Journal of Seismology 5: 329–359, 2001.© 2001 Kluwer Academic Publishers. Printed in the Netherlands.

329

Surface-rupturing history of the Bree fault scarp, Roer Valley graben:Evidence for six events since the late Pleistocene

Kris Vanneste∗,1, Koen Verbeeck1, Thierry Camelbeeck1, Etienne Paulissen2,Mustapha Meghraoui3, François Renardy4, Denis Jongmans4 & Manfred Frechen5

1Royal Observatory of Belgium, Ringlaan 5; B-1180 Brussel, Belgium; 2Lab voor Geomorfologie en RegionaleGeografie, Katholieke Universiteit Leuven, Redingenstraat 16, B-3000 Leuven, Belgium; 5EOST – Institut dePhysique du Globe, 5 rue Rene Descartes, 67084 Strasbourg cedex, France; 4LGIH, University of Liege, Bât.B19, B-4000 Liege, Belgium; 5Centre for Environmental Change & Quaternary Research, GEMRU, Francis CloseHall, Swindon Road, Cheltenham GL50 4AZ, United Kingdom; ∗Corresponding author: tel: +32-2-3730280; fax:+32-2-3730339; e-mail: kris.vanneste@oma.be

Received 28 August 1999; accepted in revised form 25 July 2000

Key words: colluvial wedge, coversand, Feldbiss fault zone, low strain rate, paleoearthquake, periglacial, trenching

Abstract

Since 1996 paleoseismological investigations have been used to develop the surface-rupturing history of the Breefault scarp, the morphologically best-defined segment of the southwestern border fault of the Roer Valley grabenin northeastern Belgium. The first studies determined that the escarpment is associated with a surface fault, andthey exposed evidence for three surface displacements since about 40 ka BP. The most recent event probablyoccurred between 1000 and 1350 yr cal BP. Geophysical and trenching studies at a new site near the southeasternend of the fault scarp reconfirmed the coincidence of the frontal escarpment with a shallow normal fault, whichdisplaces the Middle Pleistocene ‘Main Terrace’ of the Maas River, as well as overlying coversands of Saalianto late Weichselian age. Different amounts of displacement shown by the two youngest coversand units indicatetwo discrete faulting events, but primary evidence for the coseismic nature of these events is sparse. Radiocarbonand optically stimulated luminescence dating constrain the age of these events to the Holocene and between 14.0± 2.3 ka BP and 15.8 ± 2.9 ka BP, respectively. In addition, four older surface-rupturing events are inferredfrom the presence of four wedge-shaped units of reworked Main Terrace deposits that are interbedded with cov-ersand units in the hanging wall of the trench and in shallow boreholes. These wedges are interpreted as colluvialwedges, produced by accelerated slope processes in response to rejuvenation of the fault scarp, most probably ina periglacial environment. Luminescence dating indicates that five out of a total of six identified faulting eventsare younger than 136.6 ± 17.6 ka. The antepenultimate event was the largest faulting event, associated with atotal fault displacement in excess of 1 m. Thus, the newly investigated trench site represents the longest and mostcomplete record of surface rupturing recovered so far along the Bree fault scarp. This study also demonstrates theviability of the paleoseismological approach to identify past large earthquakes in areas of present-day moderate tolow seismic activity.

Introduction and seismotectonic setting

The Roer Valley graben (RVG) defines the maingraben of the Lower Rhine graben system, a series ofNW-SE-trending grabens and horsts, transecting theborderland of Germany, the Netherlands and Belgium(Figure 1). This graben system is part of the Cenozoic

rift system of western and central Europe in the fore-land of the Alpine fold belt and is subject to SW-NEextension (Ahorner, 1975; Ziegler, 1994). Althoughthis area represents one of the seismically most activezones in intraplate northwestern Europe, its present-day seismic activity (Figure 1) must be considered as

330

Figure 1. Seismotectonic map of the Lower Rhine graben area in the border zone of Belgium, Germany and the Netherlands. Map showslocation of major Quaternary faults, instrumental seismicity (from the Royal Observatory of Belgium catalog 1911–1999) and historicallydocumented earthquakes (Alexandre, 1990, 1994). Location of map area in northwestern Europe is shown in inset. BF = Bocholt fault, FF =Feldbiss fault, GBF = Grote Brogel fault, HF = Heerlerheide fault, RF = Reppel fault.

low to moderate. The seismicity appears to be mostlyconcentrated along the known major border faults, in adepth range of 5–20 km; focal mechanisms (Ahorner,1983; Camelbeeck et al., 1994) show predominantlynormal dip-slip along NW-SE-striking fault planes.Since 1755, the RVG has experienced five earthquakeswith magnitude MS > 5.0; the strongest of those oc-curred in 1756 in the vicinity of Düren, Germany andhad an epicentral intensity of I0 = VIII MSK andan estimated magnitude ML = 6.1 (Ahorner, 1994).The largest event recorded during the 20th centurywas the MW = 5.4 (I0 = VII MSK) Roermond, theNetherlands, earthquake in 1992 (Camelbeeck et al.,1994). Whereas the instrumental and historical recordsindicate a log-linear cumulative frequency-magnituderelation up to magnitude ML = 6.0 (corresponding toMS ∼ 5.4) (Ahorner, 1975), the time interval they spanis too short to assess the potential for larger, surface-rupturing earthquakes in a low-strain-rate (< 1 mm/yr)intraplate domain such as the Lower Rhine grabenarea.

In order to identify active, seismogenic faults andto find evidence (or lack thereof) for past coseis-mic surface faulting, we started a paleoseismologicalstudy in northeastern Belgium, where the southwest-ern boundary of the RVG, the Feldbiss fault zone,is located (Figure 1). Compared to the well-definedPeelrand fault (which was the likely source of theRoermond earthquake; Camelbeeck et al., 1994) onthe northeastern side of the graben, the Feldbissfault zone has a more complex geometry (Figure 1).Seismic-reflection and borehole data (Rossa, 1986;Demyttenaere and Laga, 1988) show that, from south-east to northwest, three NW-SE-trending, subparallelfault branches (the Heerlerheide, Geleen and Feld-biss faults) merge into the single Feldbiss fault proper,which has accumulated more than 400 m of throwsince the late Oligocene. Further north, in the vicin-ity of the town of Bree, the Feldbiss fault splits againinto three NNW-SSE trending faults (the Grote Bro-gel, Reppel and Bocholt faults), among which the totaldisplacement is more or less equally divided. The Qua-

331

Figure 2. Geomorphology of the Belgian part of the Roer graben area, showing position of fault scarps with respect to Middle and LatePleistocene alluvial terraces (after Paulissen, 1973 and Paulissen, 1997). Superposed topographic contours are illustrative only. BFS = Breefault scarp, BoFS = Bocholt fault scarp, GBFS = Grote Brogel fault scarp, HFS = Heerlerheide fault scarp, RFS = Reppel fault scarp.

332

ternary activity of the Feldbiss fault zone has sincelong been recognized (e.g. Briquet, 1908; Paulissen,1973) and is attested by the displacement of alluvialterraces of the rivers Maas and Rhine. The Feldbissfault marks the northeastern limit of the Campine Plat-eau (Figure 2), an elevated area covered by the MiddlePleistocene (770–350 ka BP) alluvial-fan complex ofthe Maas River (i.e. the ‘Main Terrace’). This terracedefines a relatively flat deposit over a large area andcan therefore be regarded as a reference level for re-cent tectonic activity in this region. Based on datafrom scattered borings in the Belgian Maas Valley,Beerten et al. (1999a, 1999b) report a total displace-ment of up to 40 m for the base of the Main Terraceacross the entire fault zone.

The recency of activity of the different branches ofthe Feldbiss fault zone is reflected by their surface ex-pression (Figure 2). Several topographic escarpmentshave been identified by Paulissen (1983, 1997), coin-ciding with the surface projection of faults identifiedon seismic-reflection data by Demyttenaere and Laga(1988). The most prominent morphological feature isthe NW-SE-trending lineament defining the bound-ary between the Campine Plateau and the RVG (inthis area morphologically denoted as ‘Bocholt Plain’).This escarpment extends for a distance of almost 10km between the towns of Bree and Neeroeteren, moreor less corresponding to the portion of the Feldbissfault zone that is developed as a single fault and showsa topographic offset of 10–20 m. It has been namedthe ‘Bree fault scarp’ by Camelbeeck and Meghraoui(1996). To the southeast, the Bree fault scarp bi-furcates into a main easterly and a smaller westerlystrand, both becoming less distinct over the younger,Saalian and Weichselian, alluvial terraces in the MaasValley (Figure 2). To the northwest, near Bree, thefault scarp splits into three NNW-SSE-trending es-carpments with reduced topographic offset. The twoeasterly scarps are only 2–5 m high and are relativelystraight, whereas the westernmost scarp shows a largebend, its height decreasing from 10–15 m near Bree,to less than 5 m north of the bend. All scarps disappeargradually towards the northwest.

Before the first paleoseismological studies of theBree fault scarp (Camelbeeck and Meghraoui, 1996),a possible present-day activity of Lower Rhine grabenfaults and hence their potential for generating largeearthquakes, had not been considered. In a firststep, Camelbeeck and Meghraoui (1996) were ableto identify a small, 0.5–3 m high, frontal escarp-ment along most of the Bree fault scarp (Figure 3),

which they regarded as a primary geomorphic ex-pression of active faulting. Since that time, trencheshave been excavated at three sites along the northwest-ern part of the scarp (Figure 3), providing evidenceof late Pleistocene and Holocene surface rupturing(Camelbeeck and Meghraoui, 1996; Camelbeeck andMeghraoui, 1998; Meghraoui et al., 2000). Trench1 (Figure 3) was located in the bottom of a smallvalley eroded into the Bree fault scarp, and it ex-posed 50–60 cm of net vertical offset of a presumablypost-glacial alluvial terrace and the associated Holo-cene soil. This displacement most likely representsthe most recent surface-rupturing earthquake along theBree fault scarp. The age of this event has so far onlybeen constrained by radiocarbon dates, which bracketit between 610 and 890 cal AD. Trenches 2 and 3 (Fig-ure 3) were located on higher and hence older, sectionsof the fault scarp. They show both primary (collu-vial wedges) and secondary (induced soft-sedimentdeformation) evidence for two older surface-faultingevents. The penultimate event would have occurredbetween about 27 and 14 ka BP, slightly predating theend of the Last Glacial Maximum, while the antepen-ultimate event was radiocarbon dated as younger thanabout 44 ka BP.

