Reactivated basement structures affecting the sedimentary
facies in a tectonically ‘‘quiescent’’ epicontinental basin:
an example from NW Switzerland
Andreas Wetzel *, Robin Allenbach1, Vincenzo Allia2
Geologisch-Palaontologisches Institut der Universitat Basel, Bernoullistrasse 32, CH-4056 Basel, Switzerland
Received 23 May 2001; accepted 29 April 2002
Abstract
The Jurassic deposits in the southern part of central Europe accumulated in a shallow epicontinental sea, and their deposition
has usually been believed to have corresponded to a phase of tectonic quiescence; neither on conventional seismic records nor
in outcrops obvious indications of synsedimentary tectonics were found. Nonetheless, subtle variations in lithofacies and
thickness occur above faults in the crystalline basement. In fact, preexisting structures became reactivated in the Jurassic during
major extensional phases when the Tethys and the North Atlantic opened. This reactivation led to differential subsidence and/or
to rotation of fault-bounded blocks, but the sediments were deformed mainly flexurally as vertical movements in the basement
were dissipated by Triassic salt. Thus, the depositional area was morphologically differentiated into swells and depressions. In
siliciclastic muddy environments, swells were characterized by hardbottoms and hiatus beds; depressions were filled by distal
tempestites and gravity deposits. In carbonate settings, reefs nucleated on swells; marls and muds were deposited in depressions
(many of them gravity deposits). Reactivation of faults occurred only during short episodes and was not synchronous
throughout the study area, and reactivation of individual faults was episodic and probably controlled by the palaeostress field
and the faults’ orientation. Reactivation of deep-rooted faults is also documented by hydrothermal activity which led to vein
mineralization and alteration of minerals—today exposed in nearby basement units. The chronostratigraphic ages of the
hydrothermal processes coincide with phases of enhanced subsidence during the Sinemurian, Aalenian, Bajocian/Bathonian,
and Oxfordian. In turn, facies changes of the sedimentary cover should be useful to predict basement structures when no seismic
records are available.
D 2002 Elsevier Science B.V. All rights reserved.
Keywords: Shallow water sediments; Jurassic; Switzerland; Basement; Fault reactivation
1. Introduction
Sediment accumulation in shallow-marine, epicon-
tinental settings commonly clearly responds to sea-
level changes because environmental factors such as
waves, currents, light penetration, input of clastics
and nutrients, etc. are strongly affected by water
depth or distance to coast (e.g., Johnson and Baldwin,
1996; Wright and Burchette, 1996). Consequently, the
0037-0738/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved.
doi:10.1016/S0037-0738(02)00230-0
* Corresponding author.
E-mail address: [email protected] (A. Wetzel).1 Present address: Proseis AG, Siewerdtstrasse 7, CH-8050
Zurich, Switzerland.2 Present address: Geotechnisches Institut, Hochstrasse 48, CH-
4002 Basel, Switzerland.
www.elsevier.com/locate/sedgeo
Sedimentary Geology 157 (2003) 153–172
lithology and the stratal arrangement of such deposits
is extensively discussed today in terms of sequence
stratigraphy or genetic stratigraphy (e.g., Loucks and
Sarg, 1993; Posamentier et al., 1993; Weimer and
Posamentier, 1993). The sediments within the region
on which we report in this present paper—the Swiss
Jura and adjacent areas of SW Germany—formed in
an epicontinental setting, and several authors have
successfully applied sequence-stratigraphic concepts
to them (e.g., Aigner and Bachmann, 1992; Burkhal-
ter, 1996; Gonzalez, 1996; Pittet and Strasser, 1998).
Thick epicontinental deposits only form in subsi-
ding basins. Obviously, to a first approximation, the
potential of sediments to accumulate below base level
increases with increasing subsidence rate. Except in
rifts, at young passive margins, and in other tectoni-
cally active zones, subsidence in epicontinental basins
is commonly low; hence, visible structures indicative
of synsedimentary tectonic activity are sparse or very
subtle—in the field and on seismic records. Therefore,
epicontinental shallow-marine sediments are com-
monly interpreted either to indicate tectonic quiescence
(e.g., Laubscher, 1986; Thury et al., 1994: Swiss Jura;
Schroder et al., 1997: SE Germany) or to indicate that
tectonic movements are of subordinate importance in
relation to sea-level changes (e.g., Aigner and Bach-
mann, 1992: SW Germany; Pittet and Strasser, 1998:
Swiss Jura).
The use of sequence stratigraphy emphasizes the
role of eustatic sea-level changes and subsidence on
stratal organisation and the correlatability of lithologic
units (e.g., Sarg, 1988; de Graciansky et al., 1998). In
addition, differential subsidence affecting changes in
accommodation space is considered (e.g., Robin et al.,
1998; Harris et al., 1999). In epicontinental basins,
any variation of subsidence in space and time influ-
ences sediment thickness, lithology, and/or facies
distribution (e.g., Wildi et al., 1989; Allen and Allen,
1990). When a thick sedimentary cover is forming,
preexisting structures in the basement have a high
potential to become reactivated. This has been dem-
onstrated for areas that evidently are tectonically
stretched, such as shelves or basins on or adjacent to
continental margins (e.g., Bonijoly et al., 1996; Fae-
rseth, 1996; Dromart et al., 1998). For epicontinental
basins, which are believed to be tectonically quies-
cent, the importance of preexisting structures in terms
of sediment thickness and lithofacies is increasingly
being recognized (e.g., Ziegler, 1989; Keeley, 1996;
Brandley et al., 1996).
In this study, we want to explain the basement–
cover relationship for a part of northern Switzerland
for which seismic records and some deep wells are
available. There is no visible evidence for synsedi-
mentary faults, neither on conventional 2D seismic
records nor in outcrops (e.g., Thury et al., 1994 and
references therein). However, recently performed
high-resolution 3D seismic investigations (just out-
side the study area) provided indications for small-
scale synsedimentary movements during the Jurassic
(Birkhauser et al., 2001). In outcrops subtle variations
in facies and thickness are suggestive of a syndeposi-
tionally formed relief (e.g., Wetzel and Allia, 2000).
Our study focusses on shallow-marine epicontinental
deposits and we examine when, how, and why the
reactivation of basement structures affected the lith-
ofacies of the Jurassic sedimentary cover.
2. Geological background
The investigation area is located in northern Swit-
zerland and southwestern Germany (Fig. 1). Two
time-slices in Jurassic deposits were studied in detail.
In the first one, Aalenian deposits are characterized
by about 100-m-thick terrigenous mudrocks; in the
second, Oxfordian deposits are characterized by a
transition from a carbonate platform to a muddy
basin. Both time-slices are found to be suitable as
examples of the influence of reactivation of faults in
the crystalline basement on the sedimentary cover
(see below).
The tectonic structures in the crystalline basement
formed when a mega-shear zone developed between
the Ural and the Appalachians towards the end of the
Palaeozoic (Arthaud and Matte, 1977) and strike-slip
movements led to the formation of numerous basins,
grabens, and half-grabens (e.g., Menard and Molnar,
1988; Ziegler, 1990; von Raumer, 1998). Within the
study area and its surroundings, these features in the
basement are known from seismic records (Laubscher,
1986, 1987; Diebold, 1988; Diebold et al., 1991;
Thury et al., 1994).
Some of these basins started to form prior to the
late Carboniferous (e.g., Schafer, 1986; Bruguier et
al., 1998), others during the Westphalian to early
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172154
Stephanian (Schaltegger and Corfu, 1995). During the
Late Permian, a compressional phase was deduced
from seismic records in the study area by Laubscher
(1987) who inferred an inversion of the fault-bounded
basin and subsequent erosion of basin-fill sediments.
This compressional phase is a matter of debate today
(Steck, personal communication, Lausanne, 2002),
but thermal modelling led Kempter (1987) and
Schegg and Leu (1998) to suggest that about 1500
and 1000–1200 m, respectively, were removed from
the basin fill during the Late Permian. Thereafter, the
basins continued to subside and additional basins
developed (e.g., Ziegler, 1990). Furthermore, at the
end of the Variscan Orogeny, a roughly N–S-trending
fault system and associated magmatic dykes (Metz,
1970) formed at the position of the future Rhine
Graben, the so-called ‘‘Rhenish Lineament’’ (Boigk
and Schoneich, 1974) or ‘‘Rhine Graben Lineament’’
(Ziegler, 1990). It includes a fault zone in the southern
Black Forest (e.g., Krohe, 1996).
During the Triassic, peneplanation took place and
continental and restricted marine deposits accumu-
lated. Within the context of this paper, the accumu-
lation of up to 100 m of evaporites (salt and gypsum/
anhydrite) representing parts of the Muschelkalk (Fig.
