by massive ooid-rich deposits. The interval with more *Bacinella-Lithocodium* oncoids (from 4.05 to 4.2 m) is interpreted as the maximum flooding. The transgressive deposit of the second elementary sequence consists of an ooid-dominated grainstone with cross-stratifications. Its top (at 5.05 m) reveals bioperforations and is assumed to represent the maximum-flooding surface. The two following beds consist of *Bacinella-Lithocodium* oncoid rudstones to floatstones and are interpreted as the highstand deposit. The third elementary sequence is the thickest sequence of the section and implies maximum accommodation gain. It consists of massive *Bacinella-Lithocodium* oncoid floatstones interpreted as open lagoon deposits. The fourth elementary sequence is also thick and composed of open lagoon facies: oncoid wackestones in the transgressive deposit and peloid-foraminifer wackestones in the highstand deposit. The transgressive deposit of the fifth elementary sequence consists of thin peloid wackestone beds, interpreted as open lagoonal deposits and the highstand deposit consists of more bioclastic mudstones, interpreted as more protected lagoonal facies. The general facies evolution from the third to the fifth elementary sequence points to general loss of accommodation.

The two slightly undulated beds rich in quartz and benthic foraminifera (from 12.05 to 12.2 m) and interbedded with very thin marly layers are interpreted as the small-scale SB 2. Small-scale sequence 2 presents thinner sequences than small-scale sequence 1. In the second small-scale sequence, only four elementary sequences have been examined. The first one consists of two undulated beds, interpreted as the early transgressive deposit and two more massive beds interpreted as the transgressive and highstand deposits, respectively. The second elementary sequence is relatively thin and consists of two peloid and foraminifer wackestone-mudstone beds with quartz, interpreted as the transgressive and highstand deposits, respectively. These beds are separated by a thin marly interval interpreted as the maximum flooding. The third elementary sequence consists of a nodular limestone bed and a marly interval, interpreted as the transgressive and highstand deposits, respectively. The maximum-flooding surface of this elementary is assumed to be located on the top of the nodular limestone bed. This MFS is correlated with the medium-scale MF Ox6+. The fourth elementary sequence is composed of two massive limestone beds, corresponding to the transgressive deposits and a thin marly interval that is thought to represent to the highstand.

## Synthesis of the Swiss Jura sections

Detailed facies analyses of the sections allows to sketch the platform morphology at different time intervals (cf. Chap. 7). The facies distribution and thickness variations along the six studied sections point to a complex platform morphology with high areas and depressions, where siliciclastics preferentially accumulated.

The Voyeboeuf and Savagnières sections mainly reveal low-energy facies without important clay deposits and imply relatively protected environments. The geographical position of the Voyeboeuf section suggests a proximal position in comparison to the other sections. The Vorbourg section reveals more open-marine environments with bioclastic (coral rubble, brachiopods), oncolitic, and oolitic deposits. The low thickness of the Hautes-Roches section and the dominance of high-energy ooid deposits suggests a high position on the platform. The development of ooid bars but also of coral patch-reefs and carpets in the Court section points to normal-marine conditions in open lagoonal environments. The Pertuis section is dominated by open-marine facies, characterized by *Bacinella-Lithocodium* oncoids, peloids, and bioclasts. This indicates a relative distal position on the platform. As in the Savagnières section, the deposits of small-scale sequence 2 present more protected facies probably caused by the development of a barrier in distal position.

Through all the sections, siliciclastics (quartz and clays) preferentially occur in protected lagoonal environments.

### 4.1.2 Lorraine

*Pagny-sur-Meuse* (Figs 4.16 and 4.17a-b; Pl. 15/1-7, Pl. 16/1-3)

#### Geography and stratigraphy

The Pagny-sur-Meuse section is located 33 km west of Nancy in a large quarry exploited by the Novacarb society (Fig. 1.2c and 4.16). It provides fresh accessible outcrops due to continuous quarrying. These outcrops reveal Middle to Late Oxfordian deposits from the Lorraine shallow-water carbonate platform. Such outcrop conditions are favorable for detailed sedimentological and geochemical studies
<figure><img src="image_44.png" /></figure>

**Fig. 4.16** - Location of the Pagny-sur-Meuse. Based on topographic map of Toul (33150), IGN Topo 25 series.

(e.g., VINCENT 2001; CARPENTIER 2004; OLIVIER et al. 2004a; CARPENTIER et al. 2006).

Logged in the upper part of the quarry, the section is about 10 m thick and exposes facies of the *Calcaires à polypiers de Pagny* (CPP), the *Oolithe de Saucourt inférieure* (OSI), and the *Marnes à huîtres de Pagny* (MHP) members (Fig. 4.17). A brief description of these members is given in Chap. 1. The biostratigraphic position of the Pagny-sur-Meuse section is based on the discovery of two ammonites, one at the base and the other at the top of the *Calcaires à polypiers de Pagny* (c.f. Fig. 4.17; CARPENTIER 2004). They indicate that the Bifurcatus-Bimammatum limit is positioned within the “Calcaires à polypiers de Pagny” and that the overlying deposits (the *Oolithe de Saucourt inférieure*, the *Marnes à huîtres de Pagny*, the *Oolithe de Saucourt supérieure*, and the *Marnes à serpules de Pagny*) are dated of the Bimammatum zone (CARPENTIER 2004; CARPENTIER et al. 2006).

The section has been logged through the *Oolithe de Saucourt inférieure* and the *Marnes à huîtres de Pagny*, which are situated in the lower part of the Bimammatum zone and thus are more or less time equivalent to the studied interval in the Swiss Jura.

## Sedimentological interpretation

The base of the Pagny-sur-Meuse section is composed of relatively thick dark limestones (packstone-wackestone) with bioclasts, ooids, and

locally detrital quartz. A packstone level of reworked and perforated mud pebbles occurs from 2.3 to 2.5 m. Laterally, these limestones pass into coral bioherms. These deposits represent an open lagoonal environment with local coral patch-reefs.

A decrease of accommodation space and a shallowing of the depositional environment are displayed by the marly interval (from 3.4 to 4.6 m) in the top part of the *Calcaires à polypiers de Pagny*. Marls are interbedded with argillaceous limestones, which are rich in bioclasts (e.g., oysters, echinoderms), quartz, and locally glauconite. The maximum of shallowing (4.55 m) is indicated by a reduction of faunal diversity (mainly ostracodes, gastropods, oysters) and coal debris. The following interval consists of marls (from 4.7 to 5.05 m) with lenticular bedding, commonly found in mud flats, and local coal debris (at 5 m). This interval reveals the shallowest facies corresponding to a low-energy environment (mud-dominated) with tidal currents (lenticular bedding).

The following deposits (from 5.05 to 7.65 m) consist of massive limestones rich in oysters and ooids. This abrupt lithological change points to an increase of the accommodation space during a deepening phase. The first carbonate bed has an erosive base (at 5.05 m), a wavy bed surface and consists of bioturbated (horizontal burrows) ooid- and oyster-rich grainstone-packstone. This bed presents important thickness variations and is interpreted as an ooid bar deposit. A tidal influence is demonstrated by herringbone cross-stratifications that indicate a bi-directionnal current. The following limestone beds (from 5.1 to 6.55 m) consist of oyster- and ooid-rich grainstone-packstone. Tidal influence is only observed at the base of these deposits from 5.05 to 5.45 m, suggesting a deepening of the depositional environment. At the top of a massive bed (at 6.55 m), an oyster accumulation (lumachella) probably reflects a reduction of the sedimentation rate. This level is followed by a relatively thin argillaceous oo-bioclastic (oysters, echinoderms) packstone with some quartz.

The deposition of three massive ooid- and oyster-rich packstone to grainstone beds (from 6.65 to 7.65 m), interpreted as ooid dunes (Pl. 15/6-7), points to a new increase of accommodation space and/or local migration of subtidal ooid bars. The top part of the third bed (at 7.55 m) is a wavy surface with oyster encrustations on top of the dunes and mud in depressions (Pl. 15/7). This remarkable surface is interpreted as a subtidal hardground, pointing to a period of sediment
<figure><img src="image_45.png" /><figcaption>Fig. 4.17a - Pagny-sur-Meuse section (1).</figcaption></figure>
<figure><img src="image_46.png" /><figcaption>Fig. 4.17b - Pagny-sur-Meuse section (2).</figcaption></figure>

starvation. Finally, marls, belonging to the *Marnes à huîtres de Pagny*, cap the ooid dunes. This facies change points to the turnover towards a siliciclastic-dominated sedimentary system, probably caused by a climatic change to more humid conditions.

## Sequence-stratigraphic interpretation

Marls (from 4.75 to 5.05 m) with lenticular bedding, characteristic of upper intertidal environments, represent the shallowest facies and are interpreted as

the medium-scale SB Ox6+. The base of the massive limestone is interpreted as the medium-scale TS Ox6+. This facies evolution reflects an important relative sea-level rise.

Small-scale sequence 1 consists of five elementary sequences. The first one is composed of one or two grainstone beds with herringbone cross-stratifications and a discontinuous marly level with lenticular bedding. These deposits present important lateral thickness variations. On the top of the limestone bed
(at 5.2 m), a hardground, representing an omission period, is interpreted as the maximum flooding surface. The second elementary sequence consists of two limestone beds with typical structures of tidal currents (i.e., herringbone cross-stratifications). The maximum flooding is assumed to be located between these two beds. The third elementary sequence consists of two limestone beds separated by a thin clay-rich interval, which is interpreted as the maximum flooding. A tidal influence is not recorded in the third elementary sequence, suggesting a general deepening of the depositional environment. The fourth elementary sequence comprises two limestone beds interpreted as the transgressive deposit for the first bed and the highstand deposit for the second bed. The fifth elementary sequence consists of two relatively thick limestone beds and an argillaceous packstone bed. The first two beds are interpreted as the transgressive deposit and the thick bed as the highstand deposit. The oyster lumachella (at 6.55 m) is a good candidate for the maximum-flooding surface.

In small-scale sequence 2, only three elementary sequences are investigated. Each consists of one ooid limestone bed. The sequence boundaries are assumed to be located at the base of these beds (at 6.65 m, 6.95 m, and 7.15 m), where ooids and/or quartz and lithoclasts are more abundant. The oyster-encrusted hardground at 7.65 m is considered to represent the maximum flooding of the third elementary sequence as well as of small-scale sequence 2. By lateral correlation with the Swiss Jura sections, this maximum flooding surface is interpreted as belonging to the medium-scale MF Ox6+.

<figure><img src="image_47.png" /><figcaption>**Fig. 4.18** - Location of the Latrecey-sur-Ormoy section. Based on topographic maps of Montigny-sur-Aube (3019O) and Châteauvillain (3019E), IGN Topo 25 series.</figcaption></figure>

## 4.2 DEEP PLATFORM SECTIONS

Two sections of different geographical and palaeogeographical context as well as of different deep platform facies are investigated (cf. Chap. 1).

### 4.2.1 Haute-Marne

*Latrecey-sur-Ormoy* (Figs 4.18 and 4.19a; Pl. 16/4-5)

**Geography and stratigraphy**

The Latrecey-sur-Ormoy section is located in the Haute-Marne department in an abandoned quarry ("NW quarry" in COURVILLE & VILLIER 2003) close to Latrecey-sur-Ormoy (Figs 1.2c and 4.18). The logged section is only about 3.5 m thick and exposes the "Calcaires de Latrecey" This member is characterized by homogeneous carbonate facies with quite poor and poorly diversified benthic macrofauna (COURVILLE & VILLIER 2003) and represents an outer, (relatively deep) distal platform environment (cf. Chap. 1).

Lithostratigraphy and biostratigraphy of this section were first described by COURVILLE & VILLIER (2003). *Zoophycos* trace fossils commonly occur and point to offshore environments (e.g., LORIN et al 2004). Biostratigraphy is based on the discovery, at the base of the Calcaires de Latrecey, of two ammonites (*Epipeltoceras semimammatum*) indicating the beginning of the Bimammatum zone (i.e., Semimammatum subzone; COURVILLE & VILLIER 2003).

**Sedimentological interpretation**

This section presents homogeneous facies (always mudstones) with very poor fauna (i.e., benthic foraminifera, bivalves, echinoderms and serpulids; Tab. 2.2). Bioturbation is commonly present, indicated by a difference of colors. At the base of the section (at 0.1m), a clay-rich level, surrounded by mudstone beds, contains coal debris. This level represents the shallowest facies of this section. From this level to 2.55 m, relatively thick limestone beds predominate. The deposition of a thin marly level at 2.55 m points to a change in the system. The following succession is formed by thinner limestone beds with more frequent clay levels. This facies change towards more clay-rich deposits indicates a generally higher terrigenous input.
<figure><img src="image_48.png" /><figcaption>Fig. 4.19 - Latrecey-sur-Ormoy section (1).</figcaption></figure>

## Sequence-stratigraphic interpretation

Because of low facies variations, the sequence-stratigraphic interpretation of the Latrecey section is mainly based on the bedding pattern, the bioturbation, and some sedimentary structures. The clay-rich interval at the base of the section is interpreted as the medium-scale SB Ox6+. Small-scale sequence 1 includes five elementary sequences. Each one is composed of two beds. The maximum-flooding surface of elementary sequence 1.3 is placed at the base of the marly interval (at 0.98 m). This surface is interpreted as the maximum flooding of the small-scale sequence 1 (MF1).

The marly interval at 2.55 m is considered to represent the small-scale SB 2. In small-scale sequence 2 only three elementary were investigated. The first two elementary sequences consist of one bed. The third elementary sequence is composed of a limestone bed and a more marly interval. The maximum-flooding surface of this elementary sequence is assumed to be located at the top of the limestone bed (at 3.15 m) and is interpreted to correspond to medium-scale MFOx6+.

## 4.2.2 Swabian Jura

**Balingen-Tieringen** (Figs 4.20 and 4.21a-c; Pl. 17/1-2)
<figure><img src="image_49.png" /><figcaption>Fig. 4.20 - Location of the Balingen-Tieringen section. Based on topographic maps of Balingen (7719), TK 25.</figcaption></figure>

## Geography and stratigraphy

The Balingen-Tieringen section is located in the western Swabian Jura (southern Germany; Fig. 1.2c and 4.20). The section exposes deep-water carbonate facies, characterized by marl-limestone alternations and siliceous sponge facies, probably deposited in 100-200 m water depth (OLIVIER et al. 2004b). The studied interval concerns 14.5 m of deposits dated of the Bifurcatus and Bimammatum zones (Fig. 4.21).

Sedimentology, nannofossil assemblages, and stable isotopes of this section have already been studied by PITTET & STRASSER (1998a), PITTET et al. (2000), PITTET & MATTIOLI 2002, and BARTOLINI et al. (2003). Similar sections in the Swabian Jura have also been investigated by, e.g., OLIVIER et al. (2004b).

## Sedimentological interpretation

The Balingen-Tieringen section is marl-dominated at the base (from base to 6.4 m) and limestone-dominated in the top (from 6.4 m to top). The base of the section is characterized by unfossiliferous grey to brown marls and relatively thin limestone beds. These beds are bioclastic mudstones from the base to 4.25 m and dominated by tuberoid wackestones-packstones, commonly interpreted as deposits close to sponge reefs, and thin sponge boundstones from 4.25 to 6.4 m. This marl-dominated interval points to a low carbonate input from the adjacent platform and/or to high clay input (PITTET & STRASSER 1998a; PITTET et al. 2000).

A major facies change (at 6.4 m) from marl-dominated facies to more massive limestone beds points to an increase of carbonate production and export and/or to a reduced clay input. The first two thin limestone beds (from 6.4 to 6.85 m) are tuberoid and bioclast packstone to wackestones. The two following thick and massive beds (from 6.85 to 8.6 m) are sponge boundstones with mudstone matrix. Another level of tuberoid packstone is observed between these two beds. The following bed (from 8.6 to 9.8 m) has a nodular aspect and a finely laminated bed surface. It consists of wackestones with tuberoids and sponge fragments. This bed is followed by a relatively thin tuberoid-rich wackestone to packstone, nodular in the top and a thin clay-rich level (10.1-10.15 m). PITTET & STRASSER (1998a) suggest that intervals rich in tuberoids may correspond to low sedimentation rates.

The following succession consists of tuberoid and bioclastic packstones to wackestones interbedded with thin marly limestones (from 10.15 to 11.7 m). The fifth limestone bed (from 11.45 to 11.7 m) is bioturbated, richer in bioclasts (e.g., ostracodes, Protoglobigerinids, and brachiopods) and has a wavy bed surface. A thin clay level (at 11.7 m), containing benthic foraminifera and sponge fragments, covers this bed. The following succession mainly consists of tuberoid and peloid facies that pass laterally into sponge bioerms. From 13.95 to 14.15 m, thin limestone beds rich in tuberoids and marls imply a condensed interval with a low accumulation rate. Finally, massive limestones re-appear at the top of the section and suggests an increase of carbonate input.

## Sequence-stratigraphic interpretation

Based on facies, stacking pattern and bed surfaces, a sequence-stratigraphic interpretation is proposed between SB Ox6+ and SB Ox7. It is assumed that sea-level variations play an important role in the exportation of carbonate mud from platform to deeper environments (e.g., SCHLAGER et al. 1994; cf. Chap. 5). It has been demonstrated that carbonate mud is preferentially exported from the platform during transgression when carbonate productivity is high on the platform, and during sea-level fall when progradation of the platform is forced (e.g., SCHLAGER et al. 1994; PITTET & STRASSER 1998a). Marly layers correspond to flooding of the platform or to emersion when only a small amount of carbonate mud is exported to the basin. Furthermore, pelagic carbonate
<figure><img src="image_50.png" /><figcaption>Fig. 4.21a - Balingen-Tieringen section (1).</figcaption></figure>
<figure><img src="image_51.png" /><figcaption>Fig. 4.21b - Balingen-Tieringen section (2).</figcaption></figure>
<figure><img src="image_52.png" /><figcaption>Fig. 4.21c - Balingen-Tieringen section (3).</figcaption></figure>

productivity and input of clay were certainly also modulated by climate changes.

The marl-dominated interval at the base of the section (from 0 to 6.4 m) suggests a relatively low sea level during which carbonate productivity was low on the platform. The clay-rich level at the base of the section (from 6.2 to 6.4 m) is assumed to represent the medium-scale SB O6+. The marly interval at the base of the massive limestones (at 14.25 m) is interpreted as the medium-scale SB O7. Four small-scale sequences are identified within this medium-scale sequence (between O6+ and O7).

