This section focuses on the timing of rifting, regional subsidence, and compressional events that contributed to the formation of the present-day structure of the west Iberia Margin.
Two problems must be faced in compiling information to compare the timing of tectonic episodes along the margin. The first is the fact that comparable data sets are not available from different sectors (Table 1): all have varying degrees of seismic profile coverage, but it is only in the Lusitanian Basin that profiles across the rift-basins have been calibrated using borehole data. The other problem concerns the criteria used by previous researchers to identify synrift packages on seismic sections. As shown by Wilson et al. (this volume) nearly all previously identified synrift packages in the offshore areas are probably postrift in origin.
Figure 7 is a summary of the Mesozoic and Tertiary stratigraphy of the west Iberia Margin. It shows in the right-hand column thirteen unconformity-bound sequences (UBS) defined by Wilson et al. (1989), Cunha (1992), and Pena dos Reis et al. (1992) that are related to significant events in the geological history of the margin. The discussion below summarizes the timing of tectonic events that can be inferred from this stratigraphic record.
Three main rifting episodes have affected the west Iberia Margin: (a) a Late Triassic to Early Jurassic (Hettangian) rifting event, (b) a late Oxfordian/early Kimmeridgian rifting event well documented in the Lusitanian Basin (Wilson et al., 1989), and (c) a Valanginian/ Hauterivian to Aptian rifting episode that culminated with the formation of oceanic crust and the separation of Iberia from the Grand Banks of Canada. These rifting episodes are well recorded in the stratigraphic record. Wilson et al. (1989) identified four Mesozoic unconformity-bound sequences in the Lusitanian Basin, which they related to events in the early evolution of the North Atlantic. Their paper reviewed in detail the tectonic implications of the three oldest of these sequences, and so only a brief summary of their significance is given below. This is followed by additional information concerning the Aptian–Campanian sequence.
1. Triassic-Callovian. This sequence is typical of the early riftsag successions encountered in most North Atlantic Margin basins. In the Lusitanian and Porto Basins this sequence is one to two kilometers thick. Triassic red fluvial siliciclastics are capped by Hettangian evaporites that subsequently influenced the manner in which extensional and compressional movements along Variscan basement faults affected the cover of younger sediments. Where the evaporites were thick, halokinetic structures formed above basement faults, but where they were thin, the faults propagated through younger formations. The Triassic and Hettangian sediment probably accumulated in grabens and half grabens. The later Lower and Middle Jurassic shales and carbonates exhibit simple facies geometries indicating a westward-dipping ramp, with some localized indications of contemporaneous faulting related to the Berlenga ridge. The Late Triassic–Early Jurassic rifting event resulted in thick salt deposition along the Portuguese and Grand Banks Margins (Pautot et al., 1970; Tankard and Welsink, 1987; Grant et al., 1988; Mougenot, 1988; Austin et al., 1989; Mauffret et al., 1989a,b). Soares et al. (1993) divided the Triassic–Callovian of the northern part of the Lusitanian Basin into eight megasequences bounded by discontinuities. They suggested that rifting continued into the early Liassic, whereas Wilson et al. (1989) indicated that thermal subsidence began at the beginning of the Jurassic.
2. Oxfordian–Tithonian. In the Lusitanian Basin, the base of this sequence is marked by a basin-wide late Callovian–early Oxfordian hiatus, associated with karst features. The middle Oxfordian is characterized by lacustrine and marginal marine carbonates and clastics, and some evaporites. Fully marine carbonate deposition returned to the Lusitanian Basin in the late Oxfordian, during which time apparent basement subsidence rates increased up to 200 m/m.y. (Hiscott et al., 1990). A sudden influx of coarse siliciclastic sediments at the end of the Oxfordian was accompanied by a further increase in subsidence rates (to 270 m/m.y. in one sub-basin), and the first significant diapirism occurred at this time. In the southern part of the Lusitanian Basin, half-graben basins developed in which Kimmeridgian siliciclastics clearly indicate the development of syn- and immediate postrift basin fill (Leinfelder and Wilson, 1989; in press). Rapid basement subsidence rates and a complex distribution pattern of early Kimmeridgian facies support the conclusion that this interval marks a significant transtensional rifting phase (Wilson et al., 1989). Younger sediments exhibit simpler facies distributions characteristic of a postrift episode, and carbonate shelf systems developed over Galicia Bank (Boillot, Winterer, Meyer, et al., 1987; Jansa et al., 1988), in the Porto Basin and in the southern part of the Lusitanian Basin. The complete sequence is 2–4 km thick in the Lusitanian Basin (1–3 km synrift and approximately 1 km postrift), but less than 2 km in the Porto Basin, which suggests that rift-related subsidence was less significant in the north than it was to the south.
