The structural and sedimentary evolution of the Arruda and Lower Tagus sub-basins, Portugal

Marine and Petroleum Geology 22 (2005) 427–453 www.elsevier.com/locate/marpetgeo

The structural and sedimentary evolution of the Arruda and Lower Tagus sub-basins, Portugal J. Carvalhoa,*, H. Matiasb, L. Torresa, G. Manupellaa,!, R. Pereirac, L. Mendes-Victord a

Instituto Geolo´gico e Mineiro, Estrada da Portela, Zambujal, 2720-461 Amadora, Portugal b Centro de Geofı´sica da Universidade Lisboa, 1050 Lisboa, Portugal c Partex-Oil and Gas, Av. da Republica, 50-4, 1050 Lisboa, Portugal d Instituto Inf. D. Luı´s, Rua Escola Polite´cnica 58, 1250-102 Lisboa, Portugal

Received 21 November 2003; received in revised form 9 November 2004; accepted 13 November 2004

Abstract The study discusses the Jurassic structural and sedimentary evolution of the Arruda and Lower Tagus sub-basins, located in the central and southern part of the Lusitanian Basin of Portugal. In the last five decades, thousands of kilometres of 2D seismic reflection lines were acquired for oil exploration. Reprocessing and reinterpretation of over 700 km from the study area was carried out. After the seismic to well ties based on well logs, VSP analysis, synthetic seismograms, revised petrological analysis and stratigraphic reinterpretation of the wells, nine key seismic/geologic horizons were mapped and depth converted. Reinterpretation of the seismic profiles was carried out using outcrop data and reprocessed aeromagnetic and gravimetric data of the study area that was integrated in a GIS environment. Several aspects of the tectonic and sedimentary evolution of the study area were enlightened. From the Triassic until the Late Jurassic, two rift phases were recognised, in agreement with data from other parts of the Lusitanian Basin. The initial rifting phase in the Triassic continued until the Callovian, with three prominent tectonic pulses: a first in the Late Triassic, a second in the Hettangian and a third from the Sinemurian onwards. The latter episode lasted until the Callovian in the Arruda sub-basin, and at least until the Bathonian, in the Lower Tagus sub-basin. During the Middle Jurassic, the Arruda sub-basin evolved to a NW-deepening carbonate ramp. The Lower Tagus sub-basin constituted a platform since the Early Jurassic, limited in the east by the Setu´bal-Pinhal Novo fault and in the west, by the Cadafais fault. The third rifting episode occurred in the Late Oxfordian–Early Kimmeridgian, producing a major N–S fault system already recognised by several authors, and another important NW–SE and NE–SW fault systems post-dating the former. q 2004 Elsevier Ltd. All rights reserved. Keywords: Lusitanian Basin; Arruda sub-basin; Lower Tagus sub-basin; Seismic stratigraphy; Seismic reflection profiles

1. Introduction The Lusitanian Basin is an Atlantic margin Mesozoic basin situated inwards of a marginal horst (the Berlenga Horst, Fig. 1). The study area is located in the central and southern sectors of this basin according to the definitions of tectonic sectors of several authors (in Kullberg (2000)) and

* Corresponding author. Address: Instituto Nacional de Engenharia Tecnolo´gica e Inovac¸a˜o, Estrada da Portela-Zambujal, Apartado 7586, 2721-866 Amadora, Portugal. E-mail address: [email protected] (J. Carvalho). ! Deceased. The authors are greatly indebted to his contributions to this paper. The outstanding role played by G. Manuppella in the Portuguese Jurassic stratigraphy will be felt by his colleagues for several decades. 0264-8172/$ - see front matter q 2004 Elsevier Ltd. All rights reserved. doi:10.1016/j.marpetgeo.2004.11.004

it includes the Lower Tagus Cenozoic Basin and the Arruda sub-basin (Fig. 1). Several papers have described the evolution of this basin based on stratigraphic analysis and the interpretation of seismic data from oil exploration (Montenat et al., 1988; Wilson et al., 1989; Ellis et al., 1990; Leindfelder, 1994; Leindfelder and Wilson, 1998; Rasmussen et al., 1998; Kullberg, 2000; Alves et al., 2002). Although seismic reinterpretation of the whole of the Lusitanian Basin has been published recently (Rasmussen et al., 1998), the work is based on the original seismic sections, which date from 1982 or earlier. However, despite some deterioration of the original tapes, most of the seismic data has been recovered and reformatted and is now available in digital form. Thus, a part of the available lines were reprocessed (Carvalho, 2003), in order to bring

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Fig. 1. Structural key features of the Lusitanian Basin and of different kinematic sectors. Also shown are the Late Jurassic and Lower Tagus sub-basins (after Wilson et al. (1989)). The study area includes the Arruda and Lower Tagus sub-basins. Wells discussed in the text are also represented.

new insights into the history of the two sub-basins, especially where the quality of the original lines is poor. Only the period Triassic–Early Berriasian is discussed in this paper. Well data was digitised, a new petrophysical analysis was carried out and synthetic seismograms were constructed. New seismic to well ties were produced from those results combined with VSP, log analysis and a revised lithologic interpretation of the wells. The latter was based on the original interpretation of the wells (GPEP, 1986), and the interpretations of Ramalho (1971) and Rocha et al. (1996). Potential field data—a regional aeromagnetic survey (Domsalski, 1969) and the national geodetic gravity coverage (Koll and Vasconcelos, 2000)—was useful in the interpretation of the seismic lines, by helping to resolve ambiguities

related to the presence of salt structures or magmatic intrusions and in picking the basement. To take advantage of the availability of digital data, including the 1:500,000 geological map of the study area (Oliveira et al., 1992), the different datasets were co-registered and loaded into a GIS. This environment allowed a careful control of the interpretation and 3D model building.

2. Geological setting 2.1. Tectono-sedimentary evolution Sediments deposited in the Lusitanian Basin are exposed in Mesozoic terrains in the Western Portuguese Margin

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located between the city of Aveiro, in the north, and the Arra´bida hills, 50 km south of Lisbon. Vast Mesozoic outcrops to the north of Lisbon and those of the Arra´bida hills are separated by thick Cenozoic sediments of the Lower Tagus Valley (Fig. 1). To the west, the basin is bounded by the Berlengas Horst and is limited to the east, by the Hercynian Massif. The Lusitanian Basin can be divided into three sectors, each bounded by major faults (Fig. 1). These faults were active during the formation of the basin and are probably rooted in structures formed during the Variscan orogeny (Wilson et al., 1989; Ribeiro et al., 1990; Kullberg, 2000). Since the Triassic, the Lusitanian Basin was split into a northern and a southern trough—Beira Litoral, and Estremadura, respectively—separated by the Nazare´ fault. Presently, this fault is considered the limit between the northern and central sectors of the Lusitanian Basin. The central sector of the basin is bordered to the south by the Arrife and Tagus-Gargalo faults. The southern sector is bounded by these faults to the north and by the Arra´bida fault to the south. The formation of the basin was initiated by a continuous rifting process from the Late Triassic to the Late Callovian, linked to the break-up of Pangea. This rifting phase can be sub-divided into three major tectonic pulses. The first occurred in the Late Triassic, the second in the Hettangian and the third in the Late Sinemurian onwards (Manuppella and Azeredo, 1996). From the Late Triassic until the Early Cretaceous, fault-controlled and stable regional subsidence episodes alternate to produce a thick Mesozoic cover that reaches over 4 km in some parts of the basin (Wilson et al., 1989; Watkinson, 1989; Kullberg, 2000). Informal Mesozoic lithostratigraphy of the Arruda and Lower Tagus subbasins is presented in Fig. 2. Inversion occurred in the central area of the basin during the Cenozoic (Wilson et al., 1989; Rasmussen et al., 1998). Transtensional rifting episodes affected the Estremadura Basin during the Middle Oxfordian–Late Oxfordian, originating the three sub-basins of Arruda, Turcifal and Bombarral (Leinfelder and Wilson, 1998). This episode was probably related to the Late Jurassic ocean-spreading episode in the Tagus Abyssal Plain to the west (Wilson et al., 1989). During Alpine compression in the Miocene, the central area of the Lusitanian Basin was inverted and extensional tectonics induced by the collision formed the Lower Tagus sub-basin (Ribeiro et al., 1996). The Lower Tagus Valley sub-basin belongs to the southern sector while the Arruda sub-basin is located in the central sector (Fig. 1). The Mesozoic succession of the Lusitanian Basin is subdivided into four depositional megasequences. The deposition of these megasequences is related to the opening of the Atlantic (Wilson et al., 1989): (1) Triassic– Callovian, rifting and later thermal subsidence but no ocean opening; (2) Middle Oxfordian–Early Berriasian, rifting and ocean floor spreading beneath the Tagus Abyssal Plain;

