Contrasting Styles of Siliciclastic Tidal Deposits in a Developing Thrust-Sheet-Top Basins – The Lower Eocene of the Central Pyrenees (Spain)
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.
KeywordsTidal Current Sandstone Unit Sandstone Body Delta Lobe Blind Thrust
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.
Summary of the characteristics of the underfilled foredeep and overfilled shelf stage of the Tremp-Graus-Ager Basin. See text for discussion
Early Ypresian (55.5 Ma to 51.5 Ma)
Late Ypresian to Lutetian (51.5 Ma to 43 Ma)
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)
Basin length (total marine section)
Estimated water depth in present-dayTremp-Graus and Ager Basin area
Up to 60 m on average
0 to 20 m on average
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
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.
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 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.
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)
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).
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).
18.4.2 Tidal Bars of the Alveolina Limestone
18.4.3 Tidal Bars of the Esdolomada Member
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).
18.4.5 Tidal Bars Versus Compound Tidal Dunes in the Baronia Formation
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
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)
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
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.
18.5.3 Heterolithic Tidal Lagoon Deposits
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.
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