Contrasting Styles of Siliciclastic Tidal Deposits in a Developing Thrust-Sheet-Top Basins – The Lower Eocene of the Central Pyrenees (Spain)

  • A. W. Martinius


Lower Eocene tidal deposits in the Tremp-Graus-Ager Basin in the southern Pyrenees (Spain) are well-developed and include typical examples of tidal bars, compound tidal dunes, tidal bundles and tide-dominated back-barrier lagoons as well as tidally-influenced fluvial systems. They occurred in a relatively narrow (up to 60 km) and long (up to 200 km in total) semi-enclosed sea which had an open connection to the Atlantic ocean in the west. Two groups of tidal deposits are recog­nised related to two stages of the obliquely migrating thrust-sheet-top basin which affected the position and relative dimensions of the foredeep and shelf sections. Compound tidal dune fields and large tidal bars developed mainly in the initially underfilled foredeep in relatively deep water (at least up to 40 m) during the Early Ypresian. Favourable conditions existed for basin scale tidal current circulation patterns, locally significantly amplified and modified by pronounced bathymetric variations (related to developing blind thrust related folds and blind ramps) and a variable, and probably distinct, structurally controlled, coastline morphology. Small shoal-water fan deltas and larger Gilbert-type delta(s) and associated tidal bars developed along the basin margins near, often long-lived, sediment entry points. During the Late Ypresian to Late Lutetian the basin shelf area filled-up by a rapidly developing axial east to west prograding alluvial to deltaic system. This reduced tidal amplification in the basin and shallow-water tidalites developed only in a narrow (approximately 10 km) zone, located above an oblique lateral ramp system, including the in-shore parts of the delta distributaries and the subaqueous part of the, partly barred, delta top.


Tidal Current Sandstone Unit Sandstone Body Delta Lobe Blind Thrust 
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18.1 Introduction

The central Pyrenean Eocene thrust-sheet-top basins in Spain have been well known for their siliciclastic tidal deposits since the mid 1960s when first the Dutch (Mey et al. 1968; van Eden 1970), soon followed by the Spanish, Italian and French workers, began to study the well-exposed outcrops in detail. A major step forward was the understanding of tidal bundle successions based on studies of excavations in the modern and historically well-documented Oosterschelde estuary (The Netherlands) during the late 1970s and early 1980s (Nio et al. 1980; Visser 1980; van den Berg 1981, 1982) as well as sedimentological studies in the Wadden Sea (Sha 1990; Oost 1995) that were applied to interpret analogous deposits in the Spanish Pyrenees (Donselaar and Nio 1982; Yang and Nio 1985, 1989). The tidal nature of the deposits was recognized early on, but the larger environmental setting remains the subject of discussion and re-evaluation.

The tide-influenced and tide–dominated successions discussed in this chapter form part of a number of formations of Ypresian and Lutetian age covering a continuum of almost 16 Ma. Tidal deposits of Maastrichtian age in the Aren Formation (Tremp Group; Nagtegaal et al. 1983; De Boer 1985, and a wealth of recent but unpublished data collected by Spanish workers) and of Upper Lutetian age in the Sobrarbe Formation (Campodarbe Group; Hall 1997; Dreyer et al. 1999) are also present but are not included in this chapter.

This chapter discusses examples of two different styles of tidal amplification that existed during the Ypresian and Lutetian in the central Pyrenean thrust-sheet-top Tremp-Graus-Ager Basin. The first style was controlled by the underfilled foredeep stage of basin development, the second style by the overfilled shelf stage of basin development. The two styles were closely related to significant differences in basin configuration characteristics and controlled by thrust-sheet-top basin (cf. DeCelles and Giles 1996) development and associated changes in basin paleogeography, morphology and dimensions. Therefore, usage of the terms foredeep and shelf in this chapter is strictly related to the particular development of the thrust-sheet-top basin as outlined hereafter. In addition, eustatic sea-level fluctuations influenced the degree of tidal amplification. A summary of the characteristics of the two styles is given in Table 18.1 and discussed hereafter.
Table 18.1

Summary of the characteristics of the underfilled foredeep and overfilled shelf stage of the Tremp-Graus-Ager Basin. See text for discussion


Underfilled foredeep

Overfilled shelf

Chronostratigraphic period

Early Ypresian (55.5 Ma to 51.5 Ma)

Late Ypresian to Lutetian (51.5 Ma to 43 Ma)

Basin shape

Narrow and elongate, closed in the E, open to the Atlantic Ocean in W

Narrow and elongate, closed in the E, open to the Atlantic Ocean in W

Basin width (total marine section)

50 km

40 km

Basin length (total marine section)

200 km

150 km

Estimated water depth in present-dayTremp-Graus and Ager Basin area

Up to 60 m on average

0 to 20 m on average

Sea-floor morphology

Shelf floor typified by gentle ridges and swells above blind thrust-related ramps and faults. Indented northern basin margin. Significant shelf area present. Slope located in present-day Ainsa Basin area.

Almost no shelf present (delta and alluvial plain). Present-day Ainsa Basin locus of structurally fixed, narrow and steep shelf platform-to-slope transition, and slope.

Length of (marine) shelf

Approximately 50 km

Approximately 20 km decreasing to 5 km

Nature of main basin filling processes on shelf

Dominantly shallow marine siliciclastic deposition along basin margin in inner parts, with gradual transition to central basin pelagic deposition and basinmargin fringing carbonate margins in present-day Jaca Basin area

Extensive alluvial and lower delta plain with relatively narrow delta front. Major shelf slope collapse and basin floor turbidites in present-day Ainsa and Jaca Basins with carbonate platform along margin.

Position of tidalites in basin

1) Narrow sea; 2) on platform shelf in front of (detached or attached to) deltas; 3) estuaries and embayments

Subaquous delta top along indented shoreline in inshore part of channels and embayments

Type of tidal deposits

1) Tidal dunes in embayment in front of Gilbert-type delta lobes; 2) offshore compound tidal dune fields on narrow shelf; 3) detached or attached delta-front tidal bars

1) Tide-influenced mouth bars and point bars in meandering channels; 2) interdistributary bay deposits; 3) tidal inlets and back-barrier lagoons


Alveolina Lst (Serraduy Fm), Roda Fm, Baronia Fm and Ametlla Fm

Castigaleu and Montllobar Fms (Lower Montanyana Group) and Capella and Pano Fms (Upper Montanyana Group)

18.2 Geological Framework

18.2.1 Development of Lower Eocene Thrust-Sheet-Top Basins

The Pyrenean Range (Capote et al. 2002) consists of the Aragonese-Catalan Pyrenees in the east and the Basque-Cantabrian Pyrenees in the west. The Aragonese-Catalan Pyrenees are subdivided into the Eastern, Central and West-Central Pyrenees (Fig. 18.1). The focus in this chapter will be exclusively on the thrust-sheet-top basins of the Spanish Central Pyrenees (the south Pyrenean central unit of Séguret 1972). For details about other aspects of structural evolution of the Pyrenees, one is referred to Van der Voo (1969), Puigdefàbregas and Souquet (1986), Galdeano et al. (1989), Malot (1989), Choukroune et al. (1990), Malod and Mauffret (1990), Srivastava et al. (1990), Roest and Srivastava (1991), Olivet (1996), Meigs and Burbank (1997), and Capote et al. (2002).
Fig. 18.1

Structural map of the Aragonese-Catalan Pyrenees (modified after Capote et al. 2002). The present-day Tremp-Graus and Ager Basins are located south of the axial zone in the central and eastern Pyrenees and south of a structural divergence axis in the west-central Pyrenees. The eastern boundary of the Central Pyrenees is formed by the NE-SW oriented oblique Segre transfer zone and the western boundary is formed by the NW-SE oriented oblique Boltaña anticline. The southern margin is formed by the E-W oriented frontal thrust of the Sierras Marginales (SM). B Bóixols thrust sheet, C Cadi thrust sheet, M Montsec thrust sheet, P Pedraforca thrust sheet,SM Sierras Marginales thrust sheet, AB present-day Ainsa Basin, AgB present-day Ager Basin, JB Jaca Basin, TGB present-day Tremp-Graus Basin. Inset 1: geological map of Fig. 18.10; inset 2: geological map of Fig. 18.11b>

An important phase in the development of the Pyrenees started around 84 Ma (Late Santonian) when the Iberian Plate and the African Plate collided, and subduction along the northern plate margin was initiated (Guimerá 1984, 1996). This caused conversion from a foregoing extension phase to compression with a near N-S shortening direction, and produced inversion of Mesozoic extensional rift basins and the rise of the Pyrenees (Beaumont et al. 2000; Capote et al. 2002).

Two main foreland basins, the Aquitanian basin in the north and the Pyrenean Foreland basin to the south, developed from the Late Santonian (84 Ma) to the Miocene in conjunction with the convergence from extension to compression (Cámara and Klimowitz 1985; Puigdefàbregas et al. 1992; Teixell and Muñoz 2000; Beaumont et al. 2000). Four main compressional stages are recognized in the Pyrenean Foreland Basin (Puigdefàbregas et al. 1992) of which Stage III is the focus of this chapter.

Three main thrusts developed during Stage III (Early and Middle Eocene): from north to south the Bóixols, Montsec and Sierras Marginales. They were initiated successively in time but their displacement periods overlapped; they are particularly clear in the ECORS seismic profile (Cámara and Klimowitz 1985; Choukroune et al. 1989, 1990; Figs. 18.1 and 18.2). In the central Pyrenees, the size and shape of the thrust sheets were determined by the inverted Mesozoic fault pattern which controlled the location of oblique (with respect to the thrust motion) and lateral ramps that are defined as dividing two different segments of a thrust belt and consequently the distribution of accommodation space and thrust-sheet-top basin facies (Puigdefàbregas et al. 1992; Vergés 2007).
Fig. 18.2

(a) Partially restored cross-section of the crust through the central Pyrenees at the location of the ECORS profile (Choukroune et al. 1989, 1990) illustrating the tectonic style and the three main imbricated thrust sheets (Bóixols, Montsec and Sierras Marginales). Reference frame holds the European Plate and arrows indicate estimated total convergence. NPF North Pyrenean Fault (Modified after Beaumont et al. (2000)). (b) Enlargement of the boxed area in a illu­strating the position of the present-day Ager and Tremp-Graus Basins

Initiation of the Montsec Thrust occurred at the end of the Paleocene (Cámara and Klimowitz 1985; Puigdefàbregas et al. 1986, 1992; Nijman 1998; Mascle and Puigdefàbregas 1998). The thrust underlies more than 3 km of displaced, mainly Upper Cretaceous limestones and Maastrichtian and Paleocene deposits of the Tremp Group. These were deposited in the authochtonous South Pyrenean Foreland Basin when it was not broken-up and partly displaced. The Montsec Thrust carries a Lower Eocene allochtonous ­thrust-sheet-top basin (or piggyback basin cf. Ori and Friend 1984, see also Ricci Lucchi 1986) named the (present-day Tremp-Graus Basin). The remaining autochtonous part of the Pyrenean Foreland Basin is the present-day Ebro Basin (Fig. 18.1).

As a consequence, the western part of the southward-moving Montsec Thrust developed as a northward curved oblique blind thrust with associated faults (cf. Dunne and Ferrill 1988) and acted as a lateral ramp separating shelf and slope deposition ­during Ypresian sedimentation. The formation of secondary blind thrusts, associated with the Montsec Thrust, is expressed as near-surface subtle, low-amplitude, ­gentle folding in intrabasinal areas (de Boer et al. 1991; López-Blanco et al. 2003; Clevis et al. 2004).

The southward moving Montsec Thrust sheet, inclu­ding the present-day Tremp-Graus Basin, became incorporated into the cover of the next developing thrust, the Sierras Marginales, when this was initiated in the early Lutetian and overthrusted to form several smaller imbricated thrust units and thrust-sheet-top basins; the connected Tremp-Graus, Ager and Ainsa Basins (Mascle and Puigdefàbregas 1998). The rate of thrusting was highest from the end of the Paleocene to the Lutetian (Vergés et al. 1995; Millán et al. 1995) producing a relatively deep and wide basin and the broadest expansion of marine deposits (Puigdefàbregas et al. 1992; Burbank et al. 1992). Syntectonic ­contemporaneous sequences are present in neighboring basins (Pocoví 1978; Martínez-Peña and Pocoví 1988; Vergés et al. 1995; Mascle and Puigdefàbregas 1998). Paleomagnetic data from the present-day Ainsa Basin, which rotated 30° clockwise during Montsec Thrust displacement (Poblet et al. 1998), indicates the rotation of the entire thrust-sheet-top basin (Nijman 1989; Poblet et al. 1998). This complex structural development had a profound effect on basin morpho­logy and sedimentation patterns and has influenced basin-scale tidal current patterns.

18.2.2 Lower Eocene Paleoclimate, Eustatic Sea Level, and Stratigraphy

During the upper lower Ypresian, the upper Ypresian and the Lutetian, the study area was located at a latitude of approximately 35°N and was characterized by a warm and stable, sub-tropical to tropical climate (Early Eocene Climatic Optimum; Zachos et al. 2001) and generally warm and arid to semiarid conditions in the Pyrenees (Haseldonckx 1972; Schmitz and Pujalte 2003; Pearson et al. 2007).