Small-scale sequence 1 consists of five elementary sequences. The first elementary sequence is composed of two tuberoid-rich beds (packstone and wackestone). The maximum flooding is placed in the marly interval between these two beds (at 6.55 m). The second and third elementary sequence each consists of one thick and massive sponge boundstone and one reworked sponge and tuberoid packstone-wackestone. The maximum flooding of these sequences is placed on top of the boundstones. The third elementary sequence (from

7.8 to 9.75 m) is the thickest sequence of this small-scale sequence and includes the maximum flooding of small-scale sequence 1. This probably indicates good conditions for sponge growth: low to moderate trophic level, low background sedimentation rate (e.g., OLIVIER et al. 2004b). The fourth elementary sequence is thinner and composed of two tuberoid wackestone-packstone beds separated by a clay level rich in benthic foraminifera, echinoderms, and glauconite. The top of the first bed is nodular and has a slight wavy bed surface interpreted as the maximum-flooding surface. The fifth elementary sequence is composed of a massive tuberoid wackestone-packstone bed and a marly limestone bed rich in benthic foraminifera (Miliolids) and reworked sponges. This level (from 10.75 to 10.9 m) suggests the exportation of coarse elements from the platform and thus represents a good candidate for the small-scale SB 2.

Small-scale sequence 2 consists of five elementary sequences. The first and second elementary sequence are relatively similar and composed of one massive tuberoid wackestone-packstone and one thin marly limestone. The third elementary sequence is composed
<figure><img src="image_53.png" /><figcaption>**Fig. 4.22** - Location of the Vergons section. Based on topographic map of Castellane (3442OT), IGN Topo 25 series.</figcaption></figure>

of two limestone beds separated by a clay-rich level. The first limestone bed presents bioturbation and more bioclastic facies (e.g., brachiopods, protoglobigerinids). These observations point to a maximum-flooding surface. The fourth and fifth elementary sequence are each composed of one tuberoid and peloid packstone.

Small-scale sequence 3 is thinner than sequences 1 and 2. This sequence is composed of three limestone beds relatively rich in tuberoids that pass laterally into sponge reefs.

Small-scale sequence 4 is the thinnest (around 40 cm) compared to the three other small-scale sequences and only consists of four limestone beds. This sequence corresponds to a condensed interval. By lateral correlation, the top of this small-scale sequence is interpreted as the medium-scale SB Ox7.

The small-scale sequences above the medium-scale MF Ox 6+ are thinner than those before. This difference is probably due to the highstand conditions of the large-scale trend.

## 4.3 BASIN SECTIONS

### 4.3.1 South-East France

Vergons (Figs 4.22 and 4.23a-c; Pl. 17/3-6)

**Geography and stratigraphy**

The Vergons section is located in the Alpes de Haute-Provence (France) along the road N202

between St. Julien du Verdon and Vergons (Fig. 4.22). The whole section comprises about 115 metres of basin deposits and covers the Middle Oxfordian to Early Tithonian time interval. Sedimentology and strato-nomy have already been studied by KILLIAN (1889), BEAUDOIN (1977), PELLATON & ULLRICH (1997), the geochemistry by DE RAFÉLIS SAINT-SAUVEUR (1996), PELLATON & ULLRICH (1997), the sedimentary organic matter by BOMBARDIÈRE (1998), and the clay-mineral-ogy by DECONINCK & DEBRABANT (1985).

The studied section only concerns 16.5 m of deposits, analyzed with great detail (Fig. 4.23). The bed numbers are taken from PELLATON & ULLRICH (1997). The biostratigraphy of this section is based on ammonites and partly on dinocysts (ATROPS 1982, 1984). Ammonite specimens attributed to the Bifurcatus zone are present between beds 33 and 46 and those of the Bimammatum zone are found in beds 69 and 74 (determination by F. Atrops; PELLATON & ULLRICH 1997). The section corresponds to the topmost part of the Middle to Late Jurassic "Terres Noires" facies, well known in SE France. The progressive transition from the Oxfordian Terres Noires facies to marl-limestone alternations is well exposed in the abrupt cliff on the other side of the Riou River.

**Sedimentological interpretation**

This section is characterized by pelagic to hemi-pelagic facies with marl-limestone alternations and, locally, gravity deposits (calcarenites, breccias, and slumps). Such deposits are mainly controlled by changes of planktonic productivity, carbonate mud export from the platform, and clay input, as well as by gravity flows and slope instabilities (cf. Chaps 3 and 5).

The base of the section is marl-dominated (from base to bed 52) while the top part is limestone-dominated (from bed 53 to top). In the marl-dominated part, relatively thick limestone beds occur (i.e., beds 43, 45, 47, 49; Tab. 2.3) and mark changes in sediment supply and carbonate productivity. Then, in the mid-part of the section, gravity flow deposits appear (i.e., beds 48, 50, 53) that are interbedded with marl-limestone alternations. These beds are peloidal and bioclastic packstones with some radial ooids and are interpreted as turbidites reworked by contour currents. Beds 47 and 48 show synsedimentary deformations (Plate 17/4) interpreted as slumps. Consequently, the interval from 5.5 to 7 m indicates a phase of instability of deposits.
<figure><img src="image_54.png" /><figcaption>Fig. 4.23a - Vergons section (1).</figcaption></figure>
<figure><img src="image_55.png" /><figcaption>Fig. 4.23b - Vergons section (2).</figcaption></figure>
<figure><img src="image_56.png" /><figcaption>Fig. 4.23c - Vergons section (3).</figcaption></figure>

Bed 53 is a lithoclastic-bioclastic packstone with breccia facies, interpreted as a debris flow (microfacies b6, Tab. 2.3). Bed 54 to 65 are relatively thin (10-30 cm) with more and more radiolarians, protoglobigerinids, and ammonites. The top part of the section (from bed 66) is dominated by thick limestone beds, rich in radiolarians, Protoglobigerinids, and locally allochthonous particles (bed 71). Higher up, slumps appear.

## Sequence-stratigraphic interpretation

Based on the hierarchical stacking pattern, bed surfaces, and facies, four medium-scale sequence boundaries have been identified. The first one is placed below the first thick radiolarian limestone (bed 43) in the Bifurcatus zone. According to this biostratigraphic position, it is assumed to correspond to SB Ox6 of HARDENBOL et al. (1998). The medium-scale SB Ox6+ is placed below the thick lithoclastic limestone (bed 53) between the Bifurcatus and Bimammatum zones. This interval approximately coincides with the relatively lowest manganese concentrations, which is generally associated to a low sea level (EMMANUEL & RENARD 1993). The medium-scale SB Ox7 is placed below the interval of amalgamated limestone beds at the base of bed 66, below the first ammonite indicating the Bimammatum zone. SB Ox8 is placed below slumped beds (bed 75) in the Bimammatum zone.

Medium-scale sequence 0, between Ox6 and Ox6+, consists of four small-scale sequences. The first two small-scale sequences display marl-limestone alternations. The two following sequences contain gravity flow deposits in the first part of each sequence. These deposits are interpreted as lowstand deposits because they correspond to periods of slope instability commonly occurring during low sea level (STROHMENGER & STRASSER 1993).

Medium-scale sequence 1, between Ox6+ and Ox7, consists of four small-scale sequences. The first small-scale sequence consists of five marl-limestone couplets. The first bed is a breccia and thus indicates a period of instability during a low sea level. The maximum flooding of this small-scale is placed in the thinnest marly level of this sequence (at 8.35 m), probably suggesting a condensed interval. The three following small-scale sequences each consist of five beds or marl-limestone couplets. For each small-scale sequence, the maximum flooding is placed in a marly level. The maximum flooding of medium-scale sequence 1 is placed around 9.7-9.8 m in the third elementary sequence of the third small-scale sequence (bed 60), where radiolarians and then Protoglobigerinids are abundant. This interval also coincides with the highest manganese concentrations generally associated with a high sea level (EMMANUEL & RENARD 1993).
Medium-scale sequence 2, between Ox7 and Ox8, also consists of four small-scale sequences. Sequence boundaries are placed at the base of thick massive limestones. The identification of deposits is more

difficult than in the previous medium-scale sequences mainly because of a different stacking pattern with thick massive limestones but rare marly intervals.

* * *
# 5 - STRATIGRAPHIC CORRELATIONS

For the high-resolution stratigraphic correlation of all studied sections different tools are used: ammonite biostratigraphy, lithostratigraphy, and sequence stratigraphy. While biostratigraphy and major lithological changes furnish a general framework, the highest resolution is obtained by sequence-stratigraphical correlations on the scale of elementary sequences. In one section alone, it is often difficult to place these boundaries (i.e., sequence boundaries, transgressive surfaces, maximum-flooding surfaces) unequivocally and to estimate their magnitude (i.e., if they belong to elementary, small-scale, or medium-scale sequences). Through correlation with other sections over short and long distances, however, a "best-fit solution" can be found.

This study provides a high-resolution sequence-stratigraphical correlation from shallow-water platform (Jura, Lorraine) to deeper platform (Haute-Marne, Swabian Jura) and to basin sections (SE France) during the early Bimammatum zone (Late Oxfordian). One of the targets of this study is to have a closer look at the transgression that led to the accumulation of the Hauptmumienbank and Steinebach members (Swiss Jura) with regard to its evolution in time and space (cf. Chap. 7).

No palinspastic reconstructions have been undertaken in this study. The indicated distances between the sections are measured on today's geographical maps. The distances would have to be stretched by a few kilometers between the sections along a general N-S axis. Moreover, some lateral offsets by a few kilometers between sections cannot be excluded. However, the relative proximal-distal position of the sections is not affected.

## 5.1 METHODS OF CORRELATION

### 5.1.1 Biostratigraphy

The biostratigraphic framework of the platform and basin sections was established using ammonites (ENAY et al. 1988 and GYGI 1995 for the Swiss Jura; CARPENTIER 2004 for the Lorraine; COURVILLE & VILLIER 2003 for the Haute-Marne; SCHWEIGERT 1995a, b for the Swabian Jura, and ATROPS 1982, 1984 for SE France). However, due to infrequent data on most shallow-water carbonate platforms, alternative correlation methods are demanded for the correlation of platform sections.

### 5.1.2 Lithofacies

Major lithostratigraphic changes occur regularly on the Jura platform (e.g., GYGI 1995). The base of the massive Hauptmumienbank and Steinebach beds has thus been used for regional correlations (GYGI 1995; PITTET 1996; HUG 2003). Furthermore, GYGI & PERSOZ (1986) have performed mineralostratigraphic correlations (based on quartz and clay minerals) between the biostratigraphically better dated epicontinental basins and the less well dated Jura platform.

However, lithological boundaries are not necessarily isochronous. Small-scale diachronisms and lithological heterogeneities deriving from spatial and temporal variations of facies distribution and/or lateral migration of facies belts certainly occur but are difficult to evaluate without applying facies-independent methods of correlation.
### 5.1.3 Discontinuity surfaces and depositional sequences

In the wake of sequence stratigraphy, great attention has been paid to the chronostratigraphic value of discontinuities. High-resolution correlations over long distances are usually based on diagnostic surfaces and/or intervals, which are assumed to be more or less isochronous. However, especially in marginal-marine and slope environments sequence boundaries include hiatuses caused by erosion and/ or non-deposition (Fig. 5.1a). Transgressive surfaces are generally not isochronous if there is a pre-existing morphological relief (Fig. 5.1b). Maximum-flooding surfaces or intervals have the best chance to homogenize a morphological relief in shallow-marine settings and, consequently, to be isochronous (Fig. 5.1c). Accordingly, they are considered to be the most appropriate for high-resolution correlation over long distances (e.g., SANDULLI & RASPINI 2004; TRESCH 2007).

In platform sections, the stacking pattern of depositional sequences can be used to identify trends in accommodation change. However, the comparison of thickness variations of laterally corresponding depositional sequences can only be made as a first approximation. In order to translate the thickness of a sequence into accommodation and eventually sea-level change, the sequence would first have to be decompacted according to facies and water depth for the deposition of each facies would have to be estimated (STRASSER et al. 2004; HILLGÄRTNER & STRASSER 2003). This very time-consuming procedure has not been performed in the present study.

The investigated sedimentary record lies between two well-marked discontinuities, which are attributed to the 3rd-order (large-scale) sequence boundaries Ox6 and Ox7 according to HARDENBOL et al. (1998). These sequence boundaries have been dated and define the general chronostratigraphic time framework. The cyclostratigraphical analysis can further improve the time resolution (cf. Chap. 6). Two medium-scale sequences, identified between these major sequence boundaries, create a correlation framework, which has already been established in previous studies for the Swiss Jura platform (PITTET 1996; HUG 2003). The first half of the second medium-scale sequence at the base of the long-term transgressive trend has been chosen to investigate small-scale and elementary depositional sequences in more detail.

A

<figure><img src="image_57.png" /></figure>

Time interval of the sequence boundary is isochronous, but not the physical surfaces.

B

<figure><img src="image_58.png" /></figure>

Transgressive surfaces are not isochronous if there is a pre-existent morphological relief (arrows symbolize TS)

C

<figure><img src="image_59.png" /></figure>

Maximum-flooding intervals or surfaces have the best chance to homogenize a morphological relief and therefore to be isochronous at least at a regional scale.

**Fig. 5.1** - Models illustrating diagnostic intervals or surfaces in time and space. a) Sequence boundaries; b) Transgressive surfaces; c) Maximum floodings. Modified from TRESCH (2007).

High-resolution sequence-stratigraphic correlations are presented for each sedimentary domain (i.e., shallow platform, deep platform, basin) and then are discussed together.

## 5.2 CORRELATION ON THE SHALLOW PLATFORM

The base of the Hauptmumienbank and Steinebach members for the Swiss Jura sections and the base of the Oolithe de Saucourt inférieure Member for the Lorraine section are used as regional markers. The biostratigraphic framework of the platform sections is based on ammonites (GYGI 1995; CARPENTIER 2004).
Based on these sequence-stratigraphic elements presented in Chapter 4, a bed-by-bed correlation of the six Swiss Jura sections and the Lorraine section has been attempted providing the most consistent *best-fit correlation* (Figs 5.2 and 5.3).

The Swiss Jura sections are correlated between the medium-scale SB Ox6+ and the medium-scale MF Ox6+. The sections are distant of a few kilometres to tens of kilometres. They are positioned from “proximal” (Voyeboeuf section) to “distal” (Savagnières and Pertuis sections) based on the dominant facies and on their palaeogeographical positions. However, because platform morphology has changed through time, these attributions can only indicate the general trend.

The best-fit high-resolution correlation of the Swiss Jura sections shows that sequence boundaries and maximum floodings of the small-scale and elementary sequences have a good lateral continuity (Fig. 5.2). This suggest that the formation of the small-scale sequences was controlled by an allocyclic mechanism.

Elementary sequences can be produced by allo-
and autocyclic processes. Most small-scale sequences
include five elementary sequences suggesting a
control on the sedimentation by a common external
factor such as orbitally controlled sea-level changes.
However, Fig. 5.4 shows the intervals consisting
of migrating high-energy ooid bars, controlled by
autocyclic processes (cf. Chap. 3). Nevertheless,
the good correlation of many elementary sequences
suggests an allocyclic control also for these sequences,
even if it is occasionally overprinted by autocyclicity.

When comparing the facies of time-equivalent
small-scale and elementary sequences between the
Swiss Jura sections, significant lateral facies changes
appear. For example, ooid shoals are laterally replaced
by oncoid lagoons, or tidal flats are juxtaposed to
marly depressions (Fig. 5.2). Ooid shoals do not
always occur at the same time in the studied sections,
which suggests that they migrated across the platform
through time or were stranded in one position and
later re-initiated in another one. This points to a
facies mosaic on the Swiss Jura platform. It reflects
the dynamics and complexity of sedimentary systems
where juxtaposed sub-environments evolved and
shifted through space and time (SAMANKASSOU et al.
2003; WRIGHT & BURGESS 2005; STRASSER & VÉDRINE
in press; cf. Chap. 7).

The presence of thickness variations at the scale of small-scale and elementary sequences reveals variable sediment accumulation rates. This is interpreted as resulting mainly from differential subsidence due to the activity of tectonic blocks (ALLENBACH 2002). The relief created by tectonics therefore contributed significantly to the general facies distribution. Furthermore, the irregular distribution of siliciclastics can be explained by localized depressions, which were created by differential subsidence and served as conduits (PITTET 1996; HUG 2003). In addition, carbonate production and accumulation, i.e. the preferential growth of reefs and the initiation of ooid shoals on tectonic highs, further enhanced the relief.

The Swiss Jura sections are then correlated with
an additional shallow platform section from Lorraine.
Two Swiss Jura sections, representative of proximal
and distal facies (Voyeboeuf and Pertuis, respectively),
have been selected for the correlation with the Pagny-
sur-Meuse section (Lorraine). This section is less
thick than the Swiss Jura sections, probably due to
a lower subsidence rate. The correlation of sequence
boundaries and maximum floodings of small-scale
and elementary sequences between the seven shallow
platform sections is relatively straightforward (Fig.
5.3). A similar number of elementary sequences is
counted in all sections and thus implies an external
controlling factor. In all studied shallow platform
sections, the distribution of siliciclastics shows a
similar pattern with more or less developed levels at
the medium-scale SB Ox6+, and MF Ox6+, and at the
small-scale SB 2.

5.3 CORRELATION FROM DEEP
PLATFORM TO BASIN

Based on a well-established biostratigraphic
framework, the interpreted sequence stratigraphic
boundaries are correlated at the scale of elementary
and small-scale sequences from the deep platform
sections to the basin section (Figs 5.5 and 5.6). The
larger time framework of the Balingen-Tieringen and
Vergons sections (between Ox6 and Ox8) allows to
better constrain the studied time interval.

**Fig. 5.2** - (following pages) High-resolution best-fit correlation of the one and a half small-scale sequence and elementary sequences of the six sections of the Swiss Jura platform.
SHALLOW PLATFORM

Swiss Jura

<figure><img src="image_60.png" /></figure>

Facies

<figure><img src="image_61.png" /></figure>

oolitic

<figure><img src="image_62.png" /></figure>

oncolitic

<figure><img src="image_63.png" /></figure>

bioclastic

<figure><img src="image_64.png" /></figure>

coral

<figure><img src="image_65.png" /></figure>

peloid

<figure><img src="image_66.png" /></figure>

siliciclastics
SHALLOW PLATFORM

Swiss Jura

<figure><img src="image_67.png" /></figure>
SHALLOW PLATFORM

<figure><img src="image_68.png" /><figcaption>**Fig. 5.3** - High-resolution best-fit correlation of the platform sections: two selected sections of the shallow-water Swiss Jura platform and the section of the Lorraine platform.</figcaption></figure>
<figure><img src="image_69.png" /><figcaption>Fig. 5.4 - Intervals with autocyclic deposits where the sequence-stratigraphic interpretation is uncertain in the platform sections.</figcaption></figure>
DEEP PLATFORM

Haute-Marne

Swabian Jura

BASIN

SE France

<figure><img src="image_70.png" /></figure>

a 3rd-order sequence boundaries of HARDENBOL et al. (1998)

b Medium-scale sequences (400 kyr)

c Small-scale sequences (100 kyr)

d Elementary sequences (20 kyr)

e Ammonite biostratigraphy
<figure><img src="image_71.png" /></figure>

d: Elementary sequences (20 kyr)
5E: five elementary sequences

**Fig. 5.6** - Detail of the Ox6+ medium-scale sequence with small-scale and elementary sequences identified in the Vergons section.