3. Valanginian–Early Aptian. On shore and in the Porto basin, this interval shows a relatively simple facies geometry, with largely fluvial siliciclastic sands and conglomerates interfingering with shallow water carbonates. It is relatively thin (up to 500 m), indicating relatively low subsidence rates. On the Galicia Margin, the sediments of this age are interpreted as synrift sediments that heralded ocean opening in late Aptian times between Iberia and the Grand Banks. The source of the Lower Cretaceous sediments was probably the Galicia Bank, since sediments eroded from the Iberian mainland were trapped in the Interior Basin (Winterer et al., 1988). This episode is also well documented on the Galicia Bank margin (Boillot et al., 1985, 1989b; Boillot and Winterer, 1988; Moullade et al., 1988; Murillas et al., 1990; Malod and Mauffret, 1990). However, Wilson et al. (this volume) conclude that the intervals previously recognized as synrift beneath the deep Galicia Margin are in fact immediate postrift, as defined by Prosser (1993), as no convincing reflection divergence towards fault planes can be observed on published seismic sections from this area. They suggest that depositional rates during the synrift episode were extremely low, so that sequences thick enough to image by reflection seismic methods did not accumulate. In the light of this conclusion, they suggest that the synrift episode occurred after the deposition of the Tithonian–Berriasian carbonate platform on the Galicia Margin, but before the Valanginian siliciclastic sediments previously identified as synrift. This interpretation indicates a very short-lived period of rifting off the deep Galicia Margin, spanning a few million years from the latest Berriasian to earliest Valanginian.
4. Late Aptian–Early Campanian. In the Lusitanian Basin the base of this sequence is marked by the abrupt and widespread onset of coarse-grained siliciclastic sediments. Sediment was transported southwestward from the Hesperic Massif across coalescent alluvial fans that developed in a humid environment; these merge in the same direction into coastal marine sediments, including carbonates (Berthou, 1973; Soares, 1980; Dinis and Pena dos Reis, 1989). During the Turonian, localized movement of diapiric structures is indicated by an unconformity, karstification, and locally derived sediments (Pena dos Reis, 1993). Berthou (1973), Berthou and Lauverjat (1979), and Boillot et al. (1972; 1975) suggested that reactivation of the Nazaré-Lousã Fault at the end of the Cenomanian resulted in uplift of areas to the south of the structure, causing a northward displacement of marine sedimentation. However, we believe that the main movement was later, probably during the middle Campanian, since the upper Campanian fluvial sediments show a northerly drainage pattern in contrast to the earlier southwesterly direction. The top of the upper Aptian–Campanian sequence in the northwestern margin of the Lusitanian Basin is characterized by a thick silcrete, which indicates a long period of nondeposition during the Santonian?–early Campanian (Pena dos Reis and Cunha, 1989a; b; Cunha et al., 1992).