The Ypresian and Lutetian are periods of pronounced long-term tectonic development (see above) controlling local relative sea-level changes, that are likely to have varied spatially and temporally in the Pyrenean foreland Basin. Egger et al. (2009), in contrast, concluded that the effects of Paleogene changes in eustatic sea-level exceeded the effects of regional tectonic activity in the shelves of the European and Adriatic Plates as a result of the much shorter magnitude and time scales over which these processes operate. This is illustrated by the occurrence of a widespread and relative rapid marine transgression during the lowermost Ypresian (Pujalte et al. 2009). However, in general and in comparison Alpine tectonic activity in the Pyreneen orogenic zone was much stronger than elsewhere (Vergés et al. 1995; Meigs and Burbank 1997; Capote et al. 2002).

18.3 Tidalites and Paleogeography, Bathymetry and Fill – Underfilled Foredeep Versus Overfilled Shelf

18.3.1 Underfilled Foredeep

The underfilled foredeep stage (Table 18.1, Fig. 18.3b) is stratigraphically represented by deposits of the Ager Group (Figs. 18.4 and 18.5; see next section) and occurred during the early stage of thrust-sheet-top basin development (Early Ypresian, 55.5 Ma to 51.5 Ma). As discussed hereafter, the associated and characteristic configuration of basin morphology parameters indicate that the basin was within the tidal amplification window (cf. Sztanó and de Boer 1995; Fig. 18.6a, c).
Fig. 18.3

(a) Sketch of the estimated maximum extent of the allo­chtonous south Pyrenean Foreland Basin illustrating the narrow, elongated geomorphology of the semi-enclosed sea encompassing the Jaca and T-G-A Basins (modified after Plaziat 1981). (b) Sketch of the inferred paleogeography and water depths of the T-G-A Basin sea during the underfilled foredeep stage of the Late Ilerdian (approximately 53 Ma) based on the work of the authors referred to in the text. The main siliciclastic depositional systems are indicated

Fig. 18.4

Summary correlation diagram of Ypresian stratigraphy of the present-day Tremp-Graus and Ager Basins with a focus on the Ager Group. The base of the Ypresian is estimated at 55.8 Ma and the top at 48.6 Ma (Luterbacher et al. 2004). The magnetic polarity change from chron C24 to chron C23 occurs at 52.6 Ma, and the change from chron C23 to chron C22 occurs at 50.8 Ma (Luterbacher et al. 2004). The former change coincides with a dating of 52.6 Ma for the Plateau Limestone (López-Blanco et al. 2003). Unpublished data of the top of a Turritella-dominated mudstone interval in the Esdolomada Mbr (which is on average 180 m thick) occurring 29 m above the Plateau Limestone at the stratigraphic level of the El Villar Limestone gives an age of 52.4 Ma which coincides with top P6 and a maximum flooding of the Roda Fm in the Isabéna valley (confirmed by Torricelli et al. 2006). Following this data, the Roda Sst Mbr covers approximately 0.9 Ma and the Esdolomada Mbr covers approximately 1.6 Ma. Note, however, that López-Blanco et al. (2003), using magnetostratigraphic data obtained from the Roda Fm (Bentham and Burbank 1996) and the magnetic polarity scale of Cande and Kent (1995), estimated approximately 400 Ka for the Roda Sst Mbr and approximately 600 Ka for the Esdolomada Mbr. The allostratigraphic subdivision of the Figols Group advocated by Mutti and his co-workers, based on the philosophy that comparable facies throughout the basin are combined into groups or depositional systems (for example, the Figols Group for terrigeneous facies of mainly deltaic character, the Campodarbe-Montañana Group for continental facies and the Hecho Group for basin floor facies), is not incorporated. Note that tidal deposits of Maastrichtian age in the Aren Formation (Tremp Group; Nagtegaal et al. 1983; De Boer 1985 and a wealth of recent but unpublished data collected by Spanish workers) and of Upper Lutetian age in the Sobrarbe Formation (Campodarbe Group; Hall 1997; Dreyer et al. 1999) are also present but not included in this review

Fig. 18.5

Summary correlation diagram of Ypresian and Lutetian stratigraphy of the present-day Tremp-Graus and Ager Basins with a focus on the Montanyana Group (modified after Nijman 1998)

Fig. 18.6

The “tidal amplification window” (Sztanó and de Boer 1995) includes a number of variables (but does not necessarily need to simultaneously honor all of these) such as favourable depth and length for resonant amplification, and depth and width for the development of (an) amphidromic point(s) (Pugh 1996; see also  Chap. 13), and funnelling (cf. Bay of Fundy). Basin-scale tidal amplification thus may happen during specific phases of (i) overall basin development as well as (ii) relative sea-level cycles during one phase of basin development during which the necessary requirements for resonance are met. (a) The average depth of a basin determines the celerity and the length of the propagating tidal wave. For resonant amplification the critical basin length should be an odd multiple of the quarter of the tidal wave length. The solid line indicates the relation between the average water depth and the length of the basin for one times the quarter of the tidal wave length. The two successively steeper dashed lines indicate the same relationship for three and five times (the two following odd multiples) the quarter of the tidal wave length respectively. (b) The Rossby deformation radius of the Kelvin wave describes how wide a basin should be for the development of an amphidromic system. (c) The natural period of oscillation of basins of different length depends on water depth. Cross-points of the line of the M2 tide (12.42 h) with the hyperbolical curved oscillation period time line for different basin lengths indicate basin length and depth required for conditions near resonance of the semidiurnal tide (modified after Sztanó and de Boer (1995) and based on Pugh (1987, 1996))

The present-day Tremp-Graus and Ager Basins, now separated by the Montsec Thrust, formed one entity, Tremp-Graus-Ager (T-G-A) Basin. The T-G-A Basin was elongated along an E-W line and connected to the Atlantic Ocean via the Jaca Basin (Fig. 18.7a). A subaerial topographic high, formed by a thrusted anticlinal system, bounded the T-G-A Basin in the south (Figs.18.2 and 18.3b). In the area east and southeast of Tremp, the T-G-A Basin was most likely closed and is referred to as the Gulf of Ager by some workers (Fonnesu 1984; Eichenseer 1988; Eichenseer and Luterbacher 1992; Mutti et al. 1994; Barberà et al. 1997; Waehry 1999). It included a southern limb that likely extended farther to the east than the present-day closure (Maestro-Maideu et al. 1991; Dreyer and Fält 1993; Dercourt et al. 2000; Vincent 1993; Olariu et al. 2008a; Fig. 18.3b). Any connection with the Mediterranean is uncertain and unlikely as it would generate strong tidal currents for which no support is found in the sedimentary record. A number of NW-SE oriented anticlines developed along the central northern basin margin controlled by the long-lived inverted Mesozoic fault structures. They developed either as (i) lateral or oblique ramps (or their associated hanging­wall anticlines) to S-directed upper cover-thrust sheets (cf. Muñoz 1992) or, alternatively, (ii) linked to the Bóixols frontal thrust (Cámara and Klimowitz 1985). It is suggested here that the periodic re-activation of these structures was expressed by an indented coastline along the northern basin margin locally forming estua­ries and/or embayments causing tidal amplification and ebb-flood cyclicity.
Fig. 18.7

Sketch of the inferred paleogeography and water depths of the T-G-A basin sea during the overfilled shelf stage of the Lutetian. The dotted line indicates the position of the basin axis. (a) The Montanyana system at initial progradation. Alluvial fans along the northern basin margin and a fluvial system in the eastern and central parts of the T-G-A Basin. A relatively large delta front area existed with interdistri­butary bays and mouth bars. (b) The Montanyana system at advance progradation. River systems dominated and extended out almost to the shelf break filling-up the shelf area. Only a narrow delta front area remained (modified after Nijman 1998)

The basin had an asymmetrical transverse section with the strongest subsiding parts along the northern basin margin and along the south-side of the developing Montsec Thrust. Highest Early Eocene subsidence rates occurred during the Ypresian. Sedimentation rates, however, were also highest along the northern basin margin, such that a transverse section through the basin may have been approximately symmetrical or even asymmetrical with a depositional basin axis further south. The basin had pronounced sea-floor topography due to the effects of the southward moving thrust sheets. The E-W oriented Montsec Thrust pro­bably formed a shallow-water anticlinal topographic high on the sea floor across which exchange (intermittent?) of marine waters could occur between the areas of the present-day Ager Basin and Tremp-Graus Basin. This situation was sustained by the relatively rapid eustatic sea-level rise at the start of the Early Ypresian (Pujalte et al. 2009).

As a result, the total basin length was approximately 200 km (Mutti et al. 1985a) with an approximately 50 km long and shallower (approximately 40–60 m) T-G-A Basin section east of the Montsec oblique blind thrust (Fig. 18.3). The paleo-width of the T-G-A Basin is uncertain but is estimated at up to 60 km (larger than it is at present after post-Ypresian compression, uplift and erosion).

In conclusion, the combined Jaca and T-G-A Basin was within the tidal amplification window (cf. Sztanó and de Boer 1995) with favourable average depth and length for resonant amplification (Figs. 18.3 and 18.6a, c), favorable paleobathymetric parameters, but unfavourable depth-width configuration for development of amphidromic point(s) as the basin was too narrow.

18.3.2 Overfilled Shelf

Tidalites of the overfilled shelf stage (Table 18.1, Fig. 18.7) are represented by facies of the Middle to Late Ypresian and Lutetian (51.5 Ma to 43 Ma) Montanyana Group present across the entire T-G-A Basin (Figs. 18.4 and 18.5). The T-G-A Basin most probably moved to the margin of the tidal amplification window (cf. Sztanó and de Boer 1995; Fig. 18.6a, c). As discussed hereafter, this was caused by the development of increasingly more pronounced basin-floor topographic features formed as the result of continuing thrust movement. These had a progressively more important control on basin morphology, depositional environments and tidal resonance.

The overfilled shelf stage resulted from the conversion of the (segmented) underfilled foredeep of the T-G-A Basin to a well-defined shelf and slope configuration as a result of ongoing thrusting and associated southward thrust-sheet-top basin translation along the Montsec Thrust as well as shortening along a N-S direction. The western part of the southward moving Montsec Thrust further developed into a pronounced northward curved oblique (NW-SE) blind thrust (Nijman and Nio 1975; Dunne and Ferrill 1988; Cuevas Gozalo 1989; Donselaar 1996a; Nijman 1998; Poblet et al. 1998; Clevis et al. 2004; Figs. 18.1, 18.2, 18.7 and 18.8) and acted as a lateral ramp separating shelf from slope and basin floor deposition during Late Ypresian and Lutetian sedimentation (52 Ma to 43 Ma; Cámara and Klimowitz 1985; Puigdefàbregas et al. 1992).
Fig. 18.8

Block diagram of the Montanyana delta showing alluvial fan and fluvial feeder systems combining in one delta front with break-in-slope above the lateral ramp of the underlying Montsec thrust sheet. CSPT   Central South Pyrenean Thrust system including the Montsec Thrust and its lateral ramps (modified after Marzo et al. 1988)

Associated smaller blind thrusts and associated ramps developed contemporaneously. An example of such an additional thrust is the Lascuarre reverse fault system (E of Graus; Fig. 18.9), with a NNE-SSW orientation, which acted as the most important sea-floor topographic expression (Cámara and Klimowitz 1985; Puigdefàbregas et al. 1992) east of the main oblique lateral ramp of the Montsec Thrust from 55.8 Ma to 48.6 Ma. It formed the transition from upper delta plain environments of the Montanyana Group in the T-G-A Basin on the hanging wall of the lateral ramp of the Montsec Thrust to contemporaneous lower delta plain deposition on the shelf margin and mass-flow deposition on the slope (the latter now making up a significant part of the present-day Ainsa Basin; Nijman and Nio 1975; Cuevas Gozalo 1989; Cuevas Gozalo and de Boer 1991; Donselaar 1996a; Nijman 1998; Poblet et al. 1998; Clevis et al. 2004; Figs. 18.7 and 18.8). It also prevented the formation of well-developed clinoforms (Puigdefàbregas et al. 1989). Later, from 48 Ma to 43.5 Ma (Early Lutetian), the Mediano anticline (Figs. 18.7 and 18.9) developed over the propagating Montsec oblique blind thrust, approximately 20 km to the west of the Lascuarre fault system, causing further steepening of the slope, over a time span of 5.5 Ma (Garrido-Megías 1973; Nijman and Nio 1975; Poblet et al. 1998). In between, the Isábena Depression developed (Cuevas Gozalo 1989; Fig. 18.9).
Fig. 18.9

Correlation profile of the Capella and Pano Fms in the Isábena depression and Virgen de la Collada ramp areas (modified after Cuevas Gozalo 1989, and Donselaar 1996a)

During Middle to Late Ypresian and Lutetian, the eastern and central parts of the T-G-A basin had begun to be progressively filled with alluvial and deltaic sediments from the north and east towards the west.