The first small-scale sequence shows important sediment thickness variations. The highest sediment thickness is observed in the Balingen-Tieringen section where sponge reefs develop (cf. Figs 5.5 and 4.21). The lowest sediment thickness appears in the basin section. However, the sections are not decompacted and as the Latrecey section is entirely composed of mudstones, its original thicknesses was considerably larger. If a decompaction factor of 2.5 is applied for deposits of the first small-scale sequence at Latrecey (STRASSER & SAMANKASSOU 2003), the original thickness is higher than the one of the decompacted Balingen-Tieringen section, containing wackstones, packstones, and boundstones. At Vergons, the relatively thick first small-scale sequences owes this to the debris flow at its base. Elsewhere in this section, the limestone-marl alternation had low accumulation rates. This comparison shows the different sediment

**Fig. 5.5** - (facing page) High-resolution best-fit correlation of the sections from the deep platform to the basin.

accumulation rates from deep platform to basin. Due to the disappearance of sponge boundstones and the deposition of marl-limestone couplets in the first half of the second small-scale sequence at Balingen-Tieringen, sediment thicknesses are generally lower there.

As for the shallow platform sections, five elementary sequences have been identified in the first small-scale sequence of all three sections (cf. Chap. 4; Fig. 5.6). Elementary sequence 1.3, where the maximum flooding of the small-scale sequence 1 (MF1) is inferred, is thickest in the Balingen-Tieringen section and thinnest in the Vergons section. A similar pattern is observed for the next small-scale maximum flooding, which is equivalent to the medium-scale maximum flooding (MF Ox6+) recognized on the shallow platform. However, facies analysis of the Vergons section suggests that the best expressed MF is the one of the third small-scale sequence (Fig. 4.23). This observation is supported by the highest manganese value usually interpreted as a maximum flooding. Depending on basin morphology, the best-expressed MF apparently may shift by one or two cycles from one location to another (STRASSER et al. 2000).

## 5.4 PLATFORM-TO-BASIN CORRELATION

The high-resolution sequence-stratigraphical interpretation enables for the first time a best-fit correlation on the scale of elementary sequences between sections from shallow platform to basin in the Tethyan realm during the Late Oxfordian (cf. Fig. 5.7).

The number of elementary and small-scale sequences, identified in the shallow platform sections, is equivalent to the one of the deep platform and basin sections. This suggests a common controlling factor.

## 5.5 DISCUSSION

In the studied interval, siliciclastics (quartz and/or clays) preferentially occur at three time levels in all sections: SB Ox6+, SB 2, and MF Ox6+. On the shallow platform, humidity and/or base-level changes in the hinterland act on the terrigenous run-off and thus on the distribution of siliciclastics. On the Swiss Jura platform, the occurrence of siliciclastics around sequence boundaries are mainly explained by more humid conditions in the hinterland during the relatively low sea level (PITTET & STRASSER 1998a, b). This model explains the presence of siliciclastics around SB Ox6+
<table><thead><tr><th></th><th>DEEP PLATFORM</th><th>SHALLOW PLATFORM</th></tr></thead><tbody><tr><th>Haute-Marne</th><td></td><td>Swiss Jura</td></tr><tr><th>Lorraine</th><td></td><td></td></tr></tbody></table>

<figure><img src="image_72.png" /><figcaption>Fig. 5.7 - High-resolution best-fit correlation of the platform, deep platform and basin sections for the studied interval.</figcaption></figure>
and SB 2. During the maximum floodings, relative sea-level rise is fastest and the erosion potential is reduced. Therefore, siliciclastics probably come from the remobilization of siliciclastics formerly stored on the proximal platform and are preferentially deposited in depressions on the shallow platform (Fig. 3.2; STRASSER & HILLGÄRTNER 1998).

### 5.5.1 On the shallow platform

The hierarchical stacking pattern observed in the studied sections implies that sedimentation was influenced by cyclical or periodic changes of controlling factors. The good correlation of depositional sequences over long distances between different sedimentary basins is a valuable argument that processes other than local tectonic activity or migration of sediment bodies must have been involved in the formation of these depositional sequences. It is assumed that high-frequency **sea-level fluctuations** were at least partly responsible for the repetitive stacking of depositional sequences on several scales, as observed in the studied sections.

On shallow-carbonate platforms, the thickness of a sedimentary sequence reflects as a first approximation the changing accommodation space over time. However, it has to be kept in mind that bed thickness and accommodation generally do not have a linear relationship. This is due to facies-dependent differential compaction and because sediment not always fills up the available space (STRASSER et al. 1999). The studied sections displayed here are not decompacted. Therefore, the comparison of thickness changes of laterally corresponding depositional sequences have to be interpreted with caution. Assuming that the sequence boundaries of the depositional sequences are more or less time-equivalent, the lateral changes of accommodation across the regions of investigation can be compared. The causes of these changes result from the combination of eustatic sea-level changes, tectonic movements (i.e., differential subsidence, block faulting), and the production and accumulation rates of sediment.

Additional factors have to be considered in the formation of depositional sequences on a shallow platform. The **position on the platform** is the first important parameter. Generally, restricted and/or protected facies are common in proximal positions while open-marine facies predominate in distal position (cf. Chap. 2). A **pre-existing morphology**

determines the geometry of the coast line and the distribution of depocentres on the platform. The pre-existing relief also modifies the hydrodynamic setting (current and/or wave directions and intensity), the pattern of sediment production and deposition, and ecological conditions (e.g., water depth, salinity, turbidity, nutrient distribution).

**Climate changes** also play a role on the distribution of carbonate facies on the shallow platform. Terrigenous input, controlled by humidity changes in the hinterland, mainly influences the carbonate production and accumulation (see below).

### 5.5.2 On the deep platform and in the basin

In basinal, hemipelagic to pelagic sections, the expression of environmental changes is completely different from that in shallow platform sections. Environmental changes there are indicated by variations in palaeoecology (i.e., macro-, micro-, and nanofossils), different composition of organic matter, and/or changes of mineralogy and geochemistry. These variations can be used for sequence-stratigraphic interpretations (e.g., JAN DU CHÊNE et al. 1993; PITTET & MATTIOLI 2002). Additionally, facies analyses and variations of the stacking pattern in different basinal successions have been related to environmental changes associated with relative sea-level changes (e.g., STROHMENGER & STRASSER 1993; PASQUIER & STRASSER 1997; HILLGÄRTNER 1999; COLOMBIÉ & STRASSER 2003; cf. Chap. 3.4). Different factors control the changes in stacking pattern in deep-marine depositional systems. The main factors are winnowing of fine-grained sediments by bottom currents, intensified export of platform material to more distal, basinal depositional settings (e.g., slumps and turbidites in basin floor fans, slope fans and lowstand prograding wedges, highstand shedding), fluctuations in the calcite compensation depth (CCD) and oxygen-minimum zones, and changes of climate and/or oceanic circulation patterns, which influence the planktonic productivity and the clay input and distribution (e.g., EINSELE & RICKEN 1991; HILLGÄRTNER 1999).

### 5.5.3 Platform-basin relationships: a conceptual model

For the Oxfordian, PITTET & STRASSER (1998a) and PITTET et al. (2000) propose a model for the relationships between the Swiss Jura shallow
<figure><img src="image_73.png" /><figcaption>**Fig. 5.8** - Conceptual model linking trends in relative abundance of carbonate mud and siliciclastics between the platform and the basin. Modified from PASQUIER & STRASSER (1997) and HILLGÄRTNER (1999).</figcaption></figure>
platform and the Swabian deep platform. In this model, carbonate production in shallow platform areas and export of carbonates from the platform to more distal environments are considered as the factors controlling deep platform sedimentation (i.e., marl-limestone couplets/sponge reefs). From nannofossil assemblage composition, PITTET & MATTIOLI (2002) suggest that deep platform deposition (Balingen-Tieringen section) results from cyclical export of carbonate mud from shallow platforms basinwards and, only secondarily, by changes in nannoplankton productivity. To explain the hemipelagic to pelagic Kimmeridgian deposits from the Vocontian basin, COLOMBIÉ & STRASSER (2003) propose the same controlling factors.

Based on the studied sections and their correlation, a conceptual model applicable for small-scale (100-kyr) sequences is proposed (Fig. 5.8). During **low relative sea level**, accommodation potential on the platform is reduced, leading to frequent periods of non-deposition and/or erosion. Base level is low, forcing terrigenous input, which generally coincides with more humid conditions (PITTET & STRASSER 1998a, b). Siliciclastics are mainly concentrated in shallow lagoons on the platform but sediment export, particularly clays, also occurs (cf. Fig. 5.8a). In the basin, slumps or debris flows are common due to slope instability during low sea level.

During the **(early) transgressive** period, benthic productivity on the platform increases significantly (start-up and keep-up phases, KENDALL & SCHLAGER 1981) leading to an increase of carbonate mud export (cf. Fig. 5.8b-c). In addition, clays formerly stored on the shallow platform can be remobilized and preferentially accumulated in deeper environments.

During accelerated sea-level rise and **maximum flooding**, benthic carbonate productivity on the platform decreases locally (give-up phase, KENDALL & SCHLAGER 1981) because water depth increases too rapidly for the system to keep up, and carbonate producers find themselves below the photic zone. Less humid conditions lead to reduced siliciclastic influx to platform areas and to the basin (cf. Fig. 5.8d). However, clays may also be reworked from proximal areas and spread across the platform and into the

basin. At the same time, carbonate accumulation in the basin diminishes to a degree that marly sediments predominate because of decreasing planktonic carbonate production (e.g., PASQUIER & STRASSER 1997) and reduced carbonate export from the platform (PITTET et al. 1998a). The maximum flooding is thus marked by a marly and/or condensed level.

During the **highstand** period, decreasing bathymetry on the platform successively hinders the export of clays to basinal depositional settings but increases highstand shedding of carbonate mud, resulting in thin marl intervals and thick limestone beds (e.g., SCHLAGER et al. 1994; cf. Fig. 5.8e).

### 5.5.4 Elementary sequences : a challenge

In the shallow platform sections, an elementary sequence generally consists of one to four beds including more or less developed marl intervals. In the deep platform sections, an elementary sequence generally consists of one or two limestone beds with a more or less developed marl interval. In the basin section, an elementary sequence is commonly defined by one marl-limestone couplet.

The identification of elementary sequences is based on a very detailed facies and microfacies analysis (cf. Chap. 4). In the shallow and deep platform sections, these sequences are defined based on sometimes very subtle changes of texture and fauna. Sequence boundaries are defined by more restricted or protected facies, while the maximum floodings are characterized by the most open-marine facies. The high-resolution correlation of sections, at the scale of the 20-kyr sequences, shows the great variability of facies within an elementary sequence (cf. Chap. 7), suggesting the influence of intrinsic factors (cf. Chaps 7 and 9). In the basin section, the relatively homogenous facies make the identification of sequence boundaries difficult, transgressive surfaces, or maximum floodings. Consequently, the correlation of elementary sequences from shallow platform to basin is a best-fit solution and may be improved by complementary analyses (e.g., nannofossil assemblages, stable isotopes, trace elements).

* * *
# 6 - CYCLOSTRATIGRAPHY

Cyclostratigraphy is the subdiscipline of stratigraphy that deals with the identification, correlation, and interpretation of cyclic variations in the stratigraphic record. Cycles take their origin in the orbital dynamics of the Earth-Moon-Sun system. The two main families of orbital cycles that are known to be recorded in the sedimentary deposits are the tidal and the Milankovitch cycles (HILGEN et al. 2004). This chapter focuses on the Milankovitch cyclicity records in the studied shallow platform deposits. The goal is to obtain a much more precise chronostratigraphic framework than through biostratigraphy or geochronology and thus to allow for the estimation of rates of ecological and sedimentary processes and to give a high-resolution time scale for the reconstruction of the evolution of the depositional environments.

## 6.1 CONCEPTS AND METHODOLOGY

Orbital forcing is an astronomical theory on the link between variations of the Earth's orbit and climate changes. Quasi-periodic variations of the Earth's orbital parameters modify the distribution and the intensity of insolation and engender climatic changes (MILANKOVITCH 1941; BERGER 1978; SMITH 1989; DE BOER & SMITH 1994). Insolation changes influence air temperature, growth and decay of ice sheets, atmospheric and oceanic circulation, rainfall pattern, and evaporation as well as terrestrial and oceanic ecosystems.

The main astronomical parameters are the changing eccentricity of Earth's orbit around the sun, the Earth's cycle in axial obliquity, and the cycle of precession of the equinoxes (e.g., FISCHER et al. 1990). The controlling orbital parameters (precession, obliquity, eccentricity; Fig. 6.1) are well defined in terms of cycle duration back to Mesozoic times (BERGER

et al. 1989; BERGER & LOUTRE 1994). In the geological past, the periods of precession and obliquity were shorter because of the slowing of the Earth's rotation and the increasing Earth-Moon distance (BERGER et al. 1989). In the Oxfordian, the peaks of the precession signal were approximately at 18 kyr and 22 kyr and at about 37 kyr and 48 kyr for the obliquity (BERGER et al. 1989). Eccentricity stays constant at 100 and 400 kyr.

The formation of high-frequency sedimentary cycles on shallow platforms is mainly explained by eustatic variations. During glacial periods, there were dependent on glaciation/deglaciation cycles of polar ice-caps and mountain glaciers linked to the Earth's orbital parameters. The Oxfordian is interpreted as a "greenhouse" period (HALLAM 1984). No evidence of polar ice-caps is found, but they can not be excluded, and mountain glaciers were probably present (e.g., FRAKES et al. 1992; PRICE 1999, cf. Chap. 8). Sea-level changes of a few metres amplitude can also be created by thermal expansion and retraction of the uppermost layer of the ocean water (GORNITZ et al. 1982), by thermally induced volume changes in deep-water circulation (SCHULZ & SCHÄFER-NETH 1998), and/or by water retention and release in lakes and aquifers (JACOBS & SAHAGIAN 1993).

In the sedimentary record, the 5:1 relationship between elementary and small-scale sequences and the 4:1 relationship between small-scale and medium-scale sequences reflect the hierarchy of Milankovitch cycles. The relatively good biostratigraphic control associated with the absolute dating of GRADSTEIN et al. (1995), it makes thus possible to estimate time comprised in the studied sections (including of course the error bars of radiometric dating). By counting the identified depositional sequences between dated levels (boundaries of ammonite zones or subzones,
<figure><img src="image_74.png" /><figcaption>**Fig. 6.1** - Astronomical parameters that control insolation on Earth (precession, obliquity, and eccentricity). Modified from STRASSER et al. (2007). Schematic insolation curves of the different orbital parameters. Note that the eccentricity cycles act as an amplitude modulator of the precession cycles. The obliquity cycle is not visible in the studied deposits. Based on EINSELE & RICKEN (1991).</figcaption></figure>

<figure><img src="image_75.png" /></figure>

sequence boundaries, or maximum floodings), their order of duration can be inferred (STRASSER et al. 1999) and compared with the known durations of the Milankovitch cycles.

## 6.2 ADVANTAGES AND LIMITS

Cyclostratigraphy allows obtaining a high-resolution stratigraphic framework, superior to that of biostratigraphy (PITTET 1994, 1996; HILLGÄRTNER 1999; HUG 2003; STRASSER et al. 2004; TRESCH 2007). As each sedimentary cycle represents a given duration (approximately 20, 40, 100, or 400 ka), the resulting time scale is very accurate but still floating because the age of the stage boundaries has large error margins ("floating astronomical timescale"; HILGEN et al. 2004; STRASSER et al. 2006).

The following effects or mechanisms may prevent and/or distort the record of orbital forcing in shallow platform environments and, consequently, they have to be taken into account when establishing a cyclostratigraphic time frame (STRASSER et al. 1999):

* Uncertainty about the completeness of the stratigraphic record (non-deposition or non-preservation of depositional sequences).

* In shallow carbonate platform deposits, autocyclic processes may mask or imitate orbitally forced depositional sequences.

* One cycle of environmental change can lead to the formation of multiple beds and, therefore,

no constant time span can be attributed per bed (PITTET & STRASSER 1998b).

* Bed thickness and accommodation do not have a linear relationship because of facies-dependent differential compaction. Observed bed thickness consequently does not correspond to the original accommodation.

* The estimation of accommodation potential is difficult within subtidal facies.

* Loss of record of high-frequency sea-level and/or climatic fluctuations when sedimentological and/or biological thresholds were not passed and thus no facies contrast was created.

* Interference of long-term and short-term sea-level fluctuations leading to enhancement and attenuation of high-frequency signals.

* Imprecise chronostratigraphic data that introduce large error ranges in mean cycle durations.

Despite of all these difficulties, cyclostratigraphy is the tool of choice for establishing a high-resolution time frame.

## 6.3 CYCLOSTRATIGRAPHIC TIME FRAME

A cyclostratigraphic time frame is proposed for the studied interval (Fig. 6.2). Based on the radiometric ages given by GRADSTEIN et al. (1995) for the Oxfordian and Kimmeridgian stage boundaries, HARDENBOL et al. (1998) date the 3rd-order SB Ox6 at 155.81 Ma and the 3rd-order SB Ox7 at 155.15 Ma by interpolation between the stage boundaries (Figs 1.5 and 6.2).
This implies a time span of 600 to 700 kyr between these two sequence boundaries. Between the 3rd-order sequence boundaries Ox6 and Ox7, two medium-scale sequences, each containing four small-scale sequences, were identified in the shallow platform and in the basin sections (Fig. 6.2). Consequently, it is suggested that a medium-scale sequence formed within 300 to 350 kyr and a small-scale sequence within 75 to 87.5 kyr. These values coincide well with those of the short and long eccentricity cycles (100 and 400 kyr, respectively; BERGER et al. 1989; BERGER & LOUTRE 1994). In the studied sections it is commonly observed that five elementary sequences group into a small-scale sequence. Consequently, the observed depositional sequences are the sedimentological expression of orbital cycles and are attributed to the precession of the equinoxes (20 kyr) for the elementary sequences, and to the short- and long-eccentricity cycles (100 and 400 kyr) for the small- and medium-scale sequences,

respectively. Despite the attenuated effects of the obliquity cycle (40 kyr) on the insolation in low latitudes (BERGER 1978), an influence of this cycle is possible, particularly during greenhouse periods (READ 1995). This cyclostratigraphic time frame is comparable to the one of STRASSER in press (Fig. 6.2) and now helps to improve the timing between the 3rd-order sequence boundaries Ox6 and Ox7.

In the Late Oxfordian deposits of the Jura platform, 100-kyr sequences are quite easily recognized whereas 20-kyr sequences may be locally more difficult to identify (PITTET 1996; HUG 2003). However, this study has confidently identified elementary sequences with the help of detailed facies and microfacies analyses. The obliquity cycle seems not to be expressed in the sedimentary record of the studied sections. However, from the study of more marly intervals, for example in the top part of the Röschenz Member, PITTET

<figure><img src="image_76.png" /><figcaption>**Fig. 6.2** - Cyclostratigraphic time framework proposed for the studied interval. Bio- and sequence-chronostratigraphy are from HARDENBOL et al. (1998).</figcaption></figure>
(1996) related obliquity to humidity cycles acting on carbonate production and terrigenous input and not to cycles of sea-level changes.