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(3) Valanginian–Early Aptian, rifting around north and Northwest Iberia; (4) Late Aptian–Campanian, ocean spreading around West and North Iberian Margins. A brief description of the first and second megasequences sequences follows. (1) After the Late Triassic initial rifting pulse, grabens and semi-grabens controlled by listric faults formed, creating the conditions for the infill of the basin with terrigenous, siliciclastic fluvial deposits dominating this sequence. They show facies variations from west to east, represented by evaporitic interbeds. A second and third rifting pulses occur since the Hettangian until the Callovian, the first of which is marked by volcano-sedimentary deposits south of the Tagus River that are not evidenced in the seismic stratigraphy and outcrop data in the Arruda subbasin (Manuppella and Azeredo, 1996). Marine deposits accumulated in the Beira Litoral Basin (corresponds to the northern sector in Fig. 1) during the Early and Middle Jurassic. The environment grades from open shelf to hemipelagic, in the deeper parts of the basin, with facies changes being observed to the east, during the Middle Jurassic. From the final Early Bajocian onwards, sediments are deposited in a high-energy marine shelf, which evolved into a brackish platform, during the Early Bathonian. In contrast, north of the Tagus River, in the Estremadura trough (corresponds to the central sector in Fig. 1 after Ribeiro et al. (1996)), the seismic record during the Middle–Early Jurassic reveals a hemipelagic environment in the distal depressions (Peniche) and an open shelf low energy environment, inland. From the Bajocian onwards, the depositional environment evolves into a high energy one and, during the Late Bajocian, a slow regression takes place. In the Early Callovian, a new transgression occurs and sediments are again deposited in an open shelf setting (Manuppella et al., 2000). (2) After the Late Callovian the basin experiences sub-aereal exposure and erosion. From the Middle Oxfordian till the Early Aptian a second rifting occurred on a SSW–NNE oriented Lusitanian Basin (Wilson et al., 1989), with reactivation of Variscan fractures that bring important terrigenous materials (Rey, 1992). The extension allowed the formation of sub-oceanic crust in the Tagus trough (Mauffret et al., 1989) and the emplacement of basic dikes at the Lusitanian Basin borders. The first sediments to be deposited were the lacustrine carbonates with marine influence of the Cabac¸os formation. An eustatic equilibrium was achieved with the sedimentation of the shelf carbonates of the Montejunto formation, which present strong facies and thickness variations. Halokinesis associated to salty–clay movements (Canerot et al., 1995), already manifested during the Middle Jurassic, intensify with the second rifting, originating peri-diapiric

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Fig. 2. Informal lithostratigraphy for the Lusitanian Basin (modified from Leinfelder (1994)).

reef formations such as the one at Pinhal Novo in the Lower Tagus sub-basin after the Late Oxfordian. At the beginning of the Kimmeridgian, maximum subsidence occurred (Reis et al., 1997). The basin was then invaded by terrigenous prograding sedimentation (Abadia formation). Deposits from the uppermost Upper Jurassic correspond to a siliciclastic sequence with several transgressive episodes that in the basin depocentres grade to marine carbonates.

2.2. Lithostratigraphy of the study area 2.2.1. Paleozoic–Triassic In the study area, the oldest interpreted geological units are the sediments that were highly deformed during the Hercynian orogeny and which present a variable degree of metamorphism. These units do not outcrop in the Lusitanian Basin and are deeply buried in the study area beneath the Meso-Cenozoic cover. According to the two

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Fig. 3. Seismic reflection profiles reprocessed and reinterpreted and the wells used in this study. Highlighted are the three seismic profiles and three wells used in this paper for synthetic seismograms construction.

wells in the study area that reached the basement, Sobral-1 and Montalegre-1 (Sb-1, Mt-1, Fig. 3), slaty shales and sericitic schists, respectively, were intersected. The first formation above the Paleozoic basement is the Hettangian–Carnian Silves formation. The Silves formation is composed of siliciclastic materials deposited in an alluvial fan environment, under fault-controlled subsidence (Kullberg, 2000; Rasmussen et al., 1998). This formation occupied channels of the residual relief of the Hesperian massif and faulted, tilted blocks, producing the variable thicknesses encountered in several wells located in the Lusitanian Basin. In the study area, the Silves formation is usually part of the acoustic basement. In the study area, the Silves formation was only intersected in the Samora-1 well (Sa-1, Fig. 3).

in some onshore wells (Watkinson, 1989): an uppermost member essentially composed of dolostones and anhydrite beds; an intermediate salt/dolomite member and a lowermost member where the main lithology is halite, though stringers of dolomitic shales and anhydrite can occur. The Coimbra formation consists of a succession of dolomites, dolomitic limestones and limestones deposited in a progressively deepening, marine, westerly dipping homoclinal ramp system (Kullberg, 2000). In the Arruda sub-basin the average thickness of this formation varies from 100 m in the east, to 250 m in the western part of the sub-basin (Rocha et al., 1996). According to stratigraphic interpretations of the Campelos-1 (Cp-1, Fig. 3) well (Rocha et al., 1996), the top of Lower Jurassic probably corresponds to an unconformity.

2.2.2. Dagorda and Coimbra formations (Upper Triassic–Lower Jurassic) The Dagorda formation was deposited in an alluvial plain setting with episodic marine influences, where a rapid subsidence was accompanied by the deposition of a thick column of evaporites (Kullberg, 2000). Offshore it is composed of three members, a division which is recognizable

2.2.3. Brenha and Candeeiros formations (Lower–Middle Jurassic) The Candeeiros and Brenha formations belong to Middle and Lower Jurassic and were deposited in a carbonate ramp setting of shallow and deep waters, respectively. The Brenha formation is essentially composed of marls and marly limestones while the carbonate facies predominates in

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the Candeeiros formation. In some areas, the upper part of the Brenha formation is the lateral equivalent of the Candeeiros formation. Open sea conditions prevail since the Toarcian (Mouterde et al., 1971). At the top of the Candeeiros formation significant gravitational movements occur. The thickness of the Brenha formation can reach nearly 1 km (Kullberg, 2000) and according to Manuppella et al. (1985) the thickness of Candeeiros formation exceeds 800 m. At the top of the Candeeiros formation (dated from the Callovian), a hiatus in the sedimentation occurs, from which resulted that the Cabac¸os formation (Middle Oxfordian) rests in a slight angular unconformity over the former formation. 2.2.4. Cabac¸os and Montejunto formations (Upper Jurassic) The Cabac¸os and the Montejunto formations are carbonate sequences with mudstones interbedded in the Montejunto formation and, in the case of the Cabac¸os formation, anhydrites. The Cabac¸os formation was deposited in a lacustrine marginal marine environment in the Arruda sub-basin. The Montejunto formation is characterized by a progressively deepening marine shelf environment (Leinfelder and Wilson, 1998; Alves et al., 2002). According to well data in the Lower Tagus sub-basin, the Montejunto formation presents a reefal facies at the Barreiro-4 (Br-4, Fig. 3) well location and a shallow water limestones facies at other Barreiro and Montijo wells (Ellis et al., 1990). The Cabac¸os formation was not identified in the wells drilled in Lower Tagus sub-basin. 2.2.5. The Abadia and Amaral formations (Upper Jurassic) At the top of the Montejunto formation and the base of the overlying Abadia formation, siliciclastic sediments occur (Tojeira local unit and basal Abadia) that were deposited during the rift climax (very Late Oxfordian– very Early Kimmeridgian). The local unit of Tojeira that has been recently integrated into the Montejunto formation (Manuppella et al., in press), is composed of marls and turbiditic sandstones and corresponds to the beginning of the rift climax. The basal Abadia is composed of conglomerates, coarse arkoses and olistholits. The Abadia formation corresponds to the silicilastic infill of the Arruda sub-basin after the Late-Oxfordian–Early Kimmeridgian rifting phase. The thickest succession is the Castanheira member. This member is composed of coarse grained arkosic sandstones and conglomerates that reach over 2 km thickness near Arruda dos Vinhos, in well Arruda-1 (Ar-1). Using outcrop, borehole and seismic data, Leinfelder and Wilson (1989) interpreted the Castanheira member as a submarine fan supplied from a gap in an eastern basement horst. Outside the Castanheira member, the Abadia formation is composed of siltstones and marlstones deposited in a low energy setting that correspond to the immediate post-rift phase (Leinfelder and Wilson, 1998). The unnamed upper member of the Abadia formation is composed of mudstones,

shales, siltstones and subsidiary sandstones that are part of a southerly prograding slope. Finally, the Amaral formation, that has an average thickness of several tens of meters, caps the Abadia formation. It consists of shallow-water shelf carbonates coincident with a regressive episode. Together with the Freixial, Arranho´, Sobral formations and the upper member of the Abadia formation, it constitutes the late postrift phase (Leinfelder and Wilson, 1998). 2.2.6. Sobral, Arranho´, Freixial and Torres Vedras formations (Upper Jurassic–Lower Cretaceous) The Sobral, Arranho´ and Freixial formations were deposited during the Late Kimmeridgian, Tithonian and very Early Berriasian. These three formations have been referred to in the literature as the Lourinha˜ formation (Leinfelder and Wilson, 1998; Kullberg, 2000; Alves et al., 2002). The Sobral and Arranho´ formations are composed of sandstones that grade into clayey limestones at the top. The formations were deposited in a shallow-water environment with marine influence. The Freixial formation is composed of sandstones and limestones deposited in a brackish water environment. In the Cp-1 well, located in the Bombarral sub-basin, the Sobral and Bombarral formations were traversed. The latter formation is the lateral equivalent of the Freixial formation North of the Montejunto anticline but in the study area it has a different depositional setting than that of the Feixial formation. It will therefore not be discussed here. The Freixial, Arranho´ and Sobral formations have only been drilled in the Br-4 well (Sb-1 intersects only the Sobral formation), which is located in the Lower Tagus sub-basin. The Lower Cretaceous is mainly composed of sandstones and limestones of the Torres Vedras formation, deposited in a fluvial environment (Kullberg, 2000; Rey, 2003). A transgression during the Cenomanian deposited the marine limestones of the Cace´m formation that were partially eroded in the study area. 2.2.7. Cenozoic In the Lower Tagus sub-basin, the Cenozoic is composed of interbedded clays, silts and sands, often cemented and occasionally by limestones (Antunes and Pais, 1992). These units show lateral variation of facies throughout the study area. Deposition of sandstones with some intercalation of limestones occurred during the Paleogene as alluvial fans located on the foot of faults (Barbosa, 1995). In the study area, the Miocene sediments deposited under essentially continental conditions with marine and brackish water influence (Antunes and Pais, 1992). Between Lisbon and the Arra´bida hills, deposition was predominantly marine, as opposed to the northern study area. During the Pliocene, under high-energy sea and fluvial environments, the deposition of feldspathic sandstones, coarse sandstones, gravel and conglomerates predominated.

