Evolution of the Mesozoic Central Iberian Rift System ...

4

Evolution of the Mesozoic Central Iberian Rift System and its Cainozoic inversion (lberian chain) Ramon SAIAsrn, loan GUIMERA (2), Ramón MAS r3~ Caries MARTÍN-CLOSAS (4), Alfonso MELÉNDEZ rs; & Angela ALONSO (6) M

Cameros Basin Sedimentation Area Depocentral Areas Marine influence Carbonate facies (lacustrine) Evaporitic facies (lacustrine) Siliciclastic facies (alluvial)

50 Km

Main paleocurrent trend

FIG. 12.- Schematic palaeogeographic evolution of the Cameros Basin. BU, Burgos; LO, Logroño; SO, Soria. Modified after MAS et al. (1993).

FIG. 12.- Schéma de l'évolution paléogéographique du bassin de Cameros. BU, Burgos; LO, Logroño; SO, Soria. Modifiée d'apres MAS et al. (1993).

162

RAMON SALAS ET AL.

along the basin margins by estuarine shallow-water carbonate platforms (large freshwater discharges) where carbonate production was dominated by molluscs and calcareous algae (Fig. 10). Marginal oolitic-bioclastic shoals and coralgal boundstones are abundant. The earliest Aptian (lower part of K1.8) consists of an up to 100 m thick tidally-dominated low-stand delta complex, the upper delta plain deposits of which are rich in dinosaur remains (Morella Formation). During the Aptian (Kl.8, K1.9) marine conditions prevailed again and up to 1100 m thick very expansive, rapidly prograding shallowwater carbonate platforms developed, characterized by Orbitolines, calcareous algae and rudists. The early to middle Albian sequence (K1.10) consists of a very extensive tidally-influenced delta system, up to 500 m thick, which contains thick coallayers (Escucha Formation; QUEROL et al., 1992). The Cameros Basin differs fundamentally from the Maestrát Basin in its extensional geometry and in its sedimentary fill. The latter consists predominantly of alluvial and lacustrine deposits, containing only very rare marine incursions (GÓMEZ FERNÁNDEZ, 1992; ALONSO & MAS, 1993). The Cameros Basin displays a pre-inversion synclinal geometry and is not bounded by major basement-involving faults (Fig. 11). During its evolution, depocenters migrated progressively northwards as evidenced by the northward onlap of its depositional sequences against the Jurassic pre-rift series. Palinspastic reconstructions indicated that the evolution of this basin was controlled by a south-dipping ramp along a major extensional fault that soles out in the V ariscan basement. Palinspastic reconstructions of the Cameros extensional-ramp basin suggest that the hangingwall was displaced by about 33 km to the south (Fig. 12). The Late Jurassic-Early Cretaceous syn-rift series of the Cameros Basin can be divided into six unconformity-bounded main depositional sequences which correlate with the Jl0-K1.10 sequences of the marine basins (Figs 12, 13; MAS et al., 1993). The Tithonian-Berriasian initial syn-rift sequence (JlO) consists of predominantly alluvial and lacustrine deposits that reach a maximum thickness of up to 3000 m in the eastern part of the basin, where they rest on Kimmeridgian marine strata. Towards the west, this sequence rests on progressively older Jurassic strata. Two minor discontinuities indicate, that the sequence can be subdivided into a Tithonian (JlO.l) and two Berriasian (Jl0.2, Jl0.3) subsequences. These sequences are characterized by lacustrine deposits, including shallow freshwater carbonates, which show only occasional marine influences. During subsequence Jl0.3, an evaporitic playa-lake developed in the eastern part of the basin (Figs 12, 13). The Berriasian-Valanginian sequence (K1.1) occurs only at the eastern extremity ofthe basin, where it consists of up to 200 m thick fluvial and lacustrine sediments. The V alanginian-Hauterivian sequences (K1.2 or K1.3) are absent in the central part of the basin. In its western part they consist of 100m of lacustrine carbonates, and in its eastern part of 500 m siliciclastic and mixed siliciclastic-carbonate fluvial and lacustrine sediments. The K1.4 sequence is only represented in the southernmost part of the basin (Soria sector) in a very subsident sub-basin (up to 800 m). The Barremian sequence consists of two subsequences. The lower one (K1.5, late Hauterivian?-early Barremian) displays two clearly separated depocentres in both the east and the west (where the area of sedimentation expanded), in which fluvial-lacustrine sediments accumulated. During the accumulation of the K1.6 subsequence, that is characterized by fluvial deposits, the area of sedimentation expanded and the basins started to coalesce, a process that was completed during the late Barremian-early Aptian (Kl.7, K1.8). Subsidence rates in both basins were different. In the eastern part, 1900 m of clastic fluvial deposits accumulated, grading north-eastward into 1100 m of lacustrine carbonates. In contrast in the western part, only 800 m of fluvial sediments were deposited. A marked marine influence (lagoonal deposits) is detected in this sequence (Figs 12, 13). During the deposition of the final late Aptian-middle Albian syn-rift sequences (Kl.9, Kl.lO), sedimentation was restricted to the eastern depocentre where up to 1500 m of alluvial clastics, containing rare and thin lacustrine carbonate intercalations, are preserved. An important intra-Albian erosiona! unconformity (D4, Fig. 8) constitutes the upper boundary of this sequence. The South lberian Basin is located south of the Maestrat Basin from which it is separated by the Valencia high. lt extends over 300 km and trends NW-SE. lt developed during the Berriasian to middle Albian rifting phase and contains more than 2000 m of syn-rift sediments. Their lower parts are mainly developed in a continental and lacustrine facies whereas the upper parts consist essentially of shallow marine carbonates (Fig. 14; VILAS et al., 1983). The Tithonian-Berriasian sequence (JlO), associated with the beginning of faulting and tilting of the previous carbonated ramps (sequence J9) is dominated

163

MESOZOIC CENTRAL IBERIAN RIFT SYSTEM AND ITS CAINOZOIC INVERSION

CAMEROS BASIN

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FIG. 13.- Chrono-lithostratigraphic diagram for the Cameros Basin during the Late Jurassic-Early Cretaceous rifting stage, showing depositional environments, formation names and division in depositional sequences. For legend see figure 10. FIG.