As a consequence, water depths in the early stages of infill are estimated to have been approximately 20 m in front of the Montanyana shoreline in the west (Nijman 1998) but quickly decreased during westward progradation of the system to water depths significantly less than 20 m (and up to 0). This reduction of the shelf area reduced tidal amplification. It is estimated, however, that the T-G-A Basin width and length were largely unchanged (approximately 40 km wide and 50 km long). Sediment pathways continued into contemporaneous extensive basin floor fan systems (Hecho Group; Mutti et al. 1972, 1973, 1975, 1985a) in the trough of the strongly subsiding E-W oriented Jaca Basin on the footwall. The open connection to the Atlantic Ocean still existed; no major structural barrier was present (Mutti et al. 1972, 1985a, 1988; Nijman and Nio 1975).

In conclusion, it is envisaged that the combined Jaca and T-G-A Basin was located in the margin of the tidal amplification window (cf. Sztanó and de Boer 1995). Despite the fact that average basin length was favourable for resonant amplification (Figs. 18.3 and 18.6a, c), the paleobathymetry became unfavourable as water depths on the (very) narrow shelf were (very) low and increased over (very) short (slope) distances to rela­tively deep. In addition, the depth-width configuration was unfavourable for development of amphidromic point(s) as the basin was too narrow.

18.4 Underfilled Foredeep Tidalites – The Ager Group

18.4.1 Ager Group Stratigraphy and Depositional Environments (Early Ypresian – 55.5 Ma to 51.5 Ma)

The Lower Ilerdian Alveolina Limestone is part of the Serraduy Fm (sensu Cuevas Gozalo et al. 1985; Figs. 18.4, 18.5 and 18.10) which, in general, comprises a low-relief carbonate ramp platform facies association developed along the basin margins. It is typified by a spatially and temporally complex facies architecture comprising, for example, intertidal and supratidal flats, carbonate shoreface facies, sublittoral sand bars, as well as outer ramp and open marine shelf environments (Eichenseer 1988; Eichenseer and Luterbacher 1992; Payros et al. 2000). Contemporan­eous deposition of bioclastic carbonate shales and nodular lime­stones (Metils-Milaris Fm; Figs. 18.4 and 18.5) took place in the area of relative deepest water (Fig. 18.3b).
Fig. 18.10

Simplified geological map of the present-day Tremp-Graus Basin showing the location of the outcropping Alveolina Limestone and the Roda, Capella and Pano Fms (modified after Serra-Kiel et al. 1994, who used data from Fonnesu 1984, Samsó 1988 and Tosquella 1988, and Cuevas Gozalo 1989). The box indicates the area in part covered by Fig. 19.11a

Tidal deposits of the Roda Fm (sensu Cuevas Gozalo et al. 1985; Fig. 18.4) occur along the northern margin of the Tremp-Graus Basin in the Isábena valley (Fig. 18.10) and form part of a well-developed NE-SW prograding, tide-influenced Gilbert-type delta (Nio and Siegenthaler 1978; Cuevas Gozalo et al. 1985; Yang and Nio 1985, 1989; Jimenez 1987; Eichenseer 1988; Tosquella 1988; Nio and Yang 1991; Serra-Kiel et al. 1994; Molenaar and Martinius 1990, 1996; Martinius and Molenaar 1991; Joseph et al. 1993; López-Blanco 1996a, b; López-Blanco et al. 2003; Torricelli et al. 2006; Tinterri 2007; Olariu et al. 2011; Leren et al. 2010; Michaud 2011).

In the Isábena valley, the Roda Formation is divided into the lower Roda Sandstone Mbr (approximately 120 m thick) and the upper Esdolomada Mbr (approximately 180 m thick, Nio and Yang 1991; López-Blanco et al. 2003; Fig. 18.4). The Roda Sandstone Member comprises at least six lobate-shaped sandstone wedges of a Gilbert-etype delta (López-Blanco et al. 2003; Leren et al. 2010) that displays an overall progradational (the Roda Sandstone Member) to retrogradational (lower part of the Esdolomada Mbr) pattern. The retrogradational part of the Roda Gilbert-type delta is overlain by the El Villar Limestone (Fig. 18.4). Each lobe has been subdivided into a number of smaller subunits. Each of these shows a generally lobate shape formed by large-scale (up to 15 m high) foresets with a dip angle of up to 32° (Yang and Nio 1989; Joseph et al. 1993; López-Blanco 1996a; López-Blanco et al. 2003; Tinterri 2007).

Approximately age-equivalent Roda Fm outcrops in the closed eastern part of the Tremp-Graus Basin, between the Noguera Ribagorçana and Noguera Pallaresa Rivers (Fig. 18.10), are less well described. Tidally dominated channel fills and bars formed in an embayment setting suggest NW-oriented tidal reworking along the northern and southern basin margin (Fonnesu 1984; Cuevas Gozalo et al. 1985; Mutti et al. 1994; Waehry 1999, his allostratigraphic units Figs. 18.4 and 18.5).

Contemporaneous deposition of bioclastic carbonate shales and nodular limestones (Yeba Fm; Fig. 18.5) took place in the area of relative deepest water. Oxic conditions prevailed, in places close to the lower boundary of the photic zone (Torricelli et al. 2006) with water depths reaching approximately 80 m. Near the Roda Gilbert-type delta, water depths decreased to about 40 m (Jimenez 1987).

The Baronia Fm (Mutti et al. 1972, 1973; Figs. 18.4, 18.5 and 18.11) is a tide-dominated succession of interbedded sandstones and sandy siltstones with gene­rally large-scale tabular bedding and no recognizable channels. Trace- and body-fossil assemblages throughout the formation indicate fully marine conditions (Mutti et al. 1985b; Rubino et al. 1985; Wonham 1993; Olariu et al. 2008a). The dominant paleoflow direction was towards the east, with a subordinate component towards the west. However, in some parts of the depositional system, westward oriented transport directions parallel to the orientation of the subaqueous Montsec Thrust prevailed (Rubino et al. 1985; Olariu et al. 2008a). Sediment must have been derived from the basin margin in the south and/or east and from reworking of older deposits (note the ­hiatus below the Baronia Fm; Figs. 18.4 and 18.5).
Fig. 18.11

(a) Strongly simplified geological sketch map of the eastern sector of the present-day Tremp-Graus Basin and the Ager Basin. (b) Geological sketch map of the present-day Ager Basin (modified after Mutti et al. 1985b)

Mutti et al. (1972, 1973, 1975) divided the formation into three units (lower, middle and upper). The lower and upper units consist of a series of tabular sandstone bodies characterised by an upward-coarsening grain-size profile and a sigmoidal geometry of accretion surfaces (Mutti et al. 1985 b; Olariu et al. 2008a). The upper unit of the Baronia Fm is interpreted to have been deposited in shallower water, subject to stronger tidal currents than the lower unit (Mutti et al. 1985b).

Wonham (1993) divided the Baronia Fm into two low-order sequences. The lower sequence is composed of a transgressive estuary succession with barriers at the mouth of the estuary developed above an erosional unconformity formed by lowstand fluvial incision. A tripartite estuarine facies distribution was recognized with ebb-tidal delta deposits in the distal western part of the basin and bay-head delta bars in the proximal astern part of the basin. Higher-order relative sea-level changes resulted in repeated basinward shifts of facies and a successive broadening of the estuary in time. The estuary deposits are overlain by a relatively thin (up to 12 m) succession of transgressive offshore and shelf deposits. The upper sequence is also formed by a tide-dominated estuary overlying an unconformity formed by lowstand fluvial incision (Wonham 1993). The wide estuary had no barriers at its mouth and was filled with compound cross-stratified beds formed by ebb- and flood-directed cosets containing tidal bundles; water depths were interpreted to decrease from about 20 m at the base to 5–10 m at the top. The transgressive estuary fill is overlain by a retrogradational set of lower shoreface and offshore deposits (Wonham 1993).

The Ametlla Fm (Mutti et al. 1972, 1973; Figs. 18.4, 18.5 and 18.11) rests conformably on the offshore siltstones of the Passarella Fm. It is informally ­subdivided into two members (the Pallaresa and Collada members; Dreyer and Fält 1993; Fig. 18.4) and tidal deposits occur in both members (Ghibaudo 1975). The Pallaresa member is comprised of offshore siltstone intervals which may be up to 45 m thick, and six laterally continuous tide-influenced and – dominated sandstone bodies deposited in deltaic, estuarine and tidal shelf settings which may be up to 32 m thick (Fig. 18.12). The Collada member is interpreted to have accumulated in a coastal plain to tidal flat environment (Dreyer and Fält 1993; Dreyer 1994). Sediments were shed from an extensive topographic high to the south (the developing Sierras the Marginales thrust system) and to the east. The formation was deposited during active pulsating thrust-sheet development, resulting in a number of high-frequency relative sea-level changes (Mutti et al. 1988). The highest relative subsidence occurred in the eastern part of the basin, where the Ametlla depocentre was located. Water depths, however, increased westward (Dreyer and Fält 1993). During deposition of the Ametlla Fm, the eastern closure of the T-G-A basin was (significantly) nearer to its present-day position (Dreyer and Fält 1993).
Fig. 18.12

Schematic illustration of the early transgressive stage in the inferred paleogeographical development of the T-G-A Basin during deposition of the Ametlla Fm (modified after Dreyer and Fält 1993)

18.4.2 Tidal Bars of the Alveolina Limestone

The combined effect of the relatively rapid eustatic sea level rise at the start of the Early Ypresian (Pujalte et al. 2009) and southward displacement of the Bóixols Thrust and associated lateral or oblique ramps resulted in NW-SE oriented anticlines and synclines along the central-northern margin of the T-G-A Basin. Two anticlines developed near Serraduy (Fig. 18.10; the Roda and Coll de Vent anticlines; Eichenseer 1988; Vincent 2001; López-Blanco 1996a, b) with a syncline in between (the Serraduy-Sis syncline; López-Blanco et al. 2003). This area became the locus for an approximately NE-SW oriented warm-water and southward opening narrow coastal embayment (‘Serraduy Bay’ of Unit 2 in the upper part of sequence V of Eichenseer 1988), typified by strong tidal currents, during the relative sea-level rise. Deposits formed in this embayment are part of the Alveolina Limestone Mbr of the Serraduy Fm (Fig. 18.4). Shallow-water carbonate banks and reefs developed particularly on the flank of the Roda anticline (Pool 1983; Eichenseer 1988). The embayment was dominated by WSW directed ­ebb-oriented dunes of up to 50 cm height; oppositely directed cross strata (flood currents) were subordinate. The dunes are part of thick (up to 12 m) bioclastic (grainstone) tidal bedforms, interpreted as bars (cf. Pool 1983; Eichenseer 1988; Fig. 18.13) based on the observation that cross-bedded sets show paleocurrent directions oriented obliquely to the accretion surfaces. The bars formed in tidal channels with a deeply scoured base and often covered by a lag of coralgal breccias derived from adjacent patch reefs. The bars contain several accretionary units with superposed dunes separated by discontinuity planes. Southward at short distances (200 m), small ebb tidal delta lobes with dominant WSW current directions developed, which, in places, contain double mud drapes; the existence of barrier islands, however, is not reported but inferred. The lateral equivalent of these tidal channel deposits east of Coll del Vent are formed by lagoonal to shallow-water bay deposits with occasional patch reefs. The succession is overlain by a coralgal reef and subsequently wide-spread shallow water nodular limestones of the Serraduy Fm.
Fig. 18.13

(a) Sedimentary section through Unit 2 of Cycle 5 of Eichenseer (1988) of the Alevolina Limestone in the village of Serraduy (Fig. 18.10). (b) Outcrop image of bioclastic tidal bar deposits in the upper tidal channel indicated in (a). (c) Detail of the tidal bar deposits in the upper tidal channel showing large-scale inclined stratification (master bedding surfaces) and the erosive tidal channel base. (d) Bottomsets of cross-stratified sets with mud-draped laminae formed in an inlet channel and mouthbar succession (cf. Eichenseer 1988) underneath the upper tidal channel

18.4.3 Tidal Bars of the Esdolomada Member

Tidal bars, in this case transgressive shore parallel sand­stone bodies, were developed in front of retrogradational fan-delta lobes and mouth bars of the lower part of the Esdolomada Mbr and are either attached to or detached (offshore) from the sandy delta lobe front. Underneath the El Villar Limestone (Fig. 18.4), a well-developed detached example (Fig. 18.14) is exposed near Roda de Isábena. This is formed by slightly inclined (1.6–4.6°) master bedding surfaces and contains stacked sets of high-angle (average dip angle 21°) cross-stratification up to 70 cm thick (Olariu et al. 2011). The crest of this bar is oriented sub-parallel to the tidal paleocurrent and to the nearby paleo-shoreline; the bar was built by oblique accretion, migrating transverse to the tidal currents towards the SW; it has a width to length ratio of approximately 1:10 (Olariu et al. 2011). Other examples, higher in the stratigraphy, also lack wave-generated structures and generally migrated obliquely towards the W driven by ebb-tidal currents flowing towards the NW. Tidal currents were deflected by the Gilbert-type delta lobe front along the NW-SE oriented paleoshoreline, reworking the lobe front.
Fig. 18.14

(a) A tidal bar at the base of the Esdolomada Mbr on the E side of the Isábena River SE of Roda de Isábena. Note the sheet-like geometry and well-developed slightly inclined (1.6–4.6° towards the SW) master bedding surfaces. This sandstone body migrated laterally (i.e., transverse to the tidal currents) towards the SW. (b) Detail of (a) showing stacked sets of high-angle (average 21°) cross stratification formed by dunes that migrated in a NW direction, that is approximately coast-parallel

The tidal bars appear to have developed during the transgressive phase of sedimentary cycles (López-Blanco 1996a; Olariu et al. 2011; Michaud 2011) in response to delta lobe abandonment after sediment deli­very to the delta ceased. The ensuing high-frequency (10s of millenia) relative sea-level rise initiated favou­rable conditions for a period of reworking of the Gilbert-type delta lobe front by tidal currents and the formation of tidal bars similar to headland tidal banks (cf. Michaud 2010) that subsequently drowned or, in some cases, became moribund. In the latter case, they are capped by mudstone. In the former case, when the bars became stationary, carbonate buildups developed on top (Michaud 2010; several examples in the Esdolomada Member).