The overall tectonic style of the continental margin deformation acquired during the three main rifting episodes is essentially characterized by a series of tilted fault blocks bounded by normal faults (mostly dipping to the west), which delineate a series of half-graben structures (Mauffret and Montadert, 1987; Thommeret et al., 1988; Sibuet, 1992). Off Galicia, these half-graben structures trend predominantly north-south, as shown by the available detailed seismic and seabeam coverage (Sibuet et al., 1987). In the southern Iberia Abyssal Plain and in the Tagus Abyssal Plain, however, the trend of the continental blocks is more variable (Whitmarsh et al., 1990; Pinheiro et al., 1992; Pinheiro, 1994; Whitmarsh et al., 1995). These tilted fault blocks can be continuous for a few tens of kilometers off Galicia, but they appear to be more intensely disrupted further south, being generally offset by northeast-southwest transverse faults that probably correspond to late Variscan continental strike-slip faults reactivated during rifting (Ribeiro et al., 1979; Sibuet, 1992; Ribeiro and Silva, in press). The faulting associated with rifting was accompanied by large-scale mass movements from the crests of the tilted blocks.
Off Galicia, the end of continental extension has been dated at approximately the Aptian–Albian boundary (112 Ma; Boillot et al., 1989b). However, Wilson et al. (this volume) argue that synrift sediments cannot be identified on the available seismic data and suggest that the continental extension in this segment of the margin may have ended in the late Berriasian–early Valanginian.
The mechanism of lithospheric extension is still not clear, but it seems that a composite model involving pure-shear affecting the whole lithosphere and simple-shear affecting only the upper brittle crustal levels along a low-angle detachment that soles out at the brittle/ductile transition can best explain the observations (Sibuet, 1992). A very strong seismic reflector imaged on the available seismic sections west of Galicia Bank—the S-reflector—is a good candidate for such a detachment (Boillot et al., 1989b; Sibuet, 1992; Reston et al., 1994; Sibuet et al., in press). A similar reflector was first detected in the Bay of Biscay, both on the Armorican and the Galician Margins and was called the S-reflector by de Charpal et al. (1978), who interpreted it as representing the brittle/ductile transition. Later, Boillot et al. (1989b) applied the simple shear model, initially proposed by Wernicke (1981) to explain the Basin and Range province, to the Galicia area and postulated that the S-reflector represents the contact between the thinned continental crust and serpentinized peridotite. More recently, Hoffman and Reston (1992) and Reston et al. (1994) confirmed the detachment hypothesis and Sibuet et al. (in press) have shown that this supposedly low-angle normal fault is not rooted in the mantle, although it possibly coincides with the top of the serpentinized peridotite layer in the lowermost part of the continental margin, close to the peridotite ridge. Further to the east, however, it appears to sole out in the middle crust (Sibuet et al., in press). In the Iberia Abyssal Plain, Beslier et al. (1993) have shown that a fairly strong S-like reflector, or set of reflectors, also occur locally at approximately the same depth as the S-reflector off Galicia, but their expression on the seismic sections is not as dramatic and they are far more discontinuous. No evidence of an S-like reflector has yet been found in the Tagus Abyssal Plain.
As regards the peridotite ridge, detailed structural and microstructural studies have shown that it was emplaced at seafloor level at the end of continental rifting or at the very beginning of oceanic accretion (Beslier et al., 1990). It successively experienced (1) a low degree of partial melting and strong ductile high-temperature–low-stress deformation under asthenospheric conditions during adiabatic uplift; (2) strong, but heterogeneous ductile deformation (mylonitization and ultramylonitization) under lithospheric conditions (high but decreasing temperature of 1000°–850°C and high deviatoric stress of 180 MPa); (3) brittle deformation; and (4) serpentinization by hydrothermal alteration by seawater (Girardeau et al., 1988; Beslier et al., 1990; Boillot et al., 1992; Beslier et al., 1993). The age of emplacement of the peridotite ridge off Galicia has been dated at 118–122 ± 0.9 Ma (early Aptian). This corresponds to 39Ar/40Ar dating of amphibole crystals and plagioclase neoblasts within a syntectonic dike that crosses the peridotite and which has undergone ductile deformation under a similar stress regime and physical conditions as the peridotite (Féraud et al., 1988; Boillot et al., 1992).