## 6.4 COMPARISON WITH OTHER STUDIES

Only a few studies exist on the record of Milankovitch cycles in Late Oxfordian sedimentary successions. PITTET (1996) investigated, for the first time, the orbital cycles in the deposits of the Röschenz Member and Hauptmumienbank/Steinebach members (Vellerat Formation, Swiss Jura). This author observed a hierarchy of four orders of sequences related to orbital cycles (precession, obliquity, and short and long

eccentricity). HUG (2003) studied the top part of the Vellerat Formation and the Courgenay Formation and established a precise sequence- and cyclostratigraphic framework with a 100-kyr resolution used as a starting point of this study (Fig. 6.2).

STRASSER et al. (2005) performed a sequence- and cyclostratigraphic analysis from platform to basin on sections in Spain. They identified eight small-scale sequences within the 3rd-order sequence between SB Ox6 and Ox7, leading to the same duration (800 kyr) as proposed by STRASSER (2007) for the Swiss Jura. However, a comparison at the scale of elementary sequences is not proposed in any of these studies.

* * *
# 7 - HIGH-FREQUENCY PALAEOENVIRONMENTAL CHANGES ON THE SWISS JURA PLATFORM

This chapter examines the high-frequency palaeoenvironmental changes that occurred on the Swiss Jura platform in the narrow time framework established from facies analysis coupled with sequence- and cyclostratigraphic analysis (cf. Chaps 5 and 6). The stratigraphic and lateral distribution of facies, oncoids, and benthic foraminifera is investigated within an interval of 180 kyr at the scale of elementary sequences (20 kyr), which allows to reconstruct the platform evolution through time and to interpret the palaeoecological conditions.

## 7.1 PLATFORM EVOLUTION THROUGH TIME

The shallowest facies appears at the base of all sections (Fig. 7.1a). At Vorbourg, Hautes-Roches, Court, and Pertuis, they consist of marly deposits, characterizing protected lagoonal sedimentation. In Voyeboeuf and Savagnières, the shallowest facies are made of bioclastic (mainly ostracodes and inocerams) mudstones-wackestones. These two facies types mark a major sequence boundary: **the medium-scale SB Ox6+** (see chapter 5 for the definition of this sequence boundary). Then, a rapid facies change, easily recognizable in most of the Swiss Jura deposits, occurs from marly deposits to massive limestone beds (cf. Chap. 5; Fig. 7.1a). These limestones correspond to higher energy facies and/or contain more open-marine faunal assemblages, which overlie the transgressive surface TS Ox6+. Elementary sequence 1.3 is the thickest one of the first small-scale sequence (Fig. 7.1b) suggesting maximum accommodation gain. The maximum flooding of this elementary sequence coincides with the maximum flooding of small-scale sequence 1. At the end of the first small-scale sequence, a shallowing and possibly also a climate change to

more humid conditions are indicated by a minor siliciclastic input (Fig. 7.1c) marked by relatively thin clay layers (Voyeboeuf, Hautes-Roches, and Pertuis sections), by oncoid rudstones (Vorbourg section), by bioclast mudstones with birdseyes (Court section), and by a serpulid-rich layer (Savagnières section). This interval with thin beds and more or less rich in clays is interpreted as the **small-scale SB 2**. At the top of the carbonate-dominated part of the sections, a wavy to planar surface is observed (Fig. 7.1d), which is overlain by more or less siliciclastics. In the “proximal” sections (Voyeboeuf, Vorbourg, Hautes-Roches, and Court) this wavy surface consists of a hardground, indicating sediment starvation. In the most “distal” sections (Savagnières and Pertuis), it consists of a relatively flat surface. This level belongs to the **medium-scale MF Ox6+** (for definition see chapter 5). The overlying siliciclastics, more abundant in the ‘proximal’ part of the platform, possibly indicate more humid conditions in the hinterland. However, during transgression clays may also have been remobilized and transported by currents over kilometres and preferentially deposited in depressions (PITTET 1996; STRASSER et al. 1999).

Assuming that the observed depositional sequences formed through superimposed frequencies of sea-level fluctuations controlled by orbital cycles, a precise time frame can be constructed (Fig. 7.2). Dominant facies and environments within each elementary sequence are represented in a time-space diagram (Fig. 7.3). **Siliciclastics** (clays and quartz) occur abundantly at the base (sequence 0.5 and around SB Ox6+) of the Vorbourg, Hautes-Roches, and Court sections. Then siliciclastics disappear but reappear around SB 2, especially in the Voyeboeuf and Hautes-Roches sections. Siliciclastics are flushed onto the platform by terrigenous input controlled by increased rainfall in the hinterland and/or by sea-level drop. The source
<figure><img src="image_77.png" /></figure>

SB Ox6+ and the beginning of the transgression

<figure><img src="image_78.png" /></figure>

<figure><img src="image_79.png" /></figure>

MF 1 and elementary sequence 1.3

<figure><img src="image_80.png" /><figcaption>Fig. 7.1 - Focus on major sequential limits and their concomitant elementary sequences. (a) SB Ox6+ and elementary sequence 1.1, (b) MF 1 and elementary sequence 1.3.</figcaption></figure>
SB 2 and the beginning of the transgression

<figure><img src="image_81.png" /></figure>

MF Ox 6+ and the elementary sequence 2.3

<figure><img src="image_82.png" /><figcaption>**Fig. 7.1 (continued) - Focus on major sequential limits and their concomitant elementary sequences. (c) SB 2 and the beginning of the transgression, (d) MF Ox 6+ and elementary sequence 2.3.**</figcaption></figure>
<figure><img src="image_83.png" /><figcaption>**Fig. 7.2** - High-resolution time framework used in this study related to superimposed sea-level fluctuations.</figcaption></figure>

area was located in the London-Brabant and/or Bohemian massifs, 400 km to the North (Fig. 1.3). Between the emerged lands and the shallow-carbonate platform, a fluvial plain with an important potential of siliciclastic capture and storage has probably existed. However, the climatic conditions on the platform were not necessarily the same as in the hinterland (PITTET 1996; DUPRAZ 1999; HUG 2003). For example, siliciclastics could have been mobilized when it rained in the hinterland and flushed onto the platform that experienced a semi-arid climate.

Oncoid growth is mainly controlled by the water energy, trophic level, and sediment accumulation rate (see below). **Oncoid-rich facies** appear more abundantly in the first small-scale sequence than in the second one. The Vorbourg and Pertuis sections are

essentially composed of oncolitic facies in the first small-scale sequence whereas in the Voyeboeuf, Court and Savagnières sections oncolitic facies are patchily distributed and even inexistent in the Hautes-Roches section.

The various origins of peloids (e.g., bacterial, fecal, reworked cohesive mud, internal molds of fossils, complete micritization of grains) may explain the relatively widespread distribution of the **peloid-rich facies** in almost all sections.

**Bioclastic facies** result from the accumulation of debris of organisms. The fauna and flora association reflects the environmental conditions during their life under the condition that they have been deposited in situ. However, taphonomic processes such as time averaging through bioturbation and transport through currents cannot be excluded. Bioclastic facies with semi-restricted fauna, characteristic of protected lagoonal environments, are mainly present at the base (sequences 0.5 and 1.1) of the Voyeboeuf and Savagnières sections, and around SB 2 of the Voyeboeuf, Court, and Savagnières sections. They also exist at the top of the Savagnières section around MFS Ox6+. Bioclastic facies with normal-marine fauna, representing open-lagoonal environments, are less abundant and occur only patchily (Fig. 7.3).

**Coral facies**, mainly composed of coral boundstones, reflect open-marine environments (i.e., semi-open and open lagoon in this study). Coral facies occur in the Court and Hautes-Roches sections. The first coral reefs appear in the Hautes-Roches section in sequence 1.1 and in the Court section in sequences 1.3 and 1.4. In the second small-scale sequence, coral facies are less abundant and correspond to small patchreefs.

Ooids are formed and transported by tidal currents in shallow waters. **Oolitic facies** predominate in the Hautes-Roches, Court, and Vorbourg sections. The spatial and temporal distribution is patchy. In the first small-scale sequence, the Pertuis section reveals the first and single oolitic deposit while in the Hautes-Roches and Court sections ooid accumulation starts somewhat later but then dominates especially at Hautes-Roches. The second small-scale sequence exhibits ooid deposits mainly in the Vorbourg, Hautes-Roches, and Court sections, and a few in the Savagnières section. Oolitic deposits are mainly driven by autocyclic processes (e.g., progradation or lateral migration of sedimentary bodies), which can simulate depositional sequences
or mask sequence boundaries and maximum flooding surfaces (STRASSER 1991). However, hydrodynamic changes denoted by sedimentary structures can point to relative sea-level fluctuations. Consequently, also oolitic deposits were precisely logged in order to identify potential sequential limits.

Figure 7.3 demonstrates the complexity of facies and depositional environment distribution across the platform and through time and illustrates the **facies mosaic** on the Swiss Jura platform (STRASSER & VÉDRINE in press). The observed facies distribution reveals a difference between 'proximal' and 'distal'

<figure><img src="image_84.png" /></figure>

**Fig. 7.3** - Time-space diagram illustrating the distribution of (a) major facies and (b) depositional environments of the six sections studied in the Swiss Jura Mountains (cf. Chap. 2 for details on facies and depositional environments). The sequence boundaries of the small-scale and the elementary sequences are considered as time lines delimiting the 100- and 20-kyr orbital cycles.
<figure><img src="image_85.png" /><figcaption>**Fig. 7.4** - Platform evolution through time of the Swiss Jura platform at selected time intervals (a) medium-scale SB Ox6+; (b) small-scale MF 1; (c) small-scale SB 2; (d) medium-scale MF Ox6+. Symbols as in Fig. 4.1.</figcaption></figure>
platform environments (Fig. 7.3b). Open lagoonal environments are more common in the “distal” platform positions and particularly in the Pertuis section. Inversely, protected lagoonal environments abound in Voyeboeuf, in the most “proximal” platform positions. The facies distribution of the Swiss Jura platform is thus partly linked to platform morphology. The latter is variable and was created by differential subsidence (WETZEL et al. 2003; ALLENBACH 2001). The sediment accumulation rates (carbonates and siliciclastics) accentuate or attenuate the relief of the platform. For example, during periods of coral reef growth and ooid production, equivalent to high carbonate productivity, the irregular morphology was enhanced. Reefs formed on structural highs along the platform margin but could later find themselves in rapidly subsiding areas due to tectonic inversion. Ooid bars not only occurred on the platform margin but also in the platform interior where morphological highs created favourable conditions. Locally, wide channels brought relatively open-marine conditions into more central platform positions, generating high-energy depositional environments. Behind reefs and bars, lagoons were isolated or protected from open-marine influence and from wave or current energy (Fig. 2.3; HUG 2003; STRASSER & VÉDRINE in press). Lagoonal facies with oncoids shows a patchy distribution. No systematic relationship with the other facies is recognizable, although the most persistent oncoid lagoons appear to be situated preferentially on the inner platform when compared to ooid bars and coral reefs.

The spatial facies distribution is monitored for different selected time intervals corresponding to major sequential limits of medium-scale and small-scale sequences (Fig. 7.4). At the SB Ox6+, a major sea-level fall caused the flushing of the platform by siliciclastics (Fig. 7.4a). The Voyeboeuf and Savagnières sections are dominated by bioclasts and do not present marls, suggesting that these sections were topographic highs. Then the transgression begins and leads to the disappearance of siliciclastics. The increase of accommodation favors homogenization and diversification of lagoonal facies and marks the small-scale MF 1 (Fig. 7.4b). The Voyeboeuf and Savagnières sections still present similar facies (micrite-dominated oncoid and peloidal facies). The Vorbourg and Pertuis sections show *Bacinella-Lithocodium* oncoid facies. High-energy deposits dominate in the Hautes-Roches section and a large coral patch-reef develops in the Court section. Then, sea level falls again. At this time, minor inputs of siliciclastics are noticed especially

in the Hautes-Roches and Voyeboeuf sections (Fig. 7.4c), birdseyes representing tidal flat facies are observed in the Court section, a serpulid-rich level occurs in the Savagnières section, and peloidal facies dominates in the Pertuis section. Finally, sea level rises again, creating an increase of the accommodation space (around the MF Ox6+; Fig. 7.4d). Peloid- and bioclast-rich facies exist in the Voyeboeuf, Pertuis, and Savagnières sections. As for MF 1, the Court and Hautes-Roches sections present ooid bars and coral patch-reefs while the Vorbourg section is rich in brachiopods and micrite-dominated oncoids.

## 7.2 ONCOIDS

Based on thin sections and macroscopic samples from the Hauptmumienbank Member, four types of oncoids are distinguished (cf. Chap. 2, Fig. 2.1). Based on the well-defined high-resolution time framework, stratigraphic and spatial oncoid distribution is examined and the main environmental controlling factors of oncoid growth and distribution are discussed (see also VÉDRINE et al. 2007).

### 7.2.1 Stratigraphic and spatial distribution

From the high-resolution sequence-stratigraphic framework, a time-space diagram illustrating the distribution of oncoid types is synthesized (Fig. 7.5). A proximal-distal trend is revealed by the greater abundance of type 3 oncoids in the “distal” platform sections (Pertuis and Savagnières); while in the “proximal” section (Voyeboeuf) type 2 and 1 oncoids predominate. Hautes-Roches and Court, which are mainly composed of ooid-rich deposits and patch-reefs, almost exclusively present type 2 oncoids. The different oncoid types coexist, except types 1 and 4 that occur in very different environmental settings concerning mainly water energy. Regardless of these general trends, it appears that local conditions (e.g., clay input) played an important role in oncoid development. Migrating ooid bars locally interrupted the oncoid sedimentation (Fig. 7.3).

A specific evolution of oncoid types is occasionally observed within small-scale and elementary depositional sequences. Oncoids are generally more abundant in the first small-scale sequence than in the second one (Fig. 7.5). Based on the high-resolution sequence-stratigraphic framework, a detailed description of the oncoid type distribution is given for the Vorbourg sec-
<figure><img src="image_86.png" /><figcaption>**Fig. 7.5** - Distribution of oncoid types in the elementary sequences within the transgressive deposits of the Ox 6+ medium-scale sequence for the six studied sections in the Swiss Jura. Sections are arranged from “proximal” to “distal” platform. Relative abundance (1-present, 2-common, 3-abundant) for each oncoid type is estimated from thin section analysis. Intervals detailed in Figs 7.6 and 7.7 are shaded.</figcaption></figure>

tion (Fig. 7.6). The (early) transgressive deposits of the first small-scale sequence contain abundant type 2 oncoids. Migration of ooid bars locally interrupts oncoid formation in the upper part of the transgressive deposits of elementary sequence 1.3. Around and after the maximum flooding of this elementary sequence, type 3 and 4 oncoids appear. In the two last elementary sequences, type 3 oncoids dominate and type 2 oncoids reappear. Small-scale SB 2 is marked by Miliolid foraminifera and type 3 oncoid rudstones. Within the elementary sequences of Vorbourg, no particular trend is observed.

However, in the Voyeboeuf and Savagnières sections (Fig. 7.7), some elementary sequences display

the same evolution as the small-scale sequence of Vorbourg. Elementary sequence 1.5 in Savagnières presents two beds, interpreted respectively as transgressive and highstand deposits. The transgressive bed can be subdivided into bioclastic wackestone with normal-marine fauna and oncoid-rich wackestone-floatstone. The maximum flooding is marked by type 3 and 4 oncoid mudstone to floatstone. The highstand deposit is dominated by peloidal and bioclastic wackestone. The small-scale sequence boundary is placed at the base of a serpulid bioherm with peloids and oysters. The next elementary sequence 2.1 has an early transgressive deposit rich in peloids, serpulids, oysters, Textulariid foraminifera, and a few type 2 oncoids. In Voyeboeuf, elementary sequence 1.5 is
<figure><img src="image_87.png" /><figcaption>**Fig. 7.6** - Detail of the first small-scale sequence of the Vorbourg section, showing the distribution of the four oncoid types, facies types, palaeoenvironmental interpretation, and the sequence-stratigraphic framework. For discussion refer to text. For symbols refer to Fig. 4.1.</figcaption></figure>
<figure><img src="image_88.png" /><figcaption>**Fig. 7.7** - Selected elementary sequences of the Savagnières and Voyeboeuf sections around small-scale sequence boundary SB 2, showing the distribution of the four oncoid types, facies types, the palaeoenvironmental interpretation, and the sequence-stratigraphic framework. For discussion refer to text. For symbols refer to Fig. 4.1.</figcaption></figure>
composed of ostracode and Textulariid foraminifer mudstones (Fig. 7.7b). The highstand deposit of this elementary sequence consists of clays with coal debris, detrital quartz, and serpulids, indicating relatively proximal and semi-restricted conditions, and two thin wavy limestone beds composed of ostracode and foraminifer mudstone. The small-scale sequence boundary is placed at 6.6 m in a clay-rich level. The transgressive deposit consists of foraminifer mudstone with some type 1 oncoids. The maximum flooding is placed around thin wavy beds surrounded by marls. The highstand deposit is made of a massive bed, which is rich in type 3 and 4 oncoids and benthic foraminifera at the base. In its upper part, type 3 oncoids and foraminifera (Textulariid and Miliolid) dominate. The next sequence boundary SB 2.2 is marked by peloid and foraminifer (Miliolid and less Textulariid) wackestone rich in detrital quartz. In this transgressive deposit, some type 2 oncoids are observed.

<figure><img src="image_89.png" /><figcaption>**Fig. 7.8** - Oncoid types, their preferred position in a depositional sequence, and main controlling factors. For discussion refer to text.</figcaption></figure>

This study on the spatial and stratigraphic distribution of oncoids in lagoonal deposits shows that type 1 and 2 oncoids have both a similar distribution and preferentially form around the sequence boundaries and during transgressive conditions of elementary and small-scale sequences, indicating that enough energy was furnished to roll the particles (Fig. 7.8). Type 3 and 4 oncoids have an inverse occurrence because they mainly occur around small-scale and elementary maximum-flooding intervals and in highstand deposits, characterizing periods of low-energy and increase of *Bacinella-Lithocodium* growth (Fig. 7.8).

### 7.2.2 Controlling factors

Oncoid growth and distribution are mainly controlled by sea level and climate (Figs 7.8 and 7.9). The main direct environmental parameters are water energy, water depth, trophic level, and sediment accumulation rate. Their influence is recognized in the stratigraphic and spatial distribution of oncoids in large-scale to elementary sequences as well as in the oncoid surface morphology, the biotic composition of the oncoid cortex, and the encasing sediment.

