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: andreas.wetzel@unibas.ch (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.

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