18.4.4 Tidal Bars of the Ametlla Formation

The lowermost three sandstone units of the Pallaresa mbr are dominantly composed of tide-influenced, delta-front mouth bars (Dreyer and Fält 1993). Sandstone unit 4 was studied in more detail (Dreyer 1994) and is composed of three parts. The lower part was deposited in a tide-dominated estuarine environment during flooding (early transgressive stage) of a previously created incised valley with a pronounced unconformity at the base (Dreyer 1994). Units formed by stacked cross-stratified sets are up to 4 m thick and separated by fine-grained sandstone (Dreyer and Fält 1993) resemble the compound dunes of the Baronia Fm. In the data presented, however, no specific information is provided enabling the assessment of the progradation direction of the master bedding surfaces in relation to paleoflow directions of superposed cross-stratified sets. It is therefore unclear whether the sandstone bodies represent tidal bars (sensu Mutti et al. 1985b) or compound tidal dunes (sensu Olariu et al. 2008a).

This estuarine valley fill is overlain by mouth bar deposits of a bay-head delta (middle part of unit 4) which prograded towards the NW. The delta was initially fluvially dominated and influenced by tidal processes but transformed into a tide-dominated delta that prograded out into offshore inner shelf sediments during the later stages of the transgression (upper part of unit 4). No barrier further to the west is reported that created a back-barrier lagoon into which the bay-head delta prograded. The described characteristics illustrate that a certain degree of uncertainty is associated with the published interpretations and that the tide-influenced and tide-dominated depositional setting was characterized by a high degree of spatial and temporal variability of facies and sedimentary processes (Dreyer 1994).

The overlying sandstone units 5 (Fig. 18.15a, b) and 6 were formed by tide-dominated, near-shore to inner shelf ‘sandbar complexes’ during the late transgressive stage with common large-scale sigmoidal cross-stratified sets of up to 5 m thick. Average foreset dip angle is 23° to the west (Dreyer and Fält 1993) which is comparable to the dip angle observed in the tidal bar at the base of the Esdolomada Member (Olariu et al. 2011). Compound sets, up to 30 m thick and up to 3 km long, most probably have their long axis oriented parallel to the regional tidal flow direction (Dreyer and Fält 1993). These characteristics tend to classify these bodies as tidal bars, although a lack of relevant observational data prevents the distinction between tidal bars and compound tidal dunes.
Fig. 18.15

Outcrop images of Sandstone Unit 5 of the Pallaresa mbr of the Ametlla Fm. east of the old railroad station of the town of Ametlla. (a) Overview of Sandstone Unit 5 behind the railroad station. (b) Overview of a part of the stratigraphy with Sandstone Units 4, 5 and 6 looking westward towards the railroad station. (c) Tabular cross-stratified set with mud-draped bottomsets in the outcrop of (a). (d) Tidal bundle succession at the base of Sandstone Unit 5 filling up a scour in the channel floor (2 km east of the railroad station). Note the reactivation surfaces and the neap-spring cycles. (e, f) – Details of (d) showing mud-draped foresets, reactivation surfaces and neap-spring cyclicity

18.4.5 Tidal Bars Versus Compound Tidal Dunes in the Baronia Formation

Tidal sandstone bodies of the lower part of the Baronia Fm in the T-G-A Basin (Mutti et al. 1985b; Fig. 18.16a–d) have long served as a well-documented and classical example of shelf tidal bars in front of a delta (Dalrymple 1992; Mellere and Steel 1996; Willis 2005). Recent work (Olariu et al. 2008a) on the lower unit of the Baronia Fm, however, led to the conclusion that the tidal sandstone bodies were formed by compound tidal dunes deposited in a narrow, approximately 10 km wide, sea that extended farther eastward than the eastern closure of the present-day Ager Basin. The tabular sandstone bodies are generally 4–6 m (but up to 10 m; Fig. 18.16a–c) thick and display an alignment transverse to paleoflow. They continue over 100s of m to a km, both along depositional dip and depo­sitional strike. They commonly have a gradational base and are formed by stacked siliciclastic and bioclastic cross-stratified, planar- and trough-bedded sets, ripple-laminated sandstone and highly bioturbated sandstone. No oscillatory wave produced sedimentary structures have been reported. Cross-stratified sets show unidirectional or bi-directional paleocurrent directions and have occasional mud drapes on the foresets. The rippled sandstone is finer grained and contains thicker mud drapes (Mutti et al. 1985b; Olariu et al. 2008a, b). The bedforms shingled by migrating one over the other and offlaping (Fig. 18.17d).
Fig. 18.16

(ac) Outcrop images of compound tidal dunes of the lower unit of the Baronia Fm. east of La Baronia (Fig. 18.11b); arrows in (a) indicate the base and top. Cross-stratified sets show unidirectional or bi-directional paleocurrent directions and have occasional mud drapes on the foresets. The bedforms shingled by migrating one over the other and offlaping. Single dunes in the stacked sets (compound dunes) are inclined in the same direction as the compound-dune master surfaces, that is, the surfaces on which the larger compound dune migrated by forward accretion (Olariu et al. 2008a, b). (d) Heterolithic fine-grained facies interpreted to have been formed in the distal part of compound tidal dunes and areas in between compound tidal dunes. (e) Tidal bundle succession of approximately 10 m long forming part of a compound dune in an erosional depression on the seafloor either associated with an estuarine channel (cf. Mutti et al. 1985b) or a large tidal scour (cf. Olariu et al. 2008a, b) in the upper unit of the Baronia Fm at the village of La Règula (see Fig. 18.8); the latter interpretation is considered more likely. Dune foresets and master bedding surfaces dip towards the WNW and are draped with mud (Olariu et al. 2008a, b). Successively increasing and decreasing bundle thicknesses are interpreted to represent successive neap and spring periods (cf. Mutti et al. 1985b)

Fig. 18.17

Compound tidal dune model for the Baronia Fm showing the formation of a compound dune. Note that the inferred trajectory of successive dune troughs (dashed line) caused truncation of the previous cross-strata (modified with permission after unpublished data of Cornel Olariu, University of Texas at Austin)

Importantly, single dunes in the stacked sets (compound dunes) are inclined in the same direction (eastward) as the compound-dune master surfaces, that is, the surfaces on which the larger compound dune migrated by forward accretion (Olariu et al. 2008a, b; Fig. 18.17). This observation classifies the sandstone bodies as tidal compound dunes with their crest oriented normal to the tidal currents and internal accretion surfaces that dip in the same direction as the tidal currents. This interpretation stands in contrast to the interpretation as tidal bars proposed by Mutti et al. (1985b) for sandstone bodies of the lower unit which have their long axis parallel with the tidal currents and internal accretion surfaces that migrate laterally (at a high angle to the tidal currents). Note that Wonham (1993) interpreted bedforms in the overlying upper unit (up to 6–8 m) as compound tidal dunes on the same grounds. In contrast, however, to the typical orientation of compound dunes (sensu Olariu et al. 2008a, b) inferred to have been aligned parallel to paleoflow.

At basin scale, migration directions were primarily controlled by seafloor topography (dunes migrated to fill adjacent deeper parts of the basin) and dominant tidal current directions. Additionally, migration directions of the largest compound dunes were controlled by relative sea-level changes because dunes respond to changes in water depth. Olariu et al. (2008a) suggested water depths between 25 and 36 m at a minimum average for the lower unit. The sandstone bodies are intercalated with strongly bioturbated muddy sandstones up to tens of meters thick that represent low-energy fringes of amalgamated dune fields or periods of drowning.

Wonham (1993) and Olariu et al. (2008a) compared the depositional setting of the Baronia Fm with the outer part of the San Francisco Bay where the sea floor is covered by a tidal dune field (Rubin and Hunter 1982; Barnard et al. 2006). Berné et al. (1988) used the Baronia bedforms as an ancient analogue for the modern compound dunes of Surtainville in the English Channel that are formed by strong tidal currents. It is at present, however, unclear what the geography and morphology of the south-eastern section of the T-G-A Basin was at the time of deposition of the Baronia Fm. If the basin was closed towards the southeast this closure must have been located at least 10 or more kilometres away from the location of the (preserved) Baronia bedforms.

18.4.6 Tidal Bundles

The distally located toesets of almost all progradational Gilbert-type delta lobes of the Roda Sandstone Member in the Isábena valley were modified by tidal currents during periods between fluvially-derived sediment influxes and are represented by tidal dunes. These are small (up to 20 cm) in the lowermost three lobes and increase in size in the upper three lobes (preserved thickness up to 1 m). All have mud-draped foresets (Figs. 18.18 and 18.19). Bottomsets of the Gilbert-type delta lobe that reached farthest into the basin overlie an approximately 100 m wide belt formed by an at least 5 m thick preserved succession of tidal dunes with distinct neap-spring-neap cyclicity (Yang and Nio 1985, 1989; Nio and Yang 1991). This is well displayed in an outcrop along the Isábena River close to Roda de Isábena.
Fig. 18.18

(a) Overview of the tidal bundle outcrop locality along the Isábena River. The tidal bundles are located in the lowermost 5 m of the cliff section as indicated by the box. (be) Details of the tidal bundles showing the mud draped foreset and bottomset laminae. Paleoflow from left to right (SE to NW)

Fig. 18.19

Summary of the estimated paleotidal ranges, components and maximum random deviations of the tidal bundle deposits of the Roda Fm (upper left). Paleotidal components (b), random variations (c) and longer-period variations (d, e) derived from filtering analysis (upper right) (Modified after Yang and Nio 1985). The photograph shows the tidal bundles along the Isábena River, same locality as Fig. 18.18

Based on a comparison with the thickness and characteristics of bundle successions in tidal dunes formed in channels of the modern Oosterschelde estuary, Yang and Nio (1985) estimated that the tidal bundles were formed in an estuary with water depths of about 15 m. Tidal periodicity analysis indicates that the tidal bundles were formed in a meso- to macrotidal semi-diurnal regime (M2 dominant) with an estimated mean tidal range of 3.6 m (Fig. 18.19). Large irregular deviations from the expected equality in the bundle sequence were interpreted to reflect incidental storm influence.

A universally applicable dependency between tidal current velocity and the tidal range was used to estimate the tidal range. However, only a linear relationship between the volume of water flowing through a tidal channel during the dominant tidal period and the wet cross-sectional surface of a tidal channel below mean water level has been proven (O’Brien 1931; van de Kreeke and Haring 1979; Van den Berg 1986). Although maximum current velocities vary between spring tide and neap tide, and deeper channels are associated with somewhat higher local current velocities than shallow channels, theoretically equal tidal current velocities for all channel depths can be expected if tidal current velocity is replaced with shear velocity (J.H. van den Berg, personal communication 2010). Thus, no relationship exists between tidal shear velocity and tidal range and, hence, the estimated tidal range derived from the Roda bundle succession is questionable.

López-Blanco et al. (2003, his Fig. 6) suggested that the tidal dunes migrated over the lower part of the sandy delta front as part of attached tidal bars and that they were driven generally towards the NW by tidal currents. Low-amplitude, gentle NW-SE oriented folds resulted in a seafloor topography that caused funnelling of tidal currents in a NW-SE direction; low water depth would have contributed to the effectiveness of this process (López-Blanco et al. 2003). However, from observations in the Rhine-Meuse delta, including the Oosterschelde (Siegenthaler 1982, his Fig. 1), it appears that offshore tidal currents close to the coastline follow a rotary path in contrast to inshore estuarine tidal currents that show distinct reversals of current direction approximately along a linear flow path. The latter situation requires a funnel-shaped land constriction and, consequently, it is concluded that the Roda tidal bundles, exposed along the Isábena River close to Roda de Isábena, formed in a NW-SE oriented, restricted, inshore, tide-dominated environment such as an embayment (following Nio and Siegenthaler 1978; Yang and Nio 1985). Additionally, it is doubtful whether gentle seafloor topo­graphy could cause sufficient funnelling and reversals of tidal flow. The indented coastline morpho­logy resulted from movements along the same pre-existing NW-SE oriented blind thrust and associated faults mapped by López-Blanco et al. (2003) and which also controlled the location of the ‘Serraduy Bay’ (Eichenseer 1988). The Roda Sandstone Gilbert-type delta lobes debouched into this embayment.

The uncommon occurrence of a tidal bundle succession of approximately 10 m long (Fig. 18.16e) in the Baronia Fm near the village of La Règula (Fig. 18.11) in an erosional depression on the seafloor is interpreted to be either associated with an estuarine channel and shoal (cf. Mutti et al. 1985b) or a large tidal scour filled by a forward accreting compound tidal dune (cf. Olariu et al. 2008a, b). Given the paleogeographic setting of the Baronia Fm, the latter interpretation is considered more likely.