13.- Diagramme chrono-lithostratigraphique du bassin de Cameros durant l'épisode de rifting Jurassique supérieurCrétacé inférieur. Les environnements sédimentaires, les noms des formations et la subdivison en séquences de dépots sont indiquées. Voir la légende sur la figure 10.

164

RAMON SALAS ET AL.

by shallow subtidal carbonate bars (Higueruelas Formation) and tidal-flats and deltaic siliciclastic facies (Villar del Arzobispo Formation) which reach a thickness of up to 600 m in depocentre areas. The Valanginian to Hauterivian sequence (K1.2 or K1.3) lies unconformably on the preceding sequence (JlO), and is developed only in a very narrow and elongated fault bounded trough. lt consists of up to 200 m of mixed carbonate-clastic to clastic deposits, representing lagoonal, tidal-flats and coastal alluvial plain environments. After this episode, the basin expanded laterally with the Barremian sequences (K1.5-K1.7) onlapping their substratum. These Barremian sequences locally reach thickness of up to 700 m. They consist of siliciclastic alluvial deposits (El Collado Formation) and shallow lacustrine carbonates (La Huérguina Formation) with sorne marine incursions from the SE. Similar to the Maestrat Basin, marine conditions prevailed during the Aptian (K1.8, Kl.9). The shallow-water carbonate platforms of the El Caroig Formation, up to 500 m thick, laterally interfinger with continental silicicla~tic units (i.e. Contreras Formation). The siliciclastic early to middle Albian sequence (Kl.lO) was dominated by coastal alluvial and deltaic deposits (Escucha Formation) and peritidal mixed siliciclastic-carbonate facies (Sácaras Formation) that attain a thickness of 150m. The Columbrets Basin is located in the offshore beneath the Valencia Trough (Fig. 1) and apparently developed mainly during the Late Jurassic-Early Cretaceous Rift cycle. At present, it corresponds toa NE-SW oriented syncline which resulted from its Palaeogene partial inversion. The sedimentary fill of this basin is poorly known and only defined by reflection-seismic data calibrated by a few wells. It consists ofmore than 8 km ofMesozoic and Palaeocene rocks (ROCA, 1996).

PALAEOGEOGRAPHIC EVOLUTION DURING RIFfiNG CYCLE 2 The palaeogeographic evolution of northeastern Iberia during the Late Jurassic-Early Cretaceous rifting cycle is presented in figures 15 and 16. During this rifting cycle, pre-existing fracture systems, which were partly inherited from Triassic rifting and late Hercynián wrench faulting, were reactivated. At the early Oxfordian end of the post-rift stage 1, the Soria seaway extended from southeastern Iberia along the trend of the Triassic lberian Rift to the margins of the Bay of Biscay Rift, thus connecting the Atlantic with the Tethys seas (VERA, 2001). This seaway was limited to the southwest by the Iberian meseta and to the northeast by the Ebro high. These highs were accentuated during the Late Jurassic-Early Cretaceous rifting cycle and acted as significant palaeotopographic thresholds. Towards the southern margin of the evolving rift system, a smaller palaeohigh appeared during the Early Cretaceous separating the Maestrat and South Iberian basins (Fig. 16). The Late Jur~ssic-Early Cretaceous evolution of the lberian Rift system can be divided into the following three stages: latest Oxfordian-late Hauterivian ( 146.5-117.5 Ma), latest Hauterivian-early Albian (117 .5-103 Ma), and middle Albian (103-98 Ma). At the onset of the latest Oxfordian to late Hauterivian initial rifting phase, the Soria seaway was closed in response toa latest Oxfordian regression (BULARD, 1972) which was probably of tectonic origin and related to the build-up of tectonic stress. During the early Kimmeridgian, tector.ic activity increased, leading to the re-opening of the Soria seaway in response to a relative rise in sea-level and the subsidence of tilted extensional fault blocks. This was accompanied by rapid drowning of the spongerich late Oxfordian carbonate platform in the southeastern parts of the evolving rift, such as in the Maestrat Basin. During the Tithonian, the Soria seaway was interrupted again, as evidenced by the early syn-rift series of the Cameros Basin, which formed part of a set of smaller extensional sub-basins. During Tithonian to late Hauterivian times, rapid subsidence of basins forming part of the lberian Rift System led to significant palaeogeographic changes, particularly in its north-western parts. Overall, the Berriasian-Hauterivian time interval was a regressive episode that was only interrupted by the minor late Berriasian and early Valanginian transgressive pulses that advanced from the Tethys. The South lberian Basin developed during Berriasian times into a NW-SE trending narrow trough that was filled by terrigenous sediments. In the course of Berriasian and Hauterivian, the northwestern basins of the lberian Rift system coalesced and were drained towards the Bay of Biscay Rift.

165

MESOZOIC CENTRAL IBERIAN RIFf SYSTEM AND ITS CAINOZOIC INVERSION

SOUTHWESTERN IBERIAN BASIN

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FrG. 14.- Chrono-lithostratigraphic diagram for the South Iberian basins during the Late Jurassic-Early Cretaceous rifting stage, showing depositional environments, formation names and division in depositional sequences. For legend see figure 10. FJG. 14.- Diagramme chrono-lithostratigraphique des bassins ibériques méridionaux durant l'épisode dé rifting

Jurassique supérieur-Crétacé inférieur. Les environnements sédimentaires, les noms des formations et la subdivison en séquences de dépóts sont indiquées. Voir la légende sur la figure 10.

166

RAMON SALAS ET AL.

DS J9 (LST) LATEST OXFORDIAN

OS J9 (TST) EARLY KIMMERIDGIAN ~

~146.5Ma

144

t lBM

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l

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lB: lberian Meseta, EH: Ebro high, BC: Basco-Cantabrian basin, BB: Betic basin

FIG. 15.- Palaeogeographic maps showing evolution of the lberian Rift System during the latest Oxfordian to late Hauterivian.