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Table 1 Seismic acquisition parameters for different surveys Seismic survey

Contractor

Date

Offsets (m)

Seismic source

Nominal fold

Tejo Samora Barreiro CPP Arruda 80 Arruda 81 Barreiro Bombarral

Geco-Prakla CGG Prakla-Seismos CGG CGG CGG Quest party

1978 1979 1954/1955 1980 1981/1982 1963 1979

100–1250 100–2450 20–240 100–1250 120–1530 100–2450 450–1600

Air-gun Vibroseis Dynamite Vibroseis Vibroseis Vibroseis Geosel

24 24 1 24 24 24 12

3. Data and methodology 3.1. 2D seismic data Table 1 summarises the several seismic reflection surveys used in this study and shows their different acquisition parameters. The quality of the stacked sections varies in both sub-basins. In the Lower Tagus Valley (Samora, Tejo and Barreiro CPP surveys), the signal to noise ratio of the profiles is generally reasonable, with several reflectors with good continuity until 1 s and in some cases until near 2 s. Reprocessing of these lines significantly improved the quality of the stacked sections. In the Arruda sub-basin, the Arruda 80 and 81 stacked sections vary from excellent to poor, with some of the lines not showing more than one or two reflectors. However, most of the lines are of acceptable quality. Occasionally, reprocessing produced better results, but overall the reprocessing produced inferior results compared to the original processing. This is due to the loss of some original field files caused by the deterioration of magnetic tapes and to a combination of irregularity of the acquisition geometry and steep strata dips. Over 50 seismic reflection profiles were reprocessed, corresponding to more than 700 km, and several wells were used in the reinterpretation (Fig. 3). The processing sequence includes pre-processing (geometry information, trace edits, first arrival mute, gain recovery, filtering and elevation correction), deconvolution or spectral whitening, velocity analysis, residual statics, DMO and flat datum conversion. After stacking, time-variant frequency filters, phase-shift migration, F–K and coherency filtering were applied. Due to the process of spectral whitening, a few lines present a lower quality imaging of the near surface, such as the case of profile S1 (Fig. 15). However, this process generally produced better-stacked sections at greater depths and was often used, but the stacked sections using deconvolution and the original sections were also analysed for a correct interpretation of the upper part of the profiles.

The wells in the study area, Ar-1, Mt-1 and Sa-1, provided very little velocity information and only Sb-1, located close to profile Ar10–80, had good data. Therefore, wells and associated seismic lines shot outside the main study area had to be included. The additional pairs of wells/seismic lines selected were Br-4/B3 and Cp-1/H (Fig. 3). All three wells had both sonic log (DT) and density tools (RHOB and LDT/CNL), though some parts of the wells were not sampled. The caliper (CALI) and other curves used in this paper like spontaneous potential (SP), gamma ray (GR), photoelectric factor (PEF) and neutron porosity (NPHI), suffered from the same handicap. Sonic logs were calibrated by checkshot-surveys. VSP were acquired in the Br-4 and Cp1 wells. This data was analysed here but was not reprocessed. 3.3. Potential fields Aeromagnetic data used was surveyed in 1969 from an altitude of 600 m and with a 2 km flight line average spacing. It has been published at a scale of 1: 200,000 (Domsalski, 1969). Data was digitised using the cross-lines of flight lines and contours, reprocessed and reduced-to-thepole (Carvalho, 1995). Gravity data for the study area was also available, from several surveys that were homogenised by Torres (submitted), from regional surveys by the National Geographic Institute (Koll and Vasconcelos, 2000). 3.4. Petrophysical analysis Diverse log data from Br-4, Cp-1 and Sb-1 wells was used in an attempt to define an interpreted lithological succession, calculate petrophysical parameters and properties and confirm the seismic to well tie. This analysis was made using Landmark’s PetroWorks software. Several logs were used: gamma ray, spontaneous potential, caliper, neutron porosity, sonic, bulk density and photoelectric factor. From these logs, new properties were Table 2 Baseline values and lithology

3.2. Well data A total of three wells with good velocity information were chosen to construct synthetic seismograms, in order to tie well data to the corresponding seismic reflection lines.

DT RHOB VSH

Baseline value

Lithology

!55 (ms/ft) O2.8 (g/cm3) O0.5

Limestone Anhydrite Shale

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Fig. 4. Matrix determination plot for well Sobral-1.

derived, such as shale volume (VSH), density and sonic limestone matrix porosity (respectively, DPHI and SPHI) and apparent matrix density (RhoMa), as well as the lithological indicators M, N and U (UMApp). The latter, the apparent volumetric photoelectric cross-section matrix, is basically derived from the photoelectric absorption curve, and the M and N lithological identifiers, are basically derived from the acoustic wave traveltime and neutron porosity, respectively (Schlumberger, 1988). For the available wells, two different approaches were tried, in order to characterize the lithological variation. One was made using log baseline definition using selected logs, such as DT, RHOB, and VSH (Table 2). The second was through traditional cross-plot diagrams, namely U vs. RhoMa and N vs. M (Figs. 4 and 5). Plots shown in

Fig. 6. Porosity and lithology determination from litho-density log of Sobral-1 (fresh water, liquid-filled hole, rfZ1.0).

Figs. 4 and 5 show (with VSH values below 0.5) the presence of anhydrite, as well as calcitic and dolomitic rich strata. This is consistent with the interpretation made with baseline definition (for this well) and has a very good correlation with the geological well report. A lithology determination cross-plot shows that for the Sb-1 well a major cluster of data is plotted in the dolomite/limestone and in the anhydrite areas (Fig. 6), validating other methods used above. Using this chart, it is possible to estimate matrix apparent porosity, which in this case is approximately 10%. For the wells studied, different log curves can be presented, as well as the lithological interpretation derived from the above-described methods. The results for Sb-1 well are presented in Fig. 7. At about 1380 m depth, anhydritic beds were identified, as well as limestones, sandstones and shale intercalations. For Br-4 and Cp-1 wells a succession of limestones, sandstones and shale intercalations was found. 3.5. Synthetic seismograms

Fig. 5. N vs. M mineral determination for well Sobral-1.

The synthetic seismograms are the result of reflection coefficients (RC) generated from processed sonic and density measurements made in the wells. The steps in the processing chain were the following: sonic log editing, calibration of sonic curve with checkshots, resampling and calculation of RC. These steps were followed by the determination of primaries, multiples and convolution of the RC by an input signal. All logs were also referenced in true vertical depth and time to the seismic reference datum (SRD). The sonic log was calibrated using the data from the checkshots and making use of the velocity and interval-transit time method. This method forces the integrated sonic log to match

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Fig. 7. Petrophysical interpretation for well Sobral-1.

the time/depth pairs exactly. The sonic log is first corrected with a constant velocity adjustment to each checkshot interval followed by an interval-transit time adjustment. RC were computed considering transmission losses. Multiples are obtained from these RC using the methodology described by Wuenschel (1960). Finally,

a wavelet was selected for convolution with the RC. As the frequency content of the seismic lines lies between 10 and 50 Hz it was decided to use a Ricker wavelet with a central frequency at 30 Hz. Results are presented in Figs. 8–10. Fig. 11 presents the VSP acquired for wells Br-4 and Cp-1.

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Fig. 8. Correlation between Campelos-1 and seismic profile H. Legend (tops of formations): ABA-Abadia; MON-Montejunto; CAB-Cabc¸os; CAN-Candeeiros; COI-Coimbra; DAG-Dagonda.

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Fig. 9. Corridor stacks from the VSP realised in wells Barreiro-4 and Campelos-1.