Sigmoidal and bidirectional cross-stratified beds with double mud-draped toesets and tidal bundles as well as herringbone cross-bedding and reactivation surfaces occur in sandstone bodes 4 and 5 of the Pallaresa member of the Ametlla Fm (Dreyer 1994). Large-scale sigmoidal cross-stratified sets of up to 5 m thick contain tidal bundles with well-developed double mud drapes, mostly in the toesets, and reactivation surfaces (Fig. 18.15c–f). Average foreset dip angle is 23° to the NW interpreted to have been formed by the ebb-dominant currents (Dreyer and Fält 1993).

18.5 Overfilled Shelf Tidalites – The Montanyana Group

18.5.1 Montanyana Group Stratigraphy and Depositional Environments (Late Ypresian to Late Lutetian – 51.5 Ma to43 Ma)

The Montanyana Group (Figs. 18.5, 18.8, 18.10 and 18.20) unconformably overlies shallow marine sediments of the Ager Group and had a dispersal pattern in which the main sediment transfer zones moved southward across the basin through time maintaining an E-W orientation. This was caused by the interplay between the progressive uplift in the inner parts of the thrust system along the northern basin margin and synchronous growth of transverse alluvial systems (Nijman and Nio 1975; Nijman 1998). The Montanyana system has been divided into a Lower, Middle and Upper part (van Eden 1970; Nijman and Nio 1975; Nijman 1998; Fig. 18.5). Lower delta plain facies of the Lower Montanyana Group are mapped as the brackish water facies of the Castigaleu Fm. These are contemporaneous with and interfinger with upper delta plain and alluvial facies of the Montllobat Fm. Equivalent facies of the Upper Montanyana Group are mapped as the Perrarua and Capella Fms respectively (Garrido-Megías 1968; Nijman and Nio 1975; Van der Meulen 1989; Puigdefàbregas et al. 1989; Figs. 18.5 and 18.20). The Middle Montanyana is an incised fluvial sheet sandstone that prograded rapidly westwards across the basin (the Castissent Fm; Figs. 18.5 and 18.8). The Montanyana system was fed by alluvial fans along the northern and eastern basin margin (Nijman 1998).
Fig. 18.20

Paleodrainage pattern of the Castissent Sandstone and related Corca Fm. across the underlying Castigaleu Fm. (lower and upper delta plain) and time equivalent Montllobat Fm. (fluvial); see also Fig. 18.4 (modified after Nijman 1998). Units A to C correspond to successive fluvial depositional cycles. CSPT   Central South Pyrenean Thrust system including the Montsec Thrust and its lateral ramps

The Castigaleu Fm is on average 400 m thick and is formed by a number of up to 12 m thick sandstone units intercalated with thick fine-grained (mud to very fine sand) intervals (Nijman and Nio 1975; van der Meulen 1989; Puigdefàbregas et al. 1989). Sandstone units, dominantly formed by fluvial channel fills and bars, are interpreted to generally have a meandering planform style. Commonly, sheet river-flood deposits are conglomeratic in nature. These channel fills intercalate with distributary mouth bars and interdistributary muddy, brackish bay deposits (Fig. 18.21a, e).
Fig. 18.21

(a) Outcrop impression of the alluvial valley deposits of the Gargalluda sandstone complex, a 2 km wide tidally-influenced trunk river alluvial valley exposed 2 km south of Puente de Montanyana (Fig. 18.10). (b) Lower delta plain embayment deposits into which the alluvial valley incised at this location. (1) Well-bioturbated brack-water marls and very-fine grained thin sandstone layers with oysters. (2) Non-bioturbated planar bedded and laminated very fine grained sandstone and siltstone layers showing rhythmic deposition inferred to have been controlled by tides and possibly indicating neap-spring cyclicity. (3) Medium-grained fluvial sandstone bed. (c, d) Medium-grained sandstone beds showing (i) increasing to decreasing bottomset thickness, (ii) increasing to decreasing foreset dip and shape (from convex to concave), (iii) in places increasing-decreasing organic particle concentrations, and (iv) occasional reactivation surfaces. These features interpreted as tidally-influenced fluvial accretionary bedforms resulting from successive fluctuations in flow regime conditions from lower (during flood tidal retardation) to higher (during ebb tidal drawdown) current velocities (cf. Martinius and Gowland 2010). (e) Outcrop image of a mouth bar complex associated with a lower delta-plain distributary channel in the Perarrua Fm (Fig. 18.5) directly east of the town of Salanova (8 km south of La Puebla de Roda; Fig. 18.10). (f) Mud-draped ripple-laminated (form)set forming part of the bottomset of a tabular cross-stratified bed in the upper part of the mouth bar shown in (e) and interpreted to have been formed during bankfull stage and reversal after flood combined with low current velocities in the channel (cf. Cuevas Gozalo and de Boer 1991)

Fluvial sandstone units in the distal part of the Montllobat Fm are up to 6 m thick and occur as (1) ribbons, (2) tabular bodies with lateral accretion and (3) lenticularly-bedded sheet bodies. Wide (up to 3 km) incised alluvial valley fills are formed by amalgamated sandstone complexes (up to 20 m thick) deposited in trunk rivers on the lower alluvial plain (Fig. 18.21a). In places, they are influenced by brack water and tidal processes as a consequence of having been formed in the fluvial to tidal transition zone (cf. Ghosh et al. 2005; Van den Berg et al. 2007; Martinius and Gowland 2010; Fig. 18.22).
Fig. 18.22

Schematic illustration of the fluvial to tidal transition zone showing zones of variable influence of the tides (after Martinius and Gowland 2010)

Alluvial upper and lower delta plain fluvial facies of the Capella Fm reach a thickness of approximately 1,000 m. Sediments were deposited in the depression formed by subsidence on the footwall of the Lascuarre Fault system and the lateral ramp of the Montsec Thrust (Fig. 18.9). Due to the relative high subsidence rate, a significant volume of fluvial sediment was stored in the Isábena Depression (Fig. 18.9) largely preventing progradation of the system. Phases of source area rejuvenation were characterized by an initially low-relief alluvial profile allowing tidal processes to increase their effect on the dominantly muddy lower delta plain environments (Cuevas Gozalo 1989) despite the short shelf.

The Pano Fm (uppermost Montayana Group; Fig. 18.5) was deposited as a transgressive coastal sandstone wedge forming the shallow-marine and time equivalent continuation of the uppermost part of Capella Fm (Nijman and Nio 1975; Nio and Donselaar 1978; Cuevas Gozalo 1989; Donselaar 1996a; Fig. 18.9). The Virgen de la Collada ramp, located between the Mediano anticline and the Lascuarre reverse fault system (Donselaar 1996a; Fig. 18.9), controlled sedimentation patterns and the position of internal Pano Fm facies boundaries. Only a short (approximately 10 km) shelf was present and a number of tectonically induced relative sea-level changes controlled sedimentation. These were related to a relatively high rate of subsidence alongside the growing Mediano anticline. The Pano Fm is divided in two third order sequences and each of these is further divided in a number of fourth order sequences (Cuevas Gozalo 1989; Donselaar 1996a, Fig. 1.-23).

The Ypresian and Lutetian T-G-A Basin fill is overlain by upper Eocene and Oligocene fluvial and alluvial fan deposits mainly sourced from the north (Puigdefàbregas et al. 1989).

18.5.2 Tidally-Influenced Fluvial Point-Bars and Mouth Bars

Tidal influence is encountered in fluvial sandstone bodies of the Castigaleu, Montllobar and Capella Fms and two tidally-influenced fluvial point-bar models with current reversals were proposed (Cuevas Gozalo and de Boer 1991; Fig. 18.23), highlighting the segregation of flood- and ebb-current generated structures around the meander bend.
Fig. 18.23

Two models (1 model 1, and 2 model 2) for tide-influenced fluvial point-bars with current reversals developed for the Castigaleu Fm and Capella Fm highlighting the segregation of flood- and ebb-current generated structures around the meander bend. (a)  perspective view, and (b)  plan view (modified after Cuevas Gozalo and de Boer 1991)

Tidally-influenced deposits of the Castigaleu Fm are formed in meandering channels and mouth bars associated with distributary channels (Cuevas Gozalo and de Boer 1991, their stop 4). These are typified by tabular and trough cross-stratified well-sorted sandstone beds with common reactivation surfaces and occasional herringbone structures; ripple-laminated sets occur in the top and mud drapes occur particularly in mouth bars entering brackish bays (Nijman and Nio 1975; Marzo et al. 1988; Hoornweg 1988; Fig. 18.21e, f). Some examples contain abundant brack-water to normal marine ichnofacies and oyster beds, and the top is commonly mottled. The thick fine-grained intervals between the sandstone units are deposited in inshore brack-water lagoons or shallow-water embayments based on body fossil content.

Some isolated meandering channels of the delta plain of the Montllobat Fm in the Noguera Ribagorçana River valley, originally described by Puigdefàbregas and van Vliet (1978) and Van der Meulen (1982), as well as the Gargalluda sandstone complex, a 2 km wide trunk river alluvial valley stratigraphically 30 m higher (Marzo et al. 1988), show features interpreted as tidal influence on fluvial accretionary bedforms during bankfull stage and reversal after flood combined with low current velocities in the channel (Cuevas Gozalo and de Boer 1991, their stop 9 and 11; Fig. 1.21a–d). These occur in a few stratigraphic positions indicating propagation of stronger tidal currents up into the meandering and distributary channels during certain phases of delta development while the shelf area was not completely filled yet.

The Capella Fm contains tidally-influenced fluvial channels throughout its stratigraphy (Cuevas Gozalo 1985a, b, 1989; Cuevas Gozalo and de Boer 1991). Typically, trough cross-stratified sets in the deepest part of the fluvial channels that are tidally-influenced have foresets showing mud drapes, bundling, reactivation surfaces not formed by dune overtaking and bipolar (but unevenly distributed) current directions while fluvially generated cross-stratified sets are clean, coarse and unidirectional. Tidally-influenced point bars contain large-scale avalanching foresets, undermined banks and bioturbated tops (de Boer 1998). Additionally, Cuevas Gozalo (1985a, 1989, her page 80/81) and Cuevas Gozalo and de Boer (1991) defined ebb-oriented, tidally-influenced fluvial spill-over lobes that formed in areas of fluvial flow expansion associated with fluvial channel bifurcation on the intertidal plain (Fig. 18.24). They are single sedimentary bedforms; the feeder channel shallows towards the lobe flat (Cuevas Gozalo 1985a, 1989). Furthermore, in some cases an upstream transition from a sand-dominated part of the point bar to sandy-muddy part of the point bar is observed. Cuevas Gozalo and de Boer (1991) suggest that upstream fines deposited in the upper part of the point bar were protected by flood dominated swales and/or inner parts of the channel.
Fig. 18.24

Sedimentary model for the Capella Fm for spill-over lobes in an intertidal plain. R tidally-influenced river, FEC fluvial-ebb tidal channel, FC marginal flood channel, FS flood shield (modified after Cuevas Gozalo 1985a, 1989)

18.5.3 Heterolithic Tidal Lagoon Deposits

The tidally-influenced and tidally-dominated environments in third order sequence 1 of the Pano Fm (Fig. 18.25) are formed as part of a retrogradational fourth order sequence set of a N-S to NE-SW oriented barrier and back-barrier complex succession with a shoestring geometry (Pano, Panillo and Grustán barrier chains respectively; Figs. 18.26 and 18.27; Donselaar 1996a). Each of these is associated with tidal inlets and back-barrier lagoons including tidal channels. Their formation was attributed to eustatic sea-level rise and coeval (punctuated) basin floor subsidence. In particular, sedimentation of the Pano coastal barrier complex was strongly influenced by tidal action as witnessed by the tidal channels, flood-tidal delta deposits and overall bimodal currents directions. Cross-stratified sets contain mud drapes on foresets, sigmoidal laminae shapes and convex-up reactivation surfaces. The Panillo ­barrier complex is wave dominated but the overlying Grustán barrier complex is a mixed-energy system with tidal (inlet) channels (Donselaar 1996a).
Fig. 18.25

Chrono- litho- and sequence stratigraphy of the Pano Fm in the area NW of Graus (modified after Donselaar 1996a)

Fig. 18.26

Schematic west-southeast cross-section through the Mediano and Virgin de la Collada lateral ramps showing the SE-ward stratigraphic displacement of barrier islands, inlets and tidally-dominated back-barrier environments. (Not to scale, see also Fig. 18.9. Modified after Donselaar 1996a)

Fig. 18.27

Paleogeographic reconstruction of Lutetian coastline development in the area NW of Graus. (a) Situation prior to the initiation of transgression. (b) Doming of the Mediano High caused subsidence of adjacent areas and the start of relative sea-level rise. Flooding occurred of the Capella coastal plain and the Pano barrier chain was formed. (c) Phase of maximum flooding and development of the Grustán barrier chain (3) and tide-dominated back-barrier area; the Pano (1) and Panillo (2) barrier chains are abandoned and drowned (modified after Donselaar 1996a)