In this paper, we describe the results of a newtrench (trench 4, Figure 3) that was excavated in1998 in the framework of the EC environmental pro-gram ‘PALEOSIS – Evaluation of the potential forlarge earthquakes in regions of present-day low seis-mic activity in Europe’. Studies at this site includeddetailed topographic leveling, geophysical prospectingusing various methods, trench mapping and interpret-ation, hand borings below the trench bottom and cone-penetration tests adjacent to the trench. We present thistrench separately, because it contains a paleoseismicrecord that is more complete than that of the previ-ously studied sites and also because it provided a fieldexercise for the Summer School in Active Faulting andPaleoseismology, to which this special volume is ded-icated. A more complete correlation of all trenchingstudies along the Bree fault scarp will be published atcompletion of the project.

Site reconnaissance

Location and morphology

The Bree trench site 4 is located along the southeast-ern part of the Bree fault scarp, close to the junction

333

Figure 3. Digital elevation model of the Bree fault scarp in NE Belgium, interpolated from 1/10,000-scale topographic maps of the BelgianNational Geographical Institute using the surface gridding algorithm of the Generic Mapping Tools (Smith and Wessel, 1990). Contour intervalis 1.25 m. The location is indicated of all trenches excavated across the fault scarp since 1996. BFS = Bree fault scarp, HFS = Heerlerheidefault scarp.

with the smaller scarp associated with the Heerler-heide fault and at a distance of 3.5 km from trenchsite 3, and 5 km from trench sites 1 and 2 (Fig-ure 3). At this location, the fault scarp has a broadand gentle slope, showing a maximum topographicoffset of 15 m over a distance of about 400 m. Aer-ial photographs (Figure 4) indicate the existence ofa well-defined, narrow lineament running the lengthof the main escarpment. This lineament can be tracedfor more than 1 km, mainly across farmland. A smallpasture transected diagonally by the lineament wasselected for trenching (Figure 4), also because ag-ricultural influence is thought to be minimal at thislocation. Northwest of this site, the lineament shows

a right step and converges with the edge of forestedterrain; to the southeast it eventually disappears into adry transverse valley.

In the field, the lineament appears to correspond toa subtle, but visible, topographic break, with a heightof ± 65 cm and a slope of 5◦. It is superposed on thelower half of the main escarpment, which has an over-all slope of 2.5◦. An elevation model constructed fromleveling profiles at the site (Figure 5) shows that theorientation of this scarp is identical to that of the lin-eament on the aerial photograph (Figure 4). In profile(Figure 6a), the scarp is also clearly visible. The twosurfaces it displaces are not exactly parallel, however,due to pre-existing relief (probably of eolian origin,

334

Figure 4. 1/20,000-scale aerial photograph showing lineament (indicated by white arrowheads) corresponding to frontal escarpment at trenchsite 4. In the northwestern part of the image, the lineament coincides with the edge of the forested area. Photograph reprinted with permissionof the National Geographical Institute of Belgium. Location is shown in Figure 3.

see trench stratigraphy), making it difficult to determ-ine the exact topographic offset. Because the position,orientation and lateral continuity of the small scarp arein good agreement with the known tectonic structure,it is interpreted as the frontal scarp of the main Feld-biss fault (Camelbeeck and Meghraoui, 1996). It issurprising that, in spite of the intense farming in thearea, such a small morphological feature is still quitenoticeable.

Geophysical investigation

In order to confirm the tectonic nature of the frontalescarpment prior to excavating a trench, several geo-physical prospecting methods (electrical and electro-magnetic profiling, electrical tomography and ground-penetrating radar or GPR) were applied along theplanned trench transect (Figure 6). The former threemethods all provide a measure of the ground resistiv-ity, while GPR yields a profile of energy reflected at

335

Figure 5. Shaded-relief model constructed from leveling profiles perpendicular to the fault scarp at trench site 4. Contour interval is 0.5 m. Themain topographic feature is a subtle scarp, less than 1 m high, with the same orientation as the lineament on the aerial photographs (Figure 4).The main fault that was later exposed in the trench is located at the foot of this scarp.

electric resistivity contrasts, the penetration being lim-ited by the ground resistivity. A more comprehensivesummary of the methods and results of geophysicalprospecting on all trench sites along the Bree faultscarp are presented elsewhere (Demanet et al., 2001).

Electric and electromagnetic profiles (Figure 6a)were acquired by moving a Schlumberger-like config-uration of electrodes and horizontal coils, respectively,along the profile. The half-electrode spacing was 50 m,whereas the half-coil spacing was only 10 m, resultingin a depth penetration of about 17 m and 4 m, respect-ively. The electromagnetic profile displays a small butdistinct drop of resistivity values (centered around m

60), from about 275 �.m upslope of the topographicscarp, to 200 �.m downslope. The electric profileshows a larger and broader resistivity gradient, whichis shifted c. 20 m upslope with respect to the scarplet,however. The different location and aspect of the elec-tromagnetic and electric anomalies are explained bythe integration of the latter signal over a larger lateraland vertical portion of the terrain. Towards the baseof the slope, resistivity values gradually rise again onboth profiles.

The electrical tomography profile in Figure 6bwas obtained by laying out a Wenner configurationof electrodes with a spacing of 2 m and applying a

336

Figure 6. Summary of geophysical profiles across fault scarp at trench site 4: (a) topographic, electric and electromagnetic profiles; (b)2D-inverted electrical tomography line; (c) 120 MHz ground-penetrating radar section; processing steps applied: bandpass filter (40/75–100/150MHz), automatic gain control (window 0.1 µs), static correction (velocity 90 nm/ms), no migration; detail in zoom box has different bandpassfilter settings (75/120–175/250 MHz). Horizontal scale is the same for all profiles, vertical scale varies.

337

2D-inversion algorithm (Loke and Barker, 1996) tothe measured apparent resistivities. The result is a res-istivity section extending to a depth of about 12 m.The most striking feature of this section is a narrowlateral resistivity gradient between a distance of 60and 65 m and extending over the entire depth interval.This anomaly dips 60–70◦ NE, in alignment with thefoot of the surface scarp. It also interrupts a shallow(1–6 m deep) low-resistivity layer, which appears tobe slightly thicker on the downslope side. Below thislayer is a high-resistivity layer that significantly shal-lows towards both ends of the section, much like theelectric and electromagnetic profiles.

Finally, a radar profile was recorded with a 120MHz transmitting antenna, which has a penetrationdepth not exceeding 4 m and a vertical resolution bet-ter than 0.5 m. The profile (Figure 6c) is characterizedby a clear lateral contrast in reflectivity character andpenetration depth, again more or less coinciding withthe foot of the morphological scarp near 60 m. Theupslope portion of the profile is dominated by discon-tinuous, incoherent, but high-amplitude reflectors anda small penetration depth, whereas the downslope partshows more regular reflectors extending to a greaterdepth and having lower amplitudes close to the sur-face. Using a suitable bandpass filter (see zoom boxin Figure 6c), the lateral transition between high-amplitude and predominantly low-amplitude surfacereflectors appears particularly sharp. Downslope ofthis transition, the reflectors are mostly parallel to thesurface, but near the base of the hillslope they climbto the surface and are disrupted in several places (e.g.between 100 and 105 m, Figure 6c).

In summary, all of the geophysical methods usedat this site indicate a shallow anomaly at the foot ofthe small morphological escarpment. This anomalyextends to a depth of at least 12 m and thus supportsour interpretation that the frontal scarp at trench site 4is the expression of a surface fault. In addition, there isevidence of disturbed stratification (uprising and dis-ruption of surficial layers on the GPR section, Figure6c) towards the downslope end of the profile, which issuggestive of a landslide associated with small-scalethrust faulting at its toe.

Trench investigation

The good agreement between the geomorphic andgeophysical observations helped us select a suitabletrenching profile to obtain more information about the

characteristics and timing of the deformation. Becausemovement on the Feldbiss fault is presumed to havebeen predominantly dip-slip throughout the late Qua-ternary, the excavation was oriented N 45◦E, roughlyperpendicular to the frontal scarp. The trench wasabout 90 m long, 3.5 m deep and was excavated in twolevels, separated by a horizontal platform of almost 2m wide. We also drilled hand borings every 5–10 mat the trench bottom, which extended our stratigraphicdata 2–5.5 m below the trench.

Trench stratigraphy

The stratigraphic information obtained from detailedmapping of the northwestern trench wall and from thehand borings is summarized in Figure 7. The trenchexposed two blocks with distinctly different strati-graphy, separated by a single normal fault coincidingwith the foot of the morphological scarp (at m 61in Figure 7). A detail of the fault zone is shown inFigure 8.

The footwall is dominated by red-brownish, clay-coated, gravelly, coarse sands with a distinct flu-vial character, including large-scale cross-bedding anderosional channels. These sediments are correlated tothe youngest and main member of the Maas RiverMain Terrace, the Zutendaal Gravel, which was de-posited by a braided river system in a periglacialenvironment, sometime between 770 and 350 ka BP(Paulissen, 1983). They have the same petrographiccomposition as gravel deposits on the Campine Plat-eau, but they are situated about 5 m lower due tothe throw on another important fault (the Heerlerheidefault, Figure 3) between trench 4 and the plateau.

The hanging wall mainly contains yellowish, silty,fine to medium sands with varying amounts of finegravel, occasionally concentrated in pavements. Thesedeposits show mostly parallel stratification with somesmall gullies, and are interpreted as primarily eoliansands that have been reworked by slope wash, pos-sibly fed by snowmelt (Gullentops et al., 1993) or byrainwater or water from local springs. These sandsform a continuous mantle over large parts of north-western Europe and were mainly deposited during theWeichselian and Saalian glacial periods (e.g. Paulis-sen, 1973; Kolstrup, 1980; Kasse, 1997). Becausethey tend to cover the pre-existing relief, they arecommonly referred to as ‘coversands’.

The difference in stratigraphy on either side of thefault is also highlighted by the hand borings: boreholesin the footwall were limited to a depth of 2 m due to

338

Fig

ure

7.Sc

hem

atic

draw

ing

ofno

rthw

estw

allo

ftre

nch

4an

dst

ratig

raph

icsu

mm

ary.

Not

eth

ebr

eak

betw

een

low

eran

dup

per

tren

chw

all.

Indi

cate

dar

epo

sitio

nof

mai

nfa

ult(

F),s

trat

igra

phic

units

(1–1

0)an

dda

ting

resu

ltsfo

rsa

mpl

esou

tsid

efa

ultz

one.

Ver

tical

exag

gera

tion

is2.

5.L

ocat

ion

oftr

ench

issh

own

inFi

gure

s3,

4an

d5.