#### Water energy and water depth

The oncoid-rich facies of the Hauptmumienbank Member appears at the beginning of a 2^nd^-order sea-level rise, covers the entire medium-scale Ox6+ sequence and then disappears above the 3^rd^-order SB Ox7 (Fig. 1.4). This stratigraphic distribution points to a link with changes of relative sea level. Sea-level fluctuations play an important role in opening and closing shallow lagoons behind barrier systems and thus strongly modify environmental factors (e.g., water energy, oxygenation, salinity, temperature, sediment transport). The specific evolution of oncoid types within small-scale and elementary sequences confirms the link also to high-frequency sea-level fluctuations. The sub-spherical type 1 and 2 oncoids, mainly resulting from sediment trapping by micro-organisms, preferentially occur around sequence boundaries and in transgressive deposits where higher energy conditions predominated in relatively shallow water (Fig. 7.8). This implies that these particles needed energy to be overturned and rolled. On the other hand, the lobate type 3 and 4 oncoids are characteristic of low-energy environments where microbial meshwork had time to grow. Type 3 and 4 oncoids are concentrated around maximum floodings and in highstand deposits where low-energy conditions existed in relatively deep water but still in the photic zone (Fig. 7.8). Consequently, the oncoid surface morphologies indicate the intensity of rolling on the sea floor and thus the degree of wave or current energy and, indirectly, water depth. The biotic composition and diversity in the oncoid cortex and the encasing sediment (e.g., texture, faunal association) provides additional information on water energy and water depth.

Platform morphology was created by differential subsidence of tectonic blocks and by differential sediment accumulation (e.g., PITTET 1996; ALLENBACH 2001; HUG 2003; WETZEL et al. 2003). The already
<figure><img src="image_90.png" /><figcaption>Fig. 7.9 - Interrelated factors controlling oncoid growth and distribution. For discussion refer to text.</figcaption></figure>

existing highs and depressions were further accentuated by the increase of carbonate production in reefs and on ooid shoals. Oncoids were preferentially accumulated in depressions. HALLOCK (1988) also suggests that the point sources of terrigenous input and the platform morphology increase the spatial heterogeneity of nutrient distribution and thus trophic level. Consequently, variable platform morphology was partly responsible for the patchy distribution of oncoid-rich facies and oncoid types.

In the studied sections, some oncoid-rich beds are capped by ooid bars, which may have been controlled by autocyclic processes. On shallow carbonate platforms, such processes are inherent and involve lateral migration of sediment bodies (GINSBURG 1971; PRATT & JAMES 1986; STRASSER 1991). Consequently, autocyclic processes were superimposed on the orbitally controlled sea-level fluctuations and represented an additional factor in oncoid growth and distribution.

## Accumulation rate

Carbonate accumulation rates mainly depend on the local carbonate production but can also be modified by carbonate transport (supply or exportation). On the Swiss Jura platform, the accumulation rate of oncoid-

rich lagoons has been estimated at around 0.11 mm/ year (after sediment decompaction), which is low in comparison to modern lagoonal environments (STRASSER & SAMANKASSOU 2003; STRASSER & VÉDRINE in press).

Microbial crusts commonly grow slowly. Their occurrence is considered to be diagnostic for low background sedimentation rates (LEINFELDER et al 1993). Similar requirements are probable for oncoids, which are partly or totally composed of the same microencrusters such as the *Bacinella-Lithocodium* association (e.g., PERYT 1983; DUPRAZ 1999).

## Trophic level

The deposition of the massive Hauptmumienbank limestones on top of the Röschenz marls points to a general trend toward a lower input of siliciclastics and a higher carbonate productivity. Decreasing terrigenous input may have been related to a change to less humid climate conditions in the hinterland.

In the studied sections, oncoids occur abundantly in limestones and occasionally in marls. In the limestone beds, type 1, 2, 3 and 4 oncoids are present whereas in the marly intervals only type 2 oncoids are found. The biotic composition of the oncoid cortex
and the encasing sediment imply specific trophic conditions. Type 2 oncoids enclose serpulids and Textulariid foraminifera, which tolerate meso- to eutrophic conditions (e.g., BRASIER 1995; FUGAGNOLI 2004). The concomitant growth of type 2 oncoids with such organisms and their occurrence in clay-rich environments suggest that they were adapted to low water transparency and potentially high trophic levels. Type 1 oncoids are rare in the studied sections and found in limestone beds. They show affinity with type 2 oncoids: the scarcity of associated fauna probably indicates mesotrophic or eutrophic conditions, making life difficult for carbonate producing organisms. Type

3 and 4 oncoids are constituted of light-dependent microencrusters (i.e., *Bacinella* and *Lithocodium*) implying clear water and oligotrophic conditions (DUPRAZ & STRASSER 1999). They occur in the massive limestone beds, which formed when siliciclastic and nutrient input was reduced. Growing under water on the sediment surface, oncoids thus monitor water transparency and trophic level, which are both driven by the episodic influx of terrigenous sediment (quartz, clays, organic matter, and nutrients) onto the shallow carbonate platform (Fig. 7.9). This study demonstrates that oncoids are valuable proxies for high-resolution palaeoenvironmental and palaeoecological studies.

<figure><img src="image_91.png" /><figcaption>**Fig. 7.10** - Benthic foraminifer distribution (all foraminifera, agglutinated, hyaline, porcelaneous) in the Voyagebeuf, Vorbourg, Savagnières, and Pertuis sections placed in the high-resolution time-space diagram. Specific occurrences of foraminifera in the transgressive and highstand deposits of some elementary and small-scale sequences are highlighted.</figcaption></figure>
## 7.3 BENTHIC FORAMINIFERA

Benthic foraminifera have a great potential as indicators of depositional environments in the geologic past (e.g., SMITH 1955; MARTIN 1988; CANNARIATO et al. 1999). These bottom dwelling organisms occur in a wide range of marine environments from normal-marine to hypersaline, and from deep to shallow (GEEL 2000). Temperature and salinity have been considered as the main factors controlling the spatio-temporal distribution (e.g., MITCHELL 1996). However, trophic level and oxygenation appear as more determinant factors (concept of the organic flux, e.g., VAN DER ZWAAN et al. 1999). Hence, indirectly, also climate and sea level play an important role in the distribution of benthic foraminifera.

This study proposes to investigate the role of sea level on foraminifera in the Hauptmumienbank Member. Based on the high-resolution time framework (20-kyr time resolution) established by sequence- and cyclostratigraphy, the stratigraphic and spatial distribution of foraminifera is examined in four of the six studied Swiss Jura sections. The lagoonal deposits reveal three main suborders of benthic foraminifera: Textulariina, Miliolina, and Lagenina, characterized by three different types of wall (agglutinated, hyaline, porcelaneous respectively; cf. Chap. 2). For each category, the number (absolute abundance) of foraminifera has been counted per thin section (each thin section has the same surface area). In the graphical representation of the foraminifer distribution, samples with less than ten specimens have been masked in order to remove the potential noise caused by reworking. In addition, foraminifer and oncoid distribution are compared and discussed.

### 7.3.1 Stratigraphic and spatial distribution

**Agglutinated, porcelaneous, and hyaline foraminifera**

The foraminifer assemblages of the studied sections are mainly composed of forms with agglutinated tests (Textularids) and in a lower amount of those with porcelaneous tests (Miliolids). Hyaline foraminifera (only *Lenticulina*) are much rarer (Fig. 7.10).

The distribution of the total abundance of benthic foraminifera shows a relatively patchy distribution without any "proximal-distal" trend. Due to the predominance of agglutinated foraminifera in

the studied deposits, the agglutinated foraminifer distribution is relatively similar to the one of the total number. Agglutinated foraminifera apparently occur without a clear facies dependence. The determination of genera and species would possibly give more detailed information (MURRAY 1991). This, however, is outside of the scope of this thesis. The porcelaneous and hyaline foraminifera show a correlation with the sequence-stratigraphic evolution. Miliolids are preferentially found in the highstand deposits of the elementary and/or small-scale sequences (Figs 7.10

<figure><img src="image_92.png" /><figcaption>**Fig. 7.11** - Schematic distribution of Textularids, Miliolids, and *Lenticulina* in relation with the sequence-stratigraphic evolution.</figcaption></figure>

and 7.11). The hyaline foraminifera, rare in the studied sections, are preferentially found in the transgressive deposits of elementary sequences (Figs 7.10 and 7.11). A similar distribution of benthic foraminifera has been suggested by ARNAUD-VANNEAU (1994) and was interpreted as being related to changes of accommodation and sediment and nutrient input. Consequently, the benthic foraminifera are indirectly linked to relative sea-level and indirectly also to climate changes. As for the oncoids, the variations in the foraminifer distribution are due to the combination
<figure><img src="image_93.png" /><figcaption>**Fig. 7.12** - Distribution of *Mohlerina basiliensis* and oncoids along the studied sections.</figcaption></figure>
<figure><img src="image_94.png" /><figcaption>**Fig. 7.13** - Distribution of *Mohlerina basiliensis* and oncoids from protected lagoon to basin. Published distribution according to DARGA & SCHLAGINTWEIT (1991), SCHLAGINTWEIT & EBLI (1999), SASARAN et al. (1999), VELIĆ et al. (2002), GAWLICK et al. (2003), and FLÜGEL (2004).</figcaption></figure>

of physical, chemical, and biological factors (see below).

## Co-occurrence of the benthic foraminifer *Mohlerina basiliensis* and *Bacinella-Lithocodium oncoids*

Together with Textularids, Miliolids and *Lenticulina*, the foraminifer *Mohlerina basiliensis* occurs in small amounts and is sparsely distributed in the studied sections (VÉDRINE & SPEZZAFERRI 2007; VÉDRINE 2008; Fig. 7.12). *Mohlerina basiliensis* was found in only 40 of the 265 analyzed thin sections. In the oncoid-rich Hauptmumienbank Member, *M. basiliensis* is mostly found in oncoid-rich and bioclastic (brachiopods, echinoderms) wackestones, characterizing low-energy and normal-marine conditions (Pl. 9/10-11, cf. Chap. 2). Sporadically, *M. basiliensis* appears in higher energy deposits such as ooid- and peloid-rich packstones (Pl. 9/10). Several authors found this foraminifer preferentially in high-energy facies from the internal platform to the slope, commonly associated with *Protopeneroplis* sp., *Trocholina* sp., *Pseudocyclammina* sp., *Andersolina* sp., *Nautiloculina* sp., *Lenticulina* sp. *Valvulina* sp., and Miliolids (DARGA

& SCHLAGINTWEIT 1991; SCHLAGINTWEIT & EBLI 1999; SASARAN et al. 1999; VELIĆ et al. 2002; GAWLICK et al. 2003). Recently, FLÜGEL (2004) and VELIĆ et al. (2002) described *M. basiliensis* as characteristic of upper slope facies (Fig. 7.13).

The *M. basiliensis* distribution is relatively patchy. The ‘distal’ platform facies present a higher abundance of this foraminifer (Figs 7.12 and 7.14). Its minimum abundance (only 2 specimens per thin section) is found in the relatively “proximal” Voyeboeuf section, and the maximum abundance (up to 12 specimens per thin section) in the relatively “distal” Pertuis section. In the latter, *M. basiliensis* has a more extensive stratigraphic occurrence than in the other sections (Figs 7.12 and 7.14).

The distribution of *Bacinella-Lithocodium* oncoids is also relatively patchy. However, *Bacinella-Lithocodium* oncoids are more abundant in “distal” platform facies, especially in the Pertuis section (Fig. 7.12). Micrite-dominated oncoids are present in all sections but less abundant in Pertuis. Occasionally, *Bacinella-Lithocodium* and micrite-dominated oncoids coexist.
<figure><img src="image_95.png" /><figcaption>**Fig. 7.14** - *Mohlerina basiliensis* distribution within elementary and small-scale depositional sequences. The encircled numbers indicate the first occurrence of *M. basiliensis* in each section.</figcaption></figure>

In a time-space diagram, the *M. basiliensis* distribution shows a particular evolution (Fig. 7.12). The first occurrence of the foraminifer differs between sections through time and space but follows the “proximal-distal” trend. In the most “distal” section (Pertuis), the first occurrence is coeval with the beginning of the 2nd order marine transgression (arrow 1, Fig. 7.14). In the three other sections, the first occurrence appears successively later. This illustrates the stepwise flooding and opening of the shallow platform during the transgression (Fig. 7.15).

Contrary to oncoids, the *Mohlerina basiliensis* distribution does not reveal a clear preferential occurrence within small-scale and elementary sequences or around sequence boundaries or maximum floodings (Fig. 7.14).

However, the distribution of *Mohlerina basiliensis* and *Bacinella-Lithocodium* oncoids shows a strong correlation (Fig. 7.12). *M. basiliensis* preferentially occurs when *Bacinella-Lithocodium* oncoids are
present. On the other hand, the presence of *Bacinella-Lithocodium* oncoids does not necessarily imply an occurrence of *M. basiliensis*. For example, the base of the Voyageboeuf section contains *Bacinella-Lithocodium* oncoids in three intervals but no *M. basiliensis* (elementary sequences 1.2 and 1.3; Fig. 7.12). The co-occurrence of these two components thus suggests common environmental factors.

### 7.3.2 Controlling factors

The distribution of benthic foraminifera is controlled by biological, physical, and chemical factors. Water depth and water turbidity are important parameters because they control the amount of light received by the benthos. The substrate on or in which they live influences the test morphology (MURRAY 1991). Other important factors are oxygen content and food availability, which act on the foraminifer

<figure><img src="image_96.png" /><figcaption>**Fig. 7.15** – Platform flooding illustrated by the distribution of *Mohlerina basiliensis*. The platform morphology suggested by facies analysis allows to explain the successive first occurrences of *M. basiliensis*. Concerning the Savagnières and Pertuis sections, which are relatively close, the first occurrence and the abundance of *M. basiliensis* suggest a deeper bathymetry for the Pertuis section. Refer to text for more discussion.</figcaption></figure>
diversity (e.g., JORISSEN et al. 1995; FUGAGNOLI 2004). Temperature and water chemistry also play a role in foraminifer distribution by modifying the CaCO₃ solubility in seawater, which acts on the nature of the test (GREINER 1969).

The palaeoenvironmental and palaeoecological requirements of *Mohlerina basiliensis* are not well known. The successively later occurrence of *M. basiliensis* from “distal” to “proximal” sections illustrates the stepwise flooding of the platform and thus implies a dependence on normal-marine conditions. Its occurrence in high-energy as well as in low-energy facies probably suggests a reworking by waves and currents. Good living conditions were rather

in low-energy environments, which is supported by the co-occurrence of this foraminifer with *Bacinella-Lithocodium* oncoids. Oncoids and their associated microencrusters grow on the sediment surface and thus monitor water transparency and trophic level, which are both controlled by terrigenous influx (clays and nutrients) onto the shallow carbonate platform. The *Bacinella-Lithocodium* association, included in the oncoid cortex, indicates oligotrophic, clear, oxygenated, and normal-marine waters (e.g. DUPRAZ & STRASSER 1999). Hence, the co-occurrence of *Bacinella-Lithocodium* oncoids and *M. basiliensis* suggests that this foraminifer required similar ecological conditions (VÉDRINE & SPEZZAFERRI 2007; VÉDRINE 2008).

* * *
# 8 - PALAEOCLIMATE

Jurassic climate has classically been considered as warm and equable with minimal equator-to-pole thermal gradient, extensive evaporite deposition, and temperate conditions at high latitudes without permanent ice caps (e.g., FRAKES et al. 1992; HALLAM 1993; PRICE 1999). This global greenhouse period is consistent with the recent estimations of atmospheric concentrations of the greenhouse gas carbon dioxide, which is estimated to have been 4-10 times higher than modern (pre-industrial) values (e.g., VAN AARSSEN et al. 2000; RETALLACK 2001). However, there is increasing evidence for cooler climate periods that were superimposed on this generally warm trend. For example, FRAKES et al. (1992) consider the Middle Jurassic to Early Cretaceous (Bajocian to mid-Albian; 171.6–106 Ma) as a time of cool non-glacial conditions with high-latitude seasonal ice (“cool mode”). Recent General Circulation Model results indicate that Antarctica and Australia may have sustained an ice/snow-sheet during the Oxfordian to Middle Tithonian (SELLWOOD et al. 2000). Therefore, the Late Jurassic climate may have been interspersed with cool phases, and/or with a major cold phase associated with the development of an ice/snow-sheet in the southern hemisphere (GRÖCKE et al. 2003).

Most climate information comes from land plant fossils that are the most valuable palaeoclimate indicators. During the Late Jurassic, warm and arid conditions prevailed all over the Eurasian hinterland (VAKHRAMEEV 1991) whereas chilly and humid conditions predominated in the North Pacific region (PARRISH 1992), leading to strong contrasts in the longitudinal thermal gradients in the northern hemisphere.

## 8.1 PALAEOCLIMATE DURING THE OXFORDIAN

For the northern Tethys margin, several studies have suggested a relatively cool and humid climate during the Late Callovian-Early Oxfordian, whereas the Middle and Late Oxfordian were characterized by warmer and more arid conditions (e.g., VAKHRAMEEV 1991; BAUSCH et al. 1998; ABBINK et al. 2001; DROMART et al. 2003a, b). The onset of this warming was in the earliest Middle Oxfordian (Fig. 8.1). This pattern of climate change is in agreement with the pattern of sedimentation. In the European Tethys realm, the Late Callovian and Early Oxfordian are marked by a decrease of marine carbonate sedimentation: absence of sediments, condensed and iron-rich (COLLIN et al. 2005), organic carbon-rich shales (e.g., Terres Noires, DE GRACIANSKY et al. 1999), or siliceous deposits (e.g., radiolarites, BARTOLINI et al. 1996). Then, carbonate sedimentation recovered significantly during the Middle and Late Oxfordian (DROMART et al. 1996, 2003a, b). This evolution is probably linked to oceanographic changes related to the opening of the North Atlantic (ABBINK et al. 2001; RAIS et al. 2005). However, the greater spread of limestone facies in the Late Jurassic is probably more a consequence of the greater extent of epicontinental seas at that time than of higher temperatures (HALLAM 1975).

Sedimentological, floral, faunal, and geochemical indicators for palaeoclimate in the Oxfordian will be discussed only briefly. Clay minerals and stable isotopes (C, O) have been analysed in detail and are discussed in separate chapters.
<figure><img src="image_97.png" /><figcaption>**Fig. 8.1** - General palaeoclimatic evolution (FRAKES et al. 1992; ABBINK et al. 2001), eustatic sea-level curves (HAQ et al. 1987; HALLAM 1988), composite temperature and humidity curves from palynological data (ABBINK et al. 2001), δ¹³C record (BARTOLINI et al. 1996; PADDEN et al. 2001; LOUIS-SCHMID 2006), strontium isotope record (JONES et al. 1994), and clay mineralogy of basin sections for the Oxfordian of NW and Central Europe (DECONINCK & DEBRABANT 1985; DECONINCK et al. 1985) compared to the Swiss Jura platform climatic evolution (GYGI & PERSOZ 1986; PITTET 1996; HUG 2003). Note that the limit between Middle Oxfordian and Late Oxfordian is placed between the Transversarium and Bifurcatus ammonite zone in HARDENBOL et al. (1998) and between the Bifurcatus and Bimammatum ammonite zone in GYGI (1955, 2000). The studied time interval is shaded.</figcaption></figure>
<figure><img src="image_98.png" /></figure>

**Fig. 8.2** - Coal and evaporite distribution during the Late Jurassic (from REES et al. 2004).

### 8.1.1 Sedimentological indicators

The distribution of evaporites and coals is the best means of distinguishing more humid and more arid conditions (FRAKES 1979). Substantial deposits of evaporites (notably gypsum, anhydrite, and halite) indicate conditions of both warmth and aridity, whereas coals indicate swampy conditions in generally humid regimes, though there is no particular temperature connotation. In the Late Jurassic, evaporites preferentially occur in the low to mid latitudes while coals are more present in mid to high latitudes (REES et al. 2004, cf. Fig. 8.2). On the Swiss Jura platform, HUG (2003) found evaporites in the Late Oxfordian from SB Ox7 to SB Ox8 (cf. Fig. 8.1).