2) is of importance because of their plastic behavior
and their bearing on tectonic movements. During the
Lias, a transgression occurred, and an epicontinental
Fig. 1. Location of the study area, the wells at Weiach (W), Riniken (R), and Schafisheim (S), and some late Palaeozoic basins in the subsurface
(stippled) as described by the following authors: Burgundy Basin (Boigk and Schoneich, 1974), Schramberg Basin (Boigk and Schoneich, 1974;
von Raumer, 1998), Lake Constance Basin (Boigk and Schoneich, 1974; von Raumer, 1998), North Swiss Permo–Carboniferous Basin (NSB;
Diebold, 1988; Diebold et al., 1991), Entlebuch Basin and basins south of it (Pfiffner, 1993). Not all basins south of the study area are shown. The
Variscan Rhenish Lineament (name after Boigk and Schoneich, 1974) is shown after Ziegler (1990), the associated faults after Krohe (1996).
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 155
sea covered wide parts of Europe. At that time, the area
of the developing Atlantic and Tethys oceans was
stretched and the study area was affected by an exten-
sional stress regime (e.g., Ziegler, 1990).
3. Methods
In a slowly subsiding epicontinental basin, pre-
existing tectonic structures were reported to have been
reactivated at times when subsidence was enhanced
(e.g., Brandley et al., 1996; Keeley, 1996; Shail and
Alexander, 1997; de Wet, 1998). To identify such
times, geohistory diagrams—taking into account com-
paction—were prepared by using the computer pro-
gram ’Basin Works’ by MarcoPolo Software. Using
the time scales of Menning et al. (2000) for the
Palaeozoic and Gradstein et al. (1995) for the Meso-
zoic, and the sea-level curve of Haq et al. (1987), this
software calculates isostatic subsidence, non-isostatic
tectonic subsidence, and mechanical compaction.
Besides mechanical compaction, the decrease in
porosity due to clay mineral diagenesis was calculated
following Waples and Kamata (1993) based on the
temperature history of the study area provided by
Todorov et al. (1993). Temperatures were estimated
using the computer program EASY%Ro (Sweeney
and Burnham, 1990), calibrated by the values deter-
mined from the cores at Riniken and Weiach (Matter
et al., 1987, 1988).
Isopach maps were prepared for time intervals
marked by enhanced subsidence. Isopach maps used
in combination with palaeowater depth estimates
based on wave ripples (e.g., Diem, 1985) and fossils
(e.g., Flugel, 1978) provide valuable information
about basin dynamics. If sediment thickness exceeds
water depth during deposition, the accommodation
space was provided during deposition and thickness
anomalies could reflect basin morphology. To eval-
uate this, palaeoflow directions were analysed: if
palaeoflow was towards a thickness maximum, this
probably represented a depression. The palaeomor-
phology derived in this way was compared spatially
with known basement structures. In addition, data
about vein mineralization and palaeostress field were
used to evaluate the reactivation potential of pre-
existing structures.
4. Basement structures and sedimentary cover
Sediment accumulation through time was analysed
for three drilled sections preserving a nearly complete
record of the Mesozoic sediments in the investigation
area (Figs. 1–3). Although they are located outside, at
the margin, and in the center of the Permo–Carbon-
iferous basin in northern Switzerland, they exhibit a
similar stratigraphy, albeit differing in their sediment
thicknesses except the upper Palaeozoic deposits (Fig.
3). During the Mesozoic, subsidence increased during
the Middle Triassic (Muschelkalk) and then decreased
until the end of the Lias. Thereafter, three episodes of
enhanced subsidence occurred: during the Aalenian,
the Bajocian/Bathonian, and the Early to Middle
Oxfordian, the latter slowing down towards the end
of the Jurassic. Subsidence achieved the highest rate
during the Aalenian, but only for a short period. The
Fig. 2. Simplified stratigraphic column for the study area after
Bitterli-Brunner (1987). During the Callovian/Oxfordian, a con-
densed series formed in the eastern part of the study area, and a
thick muddy sequence in the western part.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172156
largest amount of subsidence, but at a lower rate,
occurred during the Oxfordian.
Two time slices characterized by accelerated sub-
sidence and enhanced sedimentation were chosen to
exemplify the effects of basement structures on the
sedimentary cover. The first, Aalenian in age, is
represented by terrigenous mudrocks; the second,
Oxfordian in age, is represented by a carbonate-plat-
form to marl-basin transition.
4.1. Aalenian terrigenous mudrocks, case study 1
Terrigenous mudrocks, today about 100 m thick,
accumulated during the Aalenian (Fig. 2), mainly
during the opalinum ammonite subzone. Therefore,
they are called in German ‘‘Opalinuston’’, translated
as Opalinus Mudrock. This facies formed during
0.5–2 m.y., depending on the time scale used
(Gradstein et al., 1995; Harland et al., 1989). In this
study, we use a value of 1 m.y. spanning from 180 to
179 Ma.
The Opalinus Mudrock accumulated in the south-
ern part of an epicontinental sea, which covered
central Europe at that time (e.g., Ziegler, 1990). The
basin was surrounded by land, islands, or a shallow
platform. The terrigenous sediments were delivered
from Scandinavia and the Bohemian Massif, the
London–Brabant Massif and the Rhenish Massif,
and subordinately from the Alemannic Islands (Fig. 4).
These dark grey, terrigenous mudrocks have often
been considered to be monotonous, but they comprise
a variety of lithologies.
Fig. 3. Decompacted sediment thickness versus time for the standard sections drilled at Weiach, Riniken, and Schafisheim (for location, see Fig.
1). For the Upper Palaeozoic at Weiach, the effects of inversion and subsequent erosion were quantified by Kempter (1987). For Riniken, the
Palaeozoic sediment thickness was estimated from seismic records, but Carboniferous and basement were not reached. Although direct evidence
for inversion is missing, seismic records suggest a history similar to Weiach (Laubscher, 1986) being a few kilometers away. Note enhanced
sediment accumulation during the Aalenian and Oxfordian and compare the amount of Late Palaeozoic erosion to Mesozoic sediment thickness.
Inset shows sediment accumulation during the Lias at Schafisheim. Chronostratigraphy of stage boundaries after Menning et al. (2000) for the
Palaeozoic and Gradstein et al. (1995) for the Mesozoic. West: Westphalian, S: Stephanian, Rl: Rotliegend, Z: Zechstein, Bu: Buntsandstein,
Mu: Muschelkalk, Keu: Keuper, R: Rhaetian, H: Hettangian, Si: Sinemurian, P: Pliensbachian, To: Toarcian, A: Aalenian, Bj: Bajocian, B:
Bathonian, C: Callovian, O: Oxfordian, K: Kimmeridgian, and T: Tithonian.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 157
Homogeneous mudrocks: They form centimeter- to
decimeter-thick intervals and display no primary sedi-
mentary structures, but scattered faint trace fossils
may be seen.
Laminated mudrocks: These comprise laminated
pure mudrocks and silt/sand-laminated mudrocks as
well. Both types can pass laterally into each other
within some decimeters. The sand–silt laminae have
sharp bases and gradational tops; locally, they show
faint current ripples. Laminated and homogeneous
mudrocks commonly alternate.
Thick-bedded silt and sand layers: Siltstones and
sandstones can form centimeter- to decimeter-thick
continuous or discontinuous beds. They display ero-
sive features at the base, commonly associated with
sole marks; erosive surfaces are also present within
these layers. The discontinuous layers consist of
several decimeter- to meter-long trains of current
ripples or small-scale channels, which show pro-
nounced erosion at the base. The continuous sand
layers are plane-laminated, obliquely laminated, low-
angle cross-stratified, and/or cross-stratified; many of
them display the characteristics of hummocky cross-
stratification (e.g., Duke et al., 1991). Within a bed,
the various stratification types can occur, commonly
separated by scouring surfaces. Ripple cross-stratifi-
cation can be uni- or bidirectional. On top of the sand
layers, wave ripples can occur. These were used to
calculate depositional water depth using the method of
Diem (1985); the resulting values of water depth are
in the range of 20–30 m.
Nodules and nodular limestones: Carbonate nod-
ules of varying shape formed during early diagenesis
in burrows—mainly Thalassinoides—within the sul-
fate reduction zone (Wetzel and Allia, 2000). Hori-
zons rich in nodules may mark episodes of reduced
sedimentation (e.g., Spears, 1989); this is also indi-
cated by the enrichment of the sediments in ammon-
ites just above such layers. The nodules were
exhumed in some places and display features indicat-
ing exposure on the seafloor, such as borings and
encrustation.
Fig. 4. Palaeogeographic map for the Aalenian (after Ohmert and Rolf, 1994; slightly changed).