FIG. 15.- Cartes paléogéographiques montrant l'évolution du systeme de rift ibérique durant l'Oxfordien terminaJ. Hauterivien supérieur.


167

OS K1.6/K1.7 LATE BARREMIAN

OS K1.4/K 1.5 LATEST HAUTERIVIAN TO EARL Y BARREMIAN

TO EARLIEST APTIAN 115-112Ma

117.5-115 Ma

OS K1.10 EARLY TO MIOOLE ALBIAN

OS K1.9 LATE APTIAN TO EARLIEST ALBIAN 109.5-107.5 Ma

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18: lberian Meseta, EH: Ebro high, 8C: 8asco-Cantabrian basin, 88: 8etic basin

Fio. 16.- Palaeogeographic maps showing evolutioñ of the lberian Rift System during the latest Hauterivian to middle Albian. FIG. 16.- Cartes paléogéographiques montrant l'évolution du systeme de rift ibérique durant l'Hauterivien terminal-

Albien moyen.

168

RAMON SALAS ET AL.

The latest Hauterivian to early Albian rifting phase was associated with a major transgression that was interrupted by the short-lived early Aptian regression and terminated with the mid-Albian regression (Fig. 16). In tbe Maestrat and South lberian basins, these regressions are reflected by deltaic sediments. During the latest Hauterivian and Barremian, accelerated tectonic subsidence facilitated the transgression of the Atlantic and Tethys seas, resulting in a retreat of continental facies towards the southeast and northwest, respectively. At the end of early Aptian depositional sequence (109.5 Ma), synrift subsidence slowed down and terminated during the late Aptian-early Albian (Fig. 16). During middle Albian/ crustal extension was reactivated for a last time, as evidenced by the rapid subsidence of fault-controlled troughs and the accumulation of thick deltaic and lacustrine series of the carbon-rich Escucha Formation (e.g., Maestrat Basin).

LATE CRETACEOUS POST-RIFf STAGE 2 During the Late Cretaceous, Iberia formed a separate plate and moved away from Europe during Albian and Cenomanian times, undergoing a 35° counter-clockwise rotation. The lberian continental block was flanked by passive margins facing the Tethys and the North Atlantic. lts northem margin, facing the Bay of Biscay, was passive during the late Aptian to Cenomanian but was gradually converted into an active margin, starting in Cenomanian times in conjunction with the early phases of the Pyrenean orogen y. During the Palaeogene phases of the Alpine orogeny, a compressional tectonic regime govemed the evolution of the Tethys Margin of Iberia (ZIEGLER, 1988; LE VOT et al., 1996; VERA, 2001) Following the final, mid-Albian rifting pulse, the area of the lberian Basin becarne tectonically quiescent and its evolution until Maastrichtian times was controlled by thermal relaxation of. the lithosphere and eustatically rising sea-levels. As a result of the Late Cretaceous transgression, large parts of northwesten{lberia were covered by shallow marine carbonate platforms (ALONSO et al., 1993; HAQ et al., 1987). Over most of the study area, the base late Albian to Late Cretaceous megasequence rests unconformably on Aptian, Barremian, Neocomian, Jurassic, Triassic and Palaeozoic sediments as well as on metarnorphic and igneous rocks (D4, Fig. 8). This unconformity is onlapped by the fluvial siliciclastic Utrillas Formation, a homogeneous, continuous, but diachronous unit that ranges in age from late Albian to Cenomanian. During the Cenomanian transgression, an extensive carbonate platform developed, connecting the Atlantic and Tethys domains. The Turonian-Coniacian succession starts with pelagic carbonates that change to thick dolomitic facies in its upper part. The Senonian succession has a marked shallowing upwards character with an upward gradation from open marine to sublittoral, lagoon and ultimately to fresh-water carbonates. The different uhits progressively ónlap to the NW (VIALLARD, 1973; CANÉROTet al., 1982; FLOQUET etal., 1982; GARCÍA-MONDÉJAR, 1989). The Upper Cretaceous supersequence (K2, Fig. 12) can be subdivided into twelve depositional sequences (K2.1-K2.12), which can be grouped into four second-ordre cycles or megasequences (ALONSO et al., 1993). Most of these sequences can be correlated with those of the Exxon curve, although sorne significant departures are evident. During the Late Cretaceous post-rift stage, there was gentle activity along sorne of the main faults. For instance, sorne of the major basin-bounding faults of the Maestrat Basin show syndepositional displacements of up to 500 m whereas similar faults in the South lberian and Carneros basins show Late Cretaceous offsets of up to 800 m. This fault activity may be largely due to compaction of the thick Late Jurassic to Early Cretaceous syn-rift sediments. MAGMATICACTIVITY During the Triassic rifting cycle and the subsequent Jurassic post-rift phase, basaltic volcanic activity occurred along the northem and southem margins of the Iberian Rift System during Norian to Bajocian

MESOZOIC CENTRAL IBERIAN RIFf SYSTEM ANO ITS CAINOZOIC INVERSION

169

times (Fig. 8; SALAS & CASAS, 1993). Extrusive centres were associated with the main bounding faults of this rift. Contemporaneous magmatic activity is evident in the rift system of the Betic domain (VERA, 2001). Magmatic activity ended in the Bay of Biscay Rift during the Early Jurassic. In the SE parts of the Iberian chain, Triassic subvolcanic activity of pre-Hettangian age, and extrusive Jurassic volcanism, spanning Pliensbachian to Bajocian times, culminating during the Toarcian, have been studied in the area of Castelló, Teruel and Valencia (LAGO et al., 1995, 1996; MARTÍNEZ et al., 1996, 1997). According to geochemical criteria, three magma types can be distinguished which were generated under different conditions and from different sources (Table 1). T ABLE 1.- General features of the Trias sic and Juras sic volcanism in SE lberian chain. Modified after MARTÍNEZ et al. ( 1997).

TABLEA U J.- Caractéristiques générales du volcanisme triassique et jurassique dans le sud-est de la chafne ibérique. Modifié d'apres MARTÍNEZ et al. ( 1997).