On the Swiss Jura platform, siliciclastics occur during the Early Oxfordian and at the Middle-Late Oxfordian transition (Fig. 8.1). Generally, the relative abundance and distribution of siliciclastics mainly reflects the intensity of continental run-off, which is related to the abundance of rainfall in the source areas and thus to local palaeoclimatic conditions (HALLAM 1984). Figure 8.3 shows Late Oxfordian palaeogeography with the different source areas and the directions of sediment export. In the Swiss Jura, siliciclastics are observed mainly at the 100-kyr sequence boundaries and are explained by rainfall controlled by palaeolatitudinal variations of atmospheric cells in tune with orbitally induced

<figure><img src="image_99.png" /></figure>

**Fig. 8.3** - Palaeogeographic map of the study areas for the Late Oxfordian showing the different sedimentation areas, the emerged lands and the direction of sediment export (arrows). Modified from MEYER & SCHMID-KALER (1989, 1990), THIERRY et al. (2000), and HUG (2003).
insolation cycles (PITTET 1996; PITTET & STRASSER 1998b). However, tectonic uplift in the hinterland or sea-level fall may also contribute to increasing siliciclastic input into the sedimentary basins. Hence, the siliciclastic-carbonate relationship alone is not sufficient for palaeoclimatic interpretations (RUFFELL & RAWSON 1994).

Terrigenous particles such as quartz are common in mixed carbonate-siliciclastic platform systems. Their fluvial origin (GYGI & PERSOZ 1986) allows to relate the quartz proportions to the intensity of hinterland erosion. Authigenic quartz (neoformation and/or overgrowth on detrital grains) is present only in minor proportions PITTET (1996).

### 8.1.2 Floral and faunal indicators

The distribution and thermal tolerances of land plants provides one of the main sources of climatic evidence for the Late Jurassic (FRAKES et al. 1992). In mid to low latitudes, particularly in the northern hemisphere, the vegetation was dominated by Cheirolepidicean conifers, which shed *Classopollis* pollen. These conifers grew on drained slopes, often near areas of evaporite deposition. Pollen of this type is most commonly found in association with evaporitic sediments and thus the presence of this conifer family has been used as an indicator of arid climates (VAKHRAMEEV 1981). A high abundance of *Classopollis* pollen occurs at the Oxfordian in the former URSS region. Evaporites also increased in quantity at this time, there were fewer clastics, and corals spread further north in the northern hemisphere (FRAKES 1979). This coincides with a long-term transgressive trend, and it may be that the increase in area of shallow seas provided greater areas for the formation and preservation of evaporites and the entrapment of pollen, rather than reflecting a great increase of aridity (FRAKES et al. 1992).

Fluctuations in the biogeographical distribution of coral reefs during the Oxfordian suggest a latitudinal shift of climatic belts. After the cooling event of the Late Callovian-Early Oxfordian, the recovery of coral reefs in the Middle Oxfordian is limited to the European epicontinental platform between about 25° and 35°N. These reefal formations seem to indicate oligotrophic conditions at their climax stage, and hence scarce nutrient supply, suggesting an arid climate on the bordering landmass (CECCA et al. 2005). However, the maximum extent of radiolarites in the Middle

Oxfordian in the deep basin of the Mediterranean Tethys implies humid conditions marked by increased run-off and nutrient input (BARTOLINI et al. 1999). In the Late Oxfordian, coral reefs recovered completely and were distributed between 10°S to 32°N, implying generally more arid conditions and/or platform isolation (CECCA et al. 2005). These coral reefs differ from those of the Middle Oxfordian by the occurrence of microbialites suggesting mesotrophic to eutrophic conditions and implying more nutrient input from the hinterland and more humid climatic conditions (PITTET 1996; DUPRAZ & STRASSER 1999, 2002).

During the Oxfordian, two major faunal provinces existed in the northern hemisphere: the northern Boreal realm and the low-latitude Tethyan realm (HALLAM 1994; CECCA 2002 and references therein). The limits of these realms varied in response to eustatic, climatic, and tectonic changes. In the latest Callovian to Kimmeridgian, Tethyan ammonite fauna migrated from the south into the Boreal realm, following a long-term palaeotemperature increase (MALCHUS & STEUBER 2002).

### 8.1.3 Geochemical indicators

The Oxfordian is characterized by the lowest $^{87}$Sr/$^{86}$Sr values in the Phanerozoic (JENKYNS et al. 2002). This may be related to decreased input of Sr from continental erosion and/or to increased input of mantle strontium with a low $^{87}$Sr/$^{86}$Sr ratio (COGNÉ & HUMLER 2004). Generally, increased humidity (VAN AARSSEN et al. 2000; ABBINK et al. 2001) and increased siliciclastic input (GYGI & PERSOZ 1986) in the Early Oxfordian imply a high weathering rate if compared to the Middle and Late Oxfordian. Hence, low input of continental Sr is unlikely to be the cause of low $^{87}$Sr/$^{86}$Sr. JONES et al. (1994) interpreted these low values, also found in the Early-Middle Oxfordian (Cordatum-Densiplicatum zones) from the British Isles, as the result of a major pulse of hydrothermalism related to the onset of the rifting in the North and South Atlantic regions.

## 8.2 CLAY MINERALS

### 8.2.1 Introduction

The analysis of clay-mineral assemblages from marine deposits can provide significant palaeoclimatic information on the provenance, sediment sorting, and
<figure><img src="image_100.png" /></figure>

**Fig. 8.4** - Origin of clays (detrital, pedogenic, authigenic, diagenetic, metamorphic, and hydrothermal) and their link with the sedimentary cycle (uplift, erosion, and burial). Modified from ESLINGER & PEVEAR (1988).

weathering processes on the adjacent continental areas (e.g., CHAMLEY 1989; SELLWOOD & PRICE 1993; RUFFELL et al. 2002). However, clay-mineral assemblages are not only controlled by climate changes (erosion and weathering) but also by tectonics (platform morphology, uplift, burial) and eustatic sea-level changes (CHAMLEY 1989). The different types of clay minerals (detrital, pedogenic, authigenic, diagenetic, metamorphic, and hydrothermal) thus result from the interaction of climate, tectonics, and sea level (Fig. 8.4).

The formation of clay minerals in soils strongly depends on the weathering environment (SLADEN 1983). It results, either directly or indirectly, from the hydrolytic decomposition of primary aluminosilicates. Leaching potential (wash out of cations and silica oxides by percolating waters) depends on vertical water movement, weathering profiles, and parent rock and is controlled by the relationship between precipitation and evaporation (SINGER 1980). Clay minerals are formed in soils or eroded from parent rocks. Tectonic uplift favours the direct removal of materials from rocky substrates and is less favourable for pedogenesis (BLANC-VERNET et al. 1979; CHAMLEY 1989). Erosion, reworking and transport from the continental source area to marine settings by water or wind may modify the composition of clay-mineral assemblages, e.g., by mixing different horizons with different degrees of alteration (e.g., CURTIS 1990). The transformation of kaolinite in sea water may also change the clay-mineral assemblages (REX & MARTIN 1966).

The relationship between climatic parameters (particularly seasonality, average annual temperature and precipitation) and clay-mineral formation is complicated by the intervention of other factors such as topography, type of vegetation and organisms, parent rock composition, and length of time available for soil formation (e.g., SINGER 1980; SLADEN 1983; CHAMLEY 1989; VELDE 1995). The longer the transport distance, the higher is the possibility that complex alteration processes happen during intermediate sediment storage, and less the palaeoclimatic message of the deposited clay-mineral assemblage is clear (THIRY 2000).

Once deposited in the sedimentary environment, clay-mineral assemblages may be further modified by diagenetic processes (CHAMLEY 1989; KÜBLER & JABOYEDOFF 2000; DECONINCK et al. 2001) and thus the original climatic signal will be distorted. Post-depositional diagenetic changes include alteration and neoformation of clay minerals (e.g., CHAMLEY 1989; RUFFELL et al. 2002). The most common process is the transformation of smectite to illite during burial diagenesis (CHAMLEY 1989). However, this transformation may also happen under surface conditions (e.g., in hypersaline lagoons; DECONINCK et al. 1988, 2001).

In marginal-marine settings, the clay-mineral assemblages are mainly characterized by kaolinite, illite, smectite, mixed-layer minerals, and palygorskite. Kaolinite formation is supported by deep weathering
associated with perennial precipitation and a soil temperature not dropping below 15°C (e.g., HALLAM et al. 1991). Fairly high evaporation in combination with high annual precipitation in humid warm to temperate climates leads to extensive leaching and to deep weathering profiles, creating kaolinite- and gibbsite-rich acid soils (e.g., SINGER 1980; SLADEN 1983; WRIGHT 1994). Consequently, kaolinite in marine sediments is mainly considered as a detrital mineral resulting of soil reworking. **Illite** and **chlorite** may derive from the erosion of unaltered bedrock. An increase in these minerals is often observed in areas with steep relief, where rapid erosion does not allow for soil formation and maturation (DECONINCK et al. 1985; GAWENDA 1999). **Illite** is often thought to be the product of physical erosion dominating in the hinterland (CHAMLEY 1989; RUFFEL & BATTEN 1990). According to DECONINCK (1993), low relative sea level causes a high erosion rate and thus favours the development of illite-rich sediments. SINGER (1984) claims that well-crystallized illite indicates formation in either cold or dry conditions with minimum hydrolysis. Chlorite mainly derives from the alteration of crystalline (basic igneous or low-grade metamorphic) rocks. Hence, regional variations in the source-rock composition may influence the chlorite content (LAMY et al. 1998).

**Palygorskite** and **smectite** are characteristic for soils with a shallow weathering profile in seasonal, semi-arid climates (Mediterranean-type climate). Palygorskite is an excellent marker for arid to semi-arid conditions (SINGER 1980). Smectite forms preferentially under tropical and subtropical conditions that are characterized by an elevated seasonality of alternating wet and dry periods (ROBERT & KENNET 1994; GAWENDA 1999). Moreover, smectite may be entirely volcanic in origin, deriving directly from the weathering of lava and ash (bentonites; e.g., PELLENARD et al. 2003).

The **kaolinite-illite ratio (K/I)** is commonly used for palaeoclimatic reconstruction and reflects the influence of climate and relief changes on chemical weathering and soil formation. The main influences on the K/I ratio are interpreted to be changes in climate and relative sea level (THIRY 2000; CURTIS 1990). A high K/I ratio generally indicates a more humid, warm climate.

The **illite crystallinity index (IC)** gives information on the chemical weathering intensity in the source area of the mineral (SINGER 1984; CHAMLEY 1989). Well-crystallized illites indicate a low chemical weathering

activity while poor illite crystallinities reflect stronger chemical weathering intensity (e.g., SINGER 1984). However, another source of illite is the replacement of smectite during burial diagenesis (diagenetic illites IC >0.42° Δ2θ CuKα, CHAMLEY 1989) but also at surface temperature and pressure in seawater of elevated salinity under a hot and arid climate (DECONINCK et al. 1988). RAMEIL (2005) noted a correlation between illite crystallinity index and regressive-transgressive phases of superimposed large- and medium-scale sequences. During large-scale maximum flooding periods, only detrital, well-crystallized illites (IC <0.42° Δ2θ CuKα) are deposited while around large-scale sequence boundaries illite neoformation is dominant. Additionally, illite crystallinity can also serve as a proxy of proximal-distal facies trends (RAMEIL 2005).

### 8.2.2 Material and analytical methods

Clay mineral analyses were performed on 70 samples from the Savagnières, Balingen-Tieringen, and Vergons sections. Sample preparation and X-ray diffraction (XRD) analyses were carried out at the Institute of Geology of the University of Neuchâtel (Switzerland) on a SCINTAG XRD 2000 diffractometer. Sample preparation is based on the methods described by KÜBLER (1987): The crushed sediment is mixed with an ionized aqueous solution (pH 7-8) and agitated. The carbonate fraction is then eliminated by the addition of HCl 10% (1.25 N) at room temperature. During a reaction time of 20 minutes or more, physical disintegration is stimulated in an ultrasound bath every 3 minutes, until all carbonate is dissolved. The insoluble residue is then repeatedly concentrated in a centrifuge and washed (5-6 times) until obtaining a neutral suspension (pH 7-8). The separation of various particle sizes (<2 µm and 2-16 µm) is obtained using the relation between sedimentation time and the depth of pipettage based on STOKES' Law. The <2µm fraction is transferred by pipette to a glass object tray and air-dried at room temperature. Then the <2 µm fraction is analysed after saturation with ethylene glycol and subsequent air-drying.

The intensities of the peaks that characterize each clay mineral (smectite, chlorite, illite, kaolinite, interstratified) are measured in *counts per second* (cps) for the semi-quantitative estimation of the proportion of clay minerals, which is therefore given in relative percent without correction factors (Annex 6). Variations of more than 10% are considered as significant. Quartz content has been estimated from thin sections and from
X-ray diffraction analyses of bulk-rock in percent, or from the 2-16 µm fraction in counts per second when the total amount of quartz in bulk was not sufficient. The illite crystallinity index (IC) is expressed as the ratio of the width of the (001) basal peak at mid-height to the peak height above background measured on <2 µm fraction diffractograms (KÜBLER 1987).

### 8.2.3 Results and interpretations

#### Swiss Jura (shallow platform)

Only the Savagnières section has been analysed for the clay minerals and is considered as representative for the Swiss Jura sections. Clay mineral assemblages of the Hauptmumienbank Member are mainly composed of kaolinite (40-60%), illite (30-50%), mixed-layer illite-smectite (2-15%), and chlorite (2-8%).

The illite and chlorite trends are relatively similar and are both inverse to the kaolinite trend (Fig. 8.5). The mixed-layer illite-smectite trend preferentially correlates with the kaolinite trend. The scarcity of smectites and the relative abundance of mixed-layer clays are a fundamental feature of the clay mineral associations of the Oxfordian of the Swiss Jura (GYGI & PERSOZ 1986). This scarcity can be explained in two ways: (1) the absence of smectite is an original feature; (2) it is due to the transformation of smectites into mixed layer illite-smectite and illite by burial diagenesis or by transformation in soils (GYGI & PERSOZ 1986; CHAMLEY 1989).

The comparison between clay-mineral assemblages and sequence stratigraphy (Figs 8.5 and 8.6) shows that illite content commonly is higher around sequence boundaries and lower around maximum floodings essentially at the scale of small-scale (100-kyr) sequences. Chlorite content is generally higher around the small-scale and elementary (100- and 20-kyr) SBs and lower around the small-scale and elementary MFs. Mixed-layer illite-smectite content is higher around the small-scale SBs and occasionally around the elementary SBs. In addition, high content of mixed-layer illite-smectite also occurs around the small-scale MF while the lowest values exist around the elementary MF. The kaolinite trend presents low values (45%) at the base of the section (i.e., in the first two elementary sequences 0.5 and 1.1) and then increases significantly (55-60%). High kaolinite contents appear in the elementary sequences 1.2 and 1.3, in the elementary sequence 1.5, and (3) in the

elementary sequences 2.2 and 2.3 but a link with the sequence stratigraphy is not evident.

The kaolinite-illite ratio is mainly controlled by the kaolinite content. High values occur around the small-scale (100-kyr) MF 1, at the small-scale (100-kyr) SB 2, and the highest values at the medium-scale (400-kyr) MF Ox6+. High quartz content preferentially occurs around the small-scale and elementary SBs (Fig. 8.5).

Most of the samples present low IC values (<0.42° Δ2θ CuKα, well-crystallized illites) rather indicating detritic illites. This is in agreement with the low burial depth (<2 km) of the Swiss Jura suggested by TRÜMPY (1980). Five samples present higher IC values pointing to poorly crystallized illites. These samples are preferentially positioned around small-scale and elementary sequence boundaries as found by RAMEIL (2005) for large- and medium-scale sequence boundaries. However, in this study, the IC is interpreted in terms of chemical weathering in the source area rather than as an indicator for aridity and elevated evaporation (see above, SINGER 1984; RAMEIL 2005).

In the Pichoux section, an increase of the K/I ratio from the Röschenz Member to the Hauptmumienbank Member (GYGI & PERSOZ 1986) can be interpreted as a change to more humid conditions. However, this evolution is in disagreement with the general facies evolution on the Swiss Jura platform that suggests a trend towards less humid conditions (cf. Fig. 8.1; GYGI 1995, 2000).

The clay-mineral distribution is mainly explained by sea-level and humidity changes (Fig. 8.6). **Humidity changes** act on the hydrological cycle that influences the continental run-off. In the Oxfordian, the sediment supply probably came from the Rhenish, London-Brabant and/or Bohemian Massifs (Fig. 8.3). **Sea-level changes** modify the base level that acts on the erosion potential (Fig. 8.6). During low sea-level periods, the terrigenous sediments (quartz) increase pointing to an intensification of land erosion in the source area, probably also intensified by more humid conditions. In contrast, during high sea-level periods, less erosion potential and probably humid conditions are evidenced by a decrease of siliciclastics (PITTET 1996; PITTET & STRASSER 1998b).

The kaolinite distribution in the studied shallow-carbonate platform section indicates that it occurred
Savagnières

<figure><img src="image_101.png" /></figure>

a Medium-scale sequences (400 kyr)

b Small-scale sequences (100 kyr)

c Elementary sequences (20 kyr)

<figure><img src="image_102.png" /></figure>

Climatic trends

more humid

more humid ?

more humid ?

less humid
<figure><img src="image_103.png" /></figure>

<figure><img src="image_104.png" /><figcaption>**Fig. 8.6** - Models explaining the clay-mineral distribution, particularly the kaolinite formation and accumulation during a low sea level (more humid conditions) and a high sea level (less humid conditions). Refer to text for discussion.</figcaption></figure>

during low as well as during high sea-level periods (Fig. 8.7). Kaolinite occurrence during low sea level is explained by favourable conditions (i.e., more humid) for kaolinite-rich palaeosoil formation and an increase of kaolinite erosion, transport, and accumulation. DECONINCK (1993) suggests that kaolinite preferentially accumulates in proximal platform areas. During high sea-level periods, kaolinite formation is reduced due to a less humid climate. Therefore, the observed high kaolinite contents around small-scale (100-kyr) maximum floodings probably suggest the remobilization of kaolinite formerly stored in the proximal part of the shallow platform.