3.6. Depth conversion of time structural maps After the seismic to well tie (described below), nine horizons were interpreted and mapped in two-way-time. These are an intra-Neogene interface and the tops of the following formations: Paleogene, Torres Vedras, Freixial, Abadia, Montejunto, Candeeiros (or Brenha, when the first is absent), Coimbra and Paleozoic. Time structure maps were depth converted afterwards using the interval-velocity method (Marsden, 1989) and a similar approach from Japsen (1993) to calculate the interval velocities. However, velocity anomalies (Japsen, 1993) were not introduced, as they reflect lithologic variations, uplift, etc. but also errors in velocity estimations, especially in areas with poor velocity information, as is the case of the study area. Unfortunately, when velocity anomalies are not considered, an exact fit is not guaranteed at the well’s locations. However, the risk of introducing false structures in the depth-converted maps caused by the interpolation of

scarce velocity anomalies is minimised. Due to a limited number of wells in the study area, stacking velocities were introduced in some controlled points to determine the parameters of the velocity gradient for each layer. The similarity between depth-converted maps, interval velocity and two-way-time maps and the fact that the maps produced also agree with well-known geological structures and other geophysical data, show that the methodology used was adequate (Carter and Siraki, 1993). Plots of the interval velocities against semi-depths for each layer of the model have been carried out (Fig. 12). The data points (velocities) were weighted according to its quality to determine the regression line. The model is composed of 6 layers, according to its velocity variations: (1) Neogene; (2) Paleogene and Cretaceous; (3) Supra-Amaral and Abadia formations; (4) Montejunto and Cabac¸os formations; (5) Candeeiros, Brenha and Coimbra formations and finally, (6) Dagorda and Silves formations.

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Fig. 10. Tie among Barreiro-4 well and profile B3. Legend (tops of formations): ING-intra-neogene; PAL-paleogene; CRS-Lowever Cretaceous; FRE-Freixial; ABA-Abadia; MON-Montejunto; CAN-Candeeiros; COI-Coimbra.

J. Carvalho et al. / Marine and Petroleum Geology 22 (2005) 427–453 Fig. 11. Tie between well Sobral-1 and reflection profile Ar10–80. Legend (top of formations): ABA-Abadia; MON-Montejunto; CAB-Cabac¸os; CAN-Candeeiros; COR-Coimbra; DAG-Dagorda; PALPaleozoic.

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Fig. 12. Cross-plots of interval velocity and mid point depths for the 6-layer model used in depth conversion. From the regression line are obtained the parameters v0 and k of the velocity curve vZv0Ckz used inside each layer.

Depth conversion of each horizon structural map allowed an adequate analysis of the tectonic and sedimentary evolution of the study area. These results, integrated with other revised and geo-referenced geophysical data (Carvalho, 2003; Torres, submitted), geological knowledge of the basin collected over decades and new sedimentological data collected by one of the authors (Manupella et al., in press), consolidated the interpretation of the seismic data. After depth conversion was carried out by the method described above, regular grids of top of horizons depths, layer thickness, interval velocities and reflectors amplitude were generated with the triangular method of interpolation. Results are presented as plan structural maps on Figs. 13 and14 (the top of the Montejunto and Candeeiros/Brenha formations). On these are overlaid the wells from Fig. 3 and the major faults from Fig. 1.

4. Seismic to well ties 4.1. Campelos-1/H This well has a VSP together with a run of the logs needed for synthetic seismograms. However, the uppermost 1500 m does not include the caliper and density tools. Below ca. 2450 m (within the Cabac¸os and Brenha formations), areas of good quality data alternate with poor ones. The tie between Cp-1 synthetic seismogram and seismic line H is reasonable (Fig. 8). At 890 ms (1550 m) and 1275 ms (2500 m), two false reflections are observed. The false reflections are originated by the start of the density tool and poor data, respectively, that induce an improper scaling of all the synthetic reflections. Near the top of Abadia formation several weak reflections are observed in the synthetic while

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Fig. 13. Structural depth map of the top of the Montejunto formation. Compare major faults from Fig. 1 and wells shown in Fig. 3 for a better location of the map.

seismic reflection data shows two strong reflectors. The discrepancy can be explained by the reduced thickness of the Amaral formation and the different frequency content of the seismic profile and log curves. The reduced thickness of the geological units at that depth and the lower resolution of the seismic reflection data may originate composite reflections, while the superior resolution of the log curves allows their separation. The lack of the density logs at that depth may also explain the discrepancy. In the VSP (Fig. 11), acquisition of which started at the Amaral formation, the trace equalization applied in the data processing, destined to enhance the weakest reflections, may in part explain a reflector excess compared to the seismic section and synthetic. However, the major boundaries correlate (top of the Montejunto and Coimbra formations). 4.2. Barreiro-4/B3 The tie between seismic profile B3 and synthetic seismogram from Br-4 well is shown in Fig. 9. According

to the seismic section, synthetic seismogram and the corridor stack from well Br-4 (Fig. 11), two major reflectors are visible within the depth range of the well. The upper one (around 780 ms) matches very well with the top of Paleogene (890 m). This reflector probably originated in reflections from the interference of interbedded claystones, sandstones and limestones from the base of Miocene/top of Paleogene. A slight difference in time is probably due to the replacement velocity used in checkshot correction and timedepth curve. The second of the strongest events observed in the VSP, seismic profile and the Br-4 synthetic occurs at about 1220 ms. This event is in fact two or three reflectors closely spaced (1250 and 1300 ms). These reflectors correspond to (i) the transition from the limestones and sandstones of the Freixial, Arranho´ and Sobral formations to the limestones of the Amaral formation (1717 m) and (ii) the transition from the carbonates of the Amaral formation to the siliciclastic Abadia formation (1770 m). In the Br-4 area the Abadia formation is composed of brechoid para-reef limestones

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A good match was found for the most prominent reflections: the top of the Montejunto formation (600 ms), the anhydrites inside the Cabac¸os formation (690 ms), the top of the Coimbra formation (1040 ms) and the dolomitic limestones with several anhydrite interbeds of the Dagorda formation, characterized by sharp rises and drops in the log curves (1080 ms). The top of the Candeeiros formation is marked on several logs (GR, SPHI) but not on the sonic and density logs, so a reflection is not produced in the synthetic. The seismic profile presents a few moderate reflections inside the Candeeiros and Brenha formations apparently produced by lithologic changes. The top of Paleozoic (1110 ms) is characterized by a decrease in both the density and sonic curves followed by an increase in both parameters, after which remains approximately constant until the end of the well (2947 m). This horizon marks the beginning of an area without continuous reflections on the seismic profile. The three good reflectors seen inside the Paleozoic in the synthetic are probably multiples generated inside the Dagorda formation, as can be deduced from the synthetic built with and without multiples (not shown here).

5. Interpretation of the seismic reflection profiles 5.1. Basement and Silves formation (Paleozoic and Upper Triassic)

Fig. 14. Structural depth map of the top of the Candeeiros/Brenha formation.

with detritic intercalations and clayish cement rather than the typical succession of mudrocks. Below the well’s TD at 2833 m (1590 ms), the most important seismic events in the profile and VSP are seen below 1750 ms. This time marks the return of parallel reflectors in the seismic profile and, as observed in other areas, it might coincide with the top of the Coimbra formation or the top of Lower Jurassic. 4.3. Sobral-1/Ar10–80 The synthetic seismogram was not produced for the upper 1100 m due to very poor sonic data and missing density data above this depth. Below this level, two restricted areas of poor data in the logs produce a pair of artificial reflectors in the synthetic seismogram at about 88 ms (1890 m) and near 925 ms (2010 m) that are obviously not observed in the seismic profile. In spite of the scaling of all amplitudes to these reflections, apart from the pair of strong reflections on the synthetic, the correlation between seismic line Ar10–80 and Sb-1 synthetic is good (Fig. 10).

The top of the Paleozoic was drilled only by the Sb-1 and Mt-1 wells (Fig. 3). Of all these wells, only the Sb-1 had available data to build a synthetic seismogram (Fig. 10). As described above, the top of the Paleozoic corresponds to the beginning of an area of weak to moderate, irregular, discontinuous reflections. In other profiles, a similar pattern occurs beneath the Lower Jurassic reflectors, with only occasionally continuous events showing up (Figs. 15 and 16). This seismic pattern agrees with the features of a metamorphic, highly deformed geological unit as the Paleozoic is known to be. The top of Paleozoic corresponds to the top of a zone of chaotic reflections on the seismic profiles. In areas where salt or igneous masses are absent, which is the case in the majority of the study area, this seismic pattern follows the variations of the magnetic anomalies and specially the trend of the second-degree residual gravimetric anomalies. This seismic horizon was tentatively mapped in both the Lower Tagus and Arruda sub-basins. However, in the first basin, the depths may be under-estimated as it is not possible to map the Lower–Middle Jurassic horizons due to the poor quality of the seismic data. Nevertheless, depths obtained are in agreement with the aeromagnetic data interpretation (Domsalski, 1969). The Triassic was not crossed in the wells of the study area, except in the structural high of the Sa-1 well.

J. Carvalho et al. / Marine and Petroleum Geology 22 (2005) 427–453 Fig. 15. Reprocessed seismic reflection profile S1, with interpretation. Legend: SIL—Silves formation. For identification of other geological horizons see caption of Figs. 9 and 10. Gravimetry and magnetic data are also shown in the upper part of the figure.

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444 J. Carvalho et al. / Marine and Petroleum Geology 22 (2005) 427–453 Fig. 16. Interpreted stack of profile Ar9–80. For the key of geological horizons see caption of Fig. 15. Vertical scale is exaggerated twice relative to other profiles for display purposes. Gravimetry and magnetic data are also shown in the upper part of the figure.