339

Fig

ure

8.D

etai

led

map

ofm

ain

faul

tzon

eon

nort

hwes

tw

allo

ftr

ench

4,sh

owin

gam

ount

ofdi

spla

cem

ent

ofco

rrel

ativ

eho

rizo

ns,i

nfer

red

posi

tion

ofev

ent

hori

zons

(EH

)an

dda

ting

resu

lts.

Thi

n,ho

rizo

ntal

beds

offin

esa

ndw

ithin

the

wed

ge-s

hape

dun

it7

inth

eha

ngin

gw

allr

esem

ble

bent

fros

tfiss

ures

info

otw

allu

nit1

F.Se

eFi

gure

7fo

rlo

catio

n.

340

the presence of coarse gravel, whereas borings of 4m and more were easily attained in the finer-grained,hanging-wall sediments. The trench section has beensubdivided into ten sedimentary units, numbered fromoldest to youngest. A brief description of their li-thology and interpreted depositional environment isshown in Figure 7. Below we discuss only those ele-ments of the stratigraphy that are of interest in relationto the history of faulting.

Unit 1 corresponds to the coarse Maas River sands,which make up most of the footwall. The high graveland clay content of this unit are probably responsiblefor the poor penetration and lack of coherent reflect-ors in this part of the GPR section (Figure 6c). Veryhigh penetration resistances and low friction ratios ofcore penetration tests suggest that the Main Terracedeposits extend at least 9–10.5 m below the trenchbottom. We subdivide this unit into six subunits A-F,corresponding to different episodes of fluvial incisionand aggradation. Of particular interest are subunits Cand E. The former is a 10-cm-thick clay band whichis almost horizontal, indicating that the unit 1 sandshave retained the same position since their deposition.Subunit E fills a channel incising more than 4 m intounderlying subunits towards the fault zone. The entireunit 1 has been subjected to intense soil development,as indicated by its pervasive brown color and high claycontent (‘As’ Soil, Gullentops et al., 1981; Paulissen,1983). Close to the main fault, the stratification in unit1 is indistinct, particularly on the lower trench wall.

The top of unit 1 shows evidence of cryogenic dis-turbance in the form of narrow fissures filled with fineyellowish sand (m 55–60 in Figure 8). These fissuresare only a few cm wide, 50–100 cm long and typicallyhave laminations that are parallel to the fissure walls.Most of these fissures are truncated at the top of unit 1,but a few of them connect to the overlying coversandunit 8 (described below), where this unit is preservedin the footwall (see m 57–58 in Figure 8). This indic-ates that the fissures formed in at least two separatephases. Based on their shape and fill, we interpret thefissures as periglacial frost fissures formed by thermalcontraction of fine-grained sediments and filled withwind-blown sand (e.g. Van Vliet-Lanoë, 1985). Frostfissures are often referred to as sand wedges, but weavoid this term because of the possible confusion withcolluvial wedges in paleoseismology. The laminationsparallel to the fissure walls result from seasonal open-ing of the crack due to freezing of the ground andrenewed infilling with eolian sand. Thermal contrac-tion of the coarse sands of unit 1 was facilitated by

their relatively high (up to 15%) clay content. Closeto the fault zone, where the frost fissures are con-centrated, frost cracking likely exploited pre-existingcracks and weak zones related to faulting. Frost cracksusually grow perpendicular to the frozen surface. Inthe vicinity of the main fault, however, the fissures areprogressively bent and stretched in downslope direc-tion (Figure 8), from subvertical near m 52 to almosthorizontal at the fault. This deformation is evidence ofsolifluction (the slow downslope movement of surfacesediments) and has implications for the origin of somewedge-shaped units in the hanging wall, as discussedin the section describing faulting evidence.

The stratigraphic equivalent of unit 1 was not en-countered in the hanging wall of the fault. The oldesthanging-wall deposits (unit 2) are only known fromthe hand borings and consist of more than 2.5 m ofgray-colored, slightly clayey, medium to coarse sandthat is finer-grained and contains less gravel than theunit 1 deposits. We interpret them also as fluvial sands,but it is not clear if they still belong to the upper partof the Main Terrace (which has not been preservedin the footwall), or if they have been reworked fromunit 1 upslope. The difference in color between thegray unit 2 sands in the hanging wall and the brownsands of unit 1 in the footwall is probably due to theirdifferent position with respect to the water table. Theoverlying unit 3 is composed of gray to red, clayey,medium to coarse sand, grading into clayey silt whichis interfingering with fine sand. This sequence is in-terpreted to document the transition from a fluvial toan eolian environment. Unit 3 is only exposed in thedeepest part of the downslope end of the trench, butthe boreholes show that it dips southwest, toward themain fault and that it pinches out about 10 m from thefault. This reverse, southwestward dip of deeper layersis also indicated by electrical tomography (Figure 6b),GPR (Figure 6c) and CPT-tests.

The remainder of the hanging wall is primarilymade up of four coversand units (units 4, 6, 8 and 10),composed of silty, fine sand with generally upward-decreasing amounts of fine gravel. The lowermostcoversand units, 4 and 6, are very heterogeneous andcontain large amounts of gravel, indicating a largedegree of reworking, mainly by slope wash (syngen-etically), and possibly also by local streams. The latteris particularly true for unit 6, as shown by the rapidlateral facies variation from a channel-like facies con-taining as much as 50% gravel close to the main fault,to a finely laminated silt facies away from the fault.The lower part of coversand unit 8 also contains a large

341

percentage of relatively coarse gravel, but the upperpart only contains yellow, well-laminated, silty sandwith a few single-pebble horizons that probably cor-respond to intervals of deflation. Unit 8 partly extendswith a small vertical offset across the main fault ontothe footwall block.

Interbedded between coversand units 4, 6 and 8are two distinctive wedge-shaped units (units 5 and7) of dark-brown, clayey and also gravelly, coarse-grained sand. Both wedges rapidly thin downslopeinto an extended ribbon less than a few centimetersthick, but merge into a single unit, more than 2 mthick, adjacent to the main fault, where they are nearlyindistinguishable from the footwall deposits. Their in-ternal stratification, which is subhorizontal and almostparallel, appears to be different, however. The upperwedge also contains several thin, horizontal beds offine, yellow sand. Our boreholes indicate that twoadditional wedge-shaped units of similar composition(units 4B and 4D) are present below the trench bottomjust downslope of the main fault. Below we will de-scribe evidence that these wedges were derived fromthe Main Terrace deposits in the footwall (unit 1)by mass wasting, and most likely represent colluvialwedges.

The stratigraphic sequence in trench 4 is concludedby two units (units 9 and 10) which extend the entirelength of the trench. Unit 9 is a striking, 5- to 10-cm-thick gravel layer, which successively truncatescoversand units 8, 6 and 4 towards the base of theslope. We interpret it as a lag gravel that formed duringa period of extensive wind erosion. The gravel layeris more irregular in most of the footwall, where itdirectly overlies unit 1, and thus marks a large de-positional hiatus. Finally, the uppermost unit 10 isagain composed of fine, silty coversand, containingfew single-pebble horizons. The internal stratificationand original color of this unit are mostly obliterated bysoil development and bioturbation, and its top containsthe plow zone.

Dating and correlation of trench units

The late Pleistocene stratigraphy in northeastern Bel-gium is not known in great detail, which makes it diffi-cult to accurately correlate the deposits in trench 4. Asmentioned above, we correlate the coarse fluvial sandsof unit 1 in the footwall with the Zutendaal Gravelmember of the Maas River Main Terrace, which wasdeposited sometime between 350–770 ka BP (Paul-

issen, 1983). Unfortunately, we could not date thesedeposits in the trench.

Most of the hanging-wall units are composed ofcoversands. Only two coversand formations have beendefined in the area (Paulissen, 1973; Beerten et al.,1999b); these were deposited during the Saalian (c.130–185 ka BP) and the Weichselian (c. 10–120 kaBP) glacial periods, respectively. It is very hard todistinguish between these formations based on grainsize or mineralogy alone. More detailed correlationshave to rely on the Weichselian coversand stratigraphyestablished in the Netherlands (Zagwijn and van Staal-duinen, 1975; Kolstrup, 1980; see Figure 9), where thesequence contains a prominent stratigraphic marker,the Beuningen gravel bed. This gravel is a widespreadlag horizon that formed at the end of the Last GlacialMaximum, mainly between 14–19 ka BP (Kolstrup,1980). The Beuningen gravel bed is known to extendinto the Belgian coversand area as well. A good can-didate in our trench for correlation with this horizon isthe gravel bed of unit 9, which is a striking erosionalbed showing evidence of deflation. The late Weich-selian coversand sequence in the Netherlands is furtherdivided into Older Coversand I (23–27 ka BP) belowthe Beuningen gravel bed and Older Coversand II (12–14 ka BP) and Younger Coversand I and II (10–12 kaBP) above it. The coversands below and above unit 9in trench 4 have sedimentological characteristics thatare comparable with those of the Older Coversands Iand II, respectively.

Due to their generally very low organic content,radiocarbon dating of coversands is difficult, so wealso attempted to date the coversand deposits in trench4 using luminescence dating techniques. Infrared op-tically stimulated luminescence (IRSL) and thermolu-minescence (TL) have proven to be a reliable datingtool for determining the depositional age of loess de-posits (Frechen, 1999; Frechen et al., submitted) andfluvial sediments in Belgium and Germany (Frechen,1995). In this study we applied this dating method forthe first time to the coversands in the Belgian MaasValley and along the Bree fault scarp. The methodo-logy used for these analyses will be reported in detailelsewhere (Frechen et al., in press). Eight lumines-cence and eleven radiocarbon samples were analyzed;the results are presented in Tables 1a and 1b, respect-ively, while the location of the samples is shown inFigures 7 and 8. The consistent stratigraphic orderof the luminescence ages and the good agreementbetween the IRSL and TL methods, support the reli-ability of the method in this particular environment.

342

Figure 9. Graphical comparison of IRSL and radiocarbon datings of coversand units 6, 8 and 10 in trench 4 and correlation with coversandstratigraphy in the Netherlands (Kolstrup, 1980). For radiocarbon dates younger than 20 ka, 2-sigma calibrated ages have been recalculatedto ka BP. See Tables 1a and 1b for the actual dating results. Samples IVNW-C9 and IVNW-C 12 were rejected and are not shown here. Agebrackets for the faulting events E1-E3 are indicated as patterned areas.

Below we only cite the IRSL ages, because they areconsidered to be more reliable, as the reset time forIRSL samples is much shorter than for TL samples.