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172158
The sandstones and mudrocks show the character-
istics of storm deposits; the thick silt/sand beds with
HCS and/or wave ripples on top are typical of
tempestites formed at or above storm-wave base
(e.g., Duke et al., 1991). The laminated mudrocks
represent distal tempestites deposited below storm-
wave base. The homogeneous muds were mixed by
organisms, and they document episodes of hemipela-
gic rain or slow sedimentation favoring complete
mixing (e.g., Reineck, 1977); they probably accumu-
lated in water depths exceeding 20–30 m. The inter-
fingering of the mudrocks with siltstones and
sandstones is typical of storm-affected shallow-water
deposits (e.g., Aigner and Reineck, 1982; Milkert,
1994). The nodular limestone beds document local
reworking on somewhat elevated parts of the seafloor,
very likely on swells (Wetzel and Allia, 2000).
Today, the Opalinus Mudrock is about 80–120 m
thick (Fig. 5). Taking into account compaction, 100 m
of compacted sediment corresponded to about 210–
230 m when deposited (calculated by ’Basin Works’,
see above). The depositional water depth was in the
range of 30–50 m. Sea level rose only by about 10–
20 m during the Aalenian (Haq et al., 1987; Branger
and Gonnin, 1994), suggesting that accommodation
space was provided during deposition. Thus, the
question arises: did the isopach maxima represent
synsedimentary depressions or not? To answer this,
the isopach pattern was compared to that of palaeo-
flow indicators such as asymmetrical ripples and sole
marks (Fig. 5). In fact, palaeoflow tends to be directed
towards areas of maximum thickness, although it
commonly deviates by 30–60j from the direction
normal to the isopachs. These deviations may be
due to uncertainties in the isopach construction, or
they may result from oblique downwelling circulation
(e.g., Duke et al., 1991; Myrow and Southard, 1996).
The storm-sand layers document palaeoflow roughly
perpendicular to the isopachs and towards the thick-
ness maxima. Consequently, these maxima are inter-
Fig. 5. Isopach map and palaeoflow directions for the Aalenian terrigenous mudrocks in the study area (from Allia, 1996). The isopachs were
palinspastically restored according to Laubscher (1965). The location of the Late Palaeozoic basins in the subsurface is based on Diebold et al.
(1991). Sandy tempestites indicate palaeoflow towards the thickness maxima, which are spatially related to the Late Palaeozoic basins.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 159
preted as depressions that formed syngenetically. At
the first glance, for the eastern part of the study area,
the thickness maximum seems to be located above the
late Palaeozoic trough (Fig. 5). At a closer look,
however, it is seen that the isopachs are systematically
shifted to the north, where numerous faults structure
the basement (e.g., Diebold, 1988; Thury et al., 1994).
The other isopach maxima are not so well constrained,
but a similar relationship may be seen. The close
spatial relationship of depocenters to Late Palaeozoic
faults suggests their reactivation. Because 1000–1200
m of sediments were removed from the fill of the Late
Palaeozoic basin prior to the Mesozoic (see above),
compaction of the trough fill as reason for differential
subsidence of the seafloor is highly unlikely (see
below).
4.2. Oxfordian carbonate-platform to marl-basin
transition, case study 2
As a second example, Oxfordian sediments were
analysed to elucidate the relationships between the
lithofacies and structures in the basement. During the
Oxfordian, the subsidence was enhanced for a longer
time but at a lower rate than during the Aalenian (Fig.
3). The palaeogeographic situation changed compared
to the Aalenian (Figs. 4 and 6): the epicontinental
realm was flooded and the study area was now
connected via the Helvetic Shelf to the Tethys ocean
(e.g., Wildi et al., 1989; Ziegler, 1990). The northern
part of the study area was occupied by a carbonate
platform, and calcareous muds accumulated in the
southern part.
Three different lithologies can be recognized in the
study area (Fig. 7): a condensed series, marl–lime-
stone alternations, and shallow-water limestones.
Condensed series: This is about 0.5–1 m thick. It
consists of iron oolites, some stromatolites, and
wacke- to packstone, and it formed during the Early
to early Middle Jurassic in the eastern part of the study
area. At that same time, marls accumulated farther to
the west (Gygi and Persoz, 1986).
Marl– limestone alternations: These consist of
centimeter- to decimeter-thick beds which accumu-
lated during the middle and early Late Oxfordian. The
whole series is today up to 240 m thick. The carbonate
content varies within a section: this variation is
interpreted to reflect climatic and sea-level changes
(e.g., Pittet, 1994; Pittet and Strasser, 1998), some of
them being within the Milankovitch band (op. cit.).
The marls are usually bioturbated. The limestones
commonly are either burrowed or display primary
Fig. 6. Palaeogeographic map for the lower Oxfordian (after Ziegler,
1990, slightly changed).
Fig. 7. Upper part: Isopach map of Middle Oxfordian marls and limestones, based on data of Buchi et al. (1965), Bitterli (1992), and our own
observations. The isopachs were palinspastically restored according to Laubscher (1965). The location of the Late Palaeozoic basins is compiled
from Diebold et al. (1991) and Diebold and Noack (1997). Because of the poor stratigraphic resolution, the NW part of the study area is not
shown; it was occupied by a carbonate platform at that time (e.g., Gygi, 1990). Lower part: SW–NE section across the study area (after Gygi,
1969; location see above A–B). Across the faults bounding the Late Palaeozoic basins, thickness of the marl and limestone series significantly
increases. The lack of continuity in the limestone beds to the north and to the south is due to missing outcrops. However, where exposed, many
of them contain redeposited material. The reefs preferably nucleated above pre-existing faults. Note that the section was drawn by interpolating
between outcrops.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172160
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 161
sedimentary structures such as lamination or asym-
metrical ripples. These ripples are useful to determine
palaeoflow directions. Oscillatory ripples were not
found. Consequently, deposition below storm-wave
base is inferred.
Shallow-water carbonates: These formed during
the Middle and Late Oxfordian. Three main litholo-
gies occur in the study area: well-bedded limestones,
oolites and calcarenites, and reefal limestones. The
well-bedded limestones consist of mud-to-wacke-
stone-containing bioclasts of echinoderms, molluscs,
brachiopods, and corals. Oolites, oncolites, and cal-
carenites formed on the landward side of the patch-
reef belt (for details, see Gygi, 1969, 1990). The
platform margin and fringe is a geometrically and
lithologically complex system consisting of a patch-
reef belt and inter-reef mud-, wacke-, and packstones
containing a considerable amount of broken platform
organisms (e.g., Gygi, 1969, 1990; Bolliger and Burri,
1970). Behind the patch-reef belt, a lagoonal area with
small reefs, micrites, and oncolites developed.
The facies boundaries of the Lower to Middle
Oxfordian deposits coincide fairly well with the
NE–SW-trending Late Palaeozoic structures in the
subsurface of the study area. This relationship can be
deduced from isopachs in combination with transverse
sections (Fig. 7) and palaeoflow data (see below). For
example, on the section drawn by Gygi (1969) at a
time when the Permo–Carboniferous basins in the
subsurface were not known, lithology significantly
changes across the Late Palaeozoic structures (Fig. 7):
the condensed series mainly occurs above areas
underlain by the tectonically complex Late Palaeozoic
grabens. The Middle Oxfordian sediments are thickest
in areas where numerous Late Palaeozoic faults are
present in the basement. Reefs nucleated in the
vicinity of Late Palaeozoic faults.
The palaeogeographic maps published by Gygi
(1969, 1990) indicate that the facies boundaries
moved with time (Fig. 8). During the Early Oxfordian,
they were preferentially NE–SW-oriented. During the
Middle to Late Oxfordian, the platform-basin transi-
tion was oriented—as before—NE–SW in the south-
ern part of the study area. The eastern boundary of the
platform, however, shifted further to the west and
became roughly N–S-oriented in spatial relation to
the Rhenish Lineament and associated faults (Krohe,
1996), which are found in continuation along strike of
the today’s Rhine Graben eastern boundary fault
system (Allenbach, 2001, 2002). On the platform
itself, differential subsidence led to small-scale varia-
tions in facies and thickness (Bolliger and Burri,
1970; Pittet, 1994; Allenbach, 2001, 2002).
The isopach map (Fig. 9) displays local thickness
maxima that, according to palaeoflow data, probably
represent synsedimentary depressions; the palaeoflow
data indicate transport towards the depocenters, but
Fig. 8. Palaeogeographic maps of the Middle (upper part) and Upper
Oxfordian (lower part), based on Gygi (1990), but palinspastically
restored after Laubscher (1965). Note that facies boundaries were
mainly NE–SW during the Middle Oxfordian and N–S during the
Late Oxfordian. The facies boundaries are in spatial vicinity to Late
Palaeozoic structures or in continuation along strike of the (1)
Rhenish Lineament and (2) associated fault.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172162
Fig. 9. Isopachs and directions of resedimentation processes (tempestites, channels, and slides) of the Middle Oxfordian Effingen Formation.