**Fig. 8.5** - (facing page) Distribution of clay minerals, K/I ratio, quartz intensity and illite crystallinity within the transgressive part of medium-scale sequence 1 of the Savagnières section. Interpretations on climate evolution are proposed (on the right side).

## Swabian Jura (deep platform)

Clay mineral assemblages of the Impressa Mergel Formation of the Balingen-Tieringen section are mainly composed of kaolinite (30-55%), illite (20-50%), chlorite (5-20%), and mixed-layer illite-smectite (5-20%). The kaolinite trend correlates negatively with the illite, chlorite, and mixed-layer illite-smectite abundances (Fig. 8.7).

The comparison between clay-mineral assemblages and sequence stratigraphy shows a certain correlation (Figs 8.7 and 8.8): the illite content is generally higher around MFs and lower around SBs at the scale of small-scale (100-kyr) and elementary (20-kyr) sequences. Inversely, the chlorite and kaolinite contents are higher around SBs and lower around MFs of the small-scale and elementary sequences. In addition, high contents of mixed-layer illite-smectite occur around small-scale SBs and MFs.

The kaolinite-illite ratio is relatively stable except at two levels: the medium-scale SB Ox6+ and the
Balingen-Tieringen

<figure><img src="image_105.png" /></figure>

a Medium-scale sequences (400 kyr)

b Small-scale sequences (100 kyr)

c Elementary sequences (20 kyr)

<figure><img src="image_106.png" /></figure>

Climatic trends

more humid

more humid ?

less humid
<figure><img src="image_107.png" /></figure>

**Fig. 8.8** - Minima (-) and maxima (+) in the abundance of clay-minerals (kaolinite, illite, chlorite and mixed illite-smectite minerals), K/I ratio, quartz intensity, and illite crystallinity for the three studied sections (from shallow platform to basin) around the sequence boundaries and maximum floodings of 100- and 20-kyr sequences.

medium-scale MF Ox6+. These levels of high values correlate with those of the Swiss Jura section (Fig. 8.5).

The quartz content and the sequence stratigraphy do not show any clear correlation. The IC values (0.4°-0.9° Δ2θ CuKα) are representative of poorly crystallized illites. No correlation exists with the Savagnières section. As the Swiss Jura, the Swabian Jura has a low burial depth.

The trend in clay-mineral distribution of the Balingen-Tieringen section is comparable to the one of the Savagnières section. Discrepancies, particularly concerning the relative abundances of kaolinite and illite, are explained by the different palaeogeographic positions that imply different source areas and different transport distances (Fig. 8.3).

**Fig. 8.7** - (facing page) Distribution of clay minerals, K/I ratio, quartz intensity and illite crystallinity within the first two small-scale sequences 1 and 2 of the Balingen-Tieringen section. Interpretations on climate evolution are proposed (on the right side).

## SE France (basin)

Clay mineral assemblages of the Vergons section are mainly composed of illite (40-80%), chlorite (10-40%), mixed-layer illite-smectite (5-40%), and kaolinite (0-30%). In this sedimentation area, terrigenous particles probably come from the Central Massif (Fig. 8.3).

In the Vergons section, the kaolinite trend is mainly parallel to the mixed-layer illite-smectite trend (Fig. 8.9). The illite percentage shows an increase towards the top part of the section. The kaolinite trend shows a negative correlation with the chlorite trend except in the transgressive part of the second small-scale sequence where the kaolinite trend is parallel to the illite trend. According to these trends, the section can be subdivided into three parts (cf. Fig. 8.9). The first two parts correspond to studied interval in the Swiss and Swabian Jura sections.

The comparison between clay-mineral assemblages and sequence stratigraphy (Figs 8.8 and 8.9) shows that in part 1 kaolinite and mixed-layer illite-smectite contents are relatively high around small-scale (100-kyr) MFs while the illite and chlorite contents are
Vergons

a
b
c

SB OX7

1

1

SB Ox6+

0.5 m

0 m

samples

a Medium-scale sequences (400 kyr)

b Small-scale sequences (100 kyr)

c Elementary sequences (20 kyr)

<figure><img src="image_108.png" /></figure>

<figure><img src="image_109.png" /></figure>

Climatic trends

less humid

more humid

more humid ?

less humid
**Fig. 8.9** - (facing page) Distribution of clay minerals, K/I ratio, quartz intensity, illite crystallinity, and bulk mineralogy within the four small-scale sequences 1 to 4 of the Vergons section. Interpretations on climate evolution are proposed (on the right side).

relatively low. In part 2, kaolinite and illite contents are relatively high around small-scale SBs and MFs while chlorite content is low. In part 3, illite content is relatively high around small-scale SBs and low around small-scale MFs. Additionally, the highest contents in chlorite mainly occur in the highstand deposits (45%) and in the transgressive deposits (30%) of the small-scale sequences. The K/I ratio presents the highest values around the three small-scale (100-kyr) maximum floodings (Fig. 8.9). In the basin, an increase of kaolinite is commonly related to a transgression and a high value to a maximum flooding (JAN DU CHÊNE et al. 1993; DECONINCK 1993; PITTET 1996).

The quartz content and the sequence stratigraphy do not show any clear correlation. As for the Balingen-Tieringen section, the IC values of 0.4°-0.9° Δ20 CuKα) indicate poorly crystallized illites, although burial depth was relatively high (>2000 m; DECONINCK & DEBRABANT 1985).

## Comparison from shallow platform to basin: discussion

A similar medium-scale evolution of the kaolinite content and the kaolinite/illite ratio is noticed between

the Swiss and Swabian Jura sections. Low values are observed at the base of the sections, around the medium-scale SB Ox6+ and are followed by an increase of values. This trend is in agreement with the one of the Pichoux section, which shows an increase of the kaolinite content and K/I ratio at the base of the Hauptmumienbank Member (Fig. 8.1) around medium-scale SB Ox6+. Generally, high K/I values indicate more humid conditions. However, the change from siliciclastic-dominated to more carbonate facies at the base of the Hauptmumienbank/Steinebach members points to inverse conditions. Consequently, the obtained trends in the shallow and deep platform sections suggest that the distribution of clay minerals is not only controlled by climate changes but also by relative sea-level changes and platform morphology. Indeed, the base of the Hauptmumienbank/Steinebach members corresponds to a major transgressive surface, and facies distribution was influenced by the platform morphology, controlled by local/regional tectonics.

In addition, high kaolinite values and high K/I are recorded in the top part of the sections, around the medium-scale MF Ox6+. This level coincides with the marl or marly limestone interval, correlated over long distances, within the massive Hauptmumienbank limestones (cf. Chap. 5). It can be interpreted as a period of more humid conditions or more remobilization of kaolinite during the maximum flooding (Figs 8.11 and 8.6). This study allows to discuss the link between the kaolinite distribution and sea level (Fig. 8.6).

In the basin section, the trends of kaolinite content and K/I ratio are relatively similar to those of the other

<figure><img src="image_110.png" /></figure>

**Fig. 8.10** - Schematic profile showing the differential settling of clay minerals from shallow platform to basin calculated for the first small-scale sequence. Relative abundance averaged over the whole small-scale sequence. For discussion refer to text.
1 Savagnières

a b c

<figure><img src="image_111.png" /></figure>

2 Balingen-Tieringen

2

<figure><img src="image_112.png" /></figure>

3 Vergons

c

<figure><img src="image_113.png" /></figure>

<figure><img src="image_114.png" /></figure>

<figure><img src="image_115.png" /></figure>

a Medium-scale sequences (400 kyr)

b Small-scale sequences (100 kyr)

c Elementary sequences (20 kyr)

<figure><img src="image_116.png" /></figure>
**Fig. 8.11** - (facing page) Correlation of the K/I ratios from the shallow platform to the basin. A generally higher K/I ratio is observed at the medium-scale MF Ox6+.

sections around medium-scale SB Ox6+ (low values) and medium-scale MF Ox6+ (high values). However, because of diagenetic transformations of clay minerals (cf. above), palaeoclimate interpretations are not reliable.

The comparison of the coeval clay-mineral assemblages from platform to basin shows that kaolinite- and illite-dominated assemblages on the platform evolve into illite- and chlorite-dominated assemblages in the basin (Fig. 8.10). One explanation for this trend is differential settling of clay particles due to the platform morphology (i.e., epeiric platform; presence of a barrier or a reef that separated the platform from the basin; GODET et al. submitted). Moreover, this process probably was enhanced by relative sea-level changes. For example, during sea-level rise, platform area increases and the distance between the continent, where kaolinite is formed, and the basin becomes more important. Hence, this platform configuration during periods of high sea level favours trapping of kaolinite in the proximal part of the platform rather than in hemipelagic environments (Fig. 8.6).

At the scale of 100- and 20-kyr depositional sequences, the interpretation of clay-mineral distribution is more difficult. Maxima and minima of the different clay minerals occur differently between the sections and around the SBs and MFs (Fig. 8.8). For the shallow platform section, these discrepancies probably result from the complex dynamics of mixed carbonate-siliciclastic platform systems (e.g., platform morphology, facies mosaics). For the deep platform and basin sections, the potential diagenetic transformation of kaolinite in seawater (REX & MARTIN 1966), the lag time through remobilization, or the effects of burial diagenesis (CHAMLEY 1989) make the interpretation difficult.

The analysis of additional sections with a higher sampling density, especially towards the hinterland, would be helpful in order to clearly identify the role of climate, sea level, and regional/local tectonics as well as local factors in the clay-mineral distribution.

## 8.3 CARBON AND OXYGEN ISOTOPES

### 8.3.1 Introduction

Carbon and oxygen isotopes are commonly used for palaeoceanographic and palaeoclimatic interpretations (e.g., WEISSERT & ERBA 2004; WIERZBOWSKI 2002; MALCHUS & STEUBER 2002; BUONOCUNTO et al. 2002). Stable isotope data viewed in a stratigraphic context may also give useful information for interpreting depositional and diagenetic histories of limestones (e.g., JOACHIMSKI 1994; VINCENT 2001; VINCENT et al. 2004; PEARCE et al. 2005).

In shallow-marine carbonates, carbon and oxygen stable isotopes have been extensively studied (e.g., SCHOLLE & ARTHUR 1980; ALLAN & MATTHEWS 1982; JOACHIMSKI 1991, 1994; JIMENEZ DE CISNEROS & VERA 1993; VEIZER et al. 1999; RASSER & FENNINGER 2002; BARTOLINI et al. 2003; PARENTE et al. 2007). However, the complexity of sedimentological and early diagenetic phenomena in shallow environments complicates the interpretation of the isotopic record. Controlling factors include fractionation by the carbonate-producing organisms, porosity and permeability of the sediment, type and extent of vegetation cover, duration of subaerial exposure, and diagenetic alteration (JOACHIMSKI 1994).

Correlated on a global scale, $\delta^{13}C$ variations have been explained mainly in terms of changes in the rate of production and accumulation of organic and inorganic carbon (WEISSERT & MOHR 1996). Generally, positive excursions in $\delta^{13}C_{carb}$ are attributed to enhanced burial of organic carbon in sediments, leading to depletion of $^{12}C$ in the ocean carbon reservoir (e.g., KUMP & ARTHUR 1999). The Late Jurassic carbon isotope stratigraphy is marked by a positive carbon excursion of up to 3‰, dated of the early Transversarium zone in the Middle Oxfordian (e.g., BARTOLINI et al. 1996) and a distinct short-lived negative excursion in the middle Transversarium zone (PADDEN et al. 2001; Fig. 8.1). For the Middle Oxfordian positive excursion, LOUIS-SCHMID (2006) and LOUIS-SCHMID et al. (2007) propose an ocean-intrinsic process triggered by plate tectonic movements. This increased $pCO_2$ and thus modified the carbon cycle. Associated were a sea-level rise and a decrease in latitudinal temperature gradient caused by the opening of the North Atlantic. The short-lived negative excursion has been interpreted as reflecting
a sudden release of methane hydrate along the continental margins (PADDEN et al. 2001).

Platform carbonates are frequently exposed to meteoric fluids that may alter their original δ¹³C as well as δ¹⁸O (e.g., DICKSON & COLEMAN 1980; ALLAN & MATTHEWS 1982; LOHMANN 1988; MARSHALL 1992; JOACHIMSKI 1994). Carbonates, which have undergone subaerial diagenesis, may be depleted in δ¹³C due to the influence of organic ¹²C (soil gas). Below the subaerial exposure surface, there commonly is a δ¹³C decrease (ALLAN & MATTHEWS 1977; VIDETICH & MATTHEWS 1980). The intensity of the δ¹³C isotopic signature thus is a function of the subaerial exposure time and degree of pedogenetic alteration (JOACHIMSKI 1994). Moreover, it has been shown that oxidation of organic matter, photosynthesis, and respiration can cause semi-restricted water masses on carbonate platforms to develop δ¹³C values significantly different from those of the open ocean (PATTERSON & WALTER 1994; IMMENHAUSER et al. 2003).

δ¹⁸O values depend on temperature, salinity, taxon-specific fractionation, and diagenetic alteration (e.g., PODLAHA et al. 1998; ROSALES et al. 2001). In carbonate sediments, depleted δ¹⁸O either results from an increase in temperature or reflects a salinity decrease due to the interaction with isotopically depleted meteoric subsurface waters (HUDSON 1977; LOHMANN 1988). The oxygen-isotope record is commonly used for the calculation of seawater palaeotemperatures (e.g., DITCHFIELD 1997; MALCHUS & STEUBER 2002; WEISSERT & ERBA 2004). FRAKES et al. (1992) indicate up to 27°C for the Late Oxfordian ocean surface temperatures according to oxygen isotopes measured on planktonic foraminifera and belemnites. PLUNKETT (1997) calculated palaeotemperatures of 26-27°C from (modified whole rock) Middle Oxfordian samples from the Swiss Jura platform. Several studies show a strong temperature increase during the Oxfordian (e.g., FRITZ 1965; HOFFMAN et al. 1991; PODLAHA et al. 1998; ABBINK et al. 2001; POULSEN & RIDING 2003; Fig. 8.1). This change coincides with a southwestward drift of the West European crustal plate, but a causal relationship remains to be demonstrated (MALCHUS & STEUBER 2002).

Carbonates may also become enriched in ¹⁸O during subaerial exposure due to the removal of ¹⁶O by surface evaporation at the sediment-air interface (ALLAN & MATTHEWS 1982). Accordingly, relatively low δ¹⁸O values (meteoric diagenesis) as well as high δ¹⁸O values (evaporation) in a carbonate succession

may point to exposure and are commonly related to sequence boundaries (e.g., PASQUIER 1995; HILLGÄRTNER 1999).

Strong variations in δ¹³C associated with low variations in δ¹⁸O are typical for meteoric diagenesis, whereas a parallel evolution of δ¹³C and δ¹⁸O curves indicates the zone of mixed, meteoric-marine diagenesis (ALLAN & MATTHEWS 1982; PLUNKETT 1997).

## 8.3.2 Material and analytical methods

For the stable isotope analysis, 77 samples from the Vorbourg, Savagnières, and Latrecey-Ormoy sections have been used (Annex 7). The analyses were performed on whole-rock samples, preferentially on micrite (homogenization of the signal) but also on grainstones. The measurements were carried out in the stable isotope laboratory of the University of Lausanne (Switzerland). The isotopes were measured with a FINNIGAN DELTA PLUS XL mass spectrometer with an attached GASBENCH II and a PAL autosampler. Rock powders were reacted with 100 % phosphoric acid at 90° C. The produced CO₂ was introduced into the GASBENCH II, which separates CO₂ from water and other possible impurities. The treated gas was then sent in a continuous flow mode into the mass spectrometer. During the continuous flow, 10 measurements are made for each sample. The obtained raw results are regularly corrected by comparison to an internal standard (Carrara Marble: 2.05 ‰ δ¹³C and -1.75 ‰ δ¹⁸O) in between two samples. The internal standard is calibrated to the international standard NBS19 (1.95 ‰ δ¹³C and -2.20 ‰ δ¹⁸O). The mean value of the 10 corrected raw results is calculated and represents the C and O isotope values for a given sample. The error margin includes the measurement and reproducibility error and is of ± 0.5 ‰.

## 8.3.3 Results and interpretations

In this work, the carbon and oxygen isotopic records from two shallow platform sections and one deep platform section are investigated at the scale of

**Fig. 8.12** - (facing page) Trends of stables isotopes (δ¹⁸O and δ¹³C) in the transgressive part of medium-scale sequence 1 of the Latrecey (Haute-Marne, deep platform), Vorbourg, and Savagnières (Swiss Jura, shallow platform) sections. Cross-plot of the δ¹⁸O and δ¹³C values.
Vorbourg

<figure><img src="image_117.png" /></figure>

Savagnières

<figure><img src="image_118.png" /></figure>

Sequence stratigraphy

a Medium-scale sequences (400 kyr)

b Small-scale sequences (100 kyr)

c Elementary sequences (20 kyr)
<figure><img src="image_119.png" /><figcaption>**Fig. 8.13** - Trends of stable isotopes (δ¹⁸O and δ¹³C) of the three studied sections placed in the high-resolution time-space diagram.</figcaption></figure>

elementary sequences (20 kyr). Due to the absence of features indicating emersion or meteoric diagenesis in the studied sections (cf. Chap. 4), the measured values probably reflect the original sea-water composition.

### Swiss Jura (shallow platform)

In the Swiss Jura sections, the δ¹³C and δ¹⁸O curves are relatively steady with only minor fluctuations (Fig. 8.12). The δ¹³C values vary between -0.5 and +2 ‰ (Δ 2.5‰) in both studied sections, while the δ¹⁸O values vary between -5 and -4 ‰ (Δ 1‰) in the Savagnières

section and between -5.5 and -3.5 ‰ (Δ 2‰) in the Vorbourg section. The comparison of δ¹⁸O and δ¹³C values shows no significant correlation and hence suggests the preservation of the original isotopic signal (Fig. 8.12). In fact, the measured values fall well into the range indicated by HUDSON (1977) for average marine limestones.

Compared with the values of JOACHIMSKI (1994), the δ¹³C values of the Vorbourg section may be associated to more open marine environments than those in the Savagnières section.
Haute-Marne (deep platform)

In the Haute-Marne section (Latrecey), isotopic values are very homogeneous. The δ¹³C values vary between +2.0 and +2.5 ‰ (Δ 0.5‰) and the δ¹⁸O values vary between -4.5 and -3.5 ‰ (Δ 1‰). No correlation between the δ¹³C and δ¹⁸O values exists, suggesting an original signal. The steady curves are mainly explained by the equilibration of the carbonates in an open, water-buffered oxygen system (ALLAN & MATTHEWS 1982), in agreement with the palaeogeographic position on the margin of the Paris Basin.

In the studied sections, no trends have been identified within the 100- and 20-kyr sequences. The minor isotopic fluctuations of ± 1‰, particularly around the small-scale SB 2 and the medium-scale MF Ox6+, probably reflect variations of local environmental conditions (e.g., salinity, temperature, trophic level). This is related to the variability of depositional environments, evidenced by the facies mosaics on the Swiss Jura carbonate platform (cf. Figs 7.3 and 7.4; STRASSER & VÉDRINE in press). The slight variations of less than 0.5‰ at the scale of the 20-kyr elementary sequences are interpreted as noise and cannot be related to environmental changes.