The ages obtained for samples from trench units 6,8 and 10 are plotted in Figure 9. In general, lumin-escence and radiocarbon datings agree relatively wellfor the upper trench units (10 and 8), but significantlydiverge for the older unit 4. As indicated, the age datessupport our correlation of unit 9 with the Beuningengravel bed. Radiocarbon and IRSL dates of the overly-ing unit 10 are mostly bracketed between 10–15 kaBP, corresponding to Older Coversand II and YoungerCoversands. The radiocarbon age of sample IVNW-C9(Table 1b) is clearly inconsistent with the other results

and is therefore rejected. The historical ages obtainedfor some pit-shaped concentrations of organic materialat the top of unit 10 (Figure 8) suggest that these areanthropogenic features.

No organic material was retrieved from unit 8below the Beuningen gravel bed, so we only have lu-minescence dates for this unit. The two IRSL ages arebetween 14–27 ka BP and indicate that this unit prob-ably correlates with Older Coversand I in the Dutchstratigraphy; its upper part may (IRSL sample OPI-05) grade into the Beuningen horizon. The datingresults for unit 6 are more difficult to interpret. Ex-cept for sample IVNW-C 12, which we reject, theradiocarbon samples show a large age scatter between

343

Table 1a. Luminescence dating results in trench 4 (1 sigma confidence interval). TL = thermoluminescence,IRSL = infrared stimulated luminescence. TL and IRSL ages were determined from identical samples in eachcase. See Figures 7 and 8 for location of samples

Sample ID Lab ID Trench unit TL age IRSL age Stratigraphic

(kyr BP) (kyr BP) correlation

OPI-11 FRE-C1-0201 10 not datable not datable Younger Coversands –

OPI-09 FRE-C1-0195 10 8.5 ± 0.9 8.5 ± 1.6 Older Coversand II

OPI-08 FRE-C1-0194 10 10.3 ± 1.3 12.8 ± 1.4

OPI-10 FRE-C1-0196 10 12.2 ± 1.7 14.0 ± 2.3

OPI-05 FRE-C2-0113 8 12.5 ± 1.4 15.8 ± 2.9 Older Coversand I

OPI-06 FRE-C2-0114 8 17.7 ± 2.0 23.4 ± 4.1

OPI-03 FRE-C2-0112 6 53.2 ± 7.1 63.3 ± 13.1 Early (?) to Middle

Weichselian coversands

OPI-01a FRE-C1-0193 4C 128.3 ± 14.4 136.6 ± 17.6 Saalian coversands

Table 1b. AMS radiocarbon dating results in trench 4. Samples IVNW-C9 and IVNW-C12 are completely out of line with the otherresults and are therefore rejected. See Figures 7 and 8 for location of samples

Sample ID Lab Trench Type of AMS age Calibrated age (2σ ) Stratigraphic

ID unit material (yr BP) correlation

IVNW-C2b UtC-9249 black organic 230 ± 30 1630 AD (0.49) 1680 AD Historic

pits spherules 1730 AD (0.43) 1810 AD

below 1940 AD (0.02) 1950 AD

IVNW-C3b UtC-9247 plow organic 480 ± 30 1400 AD (1.00) 1470 AD

zone spherules

IVNW-C10b UtC-9411 organic 140 ± 40 1660 AD (1.00) 1960 AD

spherules

IVNW-C10c UtC-9250 charcoal 780 ± 30 1210 AD (1.00) 1290 AD

IVSE-C1 UtC-9063 10 charcoal 9050 ± 50 8450 BC (0.97) 8160 BC Young Coversands

IVNW-C8 UtC-9248 10 wood/charcoal 10780 ± 70 11200 BC (1.00) 10650 BC to

IVNW-C9 UtC-9409 10 wood 3210 ± 40 1500 BC (1.00) 1260 BC Older Coversand II

IVNW-C5 UtC-9244 6 wood/charcoal 26100 ± 280 – Early (?) to Middle

IVNW-C6 UtC-9246 6 wood/charcoal 30500 ± 500 – Weichselian

IVNW-C4 UtC-9245 6 wood 39100 ± 1400 – cover sands

IVNW-C12 UtC-9410 6 charcoal 12070 ± 200 13500 BC (1.00) 11500 BC

about 26 and 40.5 ka BP. The dating results seemto be meaningful as samples IVNW-C6 and IVNW-C4 yield ages corresponding to the Denekamp andHengelo interstadials, respectively. The ages are inreverse stratigraphic order, however, which can onlybe partly explained by severe deformation (includingsmall-scale thrust faulting) of the downslope part ofunit 6, where these samples were taken (Figure 7).Closer to the main fault, the only IRSL sample from

this unit yielded a considerably older age (63.3 ± 13.1ka BP; Figure 8). Given these relations, we interpretthe radiocarbon samples from unit 6 to be ‘intrusive’,i.e. transported from the surface downward after theunit was deposited. This would imply that coversandunit 6 is older than the age of the oldest radiocarbonsample IVNW-C4, 39,100 ± 1400 yr BP (Figure 9).Thus, unit 6 dates to the Early Weichselian, possiblyextending into the early Middle Weichselian. Finally,

344

a late Saalian to Eemian age (136.6 ± 17.6 ka BP;Table 1a) was obtained for an IRSL sample from unit4C at the bottom of the trench (Figure 7). Since cov-ersands were only deposited during glacial periods,a late Saalian age seems more likely. It is not clear,however, where in that case, the Eemian is representedwithin our trench, because we found no evidence ofthe prominent soil that is typical of this interglacialperiod.

It is obvious from this discussion that our age dateshave sufficient resolution to bracket individual faultingevents only in the upper part of the section (see nextsection). In addition, Holocene deposits are mostlymissing in the trench record. Further constraints on thetiming of events are provided, however, by secondarygeological features, such as soil horizons, post-datingdeposition of the sediments in which they are con-tained. Units 8, 9 and the lower part of unit 10 con-tain several Bt-bands (Figure 8), which are rubified,10- to 15-cm-thick, argillic horizons, formed by thedownward migration (illuviation) of fine clay particlesstained with iron hydroxides. These clays were prob-ably derived from the upper part of unit 10, which isentirely bleached (eluviation). The base of these Bt-bands is typically, but not always, parallel to the loc-ally inclined stratification of the host sediment, whiletheir tops are mostly subhorizontal and independent ofthe internal bedding. Development of these Bt-bandsprobably started during the Late Glacial and prob-ably continued through the early Holocene (Miedema,1983; Van Vliet-Lanoë et al., 1992), i.e. during or justafter deposition of unit 10. Within the upper, bleachedpart of unit 10 (i.e. the E-horizon associated with theLate Glacial – early Holocene Bt-horizon), a less pro-nounced, gray-brownish Bt-band is present (markedBt2 in Figure 8), which must be related to a youngerphase of illuviation, possibly stimulated by the onsetof farming activities and associated soil fertilization inmedieval times (R. Langohr, pers. comm.).

Evidence of faulting

Detailed mapping of the trench walls documenteda normal fault just below the subtle morphologicalscarp (near 61 m, Figure 7), which coincides withthe position of the shallow anomaly identified onthe geophysical data (Figure 6). The main deforma-tion is confined along a single fault plane oriented N124◦E and dipping 76◦ to the NE. The fault juxtaposescoarse-grained, red-brownish Main Terrace deposits inthe footwall (unit 1) against lighter-colored and finer-

grained coversand-related deposits in the hanging wall(units 4–10). The fault zone is shown in more detailin Figures 8 and 10a. On the upper trench wall, thefault plane is well defined by small, but sharp dis-placements of sediment units 8, 9 and 10, as well as ofthe soil horizons (older and younger Bt-bands) withinthese units. On the lower trench wall, the fault contactis rather obscure due to the mergence of two wedgesof Main Terrace-derived material in the hanging wall(units 5 and 7), which have a very similar color andgrain size as the adjacent unit 1 in the footwall. Thefault is defined by a straight and narrow, white-coloredtrace that is in contrast with the brown-colored sedi-ments on either side (Figure 10a). Subtle gravel bedsterminate at the fault, but we could not confidentlycorrelate specific beds across the fault to determinethe amount of displacement. Parallel to the main fault,some synthetic faults with smaller offsets could bemapped but they did not extend up to the upper trenchwall. Their white color, similar to that of the mainfault, is not due to any infill, but is the result ofin situ bleaching of parent material, probably causedby preferential infiltration of reducing meteoric wateralong these lines of weakness. This is supported bythe irregular and fuzzy edges of the bleached zonesand by the continuation of grain-size variations acrossthese edges; the grain size distribution of the bleachedmaterial is identical to that of the adjacent sediment,minus the silt and clay fractions.

Hand borings (Figure 7), as well as geophysicaldata (particularly the electrical tomography profile,Figure 6b) indicate that the main fault probably ex-tends below the trench bottom in a straight line. Incombination with the geomorphic evidence, these ob-servations emphasize the tectonic significance of theshallow fault exposed in the trench.

Detailed analysis of the fault zone provides evid-ence for as many as six separate faulting events, whichwe label E1 (youngest) to E6 (oldest). The two young-est events, E1 and E2, have small displacements andprimary evidence for their coseismic nature is lim-ited. The older faulting events, E3–E6, are inferredfrom the presence of four wedge-shaped deposits inthe hanging wall that we interpret as fault-derivedcolluvial wedges.

Most recent and penultimate faulting eventsEvidence of the two most recent fault movements intrench 4 is preserved within the youngest trench units8, 9 and 10, which are the only units that can be cor-related between footwall and hanging wall (Figure 8).

345

Fig

ure

10.

Phot

ogra

phs

ofm

ain

faul

ton

nort

hwes

tern

tren

chw

all:

(a)

expr

essi

onof

mai

nfa

ult(

F)by

disp

lace

men

tof

corr

elat

ive

hori

zons

(bas

eof

cove

rsan

dun

it8,

and

grav

elbe

dof

unit

9)on

uppe

rtr

ench

wal

l,an

das

linea

r,bl

each

edfis

sure

onlo

wer

wal

l;bl

ack

line

onho

rizo

ntal

plat

form

outli

nes

posi

tion

ofla

tera

llytr

unca

ted

fros

tfiss

ure.

Loc

atio

nis

indi

cate

din

Figu

re8;

(b)

deta

ilfr

ombo

xsa

mpl

esh

owin

gsh

arp

vert

ical

disp

lace

men

tof

Bt2

soil

hori

zon

with

inun

it10

.D

otte

dlin

esin

dica

tepo

sitio

nof

mai

nfa

ult

plan

e(F

).L

ocat

ion

corr

espo

nds

appr

oxim

atel

yto

rect

angl

ein

Figu

re10

a).

346

Figure 11. Vertically exaggerated profile of the base of trench unit 9 (Beuningen gravel bed), showing amount of vertical displacement directlyat the fault compared to the total offset between the modeled upthrown and downthrown surfaces. Dashed lines show projections of thesesurfaces.