Fig. 10. Southward-directed slide in Oxfordian marl-limestone alternations observed at the northern side of the Late Palaeozoic basin in the
subsurface (quarry near Rekingen, for location, see Fig. 7). Syngenetic growth faults at the base. Exposed sediment series is about 30 m thick.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 163
can deviate slightly from the direction normal to the
isopachs, either because of uncertainties in isopach
construction or because of oblique downwelling cir-
culation (e.g., Duke et al., 1991; Myrow and South-
ard, 1996). During a short episode of the Middle
Oxfordian, we found palaeoflow from opposite direc-
tions into an area structured by Late Palaeozoic faults.
This suggests that a depocenter was temporarily
developed there. Higher up in the succession, south-
ward-directed palaeoflows dominate, indicating that
the depocenter was filled and differential subsidence
could not maintain the depression. Furthermore, rather
large masses slid southwards. They occurred—at
some localities repeatedly—above faults bounding
the Late Palaeozoic basin (Fig. 10). Presumably due
to the general tilt of the basin to the south, relief
steepened across the pre-existing faults and induced
slides. Consequently, facies boundaries coincide spa-
tially with tectonic structures in the basement, and
synsedimentary depressions strongly affected facies
development of the sedimentary cover.
5. Hydrothermal activity and movements along
faults
During the Jurassic, differential subsidence occurred
in close spatial relationship to the Late Palaeozoic
basement structures in the study area and implies their
reactivation. Further evidence for tectonic activity is
provided by hydrothermal activity that led to mineral-
ization of veins, alteration of minerals, and illitization
(e.g., Bonhomme et al., 1983; Clauer et al., 1996), all of
which occurred along basement faults that acted as
major conduits for fluids. Data on hydrothermal activity
come from drill holes and basement rocks being
exposed nearby in the Black Forest and the Vosges.
These processes were chronostratigraphically dated
by several authors (listed in caption of Fig. 11), and
hence provide valuable information about the epi-
sodes of extension within the basement (Fig. 11).
Four phases of enhanced hydrothermal activity
occurred during the Jurassic: at about 200F 2.5 Ma
(phase I), 180F 2.5 Ma (phase II), 170F 2.5 Ma
(phase III), and 150F 2.5 Ma (phase IV). These
correspond to the Sinemurian, Aalenian, Bathonian,
and Oxfordian, respectively, according to the time
scale of Gradstein et al. (1995).
The phases of enhanced hydrothermal activity—
except the Liassic—match the phases of increased
subsidence within the study area. Hydrothermal phase
I during the Liassic is not clearly reflected by the
subsidence curve, but during the Sinemurian, sedi-
ments accumulated at a three to four times higher
rate than the other Liassic deposits. The Middle and
Late Jurassic hydrothermal phases (II to IV) coincide
with subsidence pulses (Fig. 3); one of the most
pronounced phases of subsidence (Aalenian) is syn-
chronous with hydrothermal illitization of Upper
Palaeozoic deposits in the study area at 183F 4 Ma
(Schaltegger et al., 1995).
These hydrothermal phases coincide with tectonic
extension in the Atlantic and Tethyan realms: Prealps
in western Switzerland and France, which belong
palaeogeographically to the Brianc� onnais and Sub-
brianc�onnais (Borel, 1995), Western Alps (Lemoine et
Fig. 11. Hydrothermal activity within the investigation area and its
vicinity (for exact locations, see references below) documented by
chronometric ages of vein mineralizations and alterations of
minerals, based on data of Bonhomme et al. (1983), Bouladon
and de Graciansky (1985), Brockamp et al. (1994), Hagedorn and
Lippolt (1994), Lancelot et al. (1995), Lippolt and Kirsch (1994a,b),
Lippolt and Mertz (1989), Lippolt and Siebel (1991), Mertz et al.
(1991), Schaltegger et al. (1995), Wernicke and Lippolt (1993,
1994, 1995, 1997a,b), and Zuther and Brockamp (1988). Chro-
nostratigraphy of stage boundaries after Gradstein et al. (1995).
Keu: Keuper, R: Rhaetian, H: Hettangian, Si: Sinemurian, P:
Pliensbachian, To: Toarcian, A: Aalenian, Bj: Bajocian, B:
Bathonian, C: Callovian, O: Oxfordian, K: Kimmeridgian, T:
Tithonian, Be: Berriasian, V: Valanginian, Ha: Hauterivian, Br:
Barremian, Ap: Aptian, Ab: Albian, Ce: Cenomanian, Tu: Turonian,
C: Coniacian, and S: Santonian.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172164
al., 1986), Southern Alps (Bertotti et al., 1993), and
North Sea (e.g., Faerseth, 1996). The temporal coin-
cidence of hydrothermal and tectonic activities sug-
gests that the large parts of Europe were affected by
an extensional stress regime (e.g., Ziegler, 1990).
6. Discussion
The Aalenian and Oxfordian depocenters in the
study area are found above faults related to Permo–
Carboniferous (half-)grabens in the basement. There-
fore, the question arises as to whether the increase in
thickness resulted from compaction of the Palaeozoic
deposits or from active subsidence. The geologic
history of the study area implies that compaction of
the trough fill very likely has not caused the observed
facies and isopach pattern within the Mesozoic sedi-
mentary cover mainly because of two reasons: (1)
during the Permian, 1000–1500 m of sediments of the
basin fill were eroded and hence the remaining Palae-
ozoic deposits were overconsolidated (Kempter, 1987;
Schegg and Leu, 1998). At the locality Weiach, from
the Trias to the Aalenian, 615 m of compacted (1045
m decompacted) sediment accumulated and 855 m
(1450 m decompacted) up to the Upper Oxfordian.
Therefore, the Mesozoic sediments were not suffi-
ciently thick to induce significant compaction of the
substrate. (2) Thermal analysis and modelling by
Schegg and Leu (1998) showed that heat flux
decreased since the Permian tectonic phase and hence
compaction induced by diagenesis slowed down.
Because of the preconsolidation of the trough fill,
differential compaction probably did not cause the
observed facies pattern. The factors contributing to
subsidence—compaction, isostatic subsidence, and
the remaining, so-called ‘‘non-isostatic tectonic’’ sub-
Fig. 12. Tectonic subsidence and total subsidence at the three
localities Schafisheim, Riniken, and Weiach. After the Triassic,
significant tectonic subsidence only occurred during the Aalenian,
Bajo–Bathonian, and the Early to Middle Oxfordian. Chronostra-
tigraphy of stage boundaries after Menning et al. (2000) for the
Palaeozoic and Gradstein et al. (1995) for the Mesozoic. W:
Westphalian, S: Stephanian, Rl: Rotliegend, Z: Zechstein, Bu:
Buntsandstein, Mu: Muschelkalk, Keu: Keuper, R: Rhaetian, H:
Hettangian, Si: Sinemurian, P: Pliensbachian, To: Toarcian, A:
Aalenian, Bj: Bajocian, B: Bathonian, C: Callovian, O: Oxfordian,
K: Kimmeridgian, and T: Tithonian.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 165
sidence—were quantified (see Methods) for the three
sections drilled at Riniken, Weiach, and Schafisheim,
being located in the center, at the flank, and outside
the Late Palaeozoic basin (Fig. 12). This analysis
provides evidence that non-isostatic tectonic subsi-
dence contributed about 1/3 to the total subsidence.
Furthermore, the total subsidence of the two sections
drilled within the trough area was compared to that
outside (Fig. 13). All three sections display—inde-
pendent of their position relative to the Late Palae-
ozoic basin in the subsurface—subsidence pulses at
the same time, even if they differ in the amount of
subsidence. As all sections display synchronous sub-
sidence pulses, tectonic stretching is invoked as the
reason for the observed subsidence pattern.
Hydrothermal activity is synchronous with epi-
sodes of enhanced subsidence. As large parts of
Europe were affected by an extensional stress regime,
it is suggested that hydrothermal activity and subsi-
dence resulted from extension, and pre-existing faults
became reactivated as suggested by the spatial coin-
cidence of pre-existing faults and thickness maxima.