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As explained in Chapter 6, this high is located on the footwall block of a major normal fault (Setu´bal-Pinhal Novo fault), which indicates that probably the Silves formation is also present in the lower areas of the Tagus sub-basin. Using the seismic to well tie based on the stacking velocities or other well’s velocities, the top of the Silves formation corresponds to a seismic horizon located in an area of very weak, wavy reflectors on profile S1 (Fig. 15). Away from the Samora structural high it is difficult to follow this seismic horizon, due to the presence of faults. In other areas of the Lower Tagus and Arruda sub-basins, beneath the signature of the Coimbra and Dagorda formations, a chaotic reflection zone shows up on the seismic profiles. We have seen above that this seismic pattern represents the acoustic basement, which includes Paleozoic formations. Inside the massive Dagorda salt pillows and diapirs, as in the case of profile Ar9–80 (Fig. 16), a seismic horizon can be picked above this chaotic reflections basement. This horizon may represent the top of the Silves formation or an intra-salt reflector. Whatever the case, it is not possible to follow this reflector outside the massive rock salt areas in the study area, which are quite restricted, so this horizon was not mapped. 5.2. Dagorda and Coimbra formations (Upper Triassic–Lower Jurassic) Among the three wells for which synthetic seismograms were constructed (Cp-1, Sb-1 and Br-4), the top of the Coimbra formation and the top of Lower Jurassic were only drilled at the Cp-1 and Sb-1 wells. They coincide approximately with two good reflectors in their respective synthetics (Figs. 8 and 10). However, at the depths where the top of Coimbra formation and Lower Jurassic were found in the Cp-1 well, the log data quality is poor. The correspondent reflections produced in the synthetic are therefore questionable. However, well logs quality at the Sb-1 well is good at the depths were Lower Jurassic sediments were intersected and its top is clearly visible in several logs. According to the description of the cores of well Sb-1, the top of the Coimbra formation does not represent any significant lithological change, but its presence is clearly visible in the logs (Fig. 7). The Coimbra formation and the upper member of the Dagorda formation constitute a group that is characterised by good, continuous, even parallel reflectors that are easily recognized in the Arruda sub-basin. This pattern is consistent with a stable depositional and tectonic environment. The top of this seismic pattern is generally accepted as the start of the Dagorda formation but our data suggests it is rather the top of the Coimbra formation or the top of the Lower Jurassic. In the Lower Tagus sub-basin, however, this ‘seismic package’ is not observed to the west of the Setu´bal-Pinhal Novo fault (Fig. 1), probably due to the great depth at which it is situated (over 4 km) and the corresponding loss of

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the seismic signal. To the northeast of this fault, the seismic signature of the Coimbra and Dagorda formations can be recognized in several profiles (Fig. 15). The fact that the presence of a thick salt column was found in the Pinhal Novo well and that both formations were intersected at the Montijo1 well (Mj-1, Fig. 3) also suggests that the Coimbra and Dagorda formations are present throughout the Tagus subbasin. In both sub-basins of the study area, this seismic pattern is cut by several faults. In some places, these do not affect Middle Jurassic horizons, which implies that some faulting occurred by the beginning of the Pliensbachien. Outcrop data in the Bombarral sub-basin shows syn-sedimentary faulting during the deposition of the Coimbra formation (Manuppella et al., in press). In outcrop, a growth of 20–30 m against these faults was found, but that is below the seismic resolution at the depths at which the Coimbra formation is found. Outcrop data in the Arruda sub-basin is not available, but it is possible that growth also occurred in the area, during the Sinemurian. Volcano-sedimentary materials dating from the Hettangian– Sinemurian found south of the River Tagus (Manuppella and Azeredo, 1996) also suggest this possibility. Away from the wells, it is quite difficult to determine which reflector corresponds to the top of Dagorda formation, since this group of reflectors present similar amplitude and character, so the top of the Dagorda formation was not mapped in the study area. If the average thickness of the formation found in the wells (approximately 150 m) is representative of the study area, the top of the Dagorda structural map is very similar to the top of the Coimbra formation map presented here. 5.3. Brenha and Candeeiros formations (Lower–Middle Jurassic) The upper boundary of the Brenha and Candeeiros formations do not produce seismic events. This is due to the lack of velocity and density contrasts with, respectively, the Candeeiros and Cabac¸os limestones. Generally, this thick carbonate sequence presents weak, discontinuous, subparallel reflectors in the Arruda or Lower Tagus sub-basins (Figs. 10, 15, 16, 17 and 18). However, the above mentioned strong reflector produced by interbedded anhydrites and limestones inside the Cabac¸os formation, together with the average thickness of this formation (200 m), can give an approximate depth to the top of the Candeeiros formation in a large area of the Arruda sub-basin. In the Lower Tagus area, the anhydrite/limestones interface inside the Cabac¸os formation is not present, which would make it very difficult to detect the top of the Candeeiros formation on the seismic profiles. Fortunately, this horizon is a regional unconformity in the Lusitanian Basin, which helped to detect it on the seismic profiles. For example, north of the Setu´bal-Pinhal Novo fault (Fig. 1), Upper Jurassic units are not present and the Cenozoic rests directly over Middle Jurassic carbonates. Here, therefore, the top of Candeeiros is

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Fig. 17. Final stack and interpretation of profile Ar5–81. For the key of geological horizons see caption of Fig. 15. Gravimetry and magnetic data are also shown in the upper part of the figure.

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Fig. 18. Reprocessed seismic reflection profile T16-2, with interpretation. For the key of geological horizons see caption of Fig. 15. Gravimetry and magnetic data are also shown in the upper part of the figure.

a very strong and continuous reflector (Figs. 15, 18 and 19) that can easily be followed in severely faulted areas. 5.4. Cabac¸os and Montejunto formations (Upper Jurassic) The top of the Cabac¸os formation does not produce a seismic event, which is due to the lack of velocity and density contrasts with the Montejunto limestones Inside the former formation, the carbonate/anhydrites interfaces produce a strong composite reflection that can be followed throughout the Arruda sub-basin (Figs. 10, 16 and 17). The top of the Montejunto formation, which correlates approximately with a low porosity zone in the Upper Jurassic calcareous facies in some areas of the Arruda subbasin, is characterised by a moderate to strong reflector (Figs. 8 and 10), again with the exception of the Barreiro area, where the reef limestones of the Abadia formation do not present a marked change in acoustic impedance relatively to the limestones of the Montejunto formation. However, in some areas, clastic interbedds do produce reflectors inside the Montejunto formation (Figs. 8, 10, 16 and 17). Moderate amplitude, continuous reflections that diverge towards the centre of the Arruda sub-basin are observed (Fig. 16 and 17).

As stated above, in the Lower Tagus sub-basin the Cabac¸os formation was not identified and the Montejunto formation presents a shelf carbonate facies. The thickness of the latter formation is reduced relatively to the Arruda subbasin (compare Figs. 16 and 17 with Figs. 15 and 18) and only slight thickness variations are found in the SW block of the Setu´bal-Pinhal Novo fault in the Tagus sub-basin. In the NE block of this fault, the Montejunto formation also seems to be absent in the seismic profiles and like other Upper Jurassic formations was not drilled in the Sa-1 well. 5.5. The Abadia and Amaral formations (Upper Jurassic) The top of the Amaral and Abadia formations coincide with two strong seismic reflectors that show good continuity in the seismic profiles (in the Arruda sub-basin). Within the upper part of the Abadia formation, south prograding clinoforms are reported from seismic data (Leinfelder and Wilson, 1998; Alves et al., 2002). In the seismic profiles interpreted here, only occasional evidence was found suggesting the presence of such prograding clinoforms. This agrees with outcrop data that suggest that progradation is frequent at the base of the Abadia formation (conglomerates, coarse arkoses and olistholits), rare in the upper part

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Fig. 19. Interpreted seismic reflection profile S4. For the key of geological horizons, see caption of Fig. 15. Gravimetry and magnetic data are also shown in the upper part of the figure.

of the Abadia formation and disappears at the top of the formation, where brackish water influences are found (presence of Karophitas). Below this upper part of the Abadia formation, flat, parallel and continuous reflectors are observed (indicating a stable subsidence rate) in the centre of the study area sub-basins, except in the Barreiro area, where that formation deposited in a peri-diapiric reef environment (Ellis et al., 1990) and also in the Vila Franca-Castanheira area, where it forms the submarine fan system mentioned above. In the Lower Tagus sub-basin, according to the interpretation of the Br-1–4 wells, the Abadia formation in the Barreiro area has the nature of a carbonate build-up (Ellis et al., 1990). A seismic anomaly is seen on profile B3 (Fig. 9). Here, a zone of discontinuous to chaotic reflection pattern replaces the usual seismic pattern (parallel, continuous reflectors) observed on the lower part of the Abadia formation in the Arruda sub-basin. Outside the seismic anomalous zone of Br-4 well, wherever the seismic signal is of reasonable quality,

the Abadia formation also presents parallel reflectors with good continuity, though the thickness of the formation is clearly less than in the Arruda sub-basin (Fig. 18). Furthermore, the Abadia formation thickness is relatively uniform in the Lower Tagus area. Together with the seismic character of the reflectors and Br-1–3 well data, this suggests that the Lower Tagus sub-basin was a stably subsiding platform during the Kimmeridgian, though locally (peri-diapiric zones) deep-water carbonate facies are present (Ellis et al., 1990). To the NE of the Setu´balPinhal Novo fault (Fig. 1), seismic data indicate that Upper Jurassic sediments are very thin or non-existent. This is supported by the fact that these horizons were not detected in the Sa-1 well. 5.6. Sobral, Arranho´, Freixial and Torres Vedras formations (Upper Jurassic–Lower Cretaceous) These units, which usually outcrop in the Arruda subbasin, are often very shallow in the seismic profiles and