The primary evidence of these events are the differingamounts of displacement of the coversand units aboveand below the Beuningen gravel bed (unit 9). The mostrecent movement (E1) is represented by 10–12 cm ofdisplacement of unit 9 and at least the base of unit10. Modeling of the elevation profile of the Beunin-gen gravel bed indicates a total offset of as much as0.8 m (Figure 11), which agrees better with the topo-graphic offset (Figure 6a). Because this gravel bed wasdeposited over an erosional surface that was probablyrelatively even, the additional offset is likely due to

flexure, though it cannot be excluded that part of it isinherited from a previous fault movement.

In the trench, we could trace the fault line upwardto at least 30 cm below the plow zone (Figure 8),where the original stratigraphy is obliterated by soildevelopment. A second-generation Bt-band within thebleached part of unit 10 (Bt2 in Figure 8) has thesame amount of displacement as the base of this unit(Figure l0b), suggesting continuation of the fault upto the surface. The same relation is present at the op-posite trench wall, almost 6 m away. However, careshould be taken when examining soil horizons for fault

347

displacements, because apparent displacements couldalso have developed after the fault movement by soilformation processes exploiting variations in grain sizeor mineralogy along identical stratigraphic levels oneither side of the fault. In this case, however, a post-faulting origin of the Bt2-band is unlikely because theoffset is clean and sharp, and the Bt2-band is the samethickness on either side of the fault (Figure 10b). Inthe case of non-tectonic offsets, the geometric rela-tions between Bt-bands and fault planes are usuallymore complex, as we observed in previous trenches(Vanneste et al., 1999). If our interpretation is correct,then this also implies that the offset associated withE1 is not the result of continuous, aseismic fault slip,at least not during the time interval between depositionof unit 10 and development of the youngest Bt-band.

Directly downslope of the fault, a small patch ofdark, organic-rich material is preserved just below theplow zone, resembling a faulted and buried A-horizon.Radiocarbon ages indicate a historical age for this or-ganic matter (Table 1b). The feature is most likely notrelated to faulting, however, because similar organic-rich patches are present in two other places on thefootwall (e.g. between 36–38 m in Figure 7). Ex-amination of their organic content suggests that thesepatches are instead remnants of historical fire pits (R.Langohr, pers. comm.). If an event horizon (as definedby Pantosti et al., 1993) is associated with E1, it isprobably contained within the plow layer.

It is difficult to precisely determine the age ofthe most recent fault movement. The only quantitat-ive information is provided by IRSL sample OPI-09(Table 1a and Figure 9), which yields a maximum ageof 8.5 ± 1.6 ka BP. But event E1 may be consider-ably younger, because a second-generation Bt-horizonis displaced as well. The age of this Bt-horizon isunknown, but it started to form only after depos-ition of unit 10 and completion of the main Bt-banddevelopment.

Below the Beuningen gravel bed, larger fault dis-placements are obvious. The base of the laminatedcoversand unit (unit 8) is vertically offset an additional33 cm (Figure 8), which indicates a fault movementprior to E1. This value is a minimum estimate becausethe amount of flexuring possibly associated with thisevent could not be quantified. The portion of unit 8that is preserved in the footwall appears to correlatewith the lower half of unit 8 in the hanging wall, con-straining the time of fault slip to between depositionof the upper part of unit 8 and the subsequent erosionassociated with unit 9. We cannot further refine the

timing of this possible penultimate event (E2) becausewe could not identify a related colluvial wedge. Eitherthe surface rupture from that event was too small toproduce a colluvial wedge, or the event occurred nearthe end of unit 8 deposition, and the erosion whichleft the lag gravel of unit 9 may have destroyed thecolluvial wedge associated with E2.

A complex structure of soft-sediment deformation,situated a few meters downslope from the main fault(Figure 8: m 63–66), may provide additional con-straints on the timing of E2. The structure is present onboth trench walls and consists of a set of small-scaleintraformational faults, both synthetic and antitheticto the main fault, and tightly folded laminations. Thecharacter of this deformation suggests a combinationof plastic flow and brittle failure, which often occursimultaneously during liquefaction of sands (Owen,1987). Liquefaction may also occur under certainperiglacial conditions (resulting in so-called cryoturb-ation), but we favor a tectonic origin for this structurebecause of its proximity to the main fault and its lin-ear extent parallel to the fault. In addition, grain-sizedata from earlier excavations indicate that the cover-sands along the Bree fault scarp are indeed susceptibleto earthquake-induced liquefaction (Vanneste et al.,1999). A simple graphic reconstruction (Figure 12)clearly indicates that the structure formed in two sep-arate phases, one just before deposition of unit 9 andthe other sometime after deposition of this gravel bed.This is very similar to the fault slip history, and wetherefore correlate the two phases of soft-sediment de-formation with E1 and E2. This implies that the baseof unit 9 is the event horizon related to E2, and thusthe timing of E2 is constrained between the ages ofIRSL samples OPI-05 (15.8 ± 2.9 ka BP) and OPI-10(14.0 ± 2.3 ka BP) (Figure 8). This defines a max-imum time window of 7000 years (18.7–11.7 ka BP)for this event (Figure 9), but most likely even less,considering that sample OPI-10 is located about 30 cmabove the gravel bed of unit 9. In spite of the absenceof a colluvial wedge, this is good evidence that faultslip at trench site 4 was not slow and continuous, butprobably of seismic origin. A coseismic nature of theobserved fault displacements is also supported by thepresence of some secondary deformational features inthe footwall, mainly bleached subvertical fissures thattrend parallel to the main fault (Figure 8) and whichwe interpret as extensional cracks induced by suddenfault movement.

348

Figure 12. Repeated soft-sediment deformation in coversands adjacent to main fault in trench 4 (see Figure 7 for location): (a) present situation;(b) simple graphic restitution of the gravel bed of unit 9, revealing that the main deformation occurred between deposition of units 8 and 9; thebase of unit 9 thus corresponds to an event horizon (EH). Legend as in Figure 8.

Older faulting events represented by colluvial wedges

Below unit 8, direct evidence of fault displacementsis limited due to the lack of correlative horizons. TheMain Terrace deposits of unit 1 have no known coun-terpart in the hanging wall and are, in all likelihood,downfaulted below the depth of our hand borings.This implies that their cumulative vertical displace-

ment is at least 8 m. The hanging wall does containfour wedge-like units (units 5 and 7 exposed in thetrench wall and units 4B and 4D in the boreholes, Fig-ure 7) that are composed of sediment similar to unit 1.However, several lines of evidence indicate that thesewedges are not the stratigraphic equivalents of unit1, but are instead derived from this footwall unit by

349

slope-controlled processes that were most likely ini-tiated by surface-faulting events. The wedges are notsimple downslope extensions of unit 1 that interfingerwith coversand units in the hanging wall because (i)they lack the cross-bedding that characterizes most ofunit 1, but instead they have discontinuous, paralleland subhorizontal internal bedding (Figure 8); (ii) ifthey had been deposited by the same large-scale flu-vial system that deposited unit 1, then these sedimentsshould get thicker downslope, where fluvial depositswould fill a channel; instead the wedges get thinnerdownslope; and (iii) dates of the interbedded cover-sands (Figure 7) indicate that three out of four wedgesare much younger (136.6 ± 17.6 ka BP, Table 1a)than the Maas River Main Terrace, which was depos-ited between 350 and 770 ka BP. In a cross-sectionalview, the wedges are thickest upslope where they arebounded by the main fault, but they rapidly get thinnerdownslope. This shape strongly suggests depositionby downslope sediment transport from a local sourcejust upslope of the main fault. This shape is mostlyoriginal because, as we can see for the two wedgesexposed on the trench wall, the bases of the coversandunits overlying each wedge do not appear to be signi-ficant erosional contacts. This is shown by the internalstratification in the wedges, which is not systematic-ally truncated at this contact (Figure 8) and by the factthat ribbon-like downslope extensions of the wedges,less than a few cm thick, are preserved over a distanceof more than 20 m (Figure 7, 80–105 m). We found noreason why these thin distal layers would have resistederosion while the rest of the wedges was eroded.

A key element for establishing the origin of thewedge-shaped units 5 and 7 in the hanging wall isthe presence of several horizons of fine yellow sandin the wedges. These sand horizons are only a fewmillimeters thick (Figure 13a) and are very similarin color and grain size to the eolian sand filling thefrost fissures in the upper part of unit 1 in the footwall(Figures 8 and 13b). Frost fissures normally originateas subvertical features, but here they are progress-ively bent and stretched towards the fault zone. Thiswarping is probably not tectonic in origin, becausean underlying gravel bed at the base of unit 1F ishorizontal and continuous (Figure 8). Rather, thisstyle of downslope-directed deformation, decreasingin amplitude from the surface downward, is typicalof slow downslope movement of surface sediments.Thus, the thin horizons of yellow sand in the hangingwall wedges 5 and 7 are interpreted as originally ver-tical frost fissures that have been stretched into nearly

horizontal beds. This is supported by the observationof a low-angle cross-cutting relationship of these yel-low horizons with the internal wedge stratification insome places. Consequently, we interpret the wedges ofcoarse sand in which these horizons are embedded asslope deposits that were derived from the Main Terracedeposits of unit 1 in the footwall, as shown in Figure14. The mass deficit of unit 1 is clearly shown by thenoticeable thinning of subunit 1F toward the main fault(see Figure 8, 58–61 m).

Considering that most of the wedges were de-posited during the Middle to Early Weichselian andSaalian glacial periods, as indicated by IRSL andradiocarbon dates (Tables 1a and 1b), downslopemovement of the Main Terrace deposits most likelyoccurred in a periglacial environment, which is char-acterized by highly active slope processes such as frostcreep and gelifluction. Frost creep is a slow process bywhich soil particles are lifted normal to the soil surfaceby freezing and resettle slightly further downslopeduring thawing (e.g. Van Vliet-Lanoë, 1985); gelifluc-tion is the slow downslope displacement of most of thewater-saturated thaw layer (Harris, 1988). Evidence ofthe nature of the processes responsible for depositionof the wedges of coarse sand are found in the sourcesediments (unit 1) close to the main fault. Here, a sec-ondary fault, located about 1 m upslope from the mainfault (Figure 8, near 60.5 m), shows a peculiar deform-ation. The upper part of this fault is bent downslopeto a nearly horizontal attitude, whereas the lower partis nearly vertical. The inflection point between thenear-horizontal and near-vertical parts of the fault isvery sharp and at the same level as the base of theupper wedge (unit 7) prior to faulting events E1 andE2. This bending implies en masse downslope move-ment of the entire sediment layer above this level.On the other hand, the downslope bending of thefrost fissures towards the top of unit 1 (see above)indicates that sediments close to the surface movedmuch farther downslope than deeper sediments. Thesetwo observations could be explained by a combinationof slow gelifluction and frost creep. However, thin-sections from the uppermost wedge indicate that clayskins of the redeposited material are crushed, whichis more indicative of fast flow, probably in the orderof meters per hour (B. Van Vliet-Lanoë, pers. comm.).If this is correct, then this flow must have been rel-atively laminar because the redeposited sediment wasnot completely homogenized, as indicated by the pres-ence of the fine yellow laminae within the wedges.Thus, downslope sediment transport may have been

350

Figure 13. Photographs illustrating resemblance of frost fissures filled with fine sand in footwall and fine-sand horizons in hanging wall: (a)thin, horizontal beds of fine sand (indicated by white arrows) in wedge-shaped unit 7 in the hanging wall; (b) progressive bending of originallynear-vertical frost fissures at top of footwall unit 1 towards main fault (2 m right of the picture). The former are thought to have formed fromthe latter by downslope flowing of footwall sediments in response to a surface-faulting event. Location is indicated in Figures 7 and 8. Legendas in Figure 10.

351

Figure 14. Model for colluvial-wedge development following surface rupture under periglacial conditions at trench site 4. CW = colluvialwedge.

352

the result of liquefaction of unit 1 sands, which re-quires water-saturated conditions that were frequentlyattained in a periglacial environment during thawingof the active layer or by snowmelt (nivation). The sed-iments close to the surface would have flowed fartherthan deeper sediments and may also have been sub-ject to later frost creep, which explains the observeddeformation geometry (Figure 14).

The previous discussion mainly demonstrates thecolluvial nature of the uppermost wedge units 5 andparticularly 7, but we infer a similar origin for the twoolder wedges (units 4B and 4D) that were detected inboreholes below the trench bottom (Figure 7). Fromtheir position and shape, these wedges are related tomovement of the main fault, which created the neces-sary accommodation space in the hanging wall. Thenoticeable thinning of subunit 1F towards the mainfault (see Figure 8, 58–61 m) indicates that downslopetransport of unit 1 sands was not uniform over the en-tire slope, but was most pronounced within the fewmeters upslope of the fault. This strongly suggests theexistence of a fault scarp at the time the hanging-wallwedges were deposited (Figure 14), a situation thatapparently was repeated four times. Since the wedgesalternate with coversand units of different ages, theyare well-separated in time, which demonstrates theepisodic character of the fault movement. We there-fore interpret the four wedges of Main Terrace-derivedsands in the hanging wall (units 4B, 4D, 5 and 7) to becolluvial wedges, which are related to degradation offreshly formed fault scarps (McCalpin, 1996 and ref-erences therein). The deposition of each wedge wouldhave been initiated by rejuvenation of the Bree faultscarp in response to a surface-faulting event and con-sequently, the base of each wedge represents an eventhorizon (Figure 8). Most colluvial wedges describedin literature so far are situated in dry or even arid re-gions, where slope processes are dominated by gravitytransport and slope wash (Wallace, 1977). Comparedto these typical examples, formation of the collu-vial wedges in Bree trench 4 was mainly governedby gravity-driven deformation of water-saturated sed-iment. This model of colluvial wedge formation isillustrated in Figure 14 and it also partly explains themore elongate shape of the wedges. This particulartype of colluvial wedge was not observed in any ofthe earlier trenches across the Bree fault scarp, wherethe Main Terrace deposits were much more gravelly.Their presence in trench 4 is probably linked to thetypical grain size (clayey coarse sands with less than30% gravel) of the footwall sediments, which is fa-

vorable for retaining large amounts of water. Buriedcolluvial wedges commonly show a soil profile attheir top, representing stabilization of the scarp slopebetween successive faulting events (McCalpin, 1996),but we suspect that periglacial conditions inhibited thedevelopment of soils on these colluvial wedges.

Our interpretation of units 4B, 4D, 5 and 7 as col-luvial wedges allowed us to identify four additionalsurface-faulting events, labeled E3–E6. Assuming thatthe downslope sediment transport continued until thefree face was completely leveled, we can use thethickness of the wedges adjacent to the main fault toestimate the amount of displacement for each event. Inthe simple model of fault-scarp degradation (Wallace,1977), the thickness of a colluvial wedge approxim-ately equals one-half of the scarp height. It is notcertain, however, if this model is appropriate for thesituation in trench 4, because the slope process in-volved deformation of an entire water-saturated layerrather than purely mechanical erosion. We thereforeconsider the thickness of the colluvial wedges as min-imal estimates of the offset produced in each event.Thus we estimate fault displacements of at least 85 cm,55 cm, 70 cm and 30 cm for events E3, E4, E5 and E6,respectively. E3 has the largest displacement and maytherefore correspond to the largest paleoearthquake re-corded in trench 4. It probably also generated somesmaller-scale faults with offsets of 10–30 cm on thelower trench wall (Figure 8, 59–61 m), adding ± 40cm to the total amount of displacement. The age ofthis event is poorly constrained between 23.4 ± 4.1 kaBP (IRSL sample just above colluvial wedge unit 7,see Figure 8) and 63.3 ± 13.1 ka BP (the most reli-able age obtained for the coversand unit 6 below thewedge, see Figure 7). We cannot rely on the radiocar-bon ages from unit 6 to determine the maximum ageof E3 because these samples are interpreted to havebeen transported and may thus have formed after thecolluvial wedge was emplaced. We do not have suffi-cient samples to constrain the ages of the older eventsE4-E6, but an IRSL sample from coversand unit 4Cindicates that E5 is younger than 136.6 ± 17.6 ka BP.The oldest event E6 is likely of Saalian age as welland thus younger than c. 185 ka, because there is noevidence of a soil corresponding to the Holsteinianinterglacial within the entire unit 4.

Secondary deformation at the base of the slopeNear the downslope end of the trench, units 4 to 8are severely deformed by a set of low-angle thrustfaults that reverse the dip of the stratification (Fig-

353

Figure 15. Photograph of harmonica-like folding of coversand lamination directly above low-angle thrust fault in unit 4. This deformationprobably resulted from overpressurization at the toe of a detachment slide, possibly triggered by an earthquake. Location is indicated in Figure7. Legend as in Figure 10.

ure 7). This deformation was also suggested by theelectrical tomography and GPR profiles (Figure 6). Inthe trench, this deformation involves tight, harmonica-like folding (‘stacking pattern’) of coversand laminaeand interbedded single-pebble gravel beds above alow-angle fault plane, whereas below this fault, strati-fication is undeformed and parallel (Figure 15); moreclay-rich intervals show abundant small-scale shears.We interpret the entire structure as a detachment slideand the stacking pattern as liquefaction due to over-pressurization at the toe of the slide. The masterdetachment plane is probably rooted at the top of theclay-rich unit 3. The overall shape and geometry ofthis deformation is remarkably similar to active-layerdetachment slides generated by present-day perigla-cial solifluction in Ellesmere Island, Canada (Harrisand Lewkowicz,1993; compare Figure 7 with theirFigure 3). Compared to the Ellesmere Island features,which are solely triggered by seasonal thawing, thedetachment slide in trench 4 is thicker (2.5 m vs. 0.2–

0.65 m), involves more sandy sediments (as opposedto clayey silt), and is situated on a gentle slope of2.5◦, which is at the lower end of failure angles inEllesmere Island (2◦–40◦, with a mean angle between5◦ and 20◦). Considering the present-day well-drainedcondition of the hanging wall, it is likely that thisdeformation originated when periglacial conditionsexisted at the trench site, similar to the modern con-ditions in Ellesmere Island. Surface sediments thenwere intermittently saturated due to seasonal thawingor snowmelt (Gullentops et al., 1993) and thus proneto slope failure. Sliding may have been triggered bysurface faulting and would, in that case, have oc-curred simultaneously with (or just before, becauseit is a more rapid process) formation of one or moreof the identified colluvial wedges. It is not possibleto corroborate this inference, but it is suggested bythe recognition of three separate phases of sliding thatcould correspond to individual events. A first deform-ation phase is shown by an unconformity truncating

354

faulted and folded strata in the upper part of unit 4(Figure 7, 81–85 m). The main sliding stage tookplace just before deposition of wedge unit 7 and ex-tensively deformed units 4, 5 and 6. The final stageof deformation was minor and is indicated by slightrotation of units 7 and 8. The Beuningen gravel bed(unit 9) is unaffected by the deformation and sharplytruncates the underlying tilted layers. Thus, the mainphase of sliding possibly correlates with the largestfaulting event, E3, while the final stage could havebeen contemporaneous with E2.

Older cumulative fault displacementsTable 2 summarizes the paleoseismic evidence andage constraints for six fault displacements in trench4. Borehole evidence indicates that the main fault intrench 4 was also active prior to deposition of theoldest colluvial wedge (unit 4B). For instance, whenthe displacements produced by faulting events E1-E6are geometrically reconstructed, the backtilted unit3, which probably was deposited in shallow water infront of the fault scarp, is still 1.2 m below the currenttop of the Main Terrace deposits in the footwall. Thefluvial sands of unit 2, which are also assumed to haveoriginally been at the same level as the top of unit 1,are found at an even greater depth. It is not possible,however, to resolve the faulting events responsible forthese cumulative displacements, or even to confirmwhether their origin was coseismic, but the amountof cumulative displacement suggests a long history ofQuaternary faulting on the Bree scarp.

Activity of the Bree fault scarp: discussion andconclusions

Trench 4 across the Bree fault scarp exposed thelongest, and perhaps the most complete, paleoseis-mic record obtained so far along this fault segment.Combining all of the information from trench walland boreholes, we found evidence for perhaps sixindividual surface-fault displacements since 150–185ka BP (Table 2). The youngest event, E1, producedonly 12 cm of fault slip (max. 78 cm including flex-ure) and is Holocene in age (<8.5 ± 1.6 ka BP); adisplaced Bt-horizon indicates that E1 may be signi-ficantly younger, but its age cannot be further con-strained. The penultimate event, E2, is constrainedbetween 15.8 ± 2.9 ka BP and deposition of a strik-ing gravel bed that is correlated to the Beuningengravel bed, which is dated as older than 14.0 ± 2.3

ka BP in the trench. Event E2 produced 33 cm offault offset and was associated with the formation ofextensional cracks and liquefaction-related deforma-tion of coversands close to the main fault. Four olderfaulting events (E3–E6) are inferred from the presenceof four wedges of footwall-derived sediment adjacentto the main fault in the hanging wall. These wedgesare interpreted as colluvial wedges, resulting from theinterplay between surface faulting and slope processesin a periglacial environment: each wedge was prob-ably deposited by destabilization of a water-saturatedsurface layer in response to a single faulting event thatrejuvenated the scarp. The thickness of these particularcolluvial wedges is a minimum estimate of the amountof fault slip. The antepenultimate event (E3) repres-ents the largest surface-faulting event, with a totaldisplacement in excess of 1 m. This event may alsohave caused failure of coversands in the hanging wallwhich are deformed by low-angle thrust faults andtight intraformational folds. The age of E3 is poorlyconstrained between 63.3 ± 13.1 ka and 23.4 ± 4.1 ka.

The events E1-E3 probably correlate with thesurface-faulting events identified in previous pa-leoseismic investigations further northwest along theBree fault scarp (Camelbeeck and Meghraoui, 1998;Meghraoui et al., 2000): trenches 1–3 documentedthe occurrence of three paleoearthquakes since 44,790yr BP; the most recent event was radiocarbon-datedbetween 610 and 890 cal AD, while the penultimatelarge earthquake occurred between approximately 14ka and 22 ka BP. This correlation suggests that E1,the age of which cannot be accurately constrained intrench 4, may indeed be of late Holocene or even his-torical age, as we suggested above. The older eventsE4–E6, which had minimum offsets of respectively55 cm, 70 cm and 30 cm, are identified for the firsttime in this study. The times of these events are poorlyconstrained due to the low resolution of the regionallate Pleistocene stratigraphy and the rather large datinguncertainties. Only two IRSL samples are available,which indicate that E4 is older than 63.3 ± 13.1 kaBP and E5 is younger than 136.6 ± 17.6 ka BP. Theevents E4, E5 and E6 can therefore at best be assignedto the early Weichselian, late Saalian and middle toearly Saalian, respectively.

The identification of colluvial wedges, the narrowage constraints for faulting event E2 and the secondarysoft-sediment deformations associated with E2 (andpossibly E3) support a coseismic origin for the faultdisplacements in trench 4. The long record of surfacefaulting interpreted from this trench is unique, partic-

355

Table 2. Summary of primary and secondary on-fault paleoseismic evidence, as defined by McCalpin and Nelson(1996), in trench 4

Faulting Estimated amount Age constraints Primary Secondary evidence

event of fault slip evidence

E1 12 cm < 8.5 ± 1.6 kyr BP Faulted strata Extensional fissures

(78 cm including < post-depositional (units 9, 10, and Small intraformational

flexure) (historical?) Bt-horizon young Bt-soil faults close to fault zone

horizon)

E2 33 cm > 14.0 ± 2.3 kyr BP Faulted strata Extensional fissures

< 15.8 ± 2.9 kyr BP (unit 8) Liquefaction close to

main fault

Possible correlation with

late phase of sliding

E3 85 cm (main fault) > 23.4 ± 4.1 kyr BP Colluvial wedge Possible correlation with

+ c. 40 cm < 63.3 ± 13.1 kyr BP (unit 7) main phase of detachment

(synthetic faults)

E4 55 cm > 63.3 ± 13.1 kyr BP Colluvial wedge

(unit 5)

E5 70 cm < 136.6 ± 17.6 kyr BP Colluvial wedge Possible correlation with

(unit 4D) early phase of sliding

E6 30 cm < ± 185 kyr Colluvial wedge

(onset Saalian)? (unit 4B)

ularly for areas characterized by low strain rates (< 1mm/yr), where recent fault activity is easily obscuredby surface processes modifying the landscape. In addi-tion, periglacial conditions existed at the site during alarge part of the time span covered by the trench record(e.g. Paulissen, 1973; Kolstrup, 1980; Gullentops etal., 1981; Vandenberghe, 1992), resulting in unusualsurficial processes. Past climatic conditions may thushave been a factor in the style of secondary deforma-tion (e.g. liquefaction of active layer) and in the wayfault scarps eroded (resulting in the particular type ofcolluvial wedges described above).

The paleoseismological information from trench 4allows us to loosely constrain some faulting paramet-ers such as slip per event, moment magnitude, averageslip rate, recurrence interval and time elapsed sincethe most recent paleoearthquake at the trench site. Theestimated amount of slip per surface-faulting event ap-pears to be quite variable (Table 2), but averages 54cm. This value is a minimum estimate. It includes theslip on some smaller synthetic faults adjacent to themain fault, but not flexure and possible lateral slip; theformer could not be sufficiently quantified (except forE1), the latter is considered to be minor. The estim-

ated displacement for E1 (12 cm measured directly onthe fault) is considerably smaller than the average andis also smaller than the most recent displacement atthe previous trench sites (Camelbeeck and Meghraoui,1998). We suggest, however, that a relatively largeamount (max. 66 cm) of flexure may have occurredat the site of trench 4 during this event (see Figure11). In contrast, the estimated displacement for E3is about twice as large as the average. Assuming anaverage fault displacement of 60 cm, a rupture lengthof 10 km and a seismogenic layer thickness of 17km, Camelbeeck and Meghraoui (1998) calculated amoment magnitude of MW ∼= 6.3 for the most re-cent surface-rupturing earthquake along the Bree faultscarp. However, estimated displacements of as muchas 125 cm in trench 4 suggest that even larger earth-quakes may have occurred on this active fault duringthe late Pleistocene.

Finally, the long paleoseismic record interpretedfrom trench 4 permits better estimates of fault slip rateand recurrence interval along the Bree fault scarp thanwere previously possible, though we emphasize thatthese estimates (Figure 16) are still very crude. Weinterpret the occurrence of five paleoearthquakes since

356

Fig

ure

16.

Plot

ofcu

mul

ativ

eam

ount

ofdi

spla

cem

ent

prod

uced

bysi

xpa

leoe

arth

quak

eson

the

Bre

efa

ult

scar

pve

rsus

time,

base

don

data

from

tren

chsi

te4.

Not

eth

atth

etim

eax

isis

expr

esse

din

stra

tigra

phic

units

onw

hich

rele

vant

datin

gre

sults

have

been

supe

rpos

edan

dis

thus

notl

inea

r.T

heav

erag

ere

turn

peri

odan

dfa

ults

lipra

teha

vebe

enca

lcul

ated

for

the

eart

hqua

kecy

cles

E5-

E1,

E5-

E3,

E3-

E1,

and

E2-

E1.

357

136.6 ± 17.6 ka BP. The elapsed time since the mostrecent event, E1, is not well constrained in trench 4,but we use the value of 1000–1350 yr determined byCamelbeeck and Meghraoui (1998). Thus, we calcu-late an average fault slip rate of 0.018–0.024 mm/yrfor the time period between events E5 and E1, asshown in Figure 16. The average recurrence inter-val for these events is 29,413–38,300 yr. Similarly,fault slip rates of 0.009–0.029 mm/yr and 0.021–0.088mm/yr and recurrence intervals of 21,300–67,450 yrand 8,975–37,700 yr are calculated for the cycles E5-E3 and E3-E1, respectively (Figure 16). Due to thelarge uncertainties, particularly in the age of E3, wecannot determine if the rate of fault activity has in-creased toward the present or not. The best constrainedseismic cycle in the record is the cycle between E2and E1, which had a duration of only 10,350–17,700yr. This inter-event time is significantly shorter thanthe average recurrence interval for the period E5-E1. However, the fault slip rate for the cycle E2-E1(0.019–0.032 mm/yr) is consistent with the averagefault slip rate over four cycles. It thus appears that sig-nificant variations in the duration of the seismic cyclecan occur, even if the fault slip rate remains more orless constant. If we assume a relatively constant faultslip rate (in the order of 0.02 mm/yr) since E5, wewould estimate that the age of event E3 is closer tothe upper bound than to the lower bound of its agebracket (see Table 2). If this is true, then the largestsurface-rupturing event would have been followed bya longer-than-average interseismic period. We notethat the elapsed time since the most recent surface-rupturing earthquake, E1, is considerably shorter thanthe shortest recurrence interval calculated for any ofthe older events. As a result, we speculate that theprobability of a surface-rupturing earthquake occur-ring on this section of the fault in the near future isrelatively low.

The average fault slip rate calculated from trench4 accounts for at least 40% of the middle-to-latePleistocene slip rate of the entire Feldbiss fault zoneestimated from the displacement of Maas River ter-races. Beerten et al. (1999a, 1999b) report a totaloffset of as much as 40 m for the base of the MiddlePleistocene Main Terrace across the entire fault zone,which corresponds to a slip rate of about 0.05 mm/yr.Our results are compatible, however, with boreholedata from a buried scarp, which is considered to bethe southeastern extension of the Bree fault scarp inthe Maas Valley (Figure 2; Paulissen et al., 1985).The base of the Weichselian Maas River terrace is dis-

placed 2–3 m on this scarp, equivalent to an averageslip rate of 0.017–0.025 mm/yr since the late Pleisto-cene. This indicates that additional deformation maybe taken up by other faults in the fault zone, such asthe Heerlerheide fault or possibly other strands of theFeldbiss fault zone that do not have a clear surfaceexpression.

In contrast to the consistent estimates of the long-term slip rates, leveling data in the Netherlands indic-ate a present-day relative uplift rate (not proven to beactual slip rate) of 0.8 ± 0.3 mm/yr across the Dutchportion of the Feldbiss fault zone during the last cen-tury (van den Berg et al., 1994). This finding is anorder of magnitude greater than our long-term faultslip rates estimated from geologic data. The discrep-ancy between these values remains unexplained, butcertainly should be objective of future studies.

Our investigation of the Bree fault scarp in north-eastern Belgium demonstrates that, in areas character-ized by low strain rates and low modern seismicity,paleoseismology can be a powerful method to identifyand characterize paleoearthquakes. As a result, it of-fers valuable information that is needed to evaluatethe seismic hazards in such areas. Evidence fromthis study shows that the slip on this 10-km-longsegment of the Feldbiss fault zone is most likely co-seismic in nature and that, although they are relativelyrare, surface-rupturing earthquakes have occurred re-peatedly in the recent geological past and will likelydo so in the future. Studies that are currently under-way along the northeastern boundary fault of the RoerValley graben will help determine if these conclusionsapply to the entire Lower Rhine graben system, whichcontains more than 150 km of potentially active faults.