The depocenters of Mesozoic sediments are not
systematically developed above the Late Palaeozoic
basins. The spatial relation of depocenters and base-
ment structures (Figs. 5 and 9) suggests particularly (1)
that especially the faults bounding the Late Palaeozoic
basins became reactivated, and (2) that the reactivation
of basement structures did not affect all faults at the
same time. Although a fault may be within a stress
field favorable for its reactivation, that reactivation
does not necessarily occur simply because the base-
ment consists of a mosaic of blocks interacting with
each other. However, new faults might have formed in
addition. During the Early to Middle Jurassic rifting of
the Tethyian continental margins (e.g., Lemoine et al.,
1986), the j3-direction of the palaeostress field in the
area of the today’s Jura Mountains changed from N–S
to NW–SE (Philippe et al., 1996) and became favor-
able for the reactivation of NE–SW-trending struc-
tures. Therefore, the Aalenian subsidence phase in the
study area could be linked to the change of the
palaeostress field, leading to reactivation of roughly
E–W-trending basement structures.
During the Oxfordian, another episode of enhanced
subsidence is documented in the study area and ‘‘the
whole shelf, from the Helvetics in the south to the Jura
in the north, subsided abruptly, with short-time values
as high as 70 m/my on top of the basement of the Aar
massif and in the eastern Helvetic realm, and 20–40
m/my in the southern part of the Jura’’ (Wildi et al.,
1989, p. 833). For this time, an extension process
similar to the early Middle Jurassic one is not known
from the Tethys (e.g., Wildi et al., 1989). As reason
for the strong subsidence during the Oxfordian, Wildi
et al. (1989, p. 835) stated that ‘‘plastic extension of
the lower crust in a certain distance from the centre of
brittle extension of the upper crust has been demon-
strated by small-scale models to be a possible sub-
sidence mechanism (Allemand et al., 1989)’’. This
may be a plausible explanation because areas to the
west and to the north of the Swiss Jura experienced
extension, for instance, in the Prealps (Borel, 1995) or
the southern North Sea (e.g., Karner et al., 1987).
Alternatively, Lemoine et al. (1986) invoked a thermal
origin for the Oxfordian subsidence pulse.
In addition to the NE–SW-oriented Late Palae-
ozoic basins, N–S-trending structures associated with
the Rhenish Lineament (see above) probably also
affected the sedimentary cover. The spatial coinci-
Fig. 13. Comparison of the subsidence pattern for the sections
drilled at Schafisheim (outside the Late Palaeozoic basin) and
Riniken and Weiach above it. For the latter two sites, the Palaeozoic
sediment thickness was subtracted. At all sites, a very similar
subsidence pattern is found, which implies only little influence of
the Late Palaeozoic basin fill on the Mesozoic sediment accumu-
lation. Chronostratigraphy of stage boundaries after Gradstein et al.
(1995). Mu: Muschelkalk, Keu: Keuper, R: Rhaetian, H: Hettangian,
Si: Sinemurian, P: Pliensbachian, To: Toarcian, A: Aalenian, Bj:
Bajocian, B: Bathonian, C: Callovian, O: Oxfordian, K: Kimmerid-
gian, and T: Tithonian.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172166
dence of the eastern platform-basin transition during
the Late Oxfordian with the Rhenish Lineament
suggests a reactivation of this Late Palaeozoic fault
system. The at least partial reactivation of roughly N–
S-trending ‘‘Rhenish’’ structures can be explained by
the change of the palaeostress field, as stated by
Ziegler (1990, p. 102): ‘‘As a direct consequence of
the crustal separation between Europe and Africa. . .tensional stresses related to the Tethys rift system
relaxed in Europe. This caused a fundamental reor-
ganization of the stress regimes. . .. This reorientationis expressed by the abandonment of northeast–south-
west trending grabens and troughs and the develop-
ment of new northwest–southeast trending wrench
systems.’’ From the available data, however, it is not
yet clear if the N–S-trending facies boundaries were
induced by differential subsidence across a fault zone
or by local uplift in the southern Black Forest/Vosges
area due to wrench movements. Nonetheless, the
Oxfordian platform first established around topo-
graphic elevations (Allenbach, 2001, 2002). For the
Oxfordian sediments, the high-resolution sequence-
stratigraphic frame established by Pittet (1994), Pittet
and Strasser (1998), and Allenbach (2001) may help
to further differentiate the effects of sea-level changes
and differential subsidence.
As the faults in the basement have had such an
important effect on the lithology of the sedimentary
cover, the question arises, why faults within the
Mesozoic sedimentary cover have not been observed
in outcrops or on conventional 2D seismic records. We
assume that the Triassic evaporites have dissipated the
vertical tectonic movements along the faults to form
flexure-like deformations above the evaporites (Fig.
14). This has been described, for instance, by Withjack
and Callaway (2000). As the tectonic movements are
in the order of only a few tens of meters equivalent to
decompacted thickness changes, such movements can
easily be compensated for by plastically behaving salt
100 m thick. Therefore, the Mesozoic sedimentary
cover provides no direct evidence for synsedimentary
tectonic movements. Syndepositional tectonic activity,
however, is inferred from the spatial relationship to
tectonic structures in the basement of local hard-
grounds and hiatus beds and simultaneously filled
depressions during the Aalenian (Wetzel and Allia,
2000), and of rotated blocks (Bolliger and Burri,
1970), slumps, or carbonate platforms and marly
basins (Allenbach, 2001, 2002) during the Oxfordian.
Recently performed high-resolution 3D seismic stud-
ies (just outside the study area) provided evidence for
synsedimentary tectonic movements during the Juras-
sic and hence, support our findings (Birkhauser et al.,
2001). In addition, the movements along faults were
quantified to be in the order of less than 20 m.
7. Conclusions
Up to now, there has been no direct evidence for
tectonic activity during the Mesozoic in outcrops or on
seismic records (expect the very recent high-resolution
3D seismic analysis) in the Swiss Jura and hence the
Mesozoic is usually referred to as a period of tectonic
quiescence. However, facies boundaries and thickness
maxima within the sedimentary cover occur in close
spatial relationship to structures in the basement, such
as faults, half-grabens, and grabens, which formed
Fig. 14. Schematic representation of how extension within the
basement led to changes in relief of the seafloor. Lower part—initial
situation; upper part—seafloor topography after extension. Vertical
movements within the basement were transformed into flexural
deformations by Triassic salt.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 167
during the Late Palaeozoic. Compaction of the Palae-
ozoic graben fill is ruled out as reason for the observed
lithologic changes because more than 1000 m of the
basin fill were eroded during the Late Permian. This
amount is in the range of the Mesozoic sediment
thickness. Consequently, the remaining graben fill
was significantly precompacted and probably not
affected by the Mesozoic deposits.
The reactivation of basement structures is docu-
mented not only by changes of facies and sediment
thickness, but also by hydrothermal activity. Increased
hydrothermal activity coincides with subsidence
pulses within the study area (Sinemurian, Aalenian,
Bajocian/Bathonian, and Oxfordian). These subsi-
dence pulses are synchronous with extension phases
reported from the continental margins of the Tethys
and Atlantic oceans and the North Sea. Consequently,
the epicontinental area in-between was probably
affected by an extensional stress regime.
During reactivation, blocks within the basement
moved along faults. However, the vertical movements
were dissipated by Triassic salt, and the Jurassic
seafloor experienced only a flexural deformation,
which led to the formation of swells and depressions.
In this way, direct evidence of tectonic movements
such as synsedimentary faults could not develop. In
spite of the subtle deformation of the Mesozoic
seafloor, the shallow-marine depositional systems
responded distinctly even to small-scale changes in
water depth and documented the effects of differential
subsidence especially during periods of enhanced
subsidence—the Aalenian and the Oxfordian. During
both time intervals, about 1/3 of the total subsidence
was calculated to be of non-isostatic, tectonic origin.
The thickness of the Aalenian terrigenous mudrocks
(today about 100 m) clearly exceeds the depositional
water depth (20–50 m) and, therefore, extra accom-
modation space must have been provided during dep-
osition. Isopach anomalies, therefore, reflect diffe-
rentially subsiding areas. Storm-induced currents
flowed to the depocenters, which are located above
the NE–SW-trending Late Palaeozoic basins or their
boundary faults. Hardgrounds and hiatus beds formed
locally on elevated areas.
The Lower Oxfordian to basal Middle Oxfordian
facies boundaries also followed the NE–SW-trending
Late Palaeozoic troughs, whereas the Middle to Upper
Oxfordian deposits display N–S-trending facies
boundaries that are spatially related to the Rhenish
Lineament. Its reactivation during the Late Jurassic
was probably due to the changing stress field resulting
from crustal separation between Europe and Africa.
Facies changes and isopach anomalies are useful in
predicting the existence of tectonic structures in the
subsurface, although these structures cannot be
exactly located. By the same token, basement struc-
tures known from seismic records may help to identify
facies changes in the intermediate subsurface, even
though these changes are covered by a younger sedi-
ment series.