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therefore, are not properly imaged. On profile B3 (acquired near well Br-4), in agreement with the synthetic seismogram, moderate reflectors that correspond to these formations are seen. The profile shows some irregularity and slight inclination of the seismic reflectors. Just like the Lower Cretaceous reflectors (Fig. 9), the reflectors of these Upper Jurassic formations present (in the Lower Tagus subbasin) a prograding hummocky-like configuration (Fig. 9) that indicates a shallow water environment in a prodelta or inter-deltaic position. According to sedimentary data it is admitted that this type of configuration fades out towards the top of the Jurassic succession. This agrees with a eustatic control of the sedimentation due to a very low subsidence rate proposed for this period (Rasmussen et al., 1998), in an essentially continental environment with some marine shallow water depositional periods, as deduced from sedimentary data. The Br-4 synthetic shows that the Cretaceous–Tertiary boundary is a weak to moderate reflector that is reasonably continuous on the seismic profiles in the Lower Tagus Cenozoic Basin. The Cretaceous seismic reflectors generally present a hummocky configuration. North of the Setu´bal-Pinhal Novo fault, these reflectors seem to be absent, and the Cretaceous is also not found in the Sa-1 well, just like the Upper Jurassic units. 5.7. Cenozoic The Barreiro well Br-4 was the only one with data allowing the construction of synthetic seismograms for the Cenozoic interval. The well defined stratigraphy of the Cenozoic succession produce strong reflectors, with good continuity and a parallel seismic reflection pattern (Figs. 9, 18 and 19). In the left margin of the Tagus River, at the base of Miocene/top of Paleogene, a strong amplitude reflector is located. This reflector is interpreted as the approximate top of the Oligocene. The Paleogene interval is characterised by a parallel reflection pattern with an even thickness throughout the study area corresponding to ca. 300– 400 m. This seismic configuration and sedimentary data indicate a uniform deposition rate on a uniformly subsiding shelf.

6. Discussion and implications to the basin’s structural evolution 6.1. Triassic–Callovian It is generally accepted that the first rifting episode occurred in the Lusitanian Basin during the Late Triassic (Montenat et al., 1988; Wilson et al., 1989; Rasmussen et al., 1998; Kullberg, 2000). This conclusion is based on seismic, well and geological surface data. Offshore seismic data show that mainly fluvial clastic sediments were deposited in half-grabens and grabens, separated by N–S,

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NE–SW and NW–SE predominant faults (Lomholt et al., 1995; Rasmussen et al., 1998; Alves et al., 2002). In the Arruda sub-basin however, the Silves formation has not been confirmed by drilling. Assuming that the top of Triassic is a strong reflector found at the base of massive rock salt, Rasmussen et al. (1998) mapped this horizon tentatively in the Arruda sub-basin. In the authors’ opinion, it is not possible to follow the referred seismic horizon outside salt diapirs and pillows but in these areas, the wedge-shape form of this formation defined by the hypothetical top of the Triassic and faulting affecting the horizon do indicate a fault-controlled subsidence (Figs. 16 and 17). In most of the Lower Tagus sub-basin, the quality of the seismic data below the Middle Jurassic is very weak. If we take the top of the Paleozoic as the top of the chaotic reflections zone, as explained above, the seismic data reveals a uniform thickness for the interval from the top Paleozoic to the Middle Jurassic. This assumption is further corroborated by the observation that throughout the study area the gravity anomaly data is consistent with the interpreted basement topography (Fig. 16). The extensional tectonics continued during the Lower Jurassic, producing regional subsidence. The seismic pattern of the Coimbra and upper member of the Dagorda formations is often faulted without the Middle Jurassic formations being affected. A slight thinning of these formations over the salt structures of the Runa-Montejunto anticline is also observed (Figs. 16 and 17). Since these formations do not show any growth near the faults, according to seismic data the faulting occurred after the deposition and thus, after the Sinemurian. However, as seen above, outcrop data shows that the tectonic pulse started as early as the Hettangian–Sinemurian. A thick column of carbonates was deposited in a westdeepening ramp environment during the Middle Jurassic (Watkinson, 1989; Wilson et al., 1989; Kullberg, 2000). In the study area, a relatively uniform thickness of Middle Jurassic sediments (900–1500 m) was found, except in a small area 10 km north of the Montijo well (Fig. 3) and in the NW part of the Arruda sub-basin. In the Lower Tagus sub-basin, the difficulty of detecting the top of Lower Jurassic formations probably explains the interpreted incipient depocenter located near Montijo (Carvalho, 2003). In the Arruda sub-basin, a gradual thickening of the unit is seen to the NW, where it reaches more than 2 km of sediments. Alves et al. (2002) also reported a NW thickening of the Coimbra and Brenha formations in the offshore region of the northern Lusitanian Basin. Fault-controlled subsidence associated to halokinetic movements occurred around Nazare´ (Rasmussen et al., 1998), Torres Vedras, Caldas da Rainha and Rio Maior (Montenat et al., 1988). In the study area, a slight thinning of these formations is seen over salt structures in the outskirts of the Arruda sub-basin (Figs. 16 and 17). In the depleted salt areas of the Arruda and the Lower Tagus sub-basins,

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differences in thickness found for Middle Jurassic formations are partially caused by erosion of the footwall blocks, however, due to later tectonism during the Late Jurassic (Fig. 15 and 17). The increase in thickness of Middle Jurassic succession found in the Bf-1 well (Rocha et al., 1996) is, however, not caused by erosion, but by differential subsidence on the Candeeiros-Torres Vedras and Cadafais faults. Furthermore, Middle Jurassic formations are seen to onlap earlier formations of the Sa-1 high (Fig. 15). This implies that fault-controlled subsidence during the late Early Jurassic and at least the early Middle Jurassic extended to this part of the Lusitanian Basin. 6.2. Middle Oxfordian–Berriasian In the Arruda sub-basin the Montejunto, Cabac¸os and Abadia formations show a NW–SE oriented depocenter that reaches over 1.5 km thickness for the carbonate succession between Alenquer and Arruda. The siliciclastic Abadia formation surpasses 1 km in thickness in a much wider area, and is over 2 km thick in the Ar-1 well area, where the submarine fan of the Castanheira member of the Abadia formation was deposited (Wilson et al., 1989; Leinfelder and Wilson, 1998). This fan probably started to accumulate sediments as early as the lowermost Late Jurassic or even the Middle Jurassic, since the discontinuous/chaotic seismic pattern of the Castanheira member is also observed for the Montejunto, Cabac¸os and the Candeeiros formations. The Montejunto, Cabac¸os and the base of the Abadia formation have a wedge-shaped seismic geometry with reflectors (Figs. 16 and 17) showing growth onto the major faults (Leinfelder and Wilson, 1998; Alves et al., 2002) and onlap onto the basin bounding structures (Figs. 16 and 17). This is consistent with the rift phase described in the literature for the Middle and Late Oxfordian (Rasmussen et al., 1998; Leinfelder and Wilson, 1998). Rasmussen et al. (1998) describe a N–S predominant fault pattern for the area south of Nazare´ for this period. Examples are major growth faults, like the Runa fault that separates the Turcifal–Arruda sub-basins, and the Ota fault that limit the Arruda sub-basin to the east. Additional and important NW–SE and NE–SW fault systems, post-dating the main faulting, were found in this study (Fig. 13). These fault patterns, in the case of the Arruda sub-basin, are consistent with surface geology (Oliveira et al., 1992). Salt pillows formed under the Montejunto anticlinal and under the Runa complex, depleting the central area of the Arruda sub-basin of a thick salt unit. Reactivation of basement faults during the Mesozoic extensional episodes propagated into the brittle overburden, like in other areas of the Lusitanian Basin (Rasmussen et al., 1998; Alves et al., 2002). Seismic profile Ar9–80 (Fig. 16) is a good example of salt tectonics and brittle behaviour at its northern and southern ends, respectively. The rift climax is reached at the basal Abadia (Pena dos Reis et al., 1998) and coincides with the filling of the Arruda