The tidally-influenced environments of third order sequence 2 (Fig. 18.25) are formed by a retrogradational succession (parasequence set) formed by a tidally-influenced embayment fill at the base and ensuing transgressive, up to 35 m thick flood-tidal delta deposits overlain by highstand carbonate deposits (Donselaar 1996a). Time-equivalent barrier and/or inlet deposits are not reported and are assumed absent in the studied area or not preserved. The lowermost parasequence is located in the tidally-influenced (restricted) embayment and is formed by heterolithic (mud and very fine sand) deposits with double mud drapes at the base and neap-spring cyclicity (fourth order sequence 2.1; Fig. 18.28) forming an aggradational cyclic succession (Donselaar 1996a, b). The top part of the heterolithic succession forms the transition into the toe of the overlying flood-tidal delta that formed in the embayment supposedly behind an inlet between two barrier islands. Large inclined avalanche foresets dipping into the embayment characterise the deposit. Associated feeder channels are preserved in which cross-laminated sets have an E-W bimodal flood-dominated foreset dip distribution. Small-scale planar cross-stratified sets with opposite paleocurrent directions on top of convex-up reactivation surfaces illustrate the tidal signature of the deposit (Donselaar 1996a).
Fig. 18.28

(ad) Planar bedded and laminated inclined heterolithic facies of the lowermost parasequence in sequence 2.1 (Fig. 18.25) of the Pano Fm developed above a major unconformity. Erosive surfaces separate subsequent units of inclined heterolithic strata (white arrow in a). The deposits are interpreted to have been formed in a tidally-dominated part of an (restricted) embayment (cf. Donselaar 1996a, b) and are formed by heterolithic (mud and very fine sand) deposits with double mud drapes (white arrows in b) at the base and neap-spring cyclicity (black arrows in d). Regular vertical decrease and subsequent increase of bed thickness is interpreted to indicate spring to neap tide cycles with the thinnest beds formed during neap (black arrow in b; white arrows in d). Some layers have a deformed base (white arrow in c). (After Donselaar 1996b. With permission)

18.6 Summary

Lower Eocene tidalites in the T-G-A Basin in the southern Pyrenees (Spain) were deposited in response to developing thrust related folds and blind ramps which determined the position of facies belts and focussed and enhanced tidal currents. Two distinct stages of basin configuration can be recognised which share a general basin outline typified by a relatively narrow (up to 60 km) and long (up to 200 km in total) semi-enclosed sea which had an open connection to the Atlantic ocean in the west. They differ, however, significantly in water depth distribution, basin floor topography and coastal morphology. The two stages are directly related to two different configurations of basin dimensions favourable for resonant amplification and dominantly controlled by thrust sheet development.

The underfilled foredeep stage occurred during the Early Ypresian and was a period favourable for the formation of compound tidal dune fields and large tidal bars in the foredeep in relatively deep water (at least up to 40 m). Conditions were favourable for circulating and amplified outward flowing tidal currents. No major axially draining fluvial systems existed, but instead locally fed relatively small shoal-water fan deltas and larger (a) Gilbert-type delta(s) developed along the northern basin margin, dominantly located in structurally controlled areas of the coastline. Additionally, a large offshore compound tidal dune field was present in the south-eastern part of the basin as a result of sufficiently strong and confined tidal currents flowing E-W in a narrow sea.

The overfilled shelf stage became manifest during the Late Ypresian to Late Lutetian during which modest tidal amplification occurred in shallow water (up to 10 m), in-shore parts of delta distributaries and the subaqueous part of the delta top. A shelf formed behind the developing oblique lateral ramp of the Montsec Thrust, and a distinct shelf break was located above the lateral ramp. The shallow shelf sea was relatively rapidly filled by the axial east to west prograding Montanyana alluvial to deltaic system, which restricted tidal amplification. Alluvial fans fringed the basin margin. During Lower Montanyana times the depositional shoreline to shelf break distance was approximately 15 km but only a few km’s remained during Upper Montanyana times when the fluvial system almost reached the shelf margin. A relatively short and steep slope existed westward of the shelf break (the area of the present-day Ainsa Basin) and the basin floor (the present-day Jaca Basin) was relatively deep and narrow. These areas were characterised by turbidite deposition.