JURASSIC VOLCANISM, SE IBERIAN CHAIN AGE EMPLACEMENT

PETAOGRAPHY AFFINITY GEOCHEMICAL PATTERN

Pliensbachian - Baiocian Extrusive rocks interlayered in the Jurassic series Major volcaniclastic deposits Scarce massive lavas Porphydic textura 01 + Ti-Aug + PI Ti-Mt 11m Sol Alkaline High : La, Nb, Sr, Zr, Hf, Th, Ti, P Low:Y Yb

PRE-JURASSIC VOLCANISM, SE IBERIAN CHAIN AGE EMPLACEMENT

PETAOGRAPHY AFFINITY GEOCHEMICAL PATTERN

TERUEL-CASTELLON Pre-Hettanaian Subvolcanic masses within Triassic Eoizonal sills fraamented bv diaoirs Dolerltic and subofitic texturas Pegmatoid differentiates Ti-Aua + PI + Bt + Or· Ti-Mt 11m Pv. Ao Alkaline High : Nb, Sr, Zr, Hf, Th, Ti, P Low:Y Yb

VALENCIA Pre-Hettanaian Subvolcanic masses withln Triasslc Eoizonal sills fraamented bv diaoirs Doleritic and subofitic texturas No pegmatoid differentiates Aua + PI ± Bt · Mt 11m Ao Subalkaline ltholeiitic). transitional lalkaline)_ High : La, Nb, Sr, Zr, Th (Y, Yb) Low : (Y Yb Hf), Ti P

In the Valencia Province, pre-Hettangian tholeiitic and transitional alkaline sills and dykes intruded into Triassic series during the rifting cycle l. These melts were derived by a low degree of partial melting from a slighdy metasomatized thermal bound~ layer of the mande-lithosphere. During their ascent, these melts were partly contaminated by crustal material. Similarly, the alkaline pre-Hettangian sills of the Teruel-Castelló Province also reflect a slight partial melting of a metasomatized mande. On the other hand, the Pliensbachian-Bajocian alkali-basalts, which were extruded in the SE Iberian chains during the post-rift stage 1, are characterized by a lower incompatible element content than the Triassic magmas. This suggests a low degree of partial melting of an unmetasomatized and heterogeneous mande. These data are interpreted as reflecting the following succession of events (MARTÍNEZ et al., 1997). During the pre-rift stage (Permo-Carboniferous?), an asthenospheric plume impinged on the base of the lithosphere leading to its partial melting and enrichment of its thermal boundary layer. During the Late Permian-Trias sic rifting cycle, stretching of the lithosphere resulted in decompressional partial melting of its thermal boundary layer and in the generation of subalkaline and alkaline magmas. These ascended along extensional fault systems and intruded into the syn-rift sediments. The low degree of melting is surprising and could cast doubts on the contribution of a mande plume to melting processes. On the other hand, Early and Middle Jurassic volcanism occurred during the

170

RAMON SALAS ET AL.

post-rift evolution of the Iberian Rift (SALAS & CASAS, 1993) and cannot be attributed to further extension-induced decompressional melting of the basallithosphere and the upper asthenosphere. lt may be therefore attributed toa plume-induced increase in the potential temperature of the asthenosphere. In this regard, it should be noted, that during the final rifting stage of the Central Atlantic, a majar plume impinged on the base of the lithosphere, giving rise to the development of a large magmatic province that extended from northern Brazil over the Appalachian domain and West Africa as far north as Iberia where it manifested itself in the magmatism of the Betic domain and the intrusion of dyke systems tran-secting the Iberian meseta in a NE direction. This plume-related magmatism commenced during the Late Triassic and extended well into the Middle Jurassic (WILSON, 1997). In view of this, it is likely that the Late Triassic to Middle Jurassic magmatic activity in the south-eastern parts of the Iberian chain was related to this plume and its outflowing mantle currents (ZIEGLER et al., 2001). In summary, the Triassic Teruel-Castellón alkaline magmas were probably generated at greater depths and their greater differentiation dm be attributed to a slower and/or longer ascent. In contrast, the Jurassic magmas were generated at shallower levels and ascended more rapidly, as indicated by the lack or only minar differentiation. This suggests that plume-induced progressive thinning of the mantle lithosphere persisted well into the post-rift stage l. Interestingly, the latest Jurassic-Early Cretaceous rifting cycle 2 was not accompanied by any magmatic activity despite considerable crustal stretching factors controlling the subsidence of e.g., the Maestrat and South Iberian basins.

QUANTITATIVE SUBSIDENCE ANALYSIS In arder to further assess the evolution of the Maestrat, South Iberian and Cameros basins, sorne previously published quantitative subsidence curves (SALAS & CASAS, 1992, 1993) were reviewed and new ones calculated. Back-stripping techniques were applied (STECKLER & WATTS, 1978; SCLATER & CHRISTIE, 1980; WATTS, 1981; BOND & KoMINZ, 1984) to quantify tectonic subsidence. Parametres used to calculate the subsidence for each sedimentary unit include lithology, age, present-day depth, palaeobathymetry and eustatic sea-level during deposition. The absolute age scale and sea-level curve given by HAQ et al. (1987) was used. Decompaction was calculated using porosity/depth relationships and density for each lithology. Palaeobathymetry was estimated from the interpreted depositional environments. For the Maestrat Basin, we used the logging data of three oil wells, namely Mirambell-1, Maestrazgo-! and Amposta Marino C-3 (l.ANAJA, 1987), taking into account the eroded strata. For the South Iberian Basin we analysed an outcrop section (Fig. 17). Despite similarities in the general pattern of the four calculated tectonic subsidence curves, sorne differences in their shape and the magnitude of tectonic subsidence are evident (Fig. 18). These differences are mainly due to the location of the data points in the basins and the related differences in sedimentary thicknesses. Nevertheless, all curves show four phases of tectonic subsidence, each characterized by an initial interval of rapid subsidence followed by an interval of decelerating subsidence. This model of tectonic subsidence is usually diagnostic of rifting processes, which comprise an initial period of fault-controlled rapid syn-rift subsidence, followed by a post-rift interval of asymptotically decreasing subsidence, controlled by thermal relaxation of the lithosphere (MCKENZIE, 1978). Taking into account the observed four phases of tectonic subsidence, the Structural and regional setting of the control points, the following subsidence phases can be distinguished: 1) Late Permian-late Oxfordian (255-146.5 Ma), 2) latest Oxfordian-late Hauterivian (146.5-117.5 Ma), 3) latest Hauterivianearly Albian (117.5-103 Ma), and 4) middle Albian-Maastrichtian (103-68 Ma). The subsidence curve for the sedimentary sequence in the Mirambell-1 well shows in the Late Permian-late Oxfordian interval rapid subsidence during the Late Permian to Triassic that gradually decelerates during the Jurassic. A renewed subsidence acceleration begins in the latest Oxfordian. Similarly, the curve for the Amposta Marino-C3 well shows the same initial rapid subsidence and decelerating subsidence rates during the Early-Middle Jurassic. This reflects the initial Late PermianTriassic rifting cycle 1, which is followed by the Jurassic post-rift phase l. However, in the curves for


171

4~N

~E

~ Hercynian belt

1!1 Alpine belts

Fro. 17.- Simplified geological map of the NE part of the lberian peninsula showing location of analysed wells (1, 2, 3) and stratigraphic section (4). 1, Mirambell-1; 2, Maestrazgo-!; 3, Amposta Marino C-3; 4, South lberian Basin. FIG. 17.- Carte géologique simplifiée de la partie NE de la péninsule ibérique montrant la situation des forages analysés (1, 2,

3) et de la coupe stratigraphique (4). 1, Mirambell-1; 2, Maestrazgo-]; 3, Amposta Marino C-3; 4, bassin ibérique méridional.

Maestrazgo-! and the South lberian Basin, this syn- to post-rift transition is not clearly expressed, possibly due to the wider spacing of data points. Subsidence analyses carried out by VAN WEES & STEPHENSON (1995) across the central parts of the lberian chain clearly reflect the Late PermianTriassic rifting phase and the subsequent Jurassic post-rift phase. The next cycle of accelerated subsidence corresponds to the latest Oxfordian-middle Albian time interval (146.5-98 Ma) and thus to the rifting cycle 2. As discussed above, this rifting cycle was complex and apparently comprises two phases of accelerated subsidence that were interrupted by a short phase of decelerated subsidence during the Neocomian (128.5-117.5 Ma). The late Aptian-early Albian (109.5-103 Ma) final part ofthis rifting phase was again characterized by decelerating subsidence rates. The four main phases of the Mesozoic rift-post-rift evolution identified in the Maestrat and South lberian basins are in clase agreement with the three main subsidence phases identified in the Castilian branch of the lberian chain (VAN WEES & STEPHENSON, 1995). Assuming a pure-shear, uniform stretching model (MCKENZIE, 1978), the stretching factors for each of the main subsidence phases, as well as the total Mesozoic crustal stretching factor, were calculated from the tectonic subsidence data for each well. The results are summarized in Table 2. The first phase of rapid Triassic subsidence followed by decreased subsidence during the Early and Middle Jurassic (255-146.5 Ma) can be fitted in the area of the Maestrat Basin to a crustal stretching factor of ~= 1.08-1.15, and in the area ofthe Southwest lberian Basin to ~= 1.18. The latest Oxfordian to late Hauterivian subsidence phase ( 146.5-117.5 Ma) matches a crustal stretching factor of ~= 1.03-1.14 in the Maestrat Basin and ~= 1.05 in the South lberian Basin. The Barremian-early Aptian rapid subsidence phase, which is followed by slower subsidence during the late Aptian-early Albian (117 .5103 Ma), reflects a stretching factor of ~= 1.02-1.06 for both the Maestrat and the South lberian basins. The final, though moderate middle Albian rifting phase, which was associated with strongly faultcontrolled sedimentation, and was followed by Late Cretaceous decreased subsidence (103-68 Ma), accounts for about ~= 1.015-1.02 in the Maestrat and South lberian basins.

172

RAMON SALAS ET AL.

CRETACEOUS TERTIARY

JURASSIC

o

50 1 1 11

1 1 1

:> >MIRAMBELL 1

:::: :NIA."ESTRAZGO 1

Om

1000

2000 146.5 117.5 103

255

1 2 lal

1 R1

PR1

1

R2

68

4

1

1PR21

FIG. 18.- Backstripped tectonic subsidence for the Maestrat and South Iberian basins. For locations see figure 17. Shaded areas indicate the four main phases of Mesozoic rift-postrift tectonic subsidence. For stretching factors see table 2. FIG. 18.- Subsidence tectonique des bassins du Maestrat et ibérique méridional. Voir la situation des points d'analyse sur la

figure 17. En pointillé ont été indiquées les quatre phases principales de subsidence tectonique rift-postrift mésozoi"ques. Pour lesfacteurs d'étirement voir tableau 2.

173


TABLE 2.- Stretching factors calculated for wells in the Maestrat Basin and the stratigraphic section of the South Iberian Basin.

2.- Facteurs d' extention calculés pour les forages du bassin du Maestrat et pour la coupe stratigraphique du bassin ibérique méridional.

TABLEA U

Curve l. Mirambel11 2. Maestrazgo 1 3. Amposta Marino C3 4. South lberian Basin

Phase 1 255-146.5 Ma 1.15 1.12 1.08 1.18

Phase 2 146.5-117.5 Ma 1.03 1.14 1.04 1.05

Phase 3 117.5-103 Ma 1.02 1.06

Phase 4 103-68 Ma 1.015 1.015

Total 255-68 Ma 1.22 1.37

1.04

1.02

1.31

A total Mesozoic crustal stretching factor of ~= 1.37 was determined for the Maestrazgo-! well, located in the depocentre of the Maestrat Basin, ~= 1.22 for its northern marginal area (Mirambell-1 ), and ~= 1.31 for the South lberian Basin. Thus, the average Mesozoic crustal stretching factor for the eastern part of the lberian Rift System is ~= 1.3. This translates into a total extension of about 40 km. Subsidence analysis of the Cameros Basin was based on three stratigraphic sections. Their location is shown on the restored cross-section (Fig. 19) that illustrates the stratal fan-array geometry of the Early Cretaceous sedimentary fill of this basin. Two phases of basin development are recognized and were sampled by the sections Cameros 1, 2 and 3. The Tithonian to early Barremian phase (140-115 Ma) may corresponds to the second phase of the lberian Basin (Fig. 19, sections 1 and 2). The late Barrernianearly Albian phase (115-103 Ma) can only be documented in the northern part of the basin (Fig. 19, section 3) and may be equivalent to the third phase.

CRUSTAL STRUCTURE OF THE IBERIAN CHAIN: IMPLICATIONS FOR MESOZOIC EXTENSION AND PALAEOGENE INVERSION In an attempt to better understand the crustal structure of the lberian chain and the processes controlling its Mesozoic extensional and Palaeogene compressional deformation, all available geophysical data (seismic, palaeomagnetic, aeromagnetic, gravity and heat flow) were taken into consideration.

_

o.....__...._ 10

0== [ [ ] ] Mesozoic

_,20km :1,...:!'-.,.;!:~~M~.-._;....-.+

+ + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + . + + + + + + + + + + + + + + + + + + + + + • ++++++++.,_+ .. +_.+_,+.+ + + + + + . . . . . . .

IZ2l Jurasslc

1"1 Lower Alblan

~ Tertialy ~

O

L...........J (Utrlllas Fm)

H~IMH

[[]]

Trlasalc

Lower eretaceoua

~= ~.

FJ:o. 19a

FIG. 19.- Restored cross-section through the Carneros Basin, showing its basin-scale stratal fan-array geometry (Fig. 19a) with total subsidence (ts) and tectonic subsidence (TS) curves for three stratigraphic sections (Fig. 19b).

FIG. 19.- Coupe restituée du bassin de Cameros montrant la géométrie des dépots en éventail ii l'échelle de tout le bassin (Fig. 19a) avec les courbes de subsidence totale (ts) et subsidence tectonique (TS) pour trois coupes stratigraphiques (Fig. 19b).

174

RAMON SALAS ET AL.

Cameros 1 140

135

130

125

120

115

110

100 Ma

105

o TS

1000

ts 2000 3000 140

Cameros 2 135

130

125

120

115

110

100 Ma

105

o 1000 2000

TS

3000 4000 5000

ts

6000 Cameros 3 140

135

130

125

120

115

110

100 Ma

105

o 1000

TS

2000

ts

3000

FIG.

19b


175

GEOPHYSICAL DATA A Bouguer anomaly map of the area under consideration was compiled by MEZCUA et al. (1996) (Fig. 20). This map shows that the lberian chain is associated with a regional gravity low that contrasts with the relative gravity high of the Ebro Basin and the pronounced high of the Valencia Trough. The large and negative anomalies of the lberian chain are consistent with a thickened crust. CADAVID (1977) reported Moho depths of more than 46 km under the central part of the chain from Bouguer anomalies. Nevertheless, gravity data do not reflect conditions of general isostatic equilibrium since the negative anomaly of the Serranía de Cuenca (Fig. 2ld) results from a mass deficit at depth which is not compensated by topographic elevations.

3'

41°

•

FIG.

20.- Bouguer contour anomaly map of the Iberian chain and surrounding areas. Contour intervals 10-mGal. Line shows location of the gravity profile of figure 21. Modified after MEZCUA et al. (1996). The overall structure of the lberian chain is also indicated.

FIG.

20.- Carte des anomalies de Bouguer pour la chaíne ibérique et les zones adjacentes. lntervalles de contour 10-mGal. La ligne montre la situation du profil de gravité de la figure 21. Modifiée d'apres MEZCUA et al. (1996). La structure générale de la chaíne ibérique est aussi montrée.

176

RAMON SALAS ET AL.

Three gravity profiles, which trend perpendicular to the lberian Chain, were modelled by SALAS & CASAS (1993). Density models were constrained, where possible, with geological information, well data and seismic results. The geometry of the deep crustal structure inferred from the gravity profile is shown in the fi~ure :f1a. The gravity model covers the fan-like geometry of the lberian Chain, proposed by GUIMERA & ALvARO ( 1990), with northward thrusting in the north and southward thrusting in the south (Fig. 4). Results show that a reasonable fit between the observed and calculated anomaly can only be obtained if the crusts thickens asymmetrically under the chain to a maximum of 43 km with the Moho rising gradually towards the Ebro Basin. This implies that the crust of the lberian Basin was shortened and thickened during its inversion, and that the Ebro crustal block was overthrusted by crustal fragments of the in verted lberian Chain. mGal

o

•

• •

50

100 150

o

ssw

100

50

La Mancha

Sierra de Javalambre

150

200 Muela de Montalbán

250

km .NNE

Ebro basin

="""",.,..,..,..,.,..,...,.,...,..,..,.,..,...;,· •

75km

A

Present

B

End of Mesozoic

Tertiary shortening

30

47km Mesozoic streaching

32

C

End of Palaeozolc

~ ~

Cainozoic

FIG. 21.- A, Gravity profile (dots represented observed data) and interpreted crustal section. For locations see figure 20. B, Reconstruction of section A at the end of the Mesozoic. C, Reconstruction of the section A at the end of the Palaeozoic. Modified after GurMERA et al. (1996). FIG.

21.-A, Profil de gravité le long des localités indiquées sur la figure 20 (les points représentent les données observées). B, Reconstruction de la coupe A a la fin du Mésozoi'que. C, Reconstruction de la coupe A a la fin du Paléozoi'que. Modifiée d'apres GUIMERA et al. ( 1996).


177

Deep seismic sounding data, whilst assuming a laterally heterogeneous crustal model, indicate Moho depths of about 32 km beneath the central part of the lberian chain (ZEYEN et al., 1985). This is in accordance with the results obtained by SURIÑACH & VEGAS (1988) for the lberian meseta. However, a re-interpretation of the lberian chain data, assuming a simple laterally homogeneous crustal model, gives a Moho depth of 38-40 km. As for the result of the gravity interpretation for the south-eastem parts of the chain (Fig. 21a), this implies the presence of a crustal keel beneath the northem central part of the chain, which is probably related to crustal thickening during' the inversion of the lberian Basin. Unfortunately, the relevant seismic profiles were recorded only in one direction and do not cross the lberian chain normal to its strike. Nevertheless, the hypothesis that rifting processes controlling the late Oligocene-Miocene opening of the Valencia Trough had also affected the central parts of the Iberian chain, is contradicted by gravity data. Indeed, refraction-seismic data indicate that the crust rapidly increases in thickness from the axial parts of the Valencia Trough towards coast to reach values of 40 km further inland (DAÑOBEITIA et al., 1992). The total-intensity aeromagnetic map of the area studied (ARDIZONE et aL, 1989) shows a regional NW-SE trend which extends for more than 250 km from the northwestem most part of the lberian chain to the Mediterranean coast. This trend roughly coincides with the outline of the lberian chain (SALAS & CASAS, 1993). The smooth character of this regional anomaly is in part dueto the fact that the magnetic sources are deep-seated and in part to the wide spacing of the surveyed lines (10x40 km) anda flight height of 3 km. This magnetic anomaly is tentatively interpreted as being related to a suture or boundary between two basement terranes with contrasting lithologies. Sorne strong bipolar magnetic anomalies may be associated with Mesozoic rift-related volcanics. The surface heat flow of the lberian chain and surrounding areas (Fig. 22; FERNÁ.NDEZ et al., 1998) varies from 70 mW.m- 2 in the Cameros unit, the Tertiary Almazán and Ebro basins and the northem part of the Castilian branch to the Mediterranean Margin, where it reaches values around 80-100 mw.m-2 • These values of heat flow are consistent with those of the proposed crustal and lithospheric structure of the lberian chain (SALAS & CASAS, 1993).

4z>N

FIG. 22.- Heat-flow map of the NE part of the lberian peninsula. 10 mW.m- 2 isolines. Slightly modified after FERNANDEZ et al. (1998).

FIG. 22.- Carte du flux thermique du NE de la péninsule ibérique. Lignes a 10 mW.m-2• Légerement modifiée d'apres FERNÁNDEZ et al. (1998).

4°W

z>W

111 Hercynian belt

0°

2°E

¡¡¡¡¡¡¡¡¡¡¡¡¡¡¡¡¡ Alpina belts

178

RAMON SALAS ET AL.

CRUSTAL MODELLING The crustal configuration and evolution of the south-eastern parts of the Iberian chain was investigated along a gravity profile that extends over a distance of 290 km from the Ebro Basin across the Maestrat and West Iberian basins onto the La Mancha Platform (SALAS & CASAS, 1993; Figs 20, 21). Our analysis involved crustal mass balancing, and the comparison of the observed gravity profile with a modelled one, which accounts for Mesozoic crustal extension and Palaeogene compression. According to this profile, which crosses a major negative Bouguer anomaly of the lberian chain (Fig. 20), present-day crustal thicknesses vary between 32.5 km at its northern and southem ends and 43 km in its central part. In our interpretation we assumed isostatic equilibrium for the present-day Iberian crust, and standard crustal (2.74 gr/cm3) and mantle densities (3.2 gr/cm\ The magnitude of Palaeogene crustal shortening was estimated by comparing the present-day crustal profile (Fig. 21a) with a profile that was restored to end-Mesozoic times, i.e. prior to the Tertiary compression (Fig. 21b). In view of the shallow marine to lacustrine Late Cretaceous post-rift sediments, which covered the entire lberian chain, it was assumed that the area under consideration was located at sea level, i.e. sorne 200 m above the present-day sea level (HAQ et al., 1987). Taking into account crustal densities and thicknesses, and assuming mass preservation in the restored profile and an average crustal thickness of about 30 km ±1 km at the end of the Mesozoic, the crust was shortened by 75 km ±12 km during the Palaeogene contractional phase. The shortening values given in the chapter "geological setting" refer mostly to the northem branch of the Iberian chain (Cameros unit, Aragonese branch and linking zone). A maximum shortening of 38 km is reached across the Cameros transect, with about 30 km of N-directed displacement along the North lberian thrust, whereas in the other two sections (Fig. 3b, e) the displacement of that thrust is unknown. In the southern branch of the Iberian chain (Castilian-Valencian branch) and the Altomira unit shortening is 19 km. Combined, this gives a total shortening of 57 km across the Iberian chain. As in the figure 3, only plurikilometric folds and thrusts were taken into account, giving a conservative estimate, the total shortening should be larger. Thus, shortening deduced from cross-sections and the crustal model compare fairly well. The geometry and dimensions of the Palaeogene thrusts, involving the basement of the Iberian chain, indicates the presence of an upper crustal detachment level located at a depth of 7-11 km that can account for upper crustal shortening. Therefore, shortening at deeper crustal and lithospheric levels must have been accommodated either by ductile mechanisms (GUIMERA et al., 1996), and/or by an incipient subduction mechanism, involving imbrication of the rheologically strong upper part of the mantlelithosphere (SALAS &CASAS, 1993; ZIEGLER et al., 1995). As indicated by the post-inversion crustal configuration of the area, syn-inversion deformation of the crust was probably of the heterogeneous pure-shear type (heterogeneous thickening), resulting in the development of a deep crustal keel (GUIMERA et al., 1996). Whether this keel is associated with an offset of the Moho discontinuity or not remains unsolved. The overall double arched structure of the lberian chain is reflected in the Bouguer anomaly map (Fig. 20). In the northwestem section of the chain, the arches are distinct and the anomaly map shows two NW-SE trending lows that parallel the Iberian structures; to the SE, both lows merge in the same area where the arches merge. Nevertheless, a shift between the arches and the Bouguer anohlalies is observed. This is more evident in the northem arch and in the Almazán Basin to the south of it: the relative minimum and maximum anomalies are displaced sorne 10 to 15 km to the SSW with respect to the arch and the basin. This may be related to a thickening of the lower crust by heterogeneous pure shear, while the upper crust is deformed by simple shear (thrust and fold systems). Further study of the crustal structure of Iberian Chain, including deep seismic reflection profiles is required, in order to explain this shift in upper and lower crustal thickening.


179

MESOZOIC STRETCHING Assuming an average crustal thickness of 32.5 km for Late Palaeozoic times, as currently observed in the western Iberian meseta, the transformation of a cross-sectional area of 10160 km 2 to an average crustal thickness of 30 km ±1 km at the end of the Mesozoic (Fig. 2lb) would require its extension by sorne 35 km to 59 km (P= 1.11 to 1.19). However, subsidence analyses and palinspastic restoration of basins suggest higher stretching factors for the rifted basins. ROCA et al. (1994) calculated a TriassicEarly Jurassic stretching factor of P= 1.37 for the Desert de les Palmes area (Fig. 3) by analysing restored cross-sections perpendicular to the main profile (Fig. 4). Similarly, subsidence analyses of boreholes in the Maestrat Basin indicate a cumulative Mesozoic stretching factor P= 1.22-1.37 (see above). These values imply an average syn-rift extension by 40 km and a pre-inversion crustal thickness of 26.2-23.3 km. Therefore, adding the thickness of the Mesozoic Basin fill (3.3 km and 5.8 km), the crust/mantle boundary was located at the end of the Mesozoic at a depth of 29.5 km to 29.1 km. This suggests that during the Mesozoic rifting cycles crustal thinning was not uniform but more intense beneath the most rapidly subsiding basins. Thus, the assumption of an average crustal thickness of 29.5-30 km at the end of the Mesozoic is a great simplification. This indicates that at end of the Mesozoic the crust of the Iberian Basin was considerably thinner than under its flanking, unextended areas, and that, during the Palaeogene inversion, the crust of the Iberian chain was mechanically thickened. MINERAL RESOURCES The most important mineral resources of the area under consideration are the middle Albian coal measures of the Escucha Formation of the Maestrat Basin. This formation consists mainly of deltaic sediments, of which the upper and lower delta plain series contain exploitable coal deposits in the Teruel mining district (QUEROL et al., 1992). The distribution and thickness of workable coals is controlled by the sediment supply of depositional systems and their response to tectonically controlled subsidence. These sub-bituminous coals have a high sulphur content (6 to 7.3% dry weight). The Escucha Formation contains the largest Spanish sub-bituminous coal deposit and yields about 5xl06 tons/year, i.e. more than 13% of the country's coal production. Currently, this coal is mainly used for the generation of thermoelectric power, accounting for 15% of Spain's electricity requirements. The oil crisis of the 1970's gave a considerable stimulus to coal mining which hitherto had been underground and led to the opening of op,en cast pits. At present, coal mining is facing a new crisis owing to environmental problems cau sed by the high sulphur content of the coal and Spain' s commitrnent to nuclear power. However, the continuation of coal mining in the Teruel district is ensured by the construction of fluidised bed combustion and flue gas desulphurisation units. The organic-rich Kimmeridgian Ascla Formation shales of the Maestrat Basin (Fig. 9) form a viable oil source-rock which has contributed, albeit to a minor degree, to the hydrocarbon potential of the Valencia Trough. SUMMARY AND DISCUSSION Integration of geophysical, subsurface and surface data has provided a better understanding of the Triassic and Late Jurassic-Early Cretaceous rifting history of the Iberi~n Basin, of the mechanisms of its Palaeogene inversion and of the architecture of the resulting Iberian chain. Four Mesozoic evolutionary stages can be distinguished in the Iberian basins, based on the analysis of syn-depositional Mesozoic structural features and stratigraphic sequences, quantitative subsidence modelling and analysis of seismic data (Fig. 23): - Late Permian to Hettangian rifting cycle 1; - Sinemurian to early late Oxfordian post-rift stage 1;

180

RAMON SALAS ET AL.

w

EARLY TRIASSIC

LATE JURASSIC & EARLY CRETACEOUS

(!)

~

(/')

w

1-

~ (/')

a:

!:!:

(!)

MIDDLE TRIASSIC

w

(!)

__)

~

....~

.. 1-

a:

w

(!)

~ (/')

EARLY & MIDDLE JURAS SIC

. :·.

u.

~

oo.

~/

......

·..

/

''

__j

__; D

:~

·.-.·~=.~:-~::~,: r· a: ~~m~=

1-

_,,

~

•.

u.

LATE TRIASSIC

LATE CAETACEOUS

(/')

,

~, /

'/

/

~

Sandstone Marine carbonates

a

Evaporites

A Basaltic volcanism

.... o ... ,

• o ••

Molas se

~Main

faults

/( FIG. 23.- Schematic evolution ofthe lberian Basin (lberian chain). Modified after ÁLVARO et al. (1979). FIG.

23.-Évolution schématique du bassin ibérique (chaíne ibérique). Modijiée d'apres ÁLVARO et al. (1979).

- latest Oxfordian to middle Albian rifting cycle 2. Development of the Maestrat, Cameros, Columbrets and South lberian basins during three discrete rifting pulses spanning latest Oxfordian to late Hauterivian (146.5-117.5 Ma), latest Hauterivian to early Albian (117.5-103 Ma), and middle Albian (103-98 Ma) times. Rift-induced subsidence started in the Maestrat Basin during the latest Oxfordian, and in the Cameros Basin during the early Tithonian. Both basins contain up to 5 km of synrift sediments; - late Albian to Maastrichtian post-rift cycle 2; - late Eocene (?) to middle Miocene Basin inversion; -late Oligocene to middle Miocene rifting ofValencia Trough and Teruel graben.


181

Plate tectonic considerations suggest that the Late Permian to Hettangian evolution of the rifted . lberian Basin is related to the westward and southward propagation of the Tethys and ArcticNorth Atlantic Rift systems, respectively, and to their interference in the lberian-North Atlantic domain . .The Carnian to Bajocian alkaline volcanism of the southeastem parts of the lberian Basin is possibly related to the development of the Central Atlantic super-plume (WILSON, 1997). The Late JurassicEarly Cretaceous rifting cycle of the lberian Basin coincides with the gradual opening of the North Atlantic Basin and sinