Download Geological and geophysical evidence for large

Geophys. J. Int. (1998) 132, 347–362
Geological and geophysical evidence for large palaeo-earthquakes
with surface faulting in the Roer Graben (northwest Europe)
Thierry Camelbeeck and Mustapha Meghraoui*
Royal Observatory of Belgium, av. Circulaire 3, B-1180 Bruxelles, Belgium. E-mail: [email protected]
Accepted 1997 August 15. Received 1997 August 11; in original form 1996 December 16
SU MM A RY
From the analysis of geological, geodetic and geophysical data we provide clear
evidence of seismogenic faults capable of producing large earthquakes in intraplate
Europe. Previous studies (Paulissen, Vandenberghe & Gullentops 1985; Van den Berg
et al. 1994; Geluk et al. 1994) have yielded some constraints on the rate of crustal
deformation along the Roer Valley, a graben structure crossing the Netherlands,
Belgium and Germany, and have allowed us to address the fundamental questions: can
intraplate earthquakes rupture the surface in this part of ‘stable’ continental Europe,
and if so, what is their return period? Detailed palaeoseismic investigations have been
carried out in Belgium along a 10 km long fault scarp which is the morphological
expression of the Feldbiss Fault, the southwestern border fault of the Roer Graben
(Camelbeeck & Meghraoui 1996). The scarp is multiple and the frontal fault scarp
offsets young deposits and alluvial terraces. Field investigations using geological and
geomorphological methodologies combined with geophysical prospecting provide evidence of Holocene seismic surface faulting. From 14C dating it is suggested that the
last earthquake along the fault scarp occurred between 610 AD and 890 AD. Levelling
profiles across the scarp suggest that it produced a vertical coseismic displacement of
0.5–1 m along the scarp. If we suppose that the last surface-faulting earthquake
ruptured the whole seismogenic layer (17 km thickness) over a minimum length (10 km)
corresponding to the length of the Bree scarp with an average slip of 0.6 m, its seismic
moment was at least 3.1×1018 N m (M =6.3) for an average rigidity m=3×1010 Pa.
W
By estimating the offset of the main terrace of the Maas River by slippage along the
Feldbiss Fault, we calculate the average Late Pleistocene vertical deformation rate as
0.08±0.04 mm yr−1. Palaeoseismic information combining the trench and geomorphic
observations suggests the occurrence of two surface-faulting earthquakes during the
last 20 kyr. A third dates before 28–35 kyr BP. Then, if the time distribution of
earthquakes is uniform, a return period of 12±5 ka and a vertical deformation rate of
0.06±0.04 mm yr−1 are inferred.
Key words: crustal deformation, earthquakes, normal faulting, Roer Graben.
I NT R O DU C TI O N
Continental earthquakes are qualified as large when they
rupture the whole seismogenic layer and hence produce surface
ruptures or at least measurable surface coseismic deformation
(Scholz 1990). Compared with seismic activity along plate
boundaries, the occurrence of such earthquakes in intraplate
regions is relatively rare. Worldwide, only 11 intraplate historical earthquakes which produced coseismic surface faulting are
* Now at C.N.R.–C.S. Geologia Tecnica, via Eudossiana 18, 00184
Rome, Italy.
© 1998 RAS
known (Johnston et al. 1994), and presumably none occurred
in intraplate Europe in historic times. However, due to the
higher vulnerability of developed urban areas, widespread
damage is to be expected for densely populated regions. This
was the case for the 1993 September 29 M =6.4 Killari
S
earthquake, which struck the stable Indian continental domain
(10 000 deaths). The densely populated areas of northwest
Europe, where seismic activity is apparently low and damaging
earthquakes virtually unknown, would be at high risk from
the occurrence of an earthquake similar to the Killari event.
Although historical seismic catalogues covering the last six
centuries provide a long list of moderate earthquakes (Table 1),
intraplate Europe is often considered to be a region safe from
347
348
T . Camelbeeck and M. Meghraoui
Table 1. Main historical earthquakes (M >5.0) in central and northwestern Europe and in the northwestern part of the Alps (1350–1995). All
S
the earthquakes with estimated magnitude >4.5 in the Lower Rhine Graben region are also indicated.
N°
Date
Latitude
Longitude
1
2
3*
4
5*
6
7
8*
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27*
28
29
30
31
32
33
34
13561018
135705
13820521
13950611
14490423
15040823
15750226
15800406
16010908
16400404
16820512
16901218
16920908
17330518
17551209
17551227
17560218
17600120
18280223
18460729
18731022
18780826
19050429
19111116
19130720
19260730
19310607
19321120
19380611
19460125
19510314
19780903
19831108
19920413
47.5 N
50.0 N
51.3 N
50.6 N
51.6 N
50.8 N
53.0 N
51.0 N
46.9 N
50.8 N
48.0 N
50.8 N
50.6 N
50.0 N
46.3 N
50.8 N
50.7 N
50.7 N
50.7 N
50.1 N
50.90 N
50.93 N
45.90 N
48.22 N
48.24 N
49.17 N
54.05 N
51.71 N
50.78 N
46.30 N
50.64 N
48.30 N
50.63 N
51.16 N
7.5 E
8.3 E
2.0 E
6.3 E
2.5 E
6.1 E
2.0 W
1.5 E
8.2 E
6.1 E
6.6 E
6.1 E
5.8 E
8.3 E
7.5 E
6.2 E
6.3 E
6.3 E
5.2 E
7.7 E
6.10 E
6.55 E
7.00 E
9.05 E
9.00 E
1.62 W
1.45 E
5.61 E
3.58 E
7.50 E
6.73 E
8.93 E
5.53 E
5.92 E
M
S
M
est
>6.5
6.0
5.5
6.0
5.5
5.4
4.6
4.6
5.2
5.7
6.0
5.2
5.1
5.5
4.7
5.0
6.1
5.3
5.2
4.5
5.4
I
0
IX–X
VII
VIII
VII
VI–VII
VII
VI–VII
VII–VIII
VIII
VII
VIII
VI–VII
VIII
VII
VIII–IX
VII
VIII
VII
VII–VIII
VII
VII
VIII
VIII
VIII
VII
VI–VII
V–VI
VII
VII
VIII
VII–VIII
VII
VII
VII
Reference
Vogt (1979)
Alexandre (1994)
Melville et al. (1996)
Alexandre (1994)
Melville et al. (1996)
Alexandre (1994)
Ambraseys (personal communication, 1985)
Melville et al. (1996)
Vogt (1979)
Vogt (1994)
Vogt (1979)
Alexandre (personal communication, 1996)
Alexandre (personal communication, 1996)
Vogt (1991)
Bertrand (1757)
Alexandre & Vogt (1994)
Alexandre & Vogt (1994); Ahorner, Murawski & Schneider (1970)
Alexandre & Vogt (1994)
Ambraseys (1985b); Karnik (1969); Ahorner et al. (1970)
Ambraseys (1985b); Karnik (1969); Ahorner et al. (1970)
Ambraseys (1985b); Karnik (1969); Ahorner et al. (1970)
Ambraseys (1985b); Karnik (1969); Ahorner et al. (1970)
Karnik (1969)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Ambraseys (1985a)
Camelbeeck & van Eck (1994)
M is the estimated magnitude from macroseismic information;
est
I is the maximum observed intensity (MSK scale);
0
* epicentre on sea; the earthquakes are located and referenced, respectively, with their number and date on Figs 1(a) and ( b).
large earthquakes. In this continental interior zone between
the Alps and the North Sea (Fig. 1a), earthquakes concentrate
mainly along geological structures such as the Upper Rhine
Graben and the adjacent Vosges and Black Forest mountains,
the Swabian Jura, the Lower and Middle Rhine valleys and
the Belgian zone. These moderate-sized earthquakes have
caused occasional damage associated with intensities of I=
VIII (MSK) on the Medvedev, Sponheuer and Karnik scale,
which corresponds to moderate structural damage and heavy
non-structural damage to many masonry buildings, with supposedly no related visible coseismic surface displacements. The
well known exception of a large earthquake is the Basel
earthquake of 1356 October 18, which had an estimated
epicentral intensity of IX–X (Vogt 1979).