8.4 CONCLUSIONS

On the northern Tethys margin, the studied interval corresponds to less humid conditions with a global cooling and concurs with a long-term sea-level rise (Fig. 8.1). This study on platform and basin deposits shows the large potential of a high-resolution analysis.

Clay minerals show similar medium-scale trends between the sections from shallow platform to basin particularly around the medium-scale SB Ox6+and MF Ox6+. These patterns have not only been controlled by climate changes but also by relative sea-level changes and platform morphology. For example, at the medium-scale MF Ox6+, a remobilization of clay particles by the long-term transgression may be involved (cf. Fig. 8.6). At the scale of 100- and 20-kyr sequences, the interpretation of clay-mineral distribution is more difficult. The discrepancies probably result from the complex dynamics on the shallow platform or diagenetic transformations in seawater or during burial diagenesis.

The carbon and oxygen isotopic fluctuations are minor (±1‰) and probably reflect variations of local environmental conditions (e.g., salinity, temperature, trophic level). The oxygen isotopes do not allow to demonstrate temperature changes.

Additional sections and an increased sampling density would potentially improve the interpretation of both clay minerals and stable isotopes. This would help to answer questions on the composition of source areas, clay-mineral distribution and/or diagenetic transformation.

This study shows that clay-mineral assemblages and stable isotopes can give palaeoenvironmental information on the scale of 400 kyr and 100 kyr but that at higher time resolution is difficult to obtain. This is probably due to time-averaging by sedimentological and/or diagenetic processes.

***
# 9 - CONTROLLING FACTORS OF MIXED CARBONATE-SILICICLASTIC PLATFORMS

This chapter aims to identify the controlling factors that act on carbonate platform systems. From sedimentological, geochemical, and mineralogical data obtained in this work, the main factors that controlled the hydrodynamics, the carbonate ecosystem, and the formation of depositional sequences on the Swiss Jura shallow platform are discussed.

## 9.1 THE INTERPLAY OF TECTONICS, EUSTASY, AND CLIMATE

The sedimentary record results from a large number of local, regional, and global parameters, which are interrelated to different degrees (Fig. 9.1). In the studied mixed carbonate-siliciclastic systems, the sedimentary record results from relative sea-level fluctuations, carbonate production and accumulation, and terrigenous input. On shallow platforms, the potential of carbonate production is controlled by variations of factors intrinsic to the ecosystem (e.g., water depth, turbidity, hydrodynamism, nutrient availability, oxygenation, salinity, and temperature), which depend of sea-level variations, tectonics, and climate.

### 9.1.1 Tectonics

Tectonic activity is mainly linked to plate tectonics (VAIL et al. 1991). Phases of increased tectonic activity on a global scale (duration corresponding to 1st-order of VAIL et al. 1991) are responsible for widespread orogeny and the formation of sedimentary basins. Regional tectonics (2nd-order of VAIL et al. 1991) control variations in the subsidence rate of sedimentary basins. Local tectonics induce the small-scale structural changes.

Plate tectonics influence eustatic sea-level through variations of the volume of basins. Global climate is influenced by increased CO₂ released from ocean-floor spreading and/or volcanism or by decreased CO₂ due to continental weathering of silicates. Regional climate is furthermore controlled by the position of the continents and orogeny. Tectonics also influence platform morphology and the distribution of depositional environments through bathymetric and hydrodynamic variations. Tectonics together with climate control the production, transport, and distribution of terrigenous sediments and nutrients onto the shallow platform (MATTHEWS & PERLMUTTER 1994; VAIL et al. 1991).

### 9.1.2 Eustasy

Long-term eustatic sea-level variations (tectono-eustasy, 1st- and 2nd-order of VAIL et al. 1991) are linked to variations of the volume of oceans, which mainly depend of mid-oceanic ridge activity. The volume of oceans are also modified by the collision of continents, subduction, sub-marine magmatism, and sedimentary filling (VAIL et al. 1991). Short-term eustatic sea level fluctuations (glacio-eustasy, from 3rd- to 6th-order of VAIL et al. 1991) are linked to orbital parameters. Precession, obliquity and eccentricity variations create climate changes, which act on the volume of water masses, and consequently, on sea level. Sea-level changes were created by glaciation/deglaciation cycles of polar ice-caps and mountain glaciers, by thermal expansion and retraction of the uppermost layer of the ocean water, by thermally induced volume changes in deep-water circulation, and/or by water retention and release in lakes and aquifers (cf. Chap. 6.1).
<figure><img src="image_120.png" /><figcaption>**Fig. 9.1** - Control of orbital parameters and plate tectonics on climate, eustatic sea level, and regional tectonics modifying the ecological parameters of depositional environments. The three main controlling factors are in bold. The controlling factors evidenced by the facies, oncoid, and benthic foraminifer distribution (cf. Chap. 7) are in grey. Scale, amplitudes, and frequencies may strongly vary. Refer to text for discussion.</figcaption></figure>

Sea-level variations directly influence bathymetry, hydrodynamism, nutrient availability, and temperature (e.g., HALLOCK & SCHLAGER 1986; JAMES 1997). A relative sea-level fall accentuates the platform morphology and favours the formation of restricted lagoons. Variations of base level are linked to the relative sea-level fluctuations and thus control the continental erosion and terrigenous input. Long-term eustatic sea-level variations are partly responsible of the geometry of system tracts (sensu VAIL et al. 1991; progradation, retrogradation, aggradation) and of the formation of long-term depositional sequences while short-term variations preferentially influence the formation of elementary, small-scale, and medium-scale sequences.

### 9.1.3 Climate

Global climate partly depends on orbital variations, which affect the insolation and the atmospheric circulation (MÖRNER 1994; MATTHEWS & PERLMUTTER 1994). Global climate is also controlled by greenhouse gases. Increased activity of the mid-oceanic ridges engenders an increase of atmospheric $pCO_2$ followed by an increase of temperature and rainfall (WEISSERT & MOHR 1996). Continental weathering of silicate minerals leads to a drop in $pCO_2$ and thus to a cooling. Gas hydrate dissociation releases $CH_4$, which acts as a greenhouse gas. Land-ocean distribution and elevation of mountain chains control the regional climate.
<figure><img src="image_121.png" /><figcaption>**Fig. 9.2** - Main controlling factors acting on the siliciclastic distribution. Modified from HILLGÄRTNER (1999).</figcaption></figure>

Rainfall and temperature strongly influence the ecosystems of shallow carbonate platforms and hence the potential of carbonate production. High precipitation favours the terrigenous sediment transport and distribution. The terrigenous sediment distribution thus may be related to orbital cycles defining the humidity conditions. These vary according to palaeolatitude and/or palaeoposition of the continents (MATTHEWS & PERLMUTTER 1994).

Continent-derived nutrients stimulate primary productivity, which may lead to eutrophication and turbidity increase on the shallow platform (HALLOCK & SCHLAGER 1986; DUPRAZ & STRASSER 1999). After abundant rainfalls, salinity decreases. Eutrophication, high turbidity, and low salinity are responsible for a decrease of the potential of carbonate production. Similarly, high aridity engenders, by evaporation, an increase of salinity, and hence a decrease of the carbonate production (HALLOCK & SCHLAGER 1986). Eustasy and climate are linked, through complex processes and feedback mechanisms, to variations of the orbital parameters (Fig. 9.1).

## 9.2 FACTORS CONTROLLING THE SEDIMENTATION OF THE SWISS JURA PLATFORM

Based on the study of the Swiss Jura sections (cf. Chap. 7), the facies distribution on the Swiss Jura platform is best described by a mosaic of juxtaposed depositional environments. Platform morphology, related to local and regional tectonics, was an important factor influencing the facies distribution (Fig. 9.1). The oncoid and benthic foraminifer distribution illustrates the influence of water energy and water depth (cf. Chap. 7 and Fig. 7.8), linked to relative sea-level fluctuations as well as the carbonate production

and terrigenous input controlled by climate (humidity and temperature) and platform morphology (Figs 7.9 and 9.1).

The siliciclastic distribution in the Hauptmumienbank Member is variable. Trends in abundance of quartz and clays controlled by global factors (i.e., humidity changes, tectonics) and by local factors (i.e., platform morphology, water energy; Fig. 9.2) are observed. From the study of the marly Röschenz Member deposits, PITTET (1996) proposed more humid conditions during low sea-level periods at the scale of 100-kyr sequences. This can be explained by the palaeolatitudinal variations of the atmospheric cells (MATTHEWS & PERLMUTTER 1994; PITTET & STRASSER 1998b; PITTET et al. 2000; STRASSER et al. 2005; MARTIN-GARIN 2005). However, in the overly-ing Hauptmumienbank Member, siliciclastics occur not only around sequence boundaries but also around maximum floodings of the 100-kyr sequences. This is related to the long-term transgressive trend initi-ated at the base of the Hauptmumienbank Member, which modifies the continental erosion and allows the remobilization of siliciclastics formerly stored in the proximal part of the platform. The latter process is also proposed for the clay-mineral distribution (i.e., kaoli-nite; cf. Chap. 8).

The observed hierarchical stacking pattern of the depositional sequences is interpreted to be related to eustatic sea-level fluctuations driven by the Earth's orbital perturbation in the Milankovitch frequency band (cf. Chap. 3 and 6). The good correlation across the platform of small-scale and medium-scale sequences implies a common external controlling factor such as eustatic sea-level changes. However, the variability of facies and discontinuity surfaces (cf. Chap. 7) points to additional local controlling factors such as platform morphology and terrigenous input, acting on the carbonate productivity and accumulation.
The complex interaction of factors is evidenced
particularly at the scale of elementary sequences.
These sequences present a high facies variability (cf.
Chaps 3 and 7). They are related to the precession

cycles, which increase the seasonal contrast. These
cycles acted on variations in precipitation and run-off
on a regional scale, dependent on latitude (WATERHOUSE
1999; MALLINSON et al. 2003).

***
# 10 - CONCLUSIONS AND OUTLOOK

Based on a well-established bio- and sequence-stratigraphic framework, a high-resolution sedimentological study has been performed on a narrow time window in the Late Oxfordian (early Bimammatum zone). Seven shallow platform sections (Swiss Jura, Lorraine), two deep platform sections (Haute-Marne, Swabian Jura), and one basin section (SE France) have been logged and analysed in great detail. Based on a facies and microfacies analysis, the deposits are interpreted in terms of palaeoenvironments and sequence- and cyclostratigraphy.

The main results of this study are:

* Due to their high sensitivity to sea-level and climate changes, the shallow carbonate platform deposits from the Swiss Jura sections provide a high-resolution record of palaeoenvironmental changes occurring during a marine transgression, which is a favourable period for sediment preservation due to a high accommodation gain.

* The detailed facies and microfacies analysis allows to propose depositional models for the Swiss Jura platform and the other studied areas.

* The high-resolution sequence- and cyclostratigraphic analysis allows to define hierarchically stacked depositional sequences: medium-scale, small-scale, and elementary sequences, formed through orbitally controlled sea-level changes with periodicities of 400, 100, and 20 kyr, respectively.

* In this study, the first half of a medium-scale sequence is examined. It corresponds to two small-scale sequences, each composed of five elementary sequences. In the shallow platform sections, an

elementary sequence generally consists of one to four beds including more or less developed marl intervals. In the deep platform sections, an elementary sequence generally consists of one or two limestone beds with a more or less developed marl interval. In the basin section, an elementary sequence is commonly defined by one marl-limestone couplet.

* This study provides a "best-fit solution" of high-resolution sequence-stratigraphical correlation from shallow-water platform (Jura, Lorraine platform) to deeper platform (Haute-Marne, Swabian Jura) and to basin sections (SE France). The good correlation of depositional sequences over long distances between the seven shallow platform sections and the similar number of elementary sequences in all sections are valuable arguments that allocyclic processes (i.e., eustatic sea-level fluctuations) must have been involved in the formation of these depositional sequences.

* A pre-existing morphology determined the geometry of the coast line and the distribution of depocenters on the platform. The relief also modified the hydrodynamic setting (current and/or wave directions and intensity), the pattern of sediment production and deposition, and the ecological conditions (e.g., water depth, salinity, turbidity, nutrient distribution). Carbonate production and accumulation, i.e. the preferential growth of reefs and the initiation of ooid shoals on tectonic highs further enhanced the relief on the platform. Climate changes also played a role in the facies distribution on the shallow platform. The terrigenous input, controlled by humidity changes in the hinterland and by sea level influenced carbonate production.
* Thickness variations at the scale of small-scale and elementary sequences reveals variations in the sediment accumulation rates. This is interpreted as resulting mainly from differential subsidence due to the activity of tectonic blocks. The relief created by tectonics therefore contributed significantly to the general facies distribution.

* The vertical and lateral distribution of facies, oncoids, and benthic foraminifera is investigated within elementary and small-scale sequences related to 20-kyr precession cycle and to the 100-kyr short-eccentricity cycle, respectively. When comparing the facies of time-equivalent small-scale and elementary sequences between the Swiss Jura sections, significant lateral facies changes appear (facies mosaic).

* The distribution of siliciclastics in the shallow platform sections shows a similar pattern with more or less developed marl intervals around the medium-scale SB Ox6+ and MF Ox6+ and the small-scale SB 2. Discrepancies in this distribution is explained by localized depressions, which were created by differential subsidence and served as channels, and by the remobilization of siliciclastics formerly stored in more proximal areas of the platform.

* Four types of oncoids are defined in the Hauptmumienbank Member from the Swiss Jura based on surface morphology, configuration and composition of the cortex, and the encasing sediment. Type 1 is sub-elliptical, micritic, and has a homogeneous cortex; Type 2 is elliptical, smooth, and displays micritic laminations with serpulid worm tubes and *Bullopora*; Type 3 is elliptical, lobate, and shows alternating laminae of *Bacinella-Lithocodium* and micrite; Type 4 is lobate and consists mainly of a *Bacinella-Lithocodium* meshwork.

* The stratigraphic and spatial distribution of oncoids in the six studied sections shows that each oncoid type can be attributed to a specific depositional environment. Type 1 oncoids characterize moderate-energy conditions in protected lagoons; Type 2 oncoids are related to higher energy conditions in open lagoons; Type 3 and 4 oncoids dominate in low- to very low-energy settings in open lagoons.

* Oncoid distribution is linked to relative sea-level changes. Micrite-dominated oncoids (types 1 and 2) are preferentially found around sequence boundaries and in transgressive deposits, whereas oncoids containing *Bacinella-Lithocodium* (types 3 and 4) are found rather around maximum floodings and in highstand deposits. This implies that changes in water energy and water depth were direct controlling factors. The heterogeneous oncoid distribution through space and time implies that the variable morphology of the Swiss Jura platform was also a controlling factor. Type 1 and 2 oncoids grew under meso- to eutrophic conditions, while type 3 and 4 oncoids, rich in light-dependent microencrusters (i.e., *Bacinella* and *Lithocodium*), flourished in oligotrophic waters. This study demonstrates that oncoids are valuable proxies for high-resolution palaeoenvironmental and palaeoecological studies.

* The distribution of the total abundance of benthic foraminifera shows a relatively patchy distribution without any "proximal-distal" trend. Contrary to the agglutinated foraminifera, the porcelaneous and hyaline foraminifera show a correlation with the sequence-stratigraphic evolution. Miliolids are preferentially found in the highstand deposits of some elementary and/or small-scale sequences. The hyaline foraminifera, rare in the studied sections, are preferentially found in the transgressive deposits of some elementary sequences. Consequently, the benthic foraminifera are indirectly linked to relative sea-level (accommodation changes) and to climate changes (sediment and nutrient input).

* The distribution of *Mohlerina basiliensis* and *Bacinella-Lithocodium* oncoids shows a strong correlation. *M. basiliensis* preferentially occurs when *Bacinella-Lithocodium* oncoids are present. The co-occurrence of *Bacinella-Lithocodium* oncoids and *M. basiliensis* suggests that they required similar ecological conditions.

* The successively later occurrence of *M. basiliensis* from "distal" to "proximal" sections illustrates the stepwise flooding of the platform and thus implies a dependence on normal-marine conditions.

* Clay mineral analyses have been performed on three sections from shallow platform, deep
platform, and basin. The medium-scale trends of the kaolinite content and K/I ratio between all three sections are comparable, particularly around the medium-scale SB Ox6+ and MF Ox6+. These patterns have been controlled not only by climate changes but also by relative sea-level changes and platform morphology.

* Performed on sections from the Swiss Jura and Haute-Marne, the carbon and oxygen isotope analyses show minor fluctuations (±1‰), particularly around the small-scale SB 2 and the medium-scale MF Ox6+ that probably reflect variations of local environmental conditions (e.g., salinity, temperature, trophic level). The oxygen isotopes do not allow to demonstrate temperature changes at the scale of small-scale or elementary sequences.

**This study has led to several interesting results that may be used as starting points for further research:**

The sequence- and cyclostratigraphic time frame obtained in this study may be used for performing facies analysis within an elementary sequence in order to interpret palaeoecological changes with a resolution higher than 20 kyr.

The methodology coupling high-resolution facies analysis with sequence- and cyclostratigraphic stratigraphic interpretation is an excellent tool for monitoring high-frequency palaeoenvironmental changes. This methodology may thus be applied to platform deposits from other stratigraphic periods and palaeogeographic areas, if a bio- and sequence-stratigraphic framework is already well defined.

The vertical and lateral distribution of marine oncoids may be investigated also in other sections from other palaeogeographic areas, following the same methodology than in the present study. For example, oncoids commonly occur in the Pagnysur-Meuse section (Lorraine), in the two small-scale sequences following the interval of this study. It would be interesting to define the different types and to relate them to specific depositional environments as well as sequential positions. Similar studies could also be performed on deposits containing lacustrine and/or fluvial oncoids. In Switzerland and neighboring countries, some Tertiary deposits are rich in non-marine oncoids (e.g., Upper Palaeocene-Eocene, eastern Ebro Basin, NE Spain, ZAMARREÑO et al. 1997; Oligocene, Rhineland, W Germany, LEINFELDER & HARTKOPF-FRÖDER 1990; Miocene, Digne foreland basin, SE France, BAUER 2006). Their distribution according to the sequence-stratigraphic framework and their micromorphology could be analysed.

The identification and quantification of additional genera and/or species of benthic foraminifera, particularly among the agglutinated foraminifera (Textularids), could give more precise indications on their palaeoecology.

More research on the benthic foraminifer *Mohlerina basiliensis* would be interesting for solving the problem of taxonomy. The wall microstructure of *M. basiliensis* (Middle-Late Bathonian to Early Valanginian) presents similarities to the one of Fusulinina. Most fusulinides became extinct in the Late Permian. *Tetrataxis* is one of the genera persisting until the Late Triassic (GROVES & ALTINER 2005). Consequently, the comparison of wall microstructures between *M. basiliensis* and *Tetrataxis* sp., e.g., with SEM analyses, could allow to link *Mohlerina basiliensis* to the Fusulinina suborder.

* * *
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