The Pyrenean and Alpine orogenies in Iberia were the result of collisions between Iberia and Europe, and Africa and Iberia, respectively. In Iberia and along the west Iberia Margin, the most important compressive deformation occurred during the Miocene, although Eocene deformation occurred, particularly in the Hesperic massif (Boillot et al., 1979; Ribeiro, 1988, Ribeiro et al., 1990; Malod et al., 1993; Masson et al., 1994). The Eocene deformation episode is essentially related to the formation of the Pyrenean-Cantabrian Chain and to the incipient subduction developed in the southern Bay of Biscay (Sibuet and Le Pichon, 1971; Grimaud et al., 1982; Malod et al., 1993). The Miocene event is fundamentally related to the deformation of the Betic chain in southeastern Iberia (Sibuet and Le Pichon, 1971; Murillas et al., 1990).
In the Lusitanian and Algarve Basins, the Miocene tectonic style is "thin-skinned," with the Hettangian evaporites acting as a décollement (Ribeiro, 1988; Ribeiro et al., 1990). In contrast, structures affecting basement rocks probably have deeper roots, possibly at the base of the crust. The Central Cordillera, for example, has been interpreted as a basement pop-up structure (Ribeiro, 1988; Ribeiro et al., 1990), the trend of which continues southwestward into the major zone of inverted Mesozoic rocks in the Lusitanian Basin (see Fig. 8). In this basin, the Miocene compression was directed north-north-west–south-southeast, oblique to the predominant north-northeast-south-southwest trend of Mesozoic extensional faults. This resulted in basin inversion with a strong transpressional component (Wilson et al., 1989).
The main tectonic events that affected the west Iberia Margin in the Cenozoic are well documented by several unconformity-bound sequences (UBSs; 5–13 in Fig. 7; Cunha, 1992, Pena dos Reis et al., 1992). These are discussed briefly below.
This interval consists of two major sequences: (1) an upper Campanian–Maastrichtian sequence and (2) a Paleocene–lower Lutetian sequence; these are separated by an unconformity with a stratigraphic hiatus offshore and, probably, also onshore. This interval may have been influenced by the changing movement of Iberia relative to Europe, with the beginning of the compression at the northern border of Iberia (Gräfe and Wiedmann, 1993). The reactivation of the Nazaré-Lousã Fault is signalled by a change in fluvial paleocurrent directions from southwesterly drainage during the earlier late Aptian–Campanian, to northwesterly during this interval. Significant volcanic activity occurs south of the Nazaré Fault, particularly in the Lisbon and adjacent areas. Here, subvolcanic complexes, basaltic extrusives, and associated dikes occur. Kullberg (1985), suggested that the alkaline ring complexes were emplaced along a deep north-northwest–south-southeast strike-slip zone. According to Ribeiro et al. (1985, 1990) rift migration elsewhere in the North Atlantic may have changed the stress field and deeply fractured the previously thinned continental margin.
Onshore, Maastrichtian yellowish quartzarenites and red mudstones (Antunes, 1979), interpreted as deposited by a meandering fluvial system draining to the northwest, change distally to transitional and marine environments (Pena dos Reis, 1983). Correlative diapiric reactivation built up coalescent peridiapiric alluvial fans, transverse to the north-south salt structures. Offshore, very shallow marine fine siliciclastics and dolostones are dominant.
During this interval compression intensified in the Pyrenees (Srivastava et al., 1990a,b). On the western border of the Hesperic Massif, extensional reactivation of northeast-southwest faults caused the formation of the Mondego and lower sub-basins (Cunha, 1992). In some areas, the lower boundary corresponds to an angular unconformity. A prograding conglomerate sandstone at the top of this interval is related to the uplift of the Hesperic Massif (Pena dos Reis, 1983). The sands and conglomerates show a broad northeast to south-west drainage pattern, which contrasts with the underlying Upper Cretaceous paleocurrent trend. The upper boundary is very well defined by a thick silcrete (Meyer and Pena dos Reis, 1985).
During this period, Iberia joined the Eurasian Plate and the plate boundary with Africa migrated to the Azores-Gibraltar Fracture Zone (Fig. 1F). The sedimentary record can be divided into two unconformity-bounded sequences: (1) an upper Chattian–upper Burdigalian sequence, and (2) an upper Burdigalian–lower Tortonian sequence. The lower boundaries of these sequences correspond to the Castillan and Neocastillan tectonic phases, respectively (Pena dos Reis and Cunha, 1989b). The lower sequence is only preserved beneath the continental shelf in the Tagus sub-basin, and consists of marine, brackish, and continental facies (Antunes et al., 1987). The upper sequence occurs in both the Tagus and Mondego sub-basins, and includes fine sands, plus lacustrine and shallow marine carbonates.
During this interval, three alluvial fan sequences in the Mondego sub-basin are interpreted to result from distinct pulses of uplift (UBS 11: late Tortonian–early Messinian; UBS 12: late Messinian–Zandian; UBS13; Piazencian–early Calabrian?). Carvalho et al. (1983) and Ribeiro et al. (1990) suggested that the main uplift of the Central Cordillera (Fig. 8C) occurred during the Tortonian, although stratigraphic relationships in the Serra da Arrábida indicate that the principal deformation occurred around the early middle Miocene boundary. UBSs 11–13 are affected by neotectonic structures, and precede Quaternary fluvial incision.
Offshore, there is strong evidence of Cenozoic reactivation of late Variscan faults as reverse faults or transpressive strike-slip structures, particularly during the Eocene and the Miocene (Boillot et al., 1979; Mauffret et al., 1989a; Masson et al., 1994; Pinheiro, 1994). This deformation appears to continue today, as suggested by localized presence of seafloor scarps above compressional structures within the ocean/continent transition, accompanied by deformation of recent sediments (Pinheiro, 1994). Off Galicia Bank, the peak of the deformation occurred in the Eocene (Boillot et al., 1979). In the Tagus and in the southern Iberia abyssal plains, however, the main episode of deformation occurred in the middle Miocene, and it may have overprinted the former Eocene deformation (Mougenot, 1988; Mauffret et al., 1989a,b; Masson et al., 1994; Pinheiro, 1994). The Cenozoic deformation in the abyssal domain is essentially characterized by a broad oceanward-facing monoclinal sequence that can be several tens of kilometers wide, and which invariably terminates in a tight asymmetric fold, generally associated with a steeply dipping fault or faults (Masson et al., 1994). In general, such faults do not seem to be directly rooted in any major basement faults. In both the Iberia and the northern Tagus abyssal plains, the overall pattern of the observed deformation is compressional (or transpressional). However, there is evidence of a later extensional or transtensional regime, that could be local, but which has caused a tectonic subsidence/colapse, in places, of the original compressional structures (Pinheiro, 1994). This latter episode apparently preceded the deposition of the seismic stratigraphic Unit 1A (late Miocene to Holocene) defined by Mauffret and Montadert (1988), and therefore also occurred during the middle Miocene.
As shown by Masson et al. (1994) and Pinheiro (1994), there is a close spatial link between the western termination of the deformation zone and the ocean/continent transition. In the southern Iberia Abyssal Plain, the deformation zone generally ends abruptly at the pre-sumed contact between the transitional crust that characterizes the ocean/continent transition and the thin oceanic crust. This coincides with the location of the serpentinized peridotite ridge that marks the ocean/continent transition in this area and which was drilled during this leg at Site 897. There are some seismic lines on the Tagus and on the Iberia Abyssal Plains, however, in which the deformation zone appears to terminate just west of the presumed location of the serpentinized peridotite ridge, over what has been interpreted as an area of thin oceanic crust underlain by serpentinized upper mantle (Pinheiro, 1994). Smaller, but similar, compressional (transpressional?) structures also occur much further to the east, in an area of the ocean/continent transition interpreted as thinned continental or transitional crust underlain by serpentinized upper mantle (Pinheiro, 1994). This strongly suggests that the contrasts in lithologies and rheologies that take place across the ocean/continent transition (as inferred from the interpretation of the geophysical data: see Whitmarsh et al., 1990; Pinheiro et al., 1992; Whitmarsh et al., 1993; Masson et al., 1994; Pinheiro, 1994), together with the fact that the crustal section is very thin and it is underlain by serpentinized peridotite, have probably played an important role in controlling the fold propagation and the pattern and lateral extent of the deformation zone. The existence of large volumes of serpentinized peridotite at shallow depth, in particular, appears to have been the crucial factor on the control of the fold propagation, since it is well known that these rocks, under stress, tend to deform in a fairly ductile fashion and to accommodate a significant part of the deformation (Murton, 1986; Reinen et al., 1991).
The observations discussed above suggest the reactivation, in an essentially transpressional strike-slip regime related to the stress field generated by the oblique northwest-southeast Africa-Iberia-Eurasia convergence during the Cenozoic, of the north-northeast-south-southwest normal faults that controlled the Mesozoic rifting. The change from transpression to transtension observed on some of the available seismic lines could be local and it may have been caused by small adjustments in the plate motion at the end of the middle Miocene. Other possible concurrent causes could be phenomena of local flow within the underlying serpentinite, or the formation of localized pull-apart zones in the areas of relay (or transfer) between the overlapping adjacent fault segments, caused by the interaction between the slip movement on the faults (Pinheiro, 1994).
Magmatic activity in western Iberia during the Mesozoic is present both in the Hesperic Massif (Central Iberian Zone) and in the Lusitanian Basin. The main areas of such activity in Portugal are shown in Figure 9; they include dikes of lamprophyres, dolerites, and basalts (alkaline, transitional, and tholeiitic) as well as minor hypovolcanic massifs (Ribeiro et al., 1979; Martins, 1991). The occurrences, however, are volumetrically insignificant in comparison to volcanism that preceded and accompanied continental rifting and separation between northwest Europe and Greenland.
Mesozoic igneous rocks are largely confined to the west Iberia Margin south of the Nazaré Fault, and none have been found north of Figueira da Foz. Four main phases of Mesozoic magmatic activity have been identified:
1. A Late Triassic–Early Jurassic alkaline basaltic cycle (203–224 Ma)
2. An Early Middle Jurassic tholeiitic cycle (190–160 Ma)
3. An Early Cretaceous transitional cycle (135–130 Ma)
4. A Late Cretaceous alkaline cycle (100–70 Ma)
The first three cycles are associated with the three main rifting episodes (Ribeiro et al., 1979; Martins, 1991).
The oldest magmatic phase is restricted to the Central Iberian Zone and is represented by dolerite and porphyry dikes associated with the Variscan granites, as well as by several occurrences of lamprophyres and dolerite-basaltic dikes with K-Ar ages of 224 ± 11 Ma (Late Triassic). These suggest an alkaline basaltic activity with a shoshonitic tendency (Ribeiro et al., 1979), probably related to the last phases of the Variscan orogeny in Iberia. The dikes appear to intrude along fractures possibly related to postorogenic extension (Ribeiro et al., 1979). Basaltic dikes with ages of 203 ± 3 Ma (Hettangian/Sinemurian) have also been reported in this area (Ribeiro et al., 1979; Martins, 1991). All the ages referred to in this section ("Magmatic Activity in the Mesozoic") have been taken from syntheses in Ribeiro et al. (1979) and Martins (1991).
This episode is represented in western Iberia by the intrusion of basaltic and doleritic dikes within the Hettangian–Sinemurian sedimentary formations of the Mesozoic–Cenozoic Borderland (Algarve, Bordeira, Santiago do Cacém and Sesimbra) and by the Messejana dikes along the Odemira-Ávila Fault (Fig. 9). There is also clear evidence of Early Jurassic tholeiitic magmatism in the Algarve (190 to 180 Ma; Ribeiro et al., 1979; Ferreira and Macedo, 1979). These tholeiitic volcanics characterize an area with similar sedimentological and volcanic character to the margins of the Central Atlantic (i.e., the northwest African Margin from Senegal to Morocco and the northeast United States from the Bahamas to Nova Scotia; May, 1971; Ribeiro et al., 1979; Martins, 1991). Based on the ages obtained for the intrusions along the Odemira-Ávila Fault, Martins (1991) proposed that the tholeiitic activity could extend from 190 to 160 Ma (early Middle Jurassic), with most of the magmatic activity concentrated in the period 190–180 Ma. Judging from stratigraphic relationships, however, it is possible that the magmatic activity might have started somewhat earlier, at about 205 Ma (Rhaetian–Hettangian boundary; Martins, 1991).
