethiopia, which Mohr (1983) believed to be indicative of the contemporary regional stress field. A common extension direction prevails in both earthquake focal mechanisms (Shudofsky 1985) and fault geometries (Mougenot et al. 1986) throughout East Africa and Madagascar; the seismological evidence here suggests this stress orientation persists as far north as central Sudan.

Acknowledgements--The JED and FALT software were provided by

John Young of MOD(PE) Blacknest and Bob Pearce of University College, Cardiff, respectively. Enlightening discussions were had with Mike Daly, Graham Stuart and Derek Fairhead. One of us (SEB) acknowledges financial support from the U.K. Overseas Development Agency.

REFERENCES

Ali, H. D. and Whitely, R. J. 1981. Gravity exploration for groundwater in the Bara basin, Sudan. Geoexploration 19, 127-141.

Ambraseys, N. N. and Adams, R. D. 1986. Seismicity of the Sudan. Bull. Seism. Soc. Am. 76, 483--493.

Bermingham, P. M., Fairhead, J. D. and Stuart, G. W. 1983. Gravity study of the Central African Rift System: a model of continental disruption. 2. The Darfur domal uplift and associated Cenozoic volcanism. Tectonophys. 94, 205-222.

Browne, S. E. 1984. Gravity studies in the Sudan: a tectonic interpretation of some rifted sedimentary basins. Unpubl. Ph.D. thesis, University of Leeds.

Browne, S. E. and Fairhead, J. D. 1983. Gravity study of the Central African Rift: a model of continental disruption. 1. The Ngouandere and Abu Gabra Rifts. Tectonophys. 94,187-203.

Browne, S. E., Fairhead, J. D. and Mohammed, I. I. 1985. Gravity study of the White Nile Rift, Sudan, and its regional tectonic setting. Tectonophys. 113, 123-137.

Byerley, P. 1926. The Montana earthquake of June 28 1925 GCMT. Bull. Seism. Soc. Am. 16,209-265.

Douglas, A. 1967. Joint epicentre determination. Nature 215, 47-48.


Douglas, A., Hudson, J. A. and Blarney, C. 1972. A quantitative evaluation of seismic signals at teleseismic distances---III. Com- puted P and Rayleigh wave seismograms. Geophys. J. R. astr. Soc. 28,385-410.

Fairhead, J. D. and Stuart, G. W. 1982. The seismicity of the East African Rift System and comparison with other continental rifts. In: Continental and Oceanic Rifts (Edited by Palamason, G.), pp. 41- 61. Geodynam. Set. Am. Geophys. Union, 8.

Flege, R. F. 1982. Regional gravity and magnetics in the Sudan. Application to petroleum exploration. (Abstract). 52nd Annual Int. SEG Meeting, Dallas, Texas.

Hudson, J. A. 1969a. A quantitative evaluation of seismic signals at teleseismic distances--I. Radiation from a point source. Geophys. J. R. astr. Soc. 18,233-249.

Hudson, J. A. 1969b. A quantitative evaluation of seismic signals at teleseismic distances---II. Body waves and surface waves from an extended source. Geophys. J. R. astr. Soc. 18,353-370.

Kennedy, W. Q. 1964. The structural differentiation of Africa in the Pan-African (_+ 500 m.y.) tectonic episode. Eighth Report on scien- tific results, Research Institute of African Geology, University of Leeds.

Khattab, M. M. 1975. Sedimentary basins in northeast Kordofan, Sudan, indicated by a gravity survey. EgyptJ. Geol. 19, 77-85.

Lilwall, R. C. and Douglas, A. 1970. Estimation of P-wave travel times using the joint epicentre method. Geophys. J. R. astr. Soc. 19, 165-181.

Louis, P. 1970. Contribution g6ophysique ~ la conaissance g6ologique du basin du Lac Tchad. ORSTOM Memoir 42, Paris.

Mitchell, B. J. and Landisman, M. 1969. Electromagnetic seismograph constants by least-squares inversion. Bull. seism. Soc. Am. 59, 1335-1348.

Mitwalli, M. A. 1969. Interpretation of low gravity anomaly in the north east of Kordofan, West Sudan. Bull. Geophys. Tear. Appl. 11, 119-126.

Mohr, P. 1983. Volcanotectonic aspects of Ethiopian Rift evolution. Bull. Centres Rech. Explor.-Prod. Elf-Aquitaine 7, 175-189.

Mougenot, D., Recq, M., Virlogeux, P. and Leprivier, C. 1986. Eastward extension of the East African Rift. Nature 321,599-603.

Nicod, M. A. 1982. Oil and gas developments in North Africa in 1981. Bull. Am. Ass. Petrol. Geol. 66, 2163-2250.

Nicod, M. A. 1983. Oil and gas developments in North Africa in 1982. Bull. Am. Ass. Petrol. Geol. 67, 1795-1826.

Pearce, R. G. 1977. Fault plane solutions using relative amplitudes of P and pP. Geophys. J. R. astr. Soc. 50,381-394.

Pearce, R. G. 1979. Earthquake focal mechanisms from relative amplitudes of P, pP and sP: method and computer program. AWRE Report No. O41/79. HMSO, London.

Pearce, R. G. 1980. Fault plane solutions using relative amplitudes of P and surface reflections: further studies. Geophys. J. R. astr. Soc. 60,459-487.

Pearce, R. G. 1984. The information content of relative amplitudes observed from an earthquake radiation pattern. AWRE Report No. 082/83. HMSO, London.

Pooley, C. I., Douglas, A. and Pearce, R. G. 1983. The seismic disturbance of 1976 March 20, East Kazakhstan; earthquake or explosions? Geophys. J. R. astr. Soc. 74,621-631.

Poupinet, G. 1979. On the relation between P-wave travel-time residuals and the age of continental plates. Earth Planet. Sci. Lett. 43, 149-161.

Qureshi, I. R. and Sadig, A. A. 1967. Earthquake and associated faulting in Central Sudan. Nature 215,263-265.

Savage, J. C. 1966. Radiation from a realistic model of faulting. Bull. Seism. Soc. Am. 56,577-592.

Shudofsky, G. N. 1985. Source mechanisms and focal depths of East African earthquakes using Rayleigh-wave inversion and body-wave modelling. Geophys. J. R. astr. Soc. 83, 653-714.

Stauder, W. 1962. The focal mechanism of earthquakes. Adv. Geophys. 9, 1-76.

Strojexport. 1977. Geophysical investigations of groundwater struc- tures--central and northern parts of the Upper Nile Province, 5th Stage. Prague.

Sykes, L. R. 1970. Seismicity of the Indian Ocean and a possible nascent island arc between Ceylon and Australia. J. geophys. Res. 75, 5041-5065.

Vail, J. R. 1978. Outline of the geology and mineral resources of the Democratic Republic of the Sudan and adjacent areas. Overseas Geol. Miner. Resour. 49.

Yielding, G., Jackson, J. A., King, G. C. P., Sinvhal, H., Vita- Finzi, C. and Wood, R. M. 1981. Relations between surface deformation, fault geometry, seismicity, and rupture characteristics during the E1-Asnam (Algeria) earthquake of 10 October 1980. Earth Planet. Sci. Left. 56,287-304.

the kordofan earthquakes, central sudan

Documents