  1. Barberà X, Marzo M, Reguant S, Samsó JM, Serra-Kiel J, Tosquella J (1997) Estratigrafía del Grupo Fígols (Paleógeno, Cuenca de Graus-Tremp, NE de España). Rept Soc Geol España 10:67–81Google Scholar
  2. Barnard P, Hanes DM, Rubi DM, Kvitek RG (2006) Giant sand waves at the mouth of San Francisco Bay. EOS Trans AGU 87(29):285–289CrossRefGoogle Scholar
  3. Beaumont C, Muñoz JA, Hamilton J, Fullsack P (2000) Factors controlling the Alpine evolution of the Central Pyrenees inferred from a comparison of observations and geodynamic models. J Geophys Res 105:8121–8145CrossRefGoogle Scholar
  4. Bentham P, Burbank DW (1996) Chronology of Eocene foreland basin evolution along the western oblique margin of the south-central Pyrenees. In: Friend PF, Dabrio C (eds) Tertiary basins of Spain. The stratigraphic record of crustal kinematics. Cambridge University Press, Cambridge, pp 144–152CrossRefGoogle Scholar
  5. Berné S, Auffret JP, Walker P (1988) Internal structure of subtidal sandwaves revealed by high-resolution seismic reflection. Sedimentology 35:5–20CrossRefGoogle Scholar
  6. Burbank DW, Puigdefàbregas C, Muñoz JA (1992) The chronology of the Eocene tectonic and stratigraphic development of the eastern Pyrenean foreland basin, northeast Spain. Geol Soc Am Bull 104:1101–1120CrossRefGoogle Scholar
  7. Cámara P, Klimowitz J (1985) Interpretación geodinámica de la vertiente centr-occidental surpirenaica (cuencas de Jaca y Tremp). Estudios Geológicos 41:391–404CrossRefGoogle Scholar
  8. Cande SC, Kent DV (1995) Revised calibration of the geomagnetic polarity time scale for the Late Cretaceous and Cenozoic. J Geophys Res 100:6093–6095CrossRefGoogle Scholar
  9. Capote R, Muñoz JA, Simón JL, Liesa CL, Arlegui LE (2002) Alpine tectonics I: the Alpine system north of the Betic Cordillera. In: Gibbons W, Moreno T (eds) The geology of Spain. Geological Society, LondonGoogle Scholar
  10. Choukroune P, ECORS Team (1989) The ECORS Pyrenean deep seismic profile reflection data and the overall structure of an orogenic belt. Tectonics 8:23–39CrossRefGoogle Scholar
  11. Choukroune P, Roure F, Pinet B, ECORS Team (1990) Main results of the ECORS Pyrenees profile. Tectonophysics 173:411–423CrossRefGoogle Scholar
  12. Clevis Q, de Jager G, Nijman W, de Boer PL (2004) Stratigraphic signatures of translation of thrust-sheet top basins over low-angle detachment faults. Basin Res 16:145–163CrossRefGoogle Scholar
  13. Cuevas Gozalo M (1985a) Sedimentary lobes in a tidally influenced alluvial area, Capella Formation, Tremp-Graus Basin, southern Pyrenees, Spain. Geol Mijnb 64:145–157Google Scholar
  14. Cuevas Gozalo M (1985b) Geometry and lithofacies of sediment bodies in a tidally influenced alluvial area, Middle Eocene, southern Pyrenees, Spain. Geol Mijnb 64:221–231Google Scholar
  15. Cuevas Gozalo M (1989) Sedimentary facies and sequential architecture of tide-influenced alluvial deposits: an example from the middle Eocene Capella formation. Ph.D. thesis, University of Utrecht, Geologica Ultraiectina 61, 152 pGoogle Scholar
  16. Cuevas Gozalo M, de Boer PL (1991) Tide-influenced fluvial deposits; examples from Eocene of the southern Pyrenees. In: Marzo M, Puigdefàbregas C (eds) Guidebook to the 4th international conference on fluvial sedimentology. Publicacions del Servei Geològic de Catalunya, 92 pGoogle Scholar
  17. Cuevas Gozalo M, Donselaar ME, Nio S-J (1985) Eocene clastic tidal deposits in the Tremp-Graus Basin (Provs. of Lérida and Huesca). In: Milá MD, Rosell J (eds) 6th International Association of sediment. European Regional Meeting, Lérida, Institut d’Estudis Ilerdencs, Excursion guidebook, pp 215–266Google Scholar
  18. Dalrymple RW (1992) Tidal depositional systems. In: Walker RG, James NP (eds) Facies models, response to sea level change. Geological Association of Canada, St. John’s, Canada, pp 195–218Google Scholar
  19. De Boer PL (1985) Paleozoic/Mesozoic sedimentary development. In: Donselaar MP, Geel CR (eds) Guide to the sedimentology of the Tremp-Graus Basin. Open file report, University of Utrecht, pp 19–65Google Scholar
  20. De Boer PL (1998) Intertidal sediments: composition and structure. In: Eisma D (ed) Intertidal deposits: river mouths, tidal flats, and coastal lagoons. CRC Press/LLC, Boca Raton, pp 345–361Google Scholar
  21. De Boer PL, Pragt JSJ, Oost AP (1991) Vertically persistent sedimentary facies boundaries along growth anticlines and climate-controlled sedimentation in the thrust- sheet-top south Pyreneean Tremp-Graus Foreland Basin. Basin Res 3:63–78CrossRefGoogle Scholar
  22. DeCelles PG, Giles KA (1996) Foreland basin systems. Basin Res 8:105–123CrossRefGoogle Scholar
  23. Dercourt J, Gaetani M, Vrielinck B, Barrier E, Biju-Duval B, Brunet MF, Cadet JP, Crasquin S, Sandulescu M (2000) Atlas of Peri-Tethys, Palaeogeographical Maps. Commission de la Carte Géologique de Monde (CCGM/CGMW), Gauthier-Villars, Paris, 269 p, 13 maps, 1 plCrossRefGoogle Scholar
  24. Donselaar ME (1996a) Barrier island coasts and relative sea level rise: preservation potential, facies architecture and sequence analysis. Ph.D. thesis, University of Delft, NUGI 816, 223 pGoogle Scholar
  25. Donselaar ME (1996b) Inshore heterolithic deposits. In: Johnson HD, Wonham JP, Gupta R, Donselaar ME, van de Weerd AA, Mutterlose J, Stadler A, Ruffell AH (eds) Geological characterisation of shallow marine sands for reservoir modelling and high resolution stratigraphic analysis. JOULE II final report, Section 1, pp 1–19Google Scholar
  26. Donselaar ME, Nio SD (1982) An Eocene tidal inlet/washover type barrier island complex in the south Pyrenean marginal basin, Spain. Geol Mijnb 61:343–353Google Scholar
  27. Dreyer T (1994) Architecture of an unconformity-based tidal sandstone unit in the Ametlla Formation, Spanish Pyrenees. Sediment Geol 94:21–48CrossRefGoogle Scholar
  28. Dreyer T, Fält L-M (1993) Facies analysis and high-resolution sequence stratigraphy of the Lower Eocene shallow marine Ametlla Formation, Spanish Pyrenees. Sedimentology 40:667–697CrossRefGoogle Scholar
  29. Dreyer T, Corregidor J, Arbues P, Puigdefàbregas C (1999) Architecture of the tectonically influenced Sobrarbe deltaic complex in the Ainsa Basin, northern Spain. Sediment Geol 127:127–169CrossRefGoogle Scholar
  30. Dunne M, Ferrill DA (1988) Blind thrust systems. Geology 16:33–36CrossRefGoogle Scholar
  31. Egger H, Heilmann-Clausen C, Schmitz B (2009) From shelf to abyss: record of the Paleocene/Eocene-boundary in the Eastern Alps (Austria). Geol Acta 7:215–227Google Scholar
  32. Eichenseer H (1988) Facies geology of late Maestrichtian to early Eocene coastal and shallow marine sediments (Tremp-Graus Basin, northeastern Spain). Unpublished Ph.D. thesis, University of Tübingen, 237 pGoogle Scholar
  33. Eichenseer H, Luterbacher H (1992) The marine Paleogene of the Tremp region (northeast Spain) – depositional sequences, facies history, biostratigraphy and controlling factors. Facies 27:119–152CrossRefGoogle Scholar
  34. Fonnesu F (1984) Estratigrafía física y análisis de facies de las secuencias de Fígols entre el río Noguera Pallaresa e Iscles (provs. De Lérida y Huesca). Unpublished Ph.D. thesis, Universitat de Barcelona, 317 pGoogle Scholar
  35. Galdeano A, Moreau MG, Pozzi JP, Brethou PY, Malod JA (1989) New paleomagnetic results from Cretaceous sediments near Lisboa (Portugal) and implications for the rotation of Iberia. Earth Planet Sci Lett 92:95–106CrossRefGoogle Scholar
  36. Garrido-Megías A (1968) Sobre la estratigrafía de los con­glo­merados de Campanúe (Santa Lietsra) y formaciones superiores del Eocene (extremo occidental de la cuenca Tremp-Graus, Pirineo central, provincia de Huesca). Acta Geol Hispanica 3:39–43Google Scholar
  37. Garrido-Megías A (1973) Estudio geológico y relacíon entre tectónica y sedimentación del Secundario y Terciario de la vertiente meridional pirenaica en su zona central (prov. Huesca y Lerida). Unpublished Ph.D. thesis, University of Granada, 395 pGoogle Scholar
  38. Ghibaudo G (1975) Depositi di barra di foce nel Paleogene della valle di Ager (Provincia di Lerida, Spagna). Boll Soc Geol Ital 94:2131–2154Google Scholar
  39. Ghosh SK, Chakraborty C, Chakraborty T (2005) Influence of fluvial-tidal interactions on the nature of cross-stratified packages in a deltaic setting: examples from the Barakar Coal Measure (Permian), Satpura Gondwana Basin, central India. Geol J 40:65–81CrossRefGoogle Scholar
  40. Guimerá J (1984) Paleogene evolution of deformation in the north-eastern Iberian Peninsula. Geol Mag 121:413–420CrossRefGoogle Scholar
  41. Guimerá J (1996) Cenozoic evolution of eastern Iberia: structural data and dynamic model. Acta Geol Hispánica 29:57–66Google Scholar
  42. Hall MT (1997) Sequence stratigraphy and early diagenesis: the Sobrarbe Formation, Ainsa Basin, Spain. Unpublished Ph.D. thesis, University of Manchester, 136 pGoogle Scholar
  43. Haseldonckx P (1972) The presence of Nypa palms in Europe; a solved problem. Geol Mijnb 51:645–650Google Scholar
  44. Hoornweg A (1988) Interaction between a fan delta and a fluvial system in the Eocene Castigaleu and La Rocca Formations, Isábena valley, Spanish southern Pyrenees. Unpublished M.Sc. thesis, University of Utrecht, 150 pGoogle Scholar
  45. Jimenez C (1987) Paleoecologie et Valeur Chronostrati­graphique des Foraminiferes Benthiques dans des Systemes Sedimentaires Littoraux et ltaiques. Applicationaux Series Ilerdiennes de Roda (Versant Sud-Pyreneen) et Coustouge (Corbieres). Unpublished Ph.D. thesis, University Paul-Sabatier de Toulouse, 302 pGoogle Scholar
  46. Joseph P, Hu LY, Dubrule O, Claude D, Crumeyrolle P, Lesueur JL, Soudet HJ (1993) The Roda deltaic complex (Spain): From sedimentology to reservoir stochastic modelling. In: Eschard R, Doligez B (eds) Subsurface Reservoir characterization from Outcrop Observations. Éditions Technip, Paris, pp 97–110Google Scholar
  47. Leren BL, Howell JA, Enge HD, Martinius AW (2010) Controls on stratigraphic architecture in contemporaneous delta systems from the Eocene Roda Sandstone, Tremp-Graus Basin, northern Spain. Sediment Geol 229:9–40CrossRefGoogle Scholar
  48. López-Blanco M (1996a) Estratigrafia Secuencial de Sistemas Deltaicos en Cuencas de Antepais: Ejemplos de Sant Lorenc del Mont, Montserrat y Roda (Paleogeno, Cuenca de Antepais Surpirenaica). Unpublished Ph.D. thesis, Universitat de Barcelona, 238 pGoogle Scholar
  49. López-Blanco M (1996b) Estratigrafía secuencial de sistemas deltaicos en cuencas de antepais: Ejemplos de Sant Llorenc del Munt, Montserrat y Roda (Paleogeno, Cuenca de Antepais Surpirenaica). Acta Geol Hispánica 31:91–95Google Scholar
  50. López-Blanco M, Marzo M, Muñoz JA (2003) Low-amplitude, synsedimentary folding of a deltaic complex: Roda Sandstone (lower Eocene), south-Pyrenean Foreland Basin. Basin Res 15:73–95CrossRefGoogle Scholar
  51. Luterbacher HP, Ali JR, Brinkhuis H, Gradstein FM, Hooker JJ, Monechi S, Ogg JG, Powell J, Röhl U, Sanfilippo A, Schmitz B (2004) The Paleogene period. In: Gradstein FM,t Ogg JG, Smith AG (eds) A Geologic Time Scale 2004. Cambridge University Press, Cambridge, pp 384–408Google Scholar
  52. Maestro-Maideu E, Betzler C, van den Hurk AM, Serra-Rolg J (1991) El Ilerdiense de la Serra D’Aubens. Correlacion con la Vall d’Ager. Geogaceta 10:58–61Google Scholar
  53. Malod JA, Mauffret A (1990) Iberian plate motions during the Mesozoic. Tectonophysics 184:261–278CrossRefGoogle Scholar
  54. Malot JA (1989) Ibérides et plaque ibérigue. Bull Soc Géol France 5:927–934Google Scholar
  55. Martínez-Peña MB, Pocoví A (1988) El amortiguamiento frontal de la estructura de la cobertera surpirenaica y su relación con el anticlinal de Barbastro-Balaguer. Acta Geol Hispánica 2:81–94Google Scholar
  56. Martinius AW, Gowland S (2011) Tide-influenced fluvial bedforms and tidal bore deposits (Late Jurassic Lourinhã Formation, Luisitanian Basin, Western Portugal). Sedi­men­tology 58:285–324CrossRefGoogle Scholar
  57. Martinius AW, Molenaar N (1991) A coral-mollusc (Goniaraea – Crassatella) dominated hardground community in a siliciclastic – carbonate sandstone (the Lower Eocene Roda Formation, southern Pyrenees, Spain). Palaios 6:142–155CrossRefGoogle Scholar
  58. Marzo M, Nijman W, Puigdefabregas C (1988) Architecture of the Castissent fluvial sheet sandstones, Eocene, south Pyrenees, Spain. Sedimentology 35:719–738CrossRefGoogle Scholar
  59. Mascle A, Puigdefàbregas C (1998) Tectonics and sedimentation in foreland basins: results from the Integrated Basins Studies project. In: Mascle A, Puigdefàbregas C, Luterbacher HP, Fernàndez M (eds) Cenozoic Foreland Basins of Western Europe. Geol Soc Spec Publ 134:1–28Google Scholar
  60. Meigs AJ, Burbank DW (1997) Growth of the south Pyrenean orogenic wedge. Tectonics 16:239–258CrossRefGoogle Scholar
  61. Mellere D, Steel RJ (1996) Tidal sedimentation in inner hebrides half-grabens, Scotland: the mid-Jurassic Bearreraig sandstone formation. In: De Batist M, Jacobs P (eds) Geology of siliciclastic shelf seas. Geol Soc Lond Spec Publ 117:49–79Google Scholar
  62. Mey PHW, Nagtegaal PJC, Roberti KJ, Hartevelt JJA (1968) Lithostratigraphic subdivision of post-hercynian deposits in the south-central Pyrenees, Spain. Leidse Geologische Meded 41:221–228Google Scholar
  63. Michaud K (2011) Facies architecture and stratigraphy of tidal ridges in the Eocene Roda Formation, northern Spain. M.Sc. Thesis, Queen’s University, Kingston, 128 pGoogle Scholar
  64. Millán H, Den Bezemer T, Vergés J, Zoetemeier R, Cloetingh S, Marzo M, Muñoz JA, Puigdefàbregas C, Roca A, Cerés (1995) Palaeo-elevation and effective elastic thickness evolution at mountain ranges: inferences from flexural modeling in the Eastern Pyrenees and Ebro Basin. Mar Petrol Geol 12:917–928CrossRefGoogle Scholar
  65. Molenaar N, Martinius AW (1990) Origin of nodules in mixed siliciclastic-carbonate sandstones, the Lower Eocene Roda Sandstone Member, southern Pyrenees, Spain. Sediment Geol 66:277–293CrossRefGoogle Scholar
  66. Molenaar N, Martinius AW (1996) Fossiliferous intervals and sequence boundaries in shallow marine, siliciclastic deposits (Early Eocene, fan-deltaic deposits in the southern Pyrenees, Spain). Palaeogeogr Palaeoclim Palaeoecol 121:147–168CrossRefGoogle Scholar
  67. Muñoz JA (1992) Evolution of a continental collision belt: ECORS – Pyrenees crustal balanced cross-section. In: McClay KR (ed) Thrust tectonics. Chapman and Hall, London, pp 235–246CrossRefGoogle Scholar
  68. Mutti E, Luterbacher HP, Ferrer J, Rosell J (1972) Schema stratigrafico e lineamenti di facies del Paleogene marino nella zone central sud-pirenaica tra Tremp (Catalogna) e Pamplona (Navarra). Mem Soc Geol Ital 11:391–416Google Scholar
  69. Mutti E, Obrador A, Rosell J (1973) Sedimenti deltizii di piana dim area nel Paleogene della valle de Ager (Provincia de Lérida, Spagna). Bull Soc Geol Ital 92:517–528Google Scholar
  70. Mutti E, Rosell J, Ghibaudo G, Obrador A (1975) The Paleogene of the Ager Basin. In: 9th international sedimentological congress, International Association of Sedimentologists, Nice, Excursion guidebook part B 1, pp 1–6Google Scholar
  71. Mutti E, Remacha E, Sgavetti M, Rosell J, Valloni R, Zamorrano M (1985a) Stratigraphy and facies characteristics of the Eocene Hecho Group turbidite systems, south-central Pyrenees. In: Mila MD, Rosell J (eds) Excursion guidebook: VI European Regional Meeting, IAS, Lerida, Excursion 1, pp 521–576Google Scholar
  72. Mutti E, Rosell J, Allen GP, Fonnesu F, Sgavetti M (1985b) The Eocene Baronia tide dominated delta shelf system in the Ager Basin. In: Mila MD, Rosell J (eds) Excursion guidebook: VI European Regional Meeting, IAS, Lerida, Excursion 13, pp 579–600Google Scholar
  73. Mutti E, Séguret M, Sgavetti M (1988) Sedimentation and deformation in the tertiary sequences of the southern Pyrenees. AAPG Mediterranean Basins Conference Field Trip 7, Special Publication, Institute of Geology, University of Parma, 153 pGoogle Scholar
  74. Mutti E, Sgavetti M, Waehry A, Carminatti M, Davoli G, Ghielmi M, Figoni M, Mora S (1994) Regional stratigraphy and sequence-stratigraphic aspects of the Figols Group. In: Mutti E, Davoli G, Mora S and Sgavetti M (eds) The eastern sector of the south-central folded Pyrenean Foreland: criteria for stratigraphic analysis and excursion notes. Second high-resolution sequence stratigraphic conference, Tremp, pp 37–41Google Scholar
  75. Nagtegaal PJC, van Vliet A, Brouwer J (1983) Syntectonic coastal offlap and concurrent turbidite deposition: the Upper Cretaceous Aren Sandstone in the south-central Pyrenees, Spain. Sediment Geol 34:185–218CrossRefGoogle Scholar
  76. Nijman W (1989) Thrust sheet rotation ? – the south Pyrenean Tertiary basin configuration reconsidered. Geodin Acta 3:17–42Google Scholar
  77. Nijman W (1998) Cyclicity and basin axis shift in a piggyback basin: towards modelling of the Eocene Tremp-Ager Basin, south Pyrenees, Spain. In: Mascle A, Puigdefàbregas C, Luterbache CPr, Fernàndez M (eds) Cenozoic Foreland Basins of Western Europe. Spec Publ Geol Soc 134:135–162Google Scholar
  78. Nijman W, Nio S-D (1975) The Eocene Montañana delta (Tremp-Graus Basin, provinces of Lérida and Huesca, southern Pyrenees, N. Spain). In: 9th international Sedimentological Congress, IAS, Nice, Excursion guidebook, 19, part B, 1–20 and Appendix (36 p)Google Scholar
  79. Nio S-D, Donselaar ME (1978) Field guide to transgressive siliciclastic complexes in the southern Pyrenean Basin, Spain: part Two. An Eocene transgressive barrier complex near Graus. Open file report no. 18, University of Utrecht, pp 45–58Google Scholar
  80. Nio S-D, Siegenthaler C (1978) Field guide to transgressive siliciclastic complexes in the southern Pyrenean Basin, Spain: part One. A Lower Eocene estuarine-shelf complex in the Isabena valley. Open file report no. 18, University of Utrecht, pp 1–44Google Scholar
  81. Nio S-D, Yang CS (1991) Sea-level fluctuations and the geometric variability of tide-dominated sandbodies. Sediment Geol 70:161–193CrossRefGoogle Scholar
  82. Nio S-D, van den Berg JH, Goesten M, Smulders F (1980) Dynamics and sequential analysis of a mesotidal shoal and intershoal cannel complex in the Eastern Scheldt (southwestern Netherlands). Sediment Geol 26:263–279CrossRefGoogle Scholar
  83. O’Brien MP (1931) Estuary tidal prisms related to entrance areas. Trans ASCE 1:738–739Google Scholar
  84. Olariu C, Steel RJ, Dalrymple RW, Gingras MK (2008a) Tidal deposits of Baronia sandstone, Lower Eocene, Ager Basin, Spain: compound tidal dunes and architecture of Lower Baronia. Report for BITE Joint Industry Consortium Phase 1, 46 pGoogle Scholar
  85. Olariu C, Steel RJ, Dalrymple RW, Gingras MK and Rubino J-L (2008b) Tidal dunes of the Eocene Baronia Sandstone, Ager Basin, Spain: distinguishing tidal dunes from tidal bars; Why bother? AAPG Annual Convention, San Antonio, TXGoogle Scholar
  86. Olariu MI, Olariu C, Steel RJ, Dalrymple RW, Martinius AW (2011) Anatomy of a laterally migrating tidal bar in front of a delta system: Esdolomada Member, Roda Formation, Tremp-Graus Basin, Spain. Sedimentology 58 (in press)Google Scholar
  87. Olivet JL (1996) La cinématique de la plaque ibérique. Bull Cent Rech Explor 20:131–195Google Scholar
  88. Oost AP (1995) Dynamics and sedimentary development of the Dutch Wadden Sea with emphasis on the Frisian Inlet. Ph.D. thesis, University of Utrecht, 454 pGoogle Scholar
  89. Ori GG, Friend PF (1984) Sedimentary basins formed and carried piggyback on active thrust sheets. Geology 12:475–478CrossRefGoogle Scholar
  90. Payros A, Pujalte V, Baceta JI, Bernaola G, Orue-Etxebarria X, Apellaniz E, Caballero F, Ferrandez (2000) Lithostratigraphy and sequence stratigraphy of the Upper Thanetian to Middle Ilerdian strata of the Campo section (southern Pyrenees, Spain): revision and new data, Rev Soc Geol España 13, pp 213–226Google Scholar
  91. Pearson PN, van Dongen BE, Nicholas CJ, Pancost RD, Schouten S, Singano JM, Wade BS (2007) Stable warm tropical climate through the Eocene epoch. Geology 35:211–214CrossRefGoogle Scholar
  92. Plaziat JC (1981) Late Cretaceous to Late Eocene paleo­geographic evolution of southwest Europe. Palaeogeogr Palaeoclimatol Palaeoecol 36:263–320CrossRefGoogle Scholar
  93. Poblet J, Muñoz JA, Travé A, Serra-Kiel J (1998) Quantifying the kinematics of detachment folds using three-dimensional geometry: application to the Mediano anticline (Pyrenees, Spain). Geol Soc Am Bull 110:111–125CrossRefGoogle Scholar
  94. Pocoví A (1978) Estudio geológica de las Sierras marginales catalanas. Acta Geol Hispanica 13(73):70Google Scholar
  95. Pool W (1983) Laterale en verticale faciesveranderingen in de Onder-Eocene Serraduy-, Roda Marl- en Sandstone Formation (Ager Group) tussen Serraduy en La Puebla de Roda, provincie Huesca, Spanje. Unpublished M.Sc. thesis, Open file report, University of Utrecht, 100 pGoogle Scholar
  96. Pugh DT (1987, 1996) Tides, surges and mean sea-level. Wiley, ChichesterGoogle Scholar
  97. Puigdefàbregas C, Souquet P (1986) Tectonosedimentary cycles and depositional sequences of the Mesozoic and tertiary from the Pyrenees. Tectonophysics 129:173–203CrossRefGoogle Scholar
  98. Puigdefàbregas C, van Vliet A (1978) Meandering stream deposits from the tertiary of the southern Pyrenees. In: Miall AD (ed) Fluvial sedimentology. Can Soc Petrol Geol Mem 5:469–485Google Scholar
  99. Puigdefàbregas C, Muñoz JA, Marzo M (1986) Thrust belt development in the Eastern Pyrenees and related depositional sequences in the southern foreland basin. In: Allen PA, Homewood P (eds) Foreland basins. Int Assoc Sediment Spec Publ 8:229–246Google Scholar
  100. Puigdefàbregas C, Nijman W, Muñoz JA (1989) Alluvial deposits of the successive Foreland basin stages and their relation to the Pyrenean thrust sequences. In: Marzo M, Puigdefàbregas C (eds) Guidebook to the 4th international conference on fluvial sedimentology. Publicacions del Servei Geològic de Catalunya, 176 pGoogle Scholar
  101. Puigdefàbregas C, Muñoz JA, Vergés J (1992) Thrusting and foreland basin evolution in the southern Pyrenees. In: McClay K (ed) Thrust tectonics. Chapman and Hall, London, pp 247–254CrossRefGoogle Scholar
  102. Pujalte V, Baceta JI, Schmitz B, Orue-Etxebarria X, Payros A, Bernaola G, Apellaniz E, Caballero F, Robador A, Serra_kiel J, Tosquella J (2009) Redefinition of the Ilerdian stage (early Eocene). Geol Acta 7:177–194Google Scholar
  103. Ricci Lucchi F (1986) The Oligocene to recent foreland basins of the northern Apennines. Int Assoc Sedimentol Spec Publ 8:105–139Google Scholar
  104. Roest WR, Srivastava SP (1991) Kinematics of the plate boundaries between Eurasia, Iberia and Africa in the North Atlantic from the Late Cretaceous to the present. Geology 19:613–616CrossRefGoogle Scholar
  105. Rubin DM, Hunter RE (1982) Bedform climbing in theory and nature. Sedimentology 29:121–138CrossRefGoogle Scholar
  106. Rubino JL, Leo M, Fonnesu F(1985) Detailed stratigraphy of a tidal bar complex in the Eocene Baronia sandstones, Ager Basin, south-central Pyrenees. IAS 6th European Regional Meeting, Lerida, abstract book, pp 657–660Google Scholar
  107. Samsó JM (1988) Estudi sedimentològic i biostratigràfic de la Formació St. Esteve del Mall (Eocé, conca Tremp-Graus). Unpublished M.Sc. thesis, University of Barcelona, BarcelonaGoogle Scholar
  108. Schmitz B, Pujalte V (2003) Sea-level, humidity, and land-erosion records across the initial Eocene thermal maximum from a continental-marine transect in northern Spain. Geology 31:689–692CrossRefGoogle Scholar
  109. Séguret M (1972) Étude tectonique des nappes et séries décollées de la partie centrale du versant sud des Pyrénées. Publications de l’Université de Sciences et techniques de Languedoc, série Geologie Structurale no. 2, Montpellier, 155 pGoogle Scholar
  110. Serra-Kiel J, Canudo JI, Dinarés J, Molina E, Ortoz N, Pascual JO, Samsó JM, Tosquella J (1994) Cronoestratigrafía de los sedimentos marinos del Terciario inferior de la Cuenca de Graus-Tremp (Zona Central Surpirenaica). Rev Soc Geol España 7:273–297Google Scholar
  111. Sha LP (1990) Sedimentological studies of the Ebb-Tidal Deltas along the West-Frisian Islands, the Netherlands. Ph.D. thesis, University of Utrecht, Utrecht, 159 pGoogle Scholar
  112. Siegenthaler C (1982) Tidal cross-strata and the sediment transport rate problem: a geologists approach. Mar Geol 45:227–240CrossRefGoogle Scholar
  113. Srivastava SP, Roest WR, Kovacs LC, Oakey LC, Lévesque S, Verhoef J, Macnab R (1990) Motion of Iberia since the Late Jurassic: results from detailed aeromagnetic measurements in the Newfoundland Basin. Tectonophysics 18:229–260CrossRefGoogle Scholar
  114. Sztanó O, de Boer PL (1995) Basin dimensions and morphology as controls on amplification of tidal motions (the Early Miocene North Hungary Bay). Sedimentology 42:665–682CrossRefGoogle Scholar
  115. Teixell A, Muñoz JA (2000) Evolucíon tectono-sedimentaria del Pirineo meridional durante el terciario: una síntesis basada en la transversal del río Noguera Ribagorçana. Rev Soc Geol España 13:251–264Google Scholar
  116. Tinterri R (2007) The Lower Eocene Roda Sandstone (south-central Pyrenees): an example of a flood-dominated river-delta system in a tectonically controlled basin. Riv Ital Paleontol Strat 113:223–255Google Scholar
  117. Torricelli S, Knezaurek G, Biffi U (2006) Sequence biostratigraphy and paleoenvironmental reconstruction in the Early Eocene Figols Group of the Tremp-Graus Basin (south-central Pyrenees, Spain). Palaeogeog Palaeoclim Palaeoecol 232:1–35CrossRefGoogle Scholar
  118. Tosquella J (1988) Estudi Sedimentològic I Biostratigràfic de la Formació Gresos de Roda (Eocè, Conca Tremp-Graus). Unpublished M.Sc. thesis, University of Barcelona, BarcelonaGoogle Scholar
  119. Van de Kreeke J, Haring J (1979) Equilibrium flow areas in the Rhine-Meuse Delta. Coast Engr 3:97–111CrossRefGoogle Scholar
  120. Van den Berg JH (1981) Rhythmic seasonal layering in a mesotidal channel fill sequence, Oosterschelde Mouth, the Netherlands. In: Nio S-D, Schüttenhelm RTE, van Weering TjCE (eds) Holocene marine sedimentation in the North Sea Basin. Spec Publ Int Assoc Sedimentol 5:147–159Google Scholar
  121. Van den Berg JH (1982) Migration of large-scale bedforms and preservation of crossbedded sets in highly accretional parts of tidal channels in the Oosterschelde, SW Netherlands. Geol Mijnb 61:253–263Google Scholar
  122. Van den Berg JH (1986) Aspects of sediment- and morpho­­dynamics of subtidal deposits of the Oosterschelde (the Netherlands). Ph.D. thesis, University of Utrecht, Utrecht, 126 pGoogle Scholar
  123. Van den Berg JH, Boersma JR, van Gelder A (2007) Diagnostic sedimentary structures of the fluvial-tidal transition zone. Evidence from deposits of the Rhine and Meuse. Geol Mijnb 86:287–306Google Scholar
  124. Van der Meulen S (1982) The sedimentary facies and setting of Eocene point bar deposits, Montllobat Formation, Southern Pyrenees, Spain. Geol Mijnb 61:217–227Google Scholar
  125. Van der Meulen S (1989) The distribution of Pyrenean erosional material, deposited by Eocene sheetflood systems and associated fan-deltas: A fossil record in the Montllobat and adjacent Castigaleu Formations, in the drainage area of the present Rio Noguerra Ribagorzana, provinces of Huesca and Lérida, Spain. Geologica Ultraiectina, 59, 125 pGoogle Scholar
  126. Van der Voo R (1969) Paleomagnetic evidence for the rotation of the Iberian Peninsula. Tectonophysics 7:5–56CrossRefGoogle Scholar
  127. Van Eden JG (1970) A reconnaissance of deltaic environments in the Middle Eocene of the south-central Pyrenees, Spain. Geol Mijnb 49:145–157Google Scholar
  128. Vergés J (2007) Drainage responses to oblique and lateral thrust ramps: a review. In: Nichols G, Williams EA, Paola C (eds) Sedimentary processes, environments and basins. A tribute to Peter Friend. Int Assoc Sedimentol Spec Publ 38:29–47Google Scholar
  129. Vergés J, Millán H, Roca E, Muñoz JA, Marzo M, Cirés J, den Bezemer T, Zoetemeijer R, Cloetingh S (1995) Eastern Pyrenees and related foreland basins: pre-, syn- and post-collisional crustal-scale cross-sections. Mar Petrol Geol 12:893–915CrossRefGoogle Scholar
  130. Vincent SJ (1993) Fluvial paleovalleys in mountain belts: an example from the south-central Pyrenees. Unpublished Ph.D. thesis, University of Liverpool, Liverpool, 407 pGoogle Scholar
  131. Vincent SJ (2001) The Sis palaeovalley: a record of proximal fluvial sedimentation and drainage basin development in response to Pyrenean mountain building. Sedimentology 48:1235–1276CrossRefGoogle Scholar
  132. Visser MJ (1980) Neap-spring cycles reflected in Holocene subtidal large-scale bedform deposits: a preliminary note. Geology 8:543–546CrossRefGoogle Scholar
  133. Waehry A (1999) Facies analysis and physical stratigraphy of the Ilerdian in the eastern Tremp-Graus Basin (south-central Pyrenees, Spain). Ph.D. thesis, University of Genève, Terre et Environment, 15, XII  +  191 pGoogle Scholar
  134. Willis BJ (2005) Deposits of tide-influenced river deltas. In: Giosan L, Bhattacharya JP (eds) River deltas – concepts, models and examples. SEPM Spec Publ 83:87–129Google Scholar
  135. Wonham JP (1993) Sedimentology and sequence stratigraphy of tidal sandstone bodies: implications for reservoir characterisation. Chapter 4 – controls on the facies architecture of estuarine incised valley fill sandstone bodies from the Eocene Figols Group, Ager Basin, Spain. Unpublished Ph.D. thesis, University of Liverpool, Liverpool, pp 154–226Google Scholar
  136. Yang CS, Nio S-D (1985) The estimation of palaeohydrodynamic processes from subtidal deposits using time series analysis methods. Sedimentology 32:41–57CrossRefGoogle Scholar
  137. Yang CS, Nio S-D (1989) An ebb-tide delta depositional model – a comparison between the modern Eastern Scheldt tidal basin (southwest Netherlands) and the Lower Eocene Roda Sandstone in the southern Pyrenees (Spain). Sediment Geol 64:175–196CrossRefGoogle Scholar
  138. Zachos J, Pagani M, Sloan L, Thomas E, Billups K (2001) Trends, rhythms, and aberrations in the global climate 65 Ma to present. Science 292:686–693CrossRefGoogle Scholar

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Authors and Affiliations

  1. 1.Statoil Research and DevelopmentTrondheimNorway

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