Late Jurassic–Early Cretaceous intrusives occur mainly in the Lusitanian Basin (Ribeiro et al., 1979; Martins, 1991) as north-north-east–south-southwest and east-southeast–west-northwest dikes (consisting of dolerites, gabbros, and diorites) and also as minor hypovolcanic massifs in the Leiria area. The intrusions are mostly confined to an area between two major tectonic lineaments: the Nazaré Fault (NF) and the Serra d'Aire-Montejunto lineament (SAML) (Fig. 9; Ribeiro et al., 1979). Evidence of such activity between the Oxfordian and the Hauterivian has been recognized in Soure, Codiceira, Vermoil, Leiria, and Rio Maior (Fig. 9). The basaltic dike of Cacela (126 Ma) in the Algarve could also belong to this cycle (Ribeiro et al., 1979). With the exception of Soure (159 ± 3 Ma) and Codiceira (165 ± 3 Ma), all the K-Ar ages for these intrusions range between 130 ± 3 and 144 ± 2 Ma, and concentrate mainly in the interval 135–130 Ma. Martins (1991) and Ferreira and Macedo (1987) considered the ages of 159 and 163 Ma obtained for Soure and Codiceira, respectively, as not reliable, due to alteration of the samples. Therefore, according to these authors, the magmatic activity in this period is concentrated in the interval 135–130 Ma (Valanginian/Hauterivian), although Willis (1988) reports a 140 Ma age (Berriasian) for east-southeast–west-northwest dikes to the east of Caldas da Rainha.
The main onshore occurrences related to this episode include the radial dike complex near Mafra (ca. 100 Ma), the Basaltic Complex of Lisboa (72.5±3 Ma), and the Sines, Sintra, and Monchique massifs (80 to 70 Ma) (Fig. 9). The subvolcanic massifs of Sintra, Sines, and Monchique show ages that decrease to the south (Ribeiro et al., 1979) and a similar suggestion of a decrease in age of the alkaline magmatism towards the south is also suggested by the relative ages of the basaltic complex of Lisboa (K-Ar isochron of 72.5±3 Ma), compared with the Mafra complex (about 100 Ma; Fig. 9; Ribeiro et al., 1979). Other volcanic manifestations of the same period include a few outcrops NW of the Serra d'Aire-Montejunto lineament and a few volcanic plugs (e.g., in the Nazaré area; Ribeiro et al., 1979). Evidence of alkaline volcanism is also found in the western Algarve, where K-Ar whole rock ages between 72 and 77 Ma have been reported (Martins, 1991). In the offshore, there is also evidence of alkaline volcanism. Alkaline basalts with an 40Ar/39Ar age of 74 ± 0.7 Ma were recovered from the Estremadura Spur and alkaline rocks of Maastrichtian–Paleocene age (60–74 Ma), with a similar petrology to those at the Sines and Monchique, overlie oceanic gabbros in Gorringe Bank (Féraud et al., 1977; Mougenot, 1988).
In summary, it seems clear that the Mesozoic volcanic igneous activity in western Iberia is scattered and relatively minor (Fig. 9). In fact, the synrift volcanism associated with the final rifting event that culminated with seafloor spreading is confined to a small area between the Nazaré Fault and the Serra d'Aire-Montejunto Lineament and appears to have occurred essentially in the period 135-130 Ma (Valanginian/Hauterivian). As such, the west Iberia Margin can be considered a nonvolcanic margin.