Acknowledgements

This study was carried out in the framework of theenvironment and climate project PALEOSIS of theEuropean Commission (contract ENV4-CT97-0578).We thank P. Pâquet and R. Verberien for their ongo-ing support of our research at the Royal Observatoryof Belgium. Our analysis benefited from fruitful dis-cussions in the field with B. Van Vliet-Lanoë, R.Langohr, H. Mestdagh, F. Gullentops and the scientificstaff of the Summer School in Active Faulting andPaleoseismology, among whom we wish to mentionin particular K. and C.P. Rajendran for discoveringradiocarbon sample IVSE-C1. Reviews by AnthonyCrone and Giuliana d’Addezio helped improve struc-

358

ture and content of the manuscript significantly. Wealso wish to thank J. Vandenberghe for examining oursamples for the presence of pollen, M. Van Strydonckfor advising us on radiocarbon samples and carryingout the dating and Y. Zackzek for kindly making avail-able the CPT-data. The assistance of B. Bukasa, J.-M.Delinte, A. El-Bouch, H. Martin, D. Mesmaker andF. Renders with the trench labor was greatly appreci-ated. Last but not least, we are very thankful to K.Mertens for his hospitality at the Pollismolen and to J.Koopmans for allowing the excavation on his pasture.

References

Ahorner, L., 1975, Present-day stress field and seismotectonicblock movements along major fault zones in central Europe,Tectonophysics 29, 233–249.

Ahorner, L., 1983, Historical seismicity and present-day microe-arthquake activity of the Rhenish Massif, central Europe. In:Fuchs, K. et al. (eds), Plateau Uplift, Springer-Verlag, Berlin,pp. 198–221.

Ahorner, L., 1994, Fault-plane solutions and source parameters ofthe 1992 Roermond, the Netherlands, mainshock and its strongeraftershocks from regional seismic data, Geologie en Mijnbouw73(2-4), 199–214.

Alexandre, P., 1990, Les séismes en Europe Occidentale de 324 à1259: Nouveau catalogue critique, Série Géophysique, Obser-vatoire Royal de Belgique: 267 pp. (in French).

Alexandre, P., 1994, Historical seismicity of the lower Rhine andMeuse valleys from 600 to 1525: A new critical review, Geologieen Mijnbouw 73(2-4), 431–438.

Beerten, K., Brabers, P., Bosch, P. and Gullentops, F., 1999a,The passage of the Feldbiss Bundle through the Maas Valley,Aardkundige Mededelingen 9, 153–158.

Beerten, K., Vandenberghe, N, Gullentops, F. and Paulissen,E., 1999b, Technisch verslag bij de Quartairkaart van België,Vlaams Gewest, kaartblad Maaseik (18), Ministerie van deVlaamse Gemeenschap, Administratie Natuurlijke Rijkdommenen Energie, Brussel. (in Dutch).

Briquet, A., 1908, La Meuse en aval de Sittard, Bull. Soc. belgeGéol. 25, 347–385. (in French).

Camelbeeck, T. and Meghraoui, M., 1996, Large earthquakes innorthern Europe more likely than once thought, EOS, Transac-tions, AGU 77(42), 405, 409.

Camelbeeck, T. and Meghraoui, M., 1998, Geological and geophys-ical evidence for large paleoearthquakes with surface faultingin the Roer Graben (northwest Europe), Geophys. J. Int. 132,347–362.

Camelbeeck, T., van Eck, T., Pelzing, R., Ahorner, L., Loohuis,L., Haak, H.W., Hoang-Trong, P. and Hollnack, T., 1994, The1992 Roermond earthquake, the Netherlands and its aftershocks,Geologie en Mijnbouw 73(2-4), 181–197.

Demanet, D., Renardy, F., Vanneste, K., Jongmans, D., Camel-beeck, T. and Meghraoui, M., 2001, The use of geophysical pro-specting for imaging active faults in the Roer graben, Belgium,Geophysics, 66(1), 78–89.

Demyttenaere, R. and Laga, P., 1988, Breuken- en isohypsenkaartenvan het Belgisch gedeelte van de Roerdal Slenk: eerste resultatenvan een seismisch onderzoek in het gebied van Poppel-Lommel-

Maaseik, Professional Paper Belgische Geologische Dienst, 234:20 pp. (in Dutch).

Frechen, M., 1995, Lumineszenz-Datierungen der pleistozänenTierfährten von Bottrop-Welheim, Münchner Geowiss. Abh. 27,63–80. (in German).

Frechen, M., 1999, Upper Pleistocene loess stratigraphy in SouthernGermany, Quaternary Geochron. 18, 243–269.

Frechen, M., Van Vliet-Lanoë, B. and Van den Haute, P., submit-ted, The Upper Pleistocene loess record at Harmignies/Belgium– High resolution terrestrial archive of climate forcing, J. Qua-ternary Sci.

Frechen, M., Vanneste, K., Verbeeck, K., Paulissen, E. and Camel-beeck, T., in press, The deposition history of the coversandsalong the Bree fault escarpment, NE Belgium, Geologie enMijnbouw.

Gullentops, F., Paulissen, E. and Vandenberghe, J., 1981, Fossilperiglacial phenomena in NE-Belgium, Biuletyn Peryglacjalny28, 345–365.

Gullentops, F., Janssen, J. and Paulissen, E., 1993, Saalian niva-tion activity in the Bosbeek valley, NE Belgium, Geologie enMijnbouw 72, 125–130.

Harris, C., 1988, Mechanisms of mass movement in periglacial en-vironments. In: Anderson, M.G. and Richards, K.S. (eds), SlopeStability, John Wiley, pp. 531–559.

Harris, C. and Lewkowicz, A.G., 1993, Form and internal structureof active-layer detachment slides, Fosheim Peninsula, EllesmereIsland, Northwest Territories, Canada, Can. J. Earth Sci. 30(8),1708–1714.

Kasse, C., 1997, Cold-climate aeolian sand-sheet formation innorth-western Europe (c. 14–12.4 ka); a response to permafrostdegradation and increased aridity, Permafrost and PeriglacialProcesses 8, 295–311.

Kolstrup, E., 1980, Climate and stratigraphy in northwestern Europebetween 30,000 b.p. and 13,000 b.p., with special reference tothe Netherlands, Mededelingen Rijks Geologische Dienst 32(15),181–253.

Loke, M.H. and Barker R.D., 1996, Rapid least-squares inversionof apparent resistivity pseudosections by quasi-Newton method,Geophys. Prosp. 44, 131–152.

McCalpin, J.P. (ed.), 1996, Paleoseismology, Academic Press, Inc.,San Diego, 583 pp.

McCalpin, J.P. and Nelson, A.R., 1996, Introduction to paleoseis-mology. In: McCalpin, J.P. (ed.), Paleoseismology, AcademicPress, Inc., San Diego, pp. 1–32.

Meghraoui, M., Camelbeeck, T., Vanneste, K., Brondeel, M. andJongmans, D., 2000, Active faulting and paleoseismology alongthe Bree fault zone, Lower Rhine graben (Belgium), J. Geophys.Res. 105, 13,809–13,841.

Miedema, R., 1983, Amount, characteristics and significance ofclay illuviation features in Late Weichselian Meuse terraces. In:Bullock, P. and Murphy, C.P. (eds), Soil Micromorphology, Vol.II, pp. 519–530.

Owen, G., 1987, Deformation processes in unconsolidated sands.In: Jones, M.E. and Preston, R.M.F. (eds), Deformation ofSediments and Sedimentary Rocks, Geol. Soc. Spec. Publ., 29,Oxford, pp. 11–24.

Pantosti, D., Schwartz, D.P. and Valensise, G., 1993, Paleoseismo-logy along the 1980 surface rupture of the Irpina fault: implica-tions for earthquake recurrence in the southern Apennines, Italy,J. Geophys. Res. 98, 6561–6577.

Paulissen, E., 1973, De Morfologie en de Kwartairstratigrafie van deMaasvallei in Belgisch Limburg, Verh. Kon. Acad. Wetensch.,Letteren en Schone Kunsten v. België, Kl. der Wetensch. 127,266 pp. (in Dutch).

359

Paulissen, E., 1983, Les nappes alluviales et les failles quaternairesdu Plateau de Campine. In: Robaszynski, F. and Dupuis, C. (eds),Guides Géologiques Régionaux: Belgique, Masson, Paris, pp.167–170. (in French).

Paulissen, E., 1997, Quaternary morphotectonics in the Belgian partof the Roer graben, Aardk. Med. 8, 131–134.

Paulissen, E., Vandenberghe, J. and Gullentops, F., 1985, TheFeldbiss fault in the Maas valley bottom (Limburg, Belgium),Geologie en Mijnbouw 64, 79–87.

Rossa, H.G., 1986, Upper Cretaceous and Tertiary inversion tecton-ics in the western part of the Rhenish-Westphalian coal district(FRG) and in the Campine area (N. Belgium), Ann. Soc. Géol.Belgique 109, 367–410.

Smith, W.H.F. and Wessel, P., 1990, Gridding with continuouscurvature splines in tension, Geophysics 55, 293–305.

van den Berg, M.W., Groenewoud, W., Lorenz, G.K., Lubbers, P.J.,Brus, D.J. and Kroonenberg, S.B., 1994, Patterns and velocitiesof recent crustal movements in the Dutch part of the Roer Valleyrift system, Geologie en Mijnbouw 73(2-4), 157–168.

Vandenberghe, J., 1992, Cryoturbations: a sediment structuralanalysis, Permafrost and Periglacial Processes 3, 343–352.

Vanneste, K., Meghraoui, M. and Camelbeeck, T., 1999, LateQuaternary earthquake-related soft-sediment deformation alongthe Belgian portion of the Feldbiss Fault, Lower Rhine Grabensystem, Tectonophysics 309(1-4), 57–79.

Van Vliet-Lanoë, B., 1985, Frost effects in soils. In: Boardman, J.(eds), Soils and Quaternary Landscape Evolution, John Wileyand Sons Ltd., pp. 117–158.

Van Vliet-Lanoë, B., Fagnart, J.P., Langohr, R., Munaut, A., 1992,Importance de la succession des phases écologiques anciennes etactuelles dans la différenciation des sols lessivés de la couvertureloessique d’Europe occidentale: argumentation stratigraphique etarchéologique, 1992, Science du Sol 30(2), 75–93. (in French).

Wallace, R.E., 1977, Profiles and ages of young fault scarps, north-central Nevada, Geol. Soc. Am. Bull. 88, 1267–1281.

Zagwijn, W.H. and van Staalduinen, C.J., 1975, Toelichting bij Geo-logische Overzichtskaarten van Nederland, Rijks GeologischeDienst, Haarlem, 134 pp. + 5 maps. (in Dutch).

Ziegler, P.A., 1994, Cenozoic rift system of western and centralEurope: an overview, Geologie en Mijnbouw 73(2-4), 99–127.