Some 30 years ago, Ahorner (1962) proposed a pattern of
earthquake distribution in central and northwest Europe and
inferred a detailed correlation between earthquake activity and
Quaternary structures. More recently, Vandenberghe (1982),
Van den Berg (1994), Geluk et al. (1994) and Paulissen et al.
(1985) studied the Neogene and Quaternary tectonics of the
Roer Graben. In spite of these efforts, active faults—structures
that have an established record of activity in the late
Pleistocene (in the past 125 000 years) and a possible capability
of generating major earthquakes (Boschi et al. 1996)—
remained unidentified and coseismic surface displacements
were considered to be non-existent.
The identification of active faults and the search for evidence
of past coseismic surface faulting produced by earthquakes are
important considerations for intraplate continental deformation and seismic hazard assessment in northwest Europe.
In this paper, we provide geological and geophysical evidence
of large prehistoric surface faulting of seismic origin in the
Roer Graben. Using seismotectonic analyses, we argue for
strong earthquake activity during the Late Pleistocene and the
Holocene, and for the potential of large earthquakes in the
intraplate tectonic domain of central and northwestern Europe.
Our investigation questions the traditional approach to seismic
hazard assessment based only on short-term historical and
instrumental catalogues.
R AT E O F TEC TO NI C M O VEM EN TS
The Roer Valley is a branch of the Rhine Graben system,
which belongs to the reactivated Cenozoic rift in western
Europe. The strong subsidence of the Roer Valley during the
last 150 000 years (Geluk et al. 1994), the Quaternary faults
© 1998 RAS, GJI 132, 347–362
Figure 1. (a) Seismic activity in northwest continental Europe. Instrumental (from 1900 onwards) seismicity data are taken from the International Seismological Center and the Royal Observatory
data base. Historical earthquakes with magnitude M>5 are indicated (Table 1). Main active zones of the intraplate domain are the Lower Rhine Embayment, which comprises the Roer Graben,
the Upper Rhine Graben and the Swabian Jura. Paler regions mark the central parts of rift valleys. (b) Distribution of Quaternary faults along the Roer Valley and the late Pleistocene and Holocene
deposits. Main historical earthquakes with magnitude M>4.5 (Table 1) show that fault segments can be seismogenic in this region. Note that the Bree fault scarp location is apparently in a zone
with a gap of historical seismicity.
Holocene surface faulting in the Roer Graben
© 1998 RAS, GJI 132, 347–362
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T . Camelbeeck and M. Meghraoui
and associated morphology along the flanks of the graben, and
the 0.8–2 mm yr−1 vertical rate of deformation obtained from
the comparison of levellings during the last 100 years (Van den
Berg et al. 1994; Mälzer, Hein & Zippelt 1983), combined
with the seismic activity, are the most significant elements
demonstrating recent and present-day crustal deformation.
Based on the main geological structures and Cenozoic
subsidence data, the Lower Rhine Graben can be divided into
several tectonic units (Fig. 1b): (1) the Krefeld block, which
borders the subsiding area to the northeast, (2) the Venlo
Graben and the Peel block, which have an intermediate
subsidence, (3) the Roer Graben and the Erft block, which
correspond to central valleys and sites of strong subsidence,
and (4) the Campine and South Limburg blocks of intermediate
subsidence and the Brabant block, bordering the subsiding
area to the southwest.
A detailed map of Bouguer gravity anomalies (C. Poitevin
& M. Everaerts, personal communication, 1995) reflects this
structural subdivision and shows clear limits of the different units
(Fig. 2a). The area with the largest negative anomaly correlates
remarkably well with the Roer Graben and corresponds to the
most important region of post-Miocene subsidence (Geluk
et al. 1994). The Roer Graben is bordered by two main NNW–
SSE-trending Quaternary normal fault systems, the Peel
Boundary Fault to the east and the Feldbiss fault zone to the
west. The results of modelling the gravimetric anomaly at the
western border of the graben using the method of Rasmussen
& Pedersen (1979) are in agreement with the thickness of
sediment deposited in the graben on the Palaeozoic rocks
forming the basement, which is known from wells and seismic
profiles (Fig. 3). To improve the fit between the observations
and the modelling, it would be necessary to take into account
the lateral extension of the fault zone as deduced for example
from the analysis of Demyttenaere (1989) or Demyttenaere &
Laga (1988). On the other hand, the resolution of the seismic
profiles is better for obtaining information about deposits in
the graben and fault throws than the gravity profiles, which
have one value every kilometre.
From deep seismic profiles in the Netherlands (Geluk 1990),
the Moho depth below the graben is about 28 km, which is
low compared with the values of 30–31 km determined on the
shoulders, indicating crustal thinning of 2–3 km under the
graben. The crust is thus clearly stretched, and this stretching
is related to the strong Neogene and Quaternary subsidence
and extensional tectonics.
The repartitioning of the seismic activity with depth across
the Roer Graben during the years 1980–1995 is presented in
Fig. 4. The hypocentres range from depths of about 5 to 20 km
and are clearly bounded by straight lines which are well
correlated near the surface with the location and the geometry
of the two border faults. The recent seismic activity linked to
the Peel Boundary Fault is more significant and corresponds
to the M 5.4 Roermond earthquake (1992 April 13) and its
S
aftershocks.
These data suggest that the crustal thinning can be interpreted as lithospheric extension accommodated by faulting
along large extensional faults in the upper brittle crust, which
extends here down to a depth of 20 km, and pure shear in the
lower crust and mantle (see for example the models of Kusznir,
Marsden & Egan 1991).
As is the case for many other examples (see Roberts &
Yielding 1994) in continental extensional regions, the earth-
quake data, particularly that concerning the Roermond earthquake and its aftershocks (Ahorner 1994; Camelbeeck & van
Eck 1994), suggest also that the border faults of the Roer
Graben are planar faults throughout the seismogenic layer.
Their dip is about 60° to 70°.
There is, of course, an information gap between the model
in Fig. 3( b), which extends only to 2 km depth, and the onset
of seismic activity in Fig. 4 at about 5 km depth. Nevertheless,
the similarity of the geometry, the fault slip orientation in the
seismogenic layer, the location of the faults and the evidence
of active lithospheric extension suggest that there is no reason
to think that the observed fault throws near the surface are
not the superficial expression of large crustal faults. On the
other hand, as it is suggested by Scholz (1990), the existence
of an upper cut-off in seismicity seems limited to well-developed
fault zones, which should be the case for the Feldbiss Fault, as
shown in the following section. This top of the seismogenic
layer delimits the upper region in which earthquakes can
nucleate, but of course does not bound the region in which
they can propagate (Scholz 1990).