sub-basin with coarse siliciclastic materials and olistholiths. However, in the distal parts of the sub-basin, this episode is represented on the seismic profiles by thin, good continuity parallel reflectors (Figs. 10 and 16). This sequence is followed upwards by occasional prograding clinoforms, which correspond to the late Kimmeridgian succession (Figs. 16 and 17). The latter represents the final filling of the basin. The submarine fan mapped by Leinfelder and Wilson (1998) corresponds approximately to an area of greater thickness of the Abadia formation, while a negative seconddegree residual gravity anomaly coincides with a greater depth of the Montejunto and Candeeiros carbonate succession and basement (Fig. 16). The exact area of the submarine fan is difficult to determine from seismic data. The area is near the junction of the Ota and Cadafais fault systems and is affected by intense faulting. This Oxfordian faulting was reactivated later along NW–SE strike-slip faults during Late Cretaceous (in Alves et al. (2002)) and along NE–SW to NNE–SSW transpressive faults in the Miocene (Wilson et al., 1989). Wrench zones often originate chaotic or no reflection areas (Anstey et al., 1992) and here, this type of faulting is partially responsible for the chaotic and discontinuous seismic pattern shown by the seismic profiles that is associated to the Castanheira member, as already recognised by Leinfelder and Wilson (1989). The opening between the Montalegre and Ota horsts (Fig. 14) provided space for the sediment supply that constituted the Kimmeridgian submarine fan of Castanheira. This fan probably started to accumulate sediments as early as the early Late Jurassic. From the Middle Oxfordian to the Late Kimmeridgian, the Tagus sub-basin shows reduced thickness variations in the Montejunto, Cabac¸os and Abadia formations (500– 1000 m). Together with sedimentological evidence this indicates that during this period the area was a platform, whose eastern limit was probably controlled by the Setu´balPinhal Novo fault. Upper Jurassic sediments were not encountered in the Sa-1 well, but according to unpublished geological data collected by the authors, Upper Jurassic deposits were found in olistholiths near Vila Franca de Xira (Fig. 13), which indicates that the area north of the Setu´balPinhal Novo fault and Sa-1 well was uplifted and eroded. Consequently, Upper Jurassic sedimentation extended further to the NE of this fault. This cannot be confirmed from the seismic data, due to the presence of several faults and the lack of well data, that do not allow verifying the continuation of Upper Jurassic deposits between the Setu´bal-Pinhal Novo fault and the Sa-1 well. In the Lower Tagus Valley, the thick Quaternary cover does not allow for a comparison of surface geology with seismic data. Here, several NW–SE parallel faults can be recognised on the seismic profiles (Figs. 13 and 14). The Setu´bal-Pinhal Novo fault was active during the Late Jurassic. The timing of this movement is uncertain because, as stated earlier, the seismic mapping of Upper Jurassic

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horizons NE of the fault is uncertain. Apparently, Upper Jurassic units are very thin or non-existent in the footwall side of the fault (Fig. 15 and 18), which indicates that faultcontrolled subsidence occurred during this period on the hanging-wall block of the fault. The succession above the Amaral formation (uppermost Upper Kimmeridgian–Berriasian deposits) is not properly imaged in the Arruda sub-basin due to the acquisition parameters of the seismic lines, superficiality of these geological units and the severity of the static corrections (caused by irregular topography). In the Lower Tagus Valley, hummocky clinoforms are observed on the Sobral, Arranho´, Freixial and Lower Cretaceous formations that downlap the Abadia formation in the centre of the basin. This supports a platform environment for the area. The Ota limestone, the upper Castanheira member and the unnamed prograding member of the Abadia formation represent, together with the Amaral, Sobral, Arranho´ and Freixial formations a period of very low subsidence rates. Sedimentation was controlled by eustatic sea level during this period (Leinfelder and Wilson, 1998). This is in agreement with Rasmussen et al. (1998) result that also show a period of tectonic quiescence from the Late Kimmeridgian to the Berriasian. Tectonic quiescence continued during the Late Cretaceous (Rasmussen et al., 1998). Major tectonic inversion in the central part of the Lusitanian Basin, where the Arruda sub-basin is located, occurred during Middle Miocene due to NW–SE compressive episodes of the Alpine orogeny (Ribeiro et al., 1990; Rasmussen et al., 1998). The MesoCenozoic cover was deformed in the thin-skinned style with the Dagorda formation evaporites acting as the de´collement surface. During this period, the Lower Tagus sub-basin formed as a foredeep basin created by extensional tectonics induced by a NW–SE oriented collision and over 2 km of sediments were deposited in the basin’s depocentre. A relative uniform thickness of Paleogene deposits (200–500 m) in contrast to a much more variable and thick cover of Neogene sediments confirms the timing of the Lower Tagus sub-basin formation episode (Carvalho, 2003). The Neogene–Quaternary structural evolution of the sub-basin was controlled by NNE– SSE faults. The important monoclinal-fault system of Vila Franca de Xira-Lisboa (parallel to the Cadafais fault) that separates the Arruda and Lower Tagus sub-basins (Fig. 14) is one of such faults that limit the basin to the NW. This SE verging reverse fault is most likely the result of the reactivation of one of the Arruda sub-basin extensional boundary faults that occurred during the Miocene compressional events.

7. Conclusions The results of the reprocessing of the seismic reflection profiles improved, in some areas, the structural and

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sedimentary understanding of the study area. The correlation between seismic reflection profiles and wells, based on synthetic seismograms, well logs, petrophysical analysis, VSP and seismic stratigraphy analysis, allowed the identification of nine seismic horizons that correspond to chronologic/lithologic interfaces that could be mapped, totally or partially, in the study area. The seismic/geological horizons were depth-converted using the interval-velocity method and a velocity model composed of 6 layers above the Paleozoic basement. Several consequences from the integrated interpretation presented here are significant for the understanding of the structural and sedimentary evolution of this part of the Lusitanian Basin. The analysis of synthetic seismograms and well logs suggests that the top of Candeeiros/Brenha formations cannot be determined with accuracy as it does not produce a strong reflector. Furthermore, the Callovian boundary in many areas is not an angular unconformity and is often difficult to detect seismically. The anhydrite within the Cabac¸os formation can, in many areas, provide a clear uppermost limit of the Middle Jurassic. This study also shows that the top of the Dagorda formation is difficult to determine due to the similarity and closeness of the reflectors of the ‘seismic package’ that bears its signature and to the presence of faults. However, if an approximate constant thickness is assumed for the Coimbra dolostones, an assumption that is supported by well data, the structural maps of the two horizons are quite similar. Seismic evidence in the study area supports two rifting phases with several tectonic pulses. After the first rifting pulse in the Triassic, a second extensional pulse with a low to moderate intensity occurred during or immediately after the Sinemurian. It produced some faulting and halokinesis in the Arruda and Lower Tagus sub-basins, and continued during the Middle Jurassic with similar intensity. Evidence for this rift pulse is found on other parts of the Lusitanian Basin. However, the sedimentological record shows the first rifting phase to be a continuous rather than a discontinuous process occurring from the Triassic to the Callovian. From the early Middle Jurassic until the Callovian, the Arruda sub-basin evolved to a carbonate ramp slightly deepening to the NW with an average accumulation of up to 2 km of sediments. The area between the Sobral and Montalegre wells constituted a pre-Mesozoic basement horst (Sobral–Montalegre horst) that gradually subsided since the Early Jurassic until the Middle Jurassic. The Sobral–Montalegre and Ota Horsts are clearly visible on the structural maps. These paleo-horsts were buried and covered by sediments during the Late Jurassic. During the Middle–Late Oxfordian rifting episode, the limits of the Arruda sub-basin, witnessed a mixture of fault-controlled subsidence and halokinesis. The latter dominated in the western limits of the basin (Montejunto anticlinal and Runa area) especially in the northern part

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of the Runa fault; a fault-controlled subsidence occurred on the E and SE border of the basin (Ota fault and Cadafais fault, respectively). Besides the major N–S fault system already mentioned by several authors, an important NW–SE and NE–SW fault systems, post-dating the former, were identified in this study. Probably since the Triassic or the Early Jurassic until the Paleogene, the Lower Tagus sub-basin acted as a shallow platform controlled to the east by the Setu´bal-Pinhal Novo fault. The area east of this fault was exposed during the Late Cretaceous, when Upper Jurassic and Lower Cretaceous sequences were eroded. At the western border of this platform, no variation of thickness against the Vila Fanca de Xira-Lisboa fault was found on the Mesozoic sequence. This means that the western boundary of the platform was probably the Cadafais fault, which also borders the Arruda sub-basin to the southeast and is roughly parallel to the Vila Franca de Xira-Lisboa fault system.

Acknowledgements The authors would like to thank the Nucleus for Oil Research and Exploration (NPEP) of the former Instituto Geolo´gico e Mineiro (IGM) for allowing publication of seismic and well data, the Geophysics Division of the former IGM for support to this project, Dr Julio Branco from Petroprimo, S.A. for allowing the use of facilities and synthetic seismograms software and the Portuguese Foundation for Science and Technology for supporting the SHELT (Seismic Hazard Evaluation of the Lower Tagus Valley) project PRAXIS/P/CTE/11178/1998 and supplying a PhD grant to one of the authors. We thank Rui Baptista and Jorge Carvalho for helpful discussions on the interpretation of the seismic profiles. The authors are additionally grateful to E.S. Rasmussen for his criticism and comments that greatly improved the early version of the paper. We also acknowledge the Editor D.G. Roberts for the encouragement to the publication of this paper. We thank an anonymous reviewer for his improvements to the original manuscript and Daniel de Oliveira for corrections to the English version of the paper.

References Alves, T.M., Gawthorpe, R.L., Hunt, D.W., Monteiro, J.H., 2002. Jurassic tectono-sedimentay evolution of the Northern Lusitanian Basin (offshore Portugal). Marine and Petroleum Geology 19, 727–754. Anstey, N.A., Siraki, E.S., Montgomery, S.L., 1992. Fault Interpretation. GP502, IHRDC, Boston. Antunes, M.T., Pais, J., 1992. The Neogene of Portugal. Cieˆncias da Terra, Special Issue II, 13–24. Barbosa, B.A., 1995. Alostratigrafia e litostratigrafia das unidades continentais da Bacia Tercia´ria do baixo Tejo. Relac¸o˜es com o Eustatismo e a Tecto´nica. PhD Thesis, Univ. Lisboa, p. 252.