Acknowledgements
This was made possible by financial support of the
Schweizerische Nationalfonds zur Forderung der Wis-
senschaftlichen Forschung (grants no. 21-31115.91,
20-37269.93, 21-43103.95, and 20-48252.97). Dr. W.
Muller (NAGRA,Wettingen, Switzerland) gave access
to unpublished data. J. Tipper (Freiburg, FRG) im-
proved the English. A. Strasser (Fribourg, Switzerland)
and L. Jansa (Dartmouth, Canada) as journal reviewers
provided helpful comments. S. Lauer (Basel) drew all
figures. All these contributions are gratefully acknowl-
edged.
References
Aigner, T., Bachmann, G.H., 1992. Sequence-stratigraphic frame-
work of theGermanTriassic. SedimentaryGeology 80, 115–135.
Aigner, T., Reineck, H.-E., 1982. Proximality trends in modern
storm sands from the Helgoland Bight (North Sea) and their
implications for basin analysis. Senckenbergiana Maritima 14,
183–215.
Allemand, P., Brun, J.P., Davy, P., van den Driessche, J., 1989.
Symetrie et asymetrie des rifts et mecanismes d’aminicissement
de la lithosphere. Bulletin de la Societe Geologique de France 5,
445–551.
Allen, P.A., Allen, J.R., 1990. Basin Analysis. Blackwell, Oxford,
451 pp.
Allenbach, R., 2001. Up with sea-level; down with differential sub-
sidence—a new interpretation of the Oxfordian of northern
Switzerland. Eclogae Geologicae Helvetiae 94, 265–287.
Allenbach, R.P., 2002. The ups and downs of ‘‘Tectonic Quies-
cence’’—recognizing differential subsidence in the epicontinen-
tal sea of the Oxfordian in the Swiss JuraMountains. Sedimentary
Geology 150, 323–342.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172168
Allia, V., 1996. Sedimentologie und Ablagerungsgeschichte des
Opalinustons in der Nordschweiz. Dissertationen aus dem Geo-
logisch-Palaontologischen Institut der Universitat Basel, No. 10,
185 pp.
Arthaud, F., Matte, P.H., 1977. Late Paleozoic strike-slip faulting in
southern Europe and northern Africa. Geological Society of
America Bulletin 88, 1305–1320.
Bertotti, G., Picotti, V., Bernoulli, D., Castellarin, A., 1993. From
rifting to drifting: tectonic evolution of the South-Alpine upper
crust froom the Triassic to the Early Cretaceous. Sedimentary
Geology 86, 53–76.
Birkhauser, P., Roth, P., Meier, B., Naef, H., 2001. 3D-Seismik:
Raumliche Erkundung der mesozoischen Sedimentschichten
im Zurcher Weinland. NAGRATechnischer Bericht 00-03. NA-
GRA, Wettingen, 158 pp.
Bitterli, T., 1992. Die tektonische Materialbilanz im ostlichen Fal-
tenjura (Weissenstein-, Farisberg- und Passwang-Antiklinale).
Dissertation University of Basel, 170 pp.
Bitterli-Brunner, P., 1987. Geologischer Fuhrer der Region Basel.
Veroffentlichungen aus dem Naturhistorischen Museum Basel,
vol. 19. Birkhauser, Basel, 232 pp.
Boigk, H., Schoneich, H., 1974. The Rhinegraben: geologic history
and neotectonic activity. In: Illies, J.H., Fuchs, K. (Eds.), Ap-
proaches to Taphrogenesis. Inter-Union Commission on Geo-
dynamics, Scientific Report, vol. 8. Schweizerbart, Stuttgart,
pp. 60–71.
Bolliger, W., Burri, P., 1970. Sedimentologie von Schelf-Carbonat-
en und Beckenablagerungen im Oxfordien des zentralen Schwe-
izer Jura. Beitrage zur Geologischen Karte der Schweiz, Neue
Folge, vol. 140. Schweizerische Geologische Kommission,
Bern, 96 pp.
Bonhomme, M.G., Buhmann, D., Besnus, Y., 1983. Reliability of
K/Ar-dating of clays and silifications associated with vein min-
eralizations in western Europe. Geologische Rundschau 72,
105–117.
Bonijoly, D., Perrin, J., Roure, F., Bergerat, F., Courel, L., Elmi, S.,
Mignot, A., GPF Team, 1996. The Ardeche palaeomargin of the
South-East Basin of France: Mesozoic evolution of a part of the
Tethyan continental margin (Geologie Profonde de la France
Programme). Marine and Petroleum Geology 13, 607–623.
Borel, G., 1995. Prealpes medianes romandes: courbes de subsi-
dence et implications geodynamiques. Bulletin de Societe de
Vaudois de Science Naturelle 83, 293–315.
Bouladon, J., de Graciansky, P.-C., 1985. Les mineralisations dites
de couverture (plomb, zinc, cuivre, uranium, barytine, fluo-
rine,. . .) du Trias au Pliocene, en France. Chronique des Mines
et de la Recherche Miniere 480, 17–33.
Brandley, R.T., Krause, F.F., Varsek, J.L., Thruston, J., Spratt, D.A.,
1996. Implied basement-tectonic control on deposition of lower
Carboniferous carbonate ramp, southern Cordillera, Canada.
Geology 24, 467–470.
Branger, P., Gonnin, C., 1994. Distribution des ammonites et dyna-
mique sedimentaire sur le seuil du Poitou de l’Aalenien au
Bajocien. Servizio Geologico Nazionale, Miscellanea 5, 293–
295, Poster I.
Brockamp, O., Clauer, N., Zuther, M., 1994. K–Ar dating of epi-
sodic Mesozoic fluid migrations along the fault system of
Gernsbach between the Moldanubian and Saxothuringian
(Northern Black Forest). Geologische Rundschau 83, 180–185.
Bruguier, O., Becq-Giraudon, J.F., Bosch, D., Lancelot, J.R., 1998.
Late Visean hidden basins in the internal zones of the Variscan
belt: U–Pb zircon evidence from the French Massif Central.
Geology 26, 627–630.
Buchi, U.P., Lemcke, K., Wiener, G., Zimdars, J., 1965. Geologi-
sche Ergebnisse der Erdolexplorationsbohrungen aus das Mes-
ozoikum im Untergrund des schweizerischen Molassebeckens.
Bulletin der Vereinigung schweizerischer Petroleum-Geologen
und -Ingenieure 32, 7–38.
Burkhalter, R.M., 1996. Die Passwang-Alloformation (unteres Aa-
lenie bis unteres Bajocicien) im zentralen und nordlichen
Schweizer Jura. Eclogae Geologicae Helvetiae 89, 875–934.
Clauer, N., Zwingmann, H., Chauduri, S., 1996. Isotopic (K–Ar
and oxygen) constraints on the extent and importance of the
Liassic hydrothermal activity in western Europe. Clay Minerals
31, 301–318.
de Graciansky, P.-C., Hardenbol, J., Jacquin, T., Vail, P.R., Farley,
M.B., 1998. Sequence Stratigraphy of European Basins. SEPM
(Society of Sedimentary Geology) Special Publication 60.
SEPM (Society of Sedimentary Geology), Tulsa (Oklahoma),
786 pp.
de Wet, C.B., 1998. Deciphering the sedimentological expression of
tectonics, eustasy, and climate: a basinwide study of the Coral-
lian Formation, southern England. Journal of Sedimentary Re-
search 68, 653–667.
Diebold, P., 1988. Der Nordschweizer Permokarbon-Trog und die
Steinkohlenfrage der Nordschweiz. Vierteljahrsschrift der Na-
turforschenden Gesellschaft in Zurich 133, 143–174.
Diebold, P., Noack, T., 1997. Late Palaeozoic troughs and Tertiary
structures in the eastern Folded Jura. In: Pfiffner, O.A., Lehner,
P., Heitzmann, P., Mueller, S., Steck, A. (Eds.), Deep Structure
of the Swiss Alps. Birkhauser, Basel, pp. 59–63.
Diebold, P., Naef, H., Ammann, M., 1991. Zur Tektonik der zen-
tralen Nordschweiz. NAGRA Technischer Bericht 90-04. NA-
GRA, Wettingen, 277 pp.
Diem, B., 1985. Analytical method for estimating palaeowave cli-
mate and water depth from wave ripple marks. Sedimentology
32, 705–720.
Dromart, G., Allemand, P., Quiquerez, A., 1998. Calculating rates
of syndepositional normal faulting in the western margin of the
Mesozoic Subalpine Basin (south-east France). Basin Research
10, 235–260.
Duke, W.L., Arnott, R.W.C., Cheel, R.J., 1991. Shelf sandstones
and hummocky cross-stratification: new insights on a stormy
debate. Geology 19, 625–628.
Faerseth, R.B., 1996. Interaction of Permo–Triassic and Jurassic
extensional fault-blocks during the development of the northern
North Sea. Journal of the Geological Society (London) 153,
931–944.
Flugel, E., 1978. Mikrofazielle Untersuchungsmethoden von Kalk-
en. Springer, Berlin, 454 pp.
Gonzalez, R., 1996. Response of shallow-marine carbonate facies to
third-order and high-frequency sea-level fluctuations: Hauptro-
genstein Formation, northern Swirtzerland. Sedimentary Geol-
ogy 102, 111–130.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 169
Gradstein, F.M., Agterberg, F.P., Ogg, J.G., Hardenbol, J., van
Veen, P., Thierry, J., Huang, Z., 1995. A Triassic, Jurassic and
Cretaceous time scale. In: Berggren, W.A., Kent, D.V., Aubry,
M.-P., Hardenbol, J. (Eds.), Geochronology, Time Scales and
Global Stratigraphic Correlation. SEPM (Society for Sedimen-
tary Geology) Special Publication, vol. 54. SEPM (Society for
Sedimentary Geology), Tulsa, OK, pp. 95–126.
Gygi, R.A., 1969. Zur Stratigraphie der Oxford Stufe (oberes Jura-
System) der Nordschweiz und des suddeutschen Grenzge-
bietes. Beitrage zur Geologischen Karte der Schweiz, Neue
Folge, vol. 136. Schweizerische Geologische Kommission,
Bern, 123 pp.
Gygi, R., 1990. Die Palaogeographie im Oxfordium und fruhesten
Kimmeridgium in der Nordschweiz. Jahreshefte des Geologi-
schen Landesamtes Baden-Wurttemberg 32, 207–222.
Gygi, R.A., Persoz, F., 1986. Mineralostratigraphy, litho- and bio-
stratigraphy combined in correlation of the Oxfordian (Late
Jurassic) formations of the Swiss Jura range. Eclogae Geologi-
cae Helvetiae 79, 385–454.
Hagedorn, B., Lippolt, H.J., 1994. Isotopische Alter von Zerrut-
tungszonen als Altersschranken der Freiamt-Sexau-Mineralisa-
tion (Mittlerer Schwarzwald). Geologisches Landesamt Baden-
Wurttemberg Abhandlungen 14, 205–219.
Haq, B.U., Hardenbol, J., Vail, P.R., 1987. Chronology of fluctuat-
ing sea levels since the Triassic. Science 235, 1156–1167.
Harland, W.B., Armstrong, R.L., Cox, L.E., Smith, A.G., Smith,
D.G., 1989. A Geologic Time Scale Cambridge Univ. Press,
Cambridge, 263 pp.
Harris, P.M., Saller, A.H., Simo, J.A., 1999. Advances in Carbonate
Sequence Stratigraphy: Applications to Reservoirs, Outcrops
and Models. SEPM (Society for Sedimentary Geology) Special
Publication, vol. 63. SEPM (Society for Sedimentary Geology),
Tulsa, OK, 421 pp.
Johnson, H.D., Baldwin, C.T., 1996. Shallow clastic seas. In: Read-
ing, H.R. (Ed.), Sedimentary Environments: Processes, Facies
and Stratigraphy. Blackwell, Oxford, pp. 232–280.
Karner, G.D., Lake, S.D., Dewey, J.F., 1987. The thermal and me-
chanical development of the Wessex Basin, southern England.
In: Coward, M.P., Dewey, J.F., Hancock, P.L. (Eds.), Continen-
tal Extensional Tectonics. Geological Society London Special
Publication 28, pp. 517–536.
Keeley, M.L., 1996. The Irish Variscides: problems, perspectives
and some solutions. Terra Nova 8, 259–269.
Kempter, E.H.K., 1987. Fossile Maturitat, Palaothermogradienten
und Schichtlucken in der Bohrung Weiach im Lichte von Mod-
ellberechnungen der thermischen Maturitat. Eclogae Geologicae
Helvetiae 80, 543–552.
Krohe, A., 1996. Variscan tectonics of central Europe: postaccre-
tionary intraplate deformation of weak continental lithosphere.
Tectonics 15, 1364–1388.
Lancelot, J., Brique, L., Respaut, J.-P., Clauer, N., 1995. Geochimie
isotopique des systemes U–Pb/Pb–Pb et evolution polyphasee
des gıtes d’uranium du Lodevois et du sud du Massif central.
Chronique de la Recherche Miniere 521, 3–18.
Laubscher, H.P., 1965. Ein kinematisches Modell der Jurafaltung.
Eclogae Geologicae Helvetiae 58, 231–318.
Laubscher, H., 1986. The eastern Jura: relations between thin-
skinned and basement tectonics, local and regional. Geologische
Rundschau 75, 535–553.
Laubscher, H., 1987. Die tektonische Entwicklung der Nord-
westschweiz. Eclogae Geologicae Helvetiae 80, 278–303.
Lemoine, M., Bas, T., Arnaud-Vanneau, A., Arnaud, H., Dumont,
T., Gidon, M., Bourbon, M., de Graciansky, P.-C., Rudkiewicz,
J.-L., Megard-Galli, J., Tricart, P., 1986. The continental margin
of the Mesozoic Tethys of the western Alps. Marine and Petro-
leum Geology 3, 179–199.
Lippolt, H.J., Kirsch, H., 1994a. Isotopic investigation of post-Va-
riscan plagioclase sericitization in the Schwarzwald gneiss mas-
sif. Chemie der Erde 54, 179–198.
Lippolt, H.J., Kirsch, H., 1994b. Ar40/Ar39-Untersuchungen an
serizitisierten Plagioklasen des Frankenstein-Gabbros (NW-
Odenwald) im Hinblick auf ihren Alterationszeitpunkt. Geolo-
gisches Jahrbuch Hessen 122, 123–142.
Lippolt, H.J., Mertz, D.F., 1989. Mit Magmatismus korrelierte und
unkorrelierte Hydrothermalphasen imKristallin des Sudschwarz-
waldes. European Journal of Mineralogy 1, 111 (Beiheft 1).
Lippolt, H.J., Siebel, W., 1991. Evidence for multi-stage alteration
of Schwarzwald lamprophyres. European Journal of Mineralogy
3, 587–601.
Loucks, R.G., Sarg, J.F., 1993. Carbonate sequence stratigraphy:
recent developments and applications. American Association
of Petroleum Geologists Memoir, vol. 57. American Association
of Petroleum Geologists, Tulsa, OK, 545 pp.
Matter, A., Peters, T.J., Blasi, H.-R., Ziegler, H.-J., 1987. Sondier-
bohrung Riniken. NAGRA Technischer Bericht 86-02, NA-
GRA, Wettingen, 125 pp.
Matter, A., Peters, T.J., Blasi, H.-R., Meyer, J., Ischni, H., Meyer,
C., 1988. Sondierbohrung Weiach. NAGRATechnischer Bericht
86-01. NAGRA, Wettingen, 438 pp.
Menard, G., Molnar, P., 1988. Collapse of a Hercynian Tibetan
Plateau into a late Palaeozoic European Basin and Range prov-
ince. Nature 334, 235–237.
Menning, M., Weyer, D., Drozdzewski, G., van Ameron, H.W.J.,
Wendt, I., 2000. A Carbonifreous time scale 2000: discussion
and use of geological parametres as time indicators from central
and western Europe. Geologisches Jahrbuch. Reihe A 156, 3–44.
Mertz, D.F., Karpenko, M.I., Ivanenko, V.V., Lippolt, H.J., 1991.
Evidence for Jurassic tectonism in the Schwarzwald basement
(SW Germany) by laser probe 40Ar/39Ar dating of authigenic
feldspar. Naturwissenschaften 78, 411–413.
Metz, R., 1970. Dehnungsstrukturen im Grundgebirge des Schwarz-
walds vor Beginn der Grabentektonik. In: Illies, J.H., Muller, S.
(Eds.), Graben Problems. International Upper Mantel Project,
Scientific Report, vol. 27. Schweizerbart, Stuttgart, pp. 87–90.
Milkert, D., 1994. Auswirkungen von Sturmen auf die Schlickse-
dimente der westlichen Ostsee. Berichte—Reports, Geologisch-
Palaontologisches Institut der Universitat Kiel, vol. 66. Geo-
logisch-Palaontologisches Institut der Universitat Kiel, Kiel,
153 pp.
Myrow, P.M., Southard, J.B., 1996. Tempestite deposition. Journal
of Sedimentary Research 66A, 875–887.
Ohmert, W., Rolf, C., 1994. The Aalenian boundaries at Wittnau
(Oberrhein area, south west Germany). Servizio Geologico Na-
zionale, Miscellanea 5, 33–61.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172170
Pfiffner, O.A., 1993. Palinspastic reconstruction of the pre-Triassic
basement units in the Alps: the central Alps. In: von Raumer, J.,
Neubauer, F. (Eds.), Pre-Mesozoic Geology in the Alps. Spring-
er, Berlin, pp. 29–39.
Philippe, Y., Coletta, B., Deville, E., Mascle, A., 1996. The Jura
fold-and-thrust belt: a kinematic model based on map-balancing.
In: Ziegler, P.A., Horvath, F. (Eds.), Peri-Tethys Memoir 2:
Structure and Prospects of Alpine Basins and Forelands. Mem-
oires du Museum National d’Histoire Naturelle, vol. 170. Edi-
tions du Museum, Paris, pp. 235–261.
Pittet, B., 1994. Modele d’estimation de la subsidence et des varia-
tions du niveau marin: Un exemple de l’Oxfordien du Jura
suisse. Eclogae Geologicae Helvetiae 87, 513–543.
Pittet, B., Strasser, A., 1998. Long-distance correlations by se-
quence stratigraphy and cyclostratigraphy: examples and impli-
cations (Oxfordian from the Swiss Jura, Spain, and Normandy).
Geologische Rundschau 86, 852–874.
Posamentier, H.W., Summerhayes, C.P., Haq, B.U., Allen, G.P.,
1993. Sequence Stratigraphy and Facies Associations. Interna-
tional Association of Sedimentologists Special Publication, vol.
18. Blackwell, Oxford, 644 pp.
Reineck, H.-E., 1977. Natural indicators of energy level in recent
sediments: the application of ichnology to a coastal engineering
problem. In: Crimes, T.P., Harper, J.C. (Eds.), Trace Fossils 2.
Geological Journal Special Issue, vol. 9. Seel House Press,
Liverpool, pp. 265–272.
Robin, C., Guillocheau, F., Gaulier, J.-M., 1998. Discriminating
between tectonic and eustatic controls on the stratigraphic re-
cord in the Paris basin. Terra Nova 10, 323–329.
Sarg, J.F., 1988. Carbonate sequence stratigraphy. In: Wilgus, C.K.,
Hastings, B.S., Kendall, S.C., Posamentier, H.W., Ross, C.A.,
van Wagoner, J.C. (Eds.), Sea-Level Changes: An Integrated
Approach. Society of Economic Paleontologists and Mineralo-
gists Special Publication, vol. 42. Society of Economic Paleon-
tologists and Mineralogists, Tulsa, OK, pp. 153–181.
Schafer, A., 1986. Die Sedimente des Oberkarbons und Unterrot-
liegenden im Saar-Nahe-Becken. Mainzer Geowissenschaftliche
Mitteilungen 15, 239–365.
Schaltegger, U., Corfu, F., 1995. Late Variscan ‘‘Basin and Range’’
magmatism and tectonics in the central Alps: evidence from U–
Pb geochronology. Geodinamica Acta 8, 82–98.
Schaltegger, U., Zwingmann, H., Clauer, N., Larque, P., Stille, P.,
1995. K–Ar dating of a Mesozoic hydrothermal activity in
Carboniferous to Triassic clay minerals of northern Switzerland.
Schweizerische Mineralogische und Petrographische Mitteilun-
gen 75, 163–176.
Schegg, R., Leu, W., 1998. Analysis of erosion events and palae-
ogeothermal gradients in the North Alpine Foreland Basin of
Switzerland. In: Duppenbecker, S.J., Iliffe, J.E. (Eds.), Basin
Modelling: Practice and Progress. Geological Society Special
Publications, vol. 141. Geological Society of London, London,
pp. 137–155.
Schroder, B., Ahrendt, H., Peterek, A., Wemmer, K., 1997. Post-
Variscan sedimentary record of the SW margin of the Bohemian
massif: a review. Geologische Rundschau 86, 178–184.
Shail, R.K., Alexander, A.C., 1997. Late Carboniferous to Triassic
reactivation of Variscan basement in the western English Chan-
nel: evidence from onshore exposures in south Cornwall. Jour-
nal of the Geological Society (London) 154, 163–168.
Spears, D.A., 1989. Aspects of iron incorporation into sediments
with special reference to the Yorkshire Ironstones. In: Young,
T.P., Taylor, W.E.G. (Eds.), Phanerozoic Ironstones. Geological
Society London Special Publication 46, 19–30.
Sweeney, J., Burnham, A.K., 1990. Evaluation of a simple model of
vitrinit reflectance based on chemical kinetics. American Asso-
ciation of Petroleum Geologists Bulletin 74, 1559–1570.
Thury, M., Gautschi, A., Mazurek, M., Muller, W.H., Naef, H.,
Pearson, F.J., Vomvoris, S., Wilson, W., 1994. Geology and
Hydrogeology of the Crystalline Basement of Northern Switzer-
land. NAGRA Technical Report 93-01. NAGRA, Wettingen,
424 pp.
Todorov, I., Schegg, R., Wildi, W., 1993. Thermal maturity and
modelling of Mesozoic and Cenozoic sediments in the south
of the Rhine Graben and the eastern Jura (Switzerland). Eclogae
Geologicae Helvetiae 86, 667–692.
von Raumer, J.F., 1998. The Paleozoic evolution in the Alps: from
Gondwana to Pangea. Geologische Rundschau 87, 407–435.
Waples, D.W., Kamata, H., 1993. Modelling porosity reduction as a
series of chemical and physical processes. In: Dore, A.G., Au-
gustson, J.H., Hermanrud, C., Stewart, D.J., Sylta, E. (Eds.),
Basin Modelling: Advances and Applications. Norwegian Pe-
troleum Society Special Publication, vol. 3. Norwegian Petro-
leum Society, Oslo, pp. 303–320.
Weimer, P., Posamentier, H.W., 1993. Siliciclastic sequence stratig-
raphy: recent developments and applications. American Associ-
ation of Petroleum Geologists Memoir, vol. 58. American
Association of Petroleum Geologists, Tulsa, OK, 492 pp.
Wernicke, R.S., Lippolt, H.J., 1993. Botryoidal hematite from the
Schwarzwald (Germany): heterogeneous uranium distributions
and their bearing on the helium dating method. Earth and Plan-
etary Science Letters 114, 287–300.
Wernicke, R.S., Lippolt, H.J., 1994. 4He age discordance and re-
lease behavior of a double shell botryoidal hematite from the
Schwarzwald, Germany. Geochimica et Cosmochimica Acta 58,
421–429.
Wernicke, R.S., Lippolt, H.J., 1995. Direct isotope dating of a
northern Schwarzwald qtz-ba-hem vein. Neues Jahrbuch fur
Mineralogie Monatshefte 1995, 161–172.
Wernicke, R.S., Lippolt, H.J., 1997a. (U +Th)–He evidence of Ju-
rassic continuous hydrothermal activity in the Schwarzwald
basement, Germany. Chemical Geology 138, 273–285.
Wernicke, R.S., Lippolt, H.J., 1997b. Evidence of Mesozoic multi-
ple hydrothermal activity in the basement at Nonnenmattweiher
(southern Schwarzwald) Germany. Mineralium Deposita 32,
197–200.
Wetzel, A., Allia, V., 2000. The significance of hiatus beds in shal-
low-water mudstones: an example from the Middle Jurassic of
Switzerland. Journal of Sedimentary Research 70, 170–180.
Wildi, W., Funk, H., Loup, B., Edgardo, A., Huggenberger, P.,
1989. Mesozoic subsidence history of the European marginal
shelves of the alpine Tethys (Helvetic realm, Swiss Plateau and
Jura). Eclogae Geologicae Helvetiae 82, 817–840.
Withjack, M.O., Callaway, S., 2000. Active normal faulting beneath
a salt layer: an experimental study of deformation patterns in the
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172 171
cover sequence. American Association of Petroleum Geologists
Bulletin 84, 627–651.
Wright, V.P., Burchette, T.P., 1996. Shallow-water carbonate en-
vironments. In: Reading, H.R. (Ed.), Sedimentary Environ-
ments: Processes, Facies and Stratigraphy. Blackwell, Oxford,
pp. 325–394.
Ziegler, M.A., 1989. North German Zechstein facies patterns in
relation to their substrate. Geologische Rundschau 78, 105–127.
Ziegler, P.A., 1990. Geological Atlas of Western and Central Eu-
rope. Shell Internationale Petroleum Maatschappij, Den Haag,
239 pp.
Zuther, M., Brockamp, O., 1988. The fossil geothermal system of
the Baden–Baden trough (Northern Black Forest, F.R. Ger-
many). Chemical Geology 71, 337–353.
A. Wetzel et al. / Sedimentary Geology 157 (2003) 153–172172