For the Feldbiss fault zone, which is partly located in
Belgium, the recent tectonic activity is mainly indicated at
depth by seismic profiles (Fig. 5), showing an offset of Neogene
deposits of more than 500 m (Demyttenaere & Laga 1988).
On the basis of the same profiles, we interpret the vertical
relative displacement of units at the base of the Quaternary as
being 100–120 m. The reflector bQ (=±base Quaternary) in
Fig. 5 is tentatively correlated to a horizon inside the Mol
formation (De Batist & Versteeg 1998), which is a continental
sandy unit with a Late Pliocene–Quaternary transition age.
The measured offset corresponds to an average Quaternary
rate of deformation of 0.050±0.015 mm yr−1.
The Roer Graben subsided from the Middle Miocene
onwards, and experienced an increase in tectonic activity
around 3 Ma (Geluk et al. 1994); an average long-term uplift
of 0.06 mm yr−1 in the adjacent South Limburg block was
calculated from fluvial terraces of the Maas River ( Van den
Berg 1994) (Fig. 6a).
Based on a study of the Maas Valley geomorphology
associated with an electrical prospecting interpretation,
Paulissen et al. (1985) indicated that the most recent
Quaternary activity along the southern border of the graben
occurred along the Feldbiss Fault and concluded that the most
vertical recent movements were 2–3 m during the Weichselian,
and 8–10 m since the formation of the youngest Saalian
terraces. The vertical offsets correspond to average deformation
rates of, respectively, 0.05±0.03 mm yr−1 for the last 80 000
years and 0.06±0.02 mm yr−1 for the last 150 000 years. We
interpret these as minimum values because it is likely that the
Maas River crosses the Feldbiss Fault in a transfer zone
corresponding to a lower footwall elevation. This hypothesis
is based on the three following observations.
(1) The sudden change of the Maas River course at the
Tertiary–Quaternary limit from the Eastern Maas to the
present course was probably caused by tectonic activity along
the Feldbiss Fault in the Netherlands. This activity is clearly
suggested by the data of Juvigné & Renard (1992) concerning
the Maas terraces between Liège and Maastricht.
(2) The main terrace of the Maas River, which formed
during the Cromerian (>400 ka) and now forms the top of the
uplifted block (Fig. 7), is displaced vertically by the Feldbiss
© 1998 RAS, GJI 132, 347–362
Figure 2. (a) Map of Bouguer gravity anomalies showing the main zones of subsidence (negative anomaly) along the Roer Graben, and the related deep-seated fault distribution that limits the
active basins. Location of gravity profile in Fig. 3 is indicated in white. (b) Seismotectonic map of the Roer Graben and surrounding areas. Focal mechanisms are compiled from the work of
Camelbeeck & van Eck (1994).
Holocene surface faulting in the Roer Graben
© 1998 RAS, GJI 132, 347–362
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T . Camelbeeck and M. Meghraoui
Figure 3. Modelling of the Bouguer gravimetric anomaly across the Feldbiss Fault. (a) Bouguer gravimetric anomaly observed and calculated
with the geological model presented in ( b) by the method of Rasmussen & Pedersen (1979). ( b) Geological model of Demyttenaere & Laga (1988)
for the limit between the Campine Block and the Roer Graben. Density values appropriate for this region are taken from De Sitter (1949).
(Mälzer et al. 1983) show an even higher vertical deformation
rate of the order of 1–2 mm yr−1.
In comparison to the deformation rates along the few studied
active faults in the continental interior of Australia and the
United States, the Late Pleistocene rate of vertical deformation
in the Roer Graben is one order of magnitude greater (Crone,
Machette & Bowman 1992, 1997a) and the seismic activity
appears to be representative of ongoing crustal movements
(Fig. 6b). Because seismogenic faults showing historic or prehistoric ruptures are known even in stable cratons (for
example the Australian shield) (Johnston et al. 1994; Crone
et al. 1997a,b), we believe that earthquakes along the Roer
Graben system showing coseismic surface faulting are probable
and that their signature can be documented by geological and
geomorphological analyses.
Figure 4. Depth distribution of the seismic activity in the Roer
Graben during the years 1980–1995. The hypocentres are plotted
following a section across the graben.
Fault by an estimated distance of 20±5 m. From the analysis
of Geluk et al. (1994), it is suggested that fault activity was
not continuous during this period of time and, as illustrated
in Fig. 6(a), that during the period 700–150 ka the subsidence
in the Roer Graben was interrupted. It is thus possible that
the main terrace has been displaced only during the last
150 kyr. Considering deformation of the main terrace to have
occurred between 400 and 150 ka, we infer a vertical deformation rate of 0.08±0.04 mm yr−1 greater than that estimated
in the Maas Valley bottom.
(3) Paulissen et al. (1985) did not observe any Holocene
movement, in contrast to our observations along the Bree
fault scarp.
In the Herten and Linne borehole in the Roer Graben near
Roermond, Geluk et al. (1994) found a subsidence of 25 m
during the last 150 kyr and of 15 m during the last 50 kyr,
which suggests a subsidence rate of 0.15–0.3 mm yr−1 for the
Late Pleistocene and Holocene (Fig. 6a). These late Quaternary
rates of vertical movement are comparable to the minimum
geodetic vertical rates of 0.8 mm yr−1 (±0.5 mm yr−1) measured between the Roer Graben and the South Limburg block.
Furthermore, in the southern part of the Roer graben, at the
border with the Rhenish Massif, geodetic measurements
PA LA E O SE IS M I C EVI DEN CE A L ON G T HE
BR E E FA U LT S C A R P
Repeated coseismic displacements along a fault that may
expose young deposits allow past earthquakes to be geologically recorded and display a typical morphological expression.
In the region of Bree (west flank of the Roer Graben, Fig. 1b),
along the Feldbiss fault zone, a NNW–SSE-trending fault
scarp forms a gentle but noticeable morphology. The Bree
fault separates the Campine Plateau (to the west) from the
Roer Valley (to the east), and its geomorphic signature consists
of a 10 km long sharp scarp slope with 15–20 m of topographic
offset (Fig. 7a). The scarp has already been recognized as a
Quaternary fault scarp, but it was, until now, considered
inactive and mainly attributed to erosional processes, even
though Paulissen (1973) described the scarp as displacing the
main terrace of the Maas River. The Belgian Geological Survey
executed 150 reflection seismic lines in the region for a total
distance of 1000 km (Demyttenaere 1989), of which a dozen
crossed the scarp. On different sections (Fig. 5), the scarp
coincides at the surface with the prolongation of the Feldbiss
fault zone and must be considered as the morphological
expression of the fault’s recent activity.
The Bree fault scarp affects the main terrace of the Maas
River formed during the Cromerian ( between 700 and 400 ka).
It corresponds to the northeastern border of the Plateau
Campine, which is the uplifted block of this terrace and consists
© 1998 RAS, GJI 132, 347–362
Holocene surface faulting in the Roer Graben
353
Figure 5. Seismic profile (and corresponding topography) produced by the Belgian Geological Survey across the Feldbiss Fault. Stratigraphy
slightly modified from Demyttenaere & Laga (1988): bQ=±base Quaternary; bPli=base Pliocene; bMi=base Miocene; bTe=base Tertiary.
of terrace gravels deposited by the Maas River (Zutendaal
gravels). In the downthrown block, the Zutendaal gravels have
been eroded by the Rhine, which afterwards deposited the
Bocholt sands (Paulissen 1983). These formations are the
basement on which the Maas subsequently formed its different
terraces, which are the typical landscape of the region. The
region was later covered by aeolian sands during the Riss and
Würmian (Weichselian) glacial ages, and Holocene sediments
are mainly alluvial deposits.
Detailed field work and geodetic levelling of 36 profiles
along strike, combined with the examination of aerial photographs, has revealed the existence, for a distance of greater
than 10 km, of a frontal fault scarp with 0.5–3 m of vertical
offset (Fig. 7b) in the aeolian coversands attributed to the Late
Weichselian (Paulissen 1973). Its presence can be interpreted
as a primary geomorphic expression of active faulting (Richter
1958; Stewart & Hancock 1994). The geomorphic analysis
requires that profiles must follow the line of steepest slope so
as to be parallel to the transport material across the scarp;
consequently, their strikes may vary instead of being strictly
perpendicular to the fault trend. By comparison with known
active fault scarps in arid or semi-arid climates, the preservation
of the frontal morphology of the Bree fault scarp in a humid
climate where the erosion rates are greater is an indication of
its youthfulness. At a site where a small river crosses the scarp,
the frontal scarp corresponds to the offset of the river alluvial
terrace young deposit. Owing to successive uplifts of the
terrace, the river channel has deviated from its original bed,
formed a swampy zone, and finally cut through the young
deposits. The morphology of the deformed terrace (Fig. 8) is a
reliable indicator of the recent tectonic activity along the fault
associated with palaeo-earthquakes.
(1) The river terrace, like the whole frontal scarp, is a
compound scarp, containing two or three breaks in slope, each
© 1998 RAS, GJI 132, 347–362
of which originated in a separate rupture event followed by
intervals of erosion.
(2) The formation of a swampy area on the upthrown block,
with deposits identical to that on the downthrown block, is
explained by the sudden tilting of the uplifted block during
each earthquake, which creates a depression retaining water
on the upper block. It should be noted that the situation is
similar at the other two sites where the scarp intersects
river beds.
(3) Stream incision into the upthrown block is clearly visible
(Fig. 8a) and is expected after faulting, because the base level
of the stream is suddenly lowered. On the other hand, stream
incision is very small in the downthrown block.
As the frontal scarp does not correspond to a contact between
formations and is well visible for a distance of more than
10 km, it cannot be due to erosion, local sliding or soil
consolidation. According to the levelling profile analysis
(Meghraoui et al. 1997) it is likely that the last event produced
a vertical offset ranging from 0.5 to 1.0 m (Fig. 7).
The fault offsets young deposits and alluvial terraces, and
shows a left-lateral en echelon pattern, suggesting extension of
the Roer Graben in a NNE–SSW direction, slightly oblique
to the graben trend. The earthquake focal mechanisms in the
graben (Fig. 2b) are mostly dip-slip normal faulting along
NW–SE-striking faults. The mechanism of the M 5.4
S
Roermond earthquake of 1992 April 13 indicates SW–NE
extension along a 70° dipping fault. There is a 20° discrepancy
in the direction of extension between our surface-faulting
observations and the focal mechanisms, but they represent the
deformation at different spatial scales and also along different
fault zones in the Roer Graben. The largest recorded earthquake along this section of the southwestern limit of the Roer
Graben is the M =3.7 1960 June 25, event and there is no
S
evidence for larger historical events within the last 400 years.
354
T . Camelbeeck and M. Meghraoui
Figure 6. (a) Rates of deformation obtained from geodetic levelling and geological data for the Roer Graben structure. The average long-term
uplift rate obtained by Van den Berg et al. (1994) is 0.06 mm yr−1. Geluk et al. (1994) suggested that it increased during the last 150 kyr and
reaches 0.2–0.3 mm yr−1. Repeated geodetic measurements during the last century yielded a maximum rate of 0.8 mm yr−1 ( Van den Berg et al.
1994). ( b) Comparative rates of vertical movements showing that the Bree fault scarp is amongst the most active structures of the continental
interior (as compared with intraplate zones in the central United States (Olig et al. 1994; Crone et al. 1997b).
Two trenches about 100 m in length were excavated across
the Bree fault scarp in order to investigate recent faulting
activity and related palaeo-earthquakes. Because of the large
number of observations, we present here the most important
part of the data; a palaeoseismological analysis that includes
a detailed geological description of trench exposures combined
with a geophysical and geomorphic analysis (including the
scarp age determination) is the subject of another paper
(Meghraoui et al. 1998).
In order to locate the fault precisely in the subsurface and
to extend the depth (>3 m) of the investigation, geophysical
methods including electric prospecting, electric tomography,
seismic refraction and radar measurements were performed.
These methods were chosen because they can provide information at greater depths than trench excavation. All methods
delineated the fault location (see electric profile, seismic
refraction and radar in Fig. 9). The contrast of physical
parameters (electrical resistivity and permittivity and seismic
velocity) observed between the fault blocks results from a
difference in the nature of the soils and sedimentary layers and
from a change in the water table from the upper block to the
downthrown block. The ground-penetrating radar was cor-
rected for topography with a mean velocity of 90 mm ns−1. It
was able to reveal the structure up to 4–5 m deep in some
parts of the profiles. It suggests that the sedimentary layer 2 m
deep near trench I (Fig. 9) is faulted and flexed, giving a total
vertical displacement of about 1 m, which agrees with the total
vertical offset in the geomorphic profile (Figs 7b and 8b). The
successive coseismic displacements are thus preserved in both
the morphology and the young sediments deposited by the
river.
In trench I (Fig. 10a), despite the presence of a superficial
water level at 2.5 m depth, two well-sorted gravel horizons (e
and g) intercalated with sandy clay deposits show 0.5 m of net
vertical offset, which gives a picture of the last earthquake
along the scarp.
Due to the fault scarp morphology and the offset units and
palaeosoils, no geological process other than coseismic surface
faulting can explain the block deformation of these sediments
(see also Mc Calpin 1996). Two observations in the immediate
surroundings of the fault zone provide additional evidence of
a sudden displacement: the presence of opened fissures only in
the proximity of the scarp and the deformation (and wedging
out) of palaeosol c in the trench.
© 1998 RAS, GJI 132, 347–362
Holocene surface faulting in the Roer Graben
355
Figure 7. (a) Morphotectonic map showing the main section of the Bree fault scarp, the site of trench investigations where nine levelling profiles
have been measured and the location of 27 additional levelling profiles (indicated in white) across the scarp. The geomorphic expression is
delineated by a maximum topographic offset of 15–20 m, but the fault also shows a frontal scarp of 0.5–3 m along strike. (b)–(e) Interpretation of
levelling profiles in terms of vertical offsets caused by surface-rupturing palaeo-earthquakes (for detailed explanations, see for example Mc Calpin
1996). Profiles ( b) and (c) are located on the small alluvial terrace in site I; arrows on map mark the locations of profiles (d) and (e). Note that
profile elevations are only relative, and that offset values are cumulative from bottom to top.
The trench observation is in good agreement with the
geomorphic measurements of the fault scarp, indicating the
existence at this site of at least two tilted and displaced old
surfaces that are probably due to previous coseismic displacements (Figs 7b and 8b). Palaeosoils and sedimentary deposits
with fairly high organic and carbon contents allowed 14C
dating [Accelerator Mass Spectrometry (AMS) calibrated 2s]
yielding a date for the last coseismic event ranging between
610 and 890 AD. The historical seismic catalogue (Alexandre
1990) mentions notable seismic activity between 782 and
836 AD in the region, with a large earthquake in December
803 AD reported in Aachen (Germany), which caused great
concern in Germany, Belgium and as far away as France.
However, historical information about earthquakes prior to
the 16th century needs further investigation, and the last
coseismic displacement along the Bree active fault, which could
correspond to the 803 AD historical seismic event, can be more
precisely determined with the subsequent trenching campaign.
The second trench was excavated across the fault scarp at a
distance of 300 m northwest of the first trench (Figs 7a and
10b). Here also the precise position of the fault zone has been
obtained by a geophysical prospecting campaign. Because the
upthrown block is formed of Middle Pleistocene sediments
(Paulissen 1973, 1983) and the fault separated these sediments
with supposedly reworked Late Pleistocene and Holocene
sediments, it is impossible to observe the relative offset of the
stratigraphical units of the upthrown and downthrown blocks.
In this excavation (Fig. 10b), the fault cuts through a succession of coarse gravels (g1) with an intercalated clay horizon
(g2) and a sandy unit (g3) showing a prominent flexure, and
© 1998 RAS, GJI 132, 347–362
faulted colluvial wedge deposits (e1, e2, d). At the near surface,
a unit of sandy gravel lenses (dated at 402–206 BC) is cut by
the fault, showing a vertical offset of 0.7 m, and juxtaposed
against a dark brown fine sandy clay b, representing the last
colluvium at this place. From observations of the fault zone in
three dimensions, a dip of 70°–75° has been determined.
In the downthrown block, apart from the observations of
surface faulting, additional soft-sediment deformation induced
by strong ground motion is present. Fine sands and clay units,
into which admixed coarse gravels and sandy clay sediments
form channel incisions, display minor normal faulting and
prominent liquefaction features (Fig. 11). Some of these features could also be interpreted by a mechanism other than a
seismic one. 14C dating (AMS calibrated 2s) of organic and
carbon contents of sedimentary deposits at a depth of 2 m in
the trench yield ages ranging from 34.8 to 27.9 ka. These loamy
beds, which contain peat, may correspond to the Denekamp
interstadial (Kolstrup 1980). Their presence in the trench
allows us to identify the different sedimentary units deposited
after them.
Based on the observed induced deformation, we suggest the
occurrence of two large earthquakes after the deposition of
the ‘Older Coversand I’ and a palaeosoil formed in river
deposits, which are older than about 20 kyr. A third large
earthquake dates before the Denekamp interstadial. If the time
distribution of earthquakes is presumed uniform, an average
return period of 12±5 ka and an average vertical deformation
rate of 0.06±0.04 mm yr−1 are inferred.
The palaeoseismic analysis incorporating trench, geophysical
and geomorphic observations reveals clear evidence of surface
356
T . Camelbeeck and M. Meghraoui
Figure 8. (a) Digital terrain illustration near trench I indicating the uplifted alluvial terrace with respect to the fault location in the trench exposure
(see also Fig. 7a). ( b) Levelling profiles near trench I showing the successive uplifted alluvial units in trenches and the related fault scarp. The same
geomorphic pattern is visible along 36 profiles levelled across the 10 km length of the fault and showing vertical displacements ranging from 0.5
to 3.0 m. The frontal scarp and related displaced surface S1 with 0.6 m of uplift correspond to the last coseismic movement along the fault observed
in the trenches. The uplifted surface S2 reflects the cumulative movement.
faulting with Holocene (probably historical ) coseismic displacements, and cumulative offsets probably due to older large
earthquakes.
From our observations of the colluvial wedges in trench
II and of the block deformation in trench I, it is supposed
that coseismic slip is larger than interseismic slip and that
the major part of the measured vertical offset is caused by
palaeo-earthquakes.
S E IS M IC I TY A N D FA ULTI N G
PAR A M E TER S
The instrumental and historical seismic activity is not restricted
to the Roer Graben but is rather unevenly distributed over an
active region west of the Roer Graben bordered by a region
with very low seismic activity in the eastern part of the Rhenish
Massif (Table 1 and Figs 1 and 2b).
Amongst the most important recent seismic events, the Liège
earthquake of 1983 November 8 (M =4.6) occurred at the
S
limit between the Brabant and the Ardenne massifs, outside
the graben, and had a dextral strike-slip mechanism along an
ENE–WSW-trending fault with a thrust component at a depth
of 4–6 km (Fig. 2b, Table 1). More recently, the Roermond
earthquake of 1992 April 13 (M =5.4) occurred along the Peel
S
Boundary Fault and displayed a pure normal faulting mechanism on a NW–SE-trending fault; it originated at the base of
the brittle crust at a depth of 17 km (Camelbeeck & van Eck
1994). Two other moderate-sized earthquakes occurred this
century in the Roer Graben, the Uden event of November
1932 (M =4.7) and the Euskirchen earthquake of
S
1951 March 14 (M =5.3) (Fig. 1b). From the 17th century
S
onwards, the historical seismicity provides numerous data on
the damage distribution and related effects (Alexandre 1990;
Alexandre & Vogt 1994), which allows an estimation of seismic
parameters. Within this epoch, five earthquakes with M>5
affected the graben; amongst them (Table 1, Fig. 1b), the Düren
earthquake of 1756 February 18, with an attributed magnitude
M =5.5, was designated as the strongest. A recent examinaS
tion of historical documents provides evidence that the
1692 September 8 earthquake, which struck the northern part
of the Belgian Ardennes (Fig. 1b) and caused the destruction
of castles and churches in the epicentral area and damage as
far away as England, is perhaps the largest known event in
the region (P. Alexandre, personal communication, 1996).
© 1998 RAS, GJI 132, 347–362
Holocene surface faulting in the Roer Graben
357
Figure 9. (a) Geophysical measurements (geoelectricity and seismic refraction) near trench I. (b) Simplified stratigraphy of trench I, showing offset
of gravel horizons (indicated with arrows) by surface faulting. (c) Radar profile along trench I, showing faulting of sedimentary strata at depth.
The level indicated with arrows probably corresponds to lower gravel horizon g.
Seismic activity in the Roer Graben during the period
1985–1995 indicates that the seismogenic layer concerns the
upper 15–20 km of the crust (Fig. 4). The Roermond earthquake, with a seismic moment of 1.4×1017 N m and an
average displacement of 0.3 m over a normal fault plane of
approximately 11 km2, occurred at the base of the seismogenic
layer. This indicates that it ruptured only a small part of the
brittle crust and hence that this largest recorded instrumental
seismic event has to be considered as moderate with respect
to those involving a larger fault plane and displacement at the
surface. Thus, for an earthquake to rupture the whole seismogenic layer in the Roer Graben, the seismic moment must be at
least 2–4×1018 N m (M =6.1–6.4), with a minimum fault
W
length of 10–15 km. On the other hand, this may also explain
why the known historical and instrumental activity alone
© 1998 RAS, GJI 132, 347–362
cannot account for the present-day deformation rate as well
as for the rates obtained from geological data. Considering the
known seismicity in the graben since 1600 AD and the relationship established by Ahorner (1983) between magnitude and
seismic moment, the frequency–moment statistics representing the annual cumulative number of earthquakes N with
seismic moment greater than or equal to M is log N=
0
9.43–0.67 log M (in N m). Assuming that the seismicity distri0
bution is associated with the relative motion across simple
crustal boundaries, the vertical deformation rate should be
only 0.02 mm yr−1 if the maximum seismic moment is
2×1017 N m (M =5.5). Ahorner (1975) pointed out this
W
discrepancy between seismic and observed deformation rates
but he speculated that a large part of the deformation is
possibly due to aseismic creep movements. Our palaeoseismic
358
T . Camelbeeck and M. Meghraoui
Figure 10. (a) Central section of trench I (see Figs 7a and 8a for location). Symbols as in Fig. 9. ( b) Holocene alluvial deposits are affected by the
normal fault and 0.5 m of vertical offset can be measured from the visibly displaced units g and e (coarse and fine gravel horizons, respectively).
The trench bottom corresponds to the water level. Units f and d are sandy deposits with channel incisions and minor faults. Palaeosoil c, which
wedges out against the fault, has registered the last faulting event. The fault and fissures were subsequently penetrated by roots and covered by
the b and b∞ units ( brown and reddish palaeosoils). Calibrated 14C dates (AMS, 2s) of charcoal elements provide a time range for the last coseismic
displacement. ( b) Upper section of trench II (location in Fig. 7a), displaying two clear fault branches X and XX. The last three faulting events can
be retrieved from the colluvial wedges e, d and b. In the downthrown part (with respect to X), alluvial deposits are mainly composed of coarse
gravels, except for palaeosoils e1 and a. In the upthrown block, the coarse gravel units g of Middle to Late Pleistocene (Cromerian) age are flexed
and faulted.
analysis provides evidences that a major part of the deformation rates in the Roer Graben along the Feldbiss Fault are
caused by the occurrence of large earthquakes, even though
aseismic movements along faults cannot be a priori excluded
everywhere. To explain the average Late Pleistocene–Holocene
vertical deformation rates of 0.08±0.04 mm yr−1 with the
above-mentioned frequency–moment statistics, the occurrence
of earthquakes with maximal magnitude M =6.8±0.3 correw
0.7
sponding to a seismic moment of 2×1019 N m is necessary
and the calculated return period of such an earthquake in the
graben is around 3160±2140 years. These estimations are rough
2570
due to the uncertainties in the calculations of the seismic
moment of the earthquakes. For historical earthquakes, it is a
difficult task, mainly due to incomplete macroseismic information. For instrumental data, seismic moment is often estimated from magnitude measurements that lead to systematic
errors due to the fact that the relationship between the
two parameters is not linear and hence events smaller than
M =5.0 are underestimated in deformation-rate evaluation
S
(Johnston 1996a).
IM P LI C AT IO N S FO R SE IS M I C H A ZA R D
With a thickness of the seismogenic layer of 17 km and an
average slip of 0.6 m along the 10 km length of the Bree active
fault scarp, the seismic moment of the last identified earthquake
is estimated to be a minimum of 3.1×1018 N m, from which
a moment magnitude of M =6.3 is inferred. We assume that
W
this last surface rupturing crisis consisted of a single earthquake; this is favoured by the observations in the two trenches
and by the relationship between surface rupture length and
slip for the few observed rupturing earthquakes in stable
continental regions (Johnston et al. 1994). The occurrence of
two slightly smaller earthquakes (M =6.0–6.1) clustered in a
W
very short period of time is also possible.
The length of the fault section which ruptured during a
single earthquake is fundamental for the assessment of its
magnitude. In our calculation, we used the value of 10 km
from field observations of the frontal fault scarp. This value is
a minimum because the fault scarp which separates the
Campine Plateau from the Roer Graben disappears to the
© 1998 RAS, GJI 132, 347–362
Holocene surface faulting in the Roer Graben
359
(a)
(b)
Figure 11. (a) Minor reverse faulting structure associated with liquefaction phenomenon in sandy deposits of the downthrown block (scale in
centimetres). (b) Antithetic normal faults affecting sandy deposits in the downthrown block close to the main fault of trench II (scale=50 cm).
southeast where it crosses the Maas Valley, but becomes visible
again in the Netherlands, reaching a total length of 35 km.
Near Bree, this scarp shows a reduced topographic offset
and according to the gravimetric anomaly map (Fig. 2a) its
© 1998 RAS, GJI 132, 347–362
direction changes and it divides into several fault scarps
disappearing to the north close to the Dutch border. On
seismic profiles, the fault is clearly visible (Demyttenaere &
Laga 1988) and divides into three branches that are visible in
360
T . Camelbeeck and M. Meghraoui
the morphology following the topographic slope gradients. If
we consider a rupture length of 35 km, the seismic moment
would be 1.1×1019 N m and the related magnitude M =6.6.
W
Using the relationships established by Wells & Coppersmith
(1994) for normal faulting earthquakes, the following magnitude values have been obtained.
(1) Considering a surface rupture length of 10 km, M =
S
6.18±0.34. This value has to be considered a minimum, as
discussed above.
(2) Considering a maximum slip displacement of 1 m,
M =7.32±0.34.
S
(3) Considering an average surface displacement of 0.6 m,
M =6.63±0.33.
S
The duration of the seismic cycle and recent deformation
rates are critical for the seismic hazard assessment. The elapsed
time since the last event ranges from 1110 to 1390 years. The
extrapolated geodetic rates of 0.8 mm yr−1 imply a period of
1250 years to generate 1 m of interseismic vertical deformation,
which is the maximum value measured along the Bree frontal
scarp for the last earthquake. On the other hand, the Late
Pleistocene rates measured by Geluk et al. (1994) give a value
of 0.15–0.3 mm yr−1 and correspond to an average return
period of 3000 years for a large earthquake (M >6.3).
w
Analysis of the earthquake-related soft-sediment deformation in trench II suggests the occurrence of at least two
events in the last 20 ka and of a third event just before
35–28 ka, from which an average return period of 12 000±5000
years is calculated.
With a magnitude M =6.3 for the last large earthquake
W
along the Bree fault, the frequency–moment statistics calculated
for the Lower Rhine Embayment suggest an average return
period of 1080 years for a similar event in the whole graben
and an average return period along the Bree fault of about
30 kyr. This value, between two and four times the return
period suggested by the Bree scarp study, questions the extrapolation of the regional frequency–moment statistics to estimate recurrence times of large earthquakes in the Roer Graben.
It is of course difficult to discuss this problem without having
a minimum of information about the seismogenic behaviour
of the other Quaternary faults in the Roer Graben.
The identification of surface faulting with repeated coseismic
displacements in the Roer Graben and continental Europe is
unique, and the importance of these observations lies in their
implications for the reliability of seismic hazard assessment for
the region. The map of Rosenhauer & Ahorner (1994) is often
used as a basis for estimating the seismic hazard of the Lower
Rhine Embayment and neighbouring regions. For the Bree
area, the map indicates a maximum intensity of 7.25–7.5 (MSK
scale) for a return period of 10 kyr. According to the intensity–
magnitude relationships established for the continental interior
(Johnston 1996b), the epicentral intensity of the Last Bree
event might be I=9 for M =6.4. The maximum intensities
W
proposed previously are thus probably underestimated for this
region and for other regions in ‘stable continental Europe’.
CON CLU SION S
We have presented numerous lines of evidence of Holocene
and Late Pleistocene surface faulting of seismic origin along
the Bree fault scarp, which is the morphological expression of
the Feldbiss Fault, the southwestern border fault of the Roer
Graben. Primary evidence produced by coseismic slip along
the fault resulted from a detailed geomorphic analysis and
from a study of the stratigraphy in two trenches excavated in
different geological contexts. Trench 1 cuts the Holocene
deposits of a small river crossing the scarp, in which the offset
created by the last earthquakes is measurable, whereas trench
II cuts middle Pleistocene deposits where surface faulting and
the existence of colluvial wedges are the expression of recent
palaeo-earthquakes. The concordance of the surface faulting
in the stratigraphy with the frontal fault scarp in the morphology is clearly demonstrated and indicates the extension
of surface faulting to at least the 10 km length of the visible
frontal scarp. It has been proved that the Feldbiss Fault is an
active fault, and we studied it with a methodology identical to
that used to study active faults in active zones worldwide. We
believe that seismogenic faults in continental Europe have a
behaviour comparable with other active structures in continental domains. The only difference in the mechanical behaviour
of seismogenic faults in Europe is the duration of the seismic
cycle. Our experiment suggests that it probably ranges on
average from 7 to 17 kyr along the Bree fault scarp, but we also
believe that further palaeoseismic investigations are necessary
to obtain better estimations of past earthquake parameters. In
contrast to the few other active faults studied in stable continental regions (Crone et al. 1992; Crone et al. 1997a,b), showing
a pattern of temporal clustering of earthquakes with one or
several events in a short period of time (10–20 kyr) followed
by long intervals (>100 kyr) of relatively low activity, the
tectonic activity in the Roer Graben seems relatively stable, at
least for the last 150 kyr.
The existence of earthquake-related surface ruptures in the
Roer Graben also provides a unique opportunity to characterize fault behaviour in the continental interior. Many potentially hazardous faults in stable continental regions have subtle
geomorphic expressions which are often poorly preserved,
making them less conspicuous. Nevertheless, geological and
geophysical data lead us to suggest that the surface faulting
identified in the Roer Graben is not uncommon and that
similar investigations could be undertaken in other areas
between the Alps and the North Sea. For instance, other
regions with large earthquake potential may exist in the Upper
Rhine Graben, because they show clear Quaternary and recent
tectonic activity with a total subsidence of 350 m during the
Quaternary—a value twice that measured in the Roer
Graben—and a present-day vertical deformation rate of
0.5 mm yr–1 (Zippelt & Mälzer 1981).
A CKN O W LE DG M ENT S
We are grateful to A. Crone and E. Paulissen for fruitful
discussions in the field, to N. Ambraseys, W. Aspinall and K.
Bonjer, who provided useful comments on a first version of
the paper, and to G. Bock, who helped with comments on the
preparation of the revised version. We thank D. Jongmans and
K. Vanneste for discussions, and K. Vanneste for his help
during the preparation of the manuscript and M. Everaerts
for the modelling of the gravimetric anomaly. We thank P. M.
Grootes and M. J. Nadeau (Christian-Albrechts-Universität,
Kiel) and M. Van Strydonck (Institut Royal du Patrimoine
Artistique, Bruxelles) for the quick processing of our radiocarbon samples. This work was supported by the Royal
© 1998 RAS, GJI 132, 347–362
Holocene surface faulting in the Roer Graben
Observatory of Belgium and the European Centre for
Geodynamics and Seismology (Luxembourg).
R E FER E NCE S
Ahorner, L., 1962. Untersuchungen zur quartären Bruchtektonik der
Niederrheinischen Bucht, Eiszeitalter Ggw., 13, 24–105.
Ahorner, L., 1975. Present-day stress field and seismotectonic block
movements along major fault zones in Central Europe,
T ectonophysics, 29, 233–249.
Ahorner, L., 1983. Historical seismicity and present-day microearthquake activity of the Rhenish Massif, Central Europe, in Plateau
Uplift, pp. 188–198, eds Fuchs, K., von Gehlen, K., Mälzer, H.,
Murowski, H. & Semmel, A., Springer Verlag, Berlin.
Ahorner, L., 1994. Fault-plane solutions and source parameters of the
1992 Roermond, the Netherlands, mainshock and its stronger aftershocks from regional seismic data, Geol. Mijn., 73, 199–214.
Ahorner, L., Murawski, H. & Schneider, G., 1970. The distribution of
earthquakes causing damage on the territory of the Federal Republic
of Germany. An attempt of seismic regionalisation, Zeitschrift
Geophys., 36, 313–343.
Alexandre, P., 1990. Les séismes en Europe Occidentale de 324 à 1259:
Nouveau catalogue critique, ed. Observatoire Royal de Belgique,
Série Géophysique.
Alexandre, P., 1994. Historical seismicity of the lower Rhine and
Meuse valleys from 600 to 1525: a new critical review, Geol. Mijn.,
73, 431–438.
Alexandre, P. & Vogt, J., 1994. La crise séismique de 1755–1762 en
Europe du Nord-Ouest, in Historical Investigation of European
Earthquakes, Materials of the CEC project: Review of historical
seismicity in Europe, Vol. 2, 37–76, eds Albini, P. & Moroni, A.,
CNR, Italy.
Ambraseys, N.N., 1985a. Magnitude Assessment of Northwestern
European Earthquakes, Earthq. Eng. struct. Dyn., 13, 307–320.
Ambraseys, N.N., 1985b. Intensity-attenuation and magnitudeintensity relationships for northwest European earthquakes, Earthq.
Eng. struct. Dyn., 13, 733–778.
Bertrand, M.E., 1757. Mémoires historiques et physiques sur les
tremblements de terre, La Haye.
Boschi, E. et al., 1996. New trends in active faulting studies for seismic
hazard assessment, Ann. Geofis., XXXIX, 1301–1307.
Camelbeeck, T. & van Eck, T., 1994. The Roer Valley Graben
earthquake of 13 April 1992 and its seismotectonic setting, T erra
Nova, 6, 291–300.
Camelbeeck, T. & Meghraoui, M., 1996. Large earthquakes in northern
Europe more likely than once thought, EOS, T rans. Am. geophys.
Un., 77, 405 and 409.
Crone, A.J., Machette, M.N. & Bowman, J.R., 1992. Geologic investigations of the 1988 Tennant Creek, Australia, earthquakes—implications for paleoseismicity in stable continental regions, Geol. Surv.
Am. Bull., 2032A.
Crone, A.J., Machette, M.N. & Bowman, J.R., 1997a. The episodic
nature of earthquake activity in stable continental regions revealed
by paleoseismicity studies of Australian and North American
Quaternary faults, Aus. J. Earth Sci., 44, 203–214.
Crone, A.J., Machette, M.N., Bradley, L.A. & Mahan, S.A., 1997b.
Late Quaternary surface faulting on the Cheraw fault, southeastern
Colorado, Geologic Investigations Map, MAP I-2591, US Geological
Survey, Boulder, CO.
De Batist, M. & Versteeg, W.H., 1998. Structure and seismic stratigraphy of the Mesozoic and Cenozoic in northern Belgium: results of
a high-resolution seismic survey along rivers and canals, Geol.
Mijn., submitted.
Demyttenaere, R., 1989. The post-paleozoic geological history of northeastern Belgium, Mededelingen van de Koninklijke Academie voor
Wetenschappen, L etteren en Schone Kunsten van België, 51, (4)
51–80.
© 1998 RAS, GJI 132, 347–362
361
Demyttenaere, R. & Laga, P., 1988. Breuken- en isohypsenkaarten
van het Belgisch gedeelte van de Roerdal slenk, Belg. Geol. Dienst
Prof. Paper, 234.
De Sitter, L.U., 1949. Eindverslag van het geophysische onderzoek in
ZO.-Nederland, Mededelingen van het Geologische Stichting, serie
C-I-3-N°1.
Geluk, M.C., 1990. The Cenozoic Roer Valley Graben, Southern
Netherlands, Mededelingen Rijks Geologische Dienst, 44–4, 65–72.
Geluk, M.C., Duin, E.J., Dusar, M., Rijkers, R., van den Berg, M.W.
& van Rooijen, P., 1994. Stratigraphy and tectonics of the Roer
Valley Graben, Geol. Mijn., 73 (2–4), 129–141.
Johnston, A.C., 1996a. Seismic moment assessment of earthquakes in
stable continental regions—I. Instrumental seismicity, Geophys.
J. Int., 124, 381–414.
Johnston, A.C., 1996b. Seismic moment assessment of earthquakes in
stable continental regions—II. Historical seismicity, Geophys. J. Int.,
125, 639–678.
Johnston, A.C., Coppersmith, K.J., Kanter, L.R. & Cornell, C.A., 1994.
The earthquakes of stable continental regions, Electric Power Res.
Inst. rpt., TR-102261.
Juvigné, E. & Renard, F., 1992. Les terrasses de la Meuse de Liège à
Maastricht, Ann. Soc. géol. Belgique, 115, 167–186.
Karnik, V., 1969. Seismicity of the European Area, Reidel, Dordrecht.
Kolstrup, E., 1980. Climate and stratigraphy in northwestern Europe
between 30 000 B.P. and 13 000 B.P., with special reference to the
Netherlands, Mededelingen Rijks Geologische Dienst, 32, 181–253.
Kusznir, N.J., Marsden, G. & Egan, S.S., 1991. A flexural cantilever
simple-shear/pure-shear model of continental lithosphere extension:
application to the Jeanne d’Arc Basin, Grand Banks and Viking
Graben, North Sea, in T he Geometry of Normal Faults, eds Roberts,
A.M., Yielding, G. & Freeman, B., Spec. Publ. Geol. Soc. L ond.,
56, 41–60.
Mc Calpin, J.P., 1996. Paleoseismology, Academic Press,
Mälzer, H., Hein, G. & Zippelt, K., 1983. Height changes in the
Rhenish Massif: determination and analysis, in Plateau Uplift,
pp. 164–176, eds Fuchs, K., von Gehlen, K., Mälzer, H.,
Murowski, H. & Semmel, A., Springer Verlag, Berlin.
Meghraoui, M., Camelbeeck, T., Vanneste, K., Brondeel, M. &
Jongmans, D., 1998. Active tectonics and paleoseismology along the
Bree fault zone (Lower Rhine Graben, Belgium), J. geophys. Res.,
submitted.
Melville, C., Levret, A., Alexandre, P., Lambert, J. & Vogt, J., 1996.
Historical seismicity of the Strait of Dover–Pas de Calais, T erra
Nova, 8, 626–647.
Olig S.S., Gorton, A., Bott, D., Wong, I., Kruepfer, P., Fouman, S. &
Smith, R., 1994. Preliminary interpretations of temporal surfacefaulting patterns on the southern Lost River fault zone, northeastern
Basin and Range Province, USA. US Geological Survey open file
rept, 94–568, 145–147.
Paulissen, E., 1973. De morfologie en de kwartairstratigrafie van de
Maasvallei in Belgisch Limburg, Verhandelingen van de Koninklijke
V laamse Academie voor Wetenschappen, L etteren en Schone Kunsten
van België, jaargang XXXV, Nr 127.
Paulissen, E., 1983. Les nappes alluviales et les failles quaternaires du
Plateau de Campine, in Belgique—Guides Géologiques Régionaux,
pp. 167–170, eds Robaszynski, F. & Dupuis, C., Masson, Paris.
Paulissen, E., Vandenberghe, J. & Gullentops, F., 1985. The Feldbiss
fault in the Maas valley bottom (Limburg, Belgium), Geol. Mijn.,
64, 79–87.
Rasmussen, R. & Pedersen, L.B., 1979. End corrections in potential
field modeling, Geophys. Prospect., 27, 749–760.
Richter, C.F., 1958. Elementary Seismology, W. H. Freeman, San
Francisco, CA.
Roberts, A. & Yielding, G., 1994. Continental extensional tectonics, in
Continental Deformation, pp. 223–250, ed. Hancock, P.L., Pergamon
Press, Oxford.
362
T . Camelbeeck and M. Meghraoui
Rosenhauer, W. & Ahorner, L., 1994. Seismic hazard assessment for
the Lower Rhine Embayment before and after the 1992 Roermond
earthquake, Geol. Mijn., 73 (2–4), 415–424.
Scholz, C.H., 1990. T he Mechanics of Earthquakes and Faulting,
Cambridge University Press.
Stewart, I.S. & Hancock, P.L., 1994. Neotectonics, in Continental
Deformation, pp. 370–409, ed. Hancock, P.L., Pergamon Press,
Oxford.
Van den Berg, M.W., 1994. Neotectonics of the Roer Valley rift system.
Style and rate of crustal deformation inferred from syn-tectonic
sedimentation, Geol. Mijn., 73 (2–4), 143–156.
Van den Berg, M.W., Groenewoud, W., Lorenz, G.K., Lubbers, P.J.,
Brus, D.J. & Kroonenberg, S.B., 1994. Patterns and velocities of
recent crustal movements in the Dutch part of the Roer Valley rift
system, Geol. Mijn., 73 (2–4), 157–168.
Vandenberghe, J., 1982. Geoelectric investigations of a fault system in
Quaternary deposits, Geophys. Prospect., 30, 879–897.
Vogt, J., 1979. Les tremblements de Terre en France, Mém. B.R.G.M.,
96, Orléans.
Vogt, J., 1991. Die Erdbebenfolge vom Mai 1733 im ‘Rheingebiet’,
Mainzer Naturwissenschaftliches Archiv., 29, 65–69.
Vogt, J., 1994. Glimpses at the 1640 earthquake in north-western
Europe, Historical Investigation of European Earthquakes, 2,
77–88.
Wells, D.L. & Coppersmith, K.J., 1994. Empirical relationships among
magnitude, rupture length, rupture area, and surface displacement,
Bull. seism. Soc. Am., 84, 974–1002.
Zippelt, K. & Mälzer, H., 1981. Recent height changes in the central
segment of the Rhinegraben and its adjacent shoulders,
T ectonophysics, 73, 119–123.
© 1998 RAS, GJI 132, 347–362