Canerot, J., Rey, J., Baptista, R., Manuppella, G., Peyberne´s, B., 1995. Nouvelles interpre´tations structuralles et geodynamiques de la marge atlantique portuguaise dans le sector de Caldas da Raı´nha (Portugal). C.R. Acad. Sci. Paris 320, 523–530. Carter, M.D., Siraki, E.S., 1993. Velocity Interpretation and Depth Conversion. GP504, IHRDC, Boston, p. 141. Carvalho, J.P.G., 1995. Estudo da Zona Sul Portuguesa e da Margem Atlaˆntica Adjacente a partir de Dados Geofı´sicos. MSc Thesis, Univ. Lisboa, p. 154. Carvalho, J.P.G., 2003. Sı´smica de Alta Resoluc¸a˜o Aplicada a` Prospecc¸a˜o, Geotecnia e Risco Sı´smico. PhD Thesis, Univ. Lisboa, p. 250. Domsalski, W., 1969. Interpretation of an Aeromagnetic Survey Offshore Portugal (1:200,000). Fairey Surveys Limited Report. Ellis, P.M., Wilson, R.C.L., Leinfelder, R.R., 1990. Controls on Upper Jurassic carbonate buildup development in the Lusitanian BasinPortugal, In: Tucker, M.E., Wilson, J.L., Crevello, P.D., Sarg, J.R., Read, J.F. (Eds.), International Associations of Sedimentologists Special Publications, 9, pp. 169–202. GPEP (Gabinete para a Pesquisa e Explorac¸a˜o de Peto´leos), 1986. Petroleum Potential of Portugal. Lisboa, p. 62. Japsen, P., 1993. Influence of lithology and Neogene uplift on seismic velocities in Denmark: implications for depth conversion of maps. American Association of Petroleum Geologists Bulletin 77 (2), 194–211. Koll, H., Vasconcelos, M., 2000. Adensamento Gravime´trico: Estudos e Aplicac¸o˜es. Livro de Resumos da 2a Assembleia Luso-Espanhola de Geodesia e Geofı´sica, 75–76. Kullberg, J.C.R., 2000. Evoluc¸a˜o Tecto´nica Mesozo´ica da Bacia Lusitaniana. PhD Thesis, Univ. Lisboa, p. 280. Leinfelder, R.R., 1994. Carbonate Platforms and Coral Reefs within a Siliciclastic Setting (Upper Jurassic, Lusitanian basin, Portugal). Profil 6, 1–207. Stuttgart University. Leinfelder, R.R., Wilson, R.C.L., 1989. Seismic and sedimentologic features of the Oxfordian–Kimmeridgian syn-rift sediments on the eastern margin of the Lusitanian Basin. Geologische Rundschau 78, 81–104. Leinfelder, R.R., Wilson, R.C.L., 1998. Third-order sequences in an Upper Jurassic Rift-related second-order sequence, Central Lusitanian Basin, Portugal. In: Graciansky, P.C., Hardenbol, J., Jacquin, T., Vail, P.R. (Eds.), Mesozoic and Cenozoic Sequence Stratigraphy of European Basins SEPM Special Publication, 60, pp. 507–525. Lomholt, S., Rasmussen, E.S., Andersen, C., Vejbaek, O.V., Madsen, L., Steinhardt, H., 1995. Seismic interpretation and mapping of the Lusitanian Basin, Portugal. Final Report, Proj. MILUPOBAS, Contract No. JOU-CT94-0348, GEUS, Copenhagen, p. 70. Manuppella, G. (coord.), et al., 2000. Carta Geolo´gica de Portugal na escala de 1/50,000-Notı´cia explicativa da folha 27-A (Vila Nova de Oure´m). Instituto Geolo´gico e Mineiro. Manuppella, G. (coord.), et al., in press. Carta Geolo´gica de Portugal na escala de 1/50,000-Notı´cia explicativa da folha 30-B (Bombarral). Inst. Nac. Eng. Tecn. Inov. Manuppella, G., Azeredo, A.C., 1996. Contribuic¸a˜o para o conhecimento da regia˜o de Sesimbra. Comun. Inst. Geol. Min. 82, 37–50. Manuppella, G., Balaco´, J.C., Grac¸a e Costa, J.R., Crispim, J.A., 1985. Calca´rios e Dolomitos do Macic¸o calca´rio Estremenho. Estudos Notas e Trabalhos do Servic¸o de Fomento Mineiro, Tomo 27, 3–48. Marsden, D., 1989. Layer cake depth conversion. The Leading Edge 1, 10–14. Mauffret, A., Mougenot, D., Miles, P.R., Malot, J.A., 1989. Cenozoic deformation and Mesozoic abandoned spreading center in the Tagus Abyssal Plain (West Portugal): results of a multichannel seismic survey. Can. J. Earth Sci. 26 (6), 1101–1123. Montenat, C., Guery, F., Jamet, M., 1988. Mesozoic evolution of the Lusitanian Basin: comparison with the adjacent margin. In: Boillot, G., Winterer, E.L. (Eds.), Proceedings of the Ocean Drilling Program, Scientific Results, Washington, 103. Ocean Drilling Program, College Station, pp. 757–775.

J. Carvalho et al. / Marine and Petroleum Geology 22 (2005) 427–453 Mouterde, R., Ramalho, M., Rocha, R.B., Ruget, C., Tintinant, H., 1971. Le Jurassique du Portugal. Esquisse stratigraphique et zonale. Bol. Soc. Geol. Portugal 18, 73–104. Oliveira, J.T. (coord.), et al., 1992. Carta Geolo´gica de Portugal (escala 1/500,000). Servic¸os Geolo´gicos de Portugal. Pena dos Reis, R., et al., 1998. Excursa˜o 1—O Mesozo´ico da Bacia Lusitaˆnica, In: Oliveira, T., Dias, R. et al. (Eds.), O Livro Guia das Excurso˜es, V Congresso Nacional de Geologia. Instituto Geolo´gico e Mineiro, , pp. 9–76. Ramalho, M.M., 1971. Contribuition a` l’e´tude micropale´ontologique et stratigraphique du Jurassique supe´rieure et du Cre´tace infe´rieure des environs de Lisbonne (Portugal). Mem. Serv. Geol. Portugal 19, 112. Rasmussen, E.S., Lomholt, S., Anderson, C., Vejbaek, O., 1998. Aspects of the structural evolution of the Lusitanian Basin in Portugal and the shelf and slope area offshore Portugal. Tectonophysics 300, 199–225. Reis, R.P., Proenc¸a Cunha, P., Dinis, J.L., 1997. Hipersubsient depositional event associated with a rift climax in late Jurassic of Lusitanian Basin (W Portugal). IV Congreso Jura´sico de Espan˜a, Alcaniz, 101–103. Rey, J., 1992. Les unite´s lithostratigraphiques du Cre´tace infe´rieur de la re´gion de Lisbonne. Com. Serv. Geol. Portugal 78 (2), 103–124. Rey, J., 2003. Les se´quences de de´pot dans le Cre´tace´ Infe´rieur du Bassin Lusitanine. Com. Serv. Geol. Portugal 90, 15–42. Ribeiro, A., Kullberg, M.C., Kullberg, J.C., Manupella, G., Phipps, S., 1990. A review of Alpine tectonics in Portugal: Foreland detachment in basement and cover rocks. Tectonophysics 184, 357–366.

453

Ribeiro, A., Silva, J.B., Cabral, J., Dias, R., Fonseca, P., Kullberg, M.C., Terrinha, P., Kullberg, J.C., 1996. Tectonics of the Lusitanian Basin. Final Report, Proj. MILUPOBAS, Contract No. JOU-CT94-0348, ICTE/GG/GeoFCUL, Lisboa, p. 126. Rocha, R.B., Marques, B., Kullberg, J.C., et al., 1996. The 1st and 2nd Rifting Phases of the Lusitanian Basin: Stratigraphy, Sequence Analysis and Sedimentary Evolution. Final Report, Proj. MILUPOBAS, Contract No. JOU-CT94-0348, UNLIS, Lisboa, p. 80. Schlumberger, 1988. Log interpretation charts. Schlumberger Educational Services, USA, p. 151. Torres, L.M., submitted. Contribuic¸a˜o dos Me´todos Potenciais para o Estudo Estrutural da Zona Sul Portuguesa. PhD Thesis, Univ. Lisboa. Watkinson, M.P., 1989. Triassic to Middle Jurassic sequences from the Lusitanian Basin Portugal, and their equivalents in other North Atlantic margin basins. PhD Thesis, Open University, Milton Keynes, p. 390. Wilson, R.C.L., Hiscott, R.N., Willis, M.G., Gradstein, F.M., 1989. The Lusitanian Basin of west central Portugal: Mesozoic and Tertiary tectonics, stratigraphy and subsidence history. In: Tankard, A.J., Balkwill, H. (Eds.), Extensional Tectonics and Stratigraphy of the North Atlantic Margins American Association Petroleum Geologists Memoir, 46, pp. 341–361. Wuenschel, P.C., 1960. Seismogram synthesis including multiples and transmission coefficients. Geophysics 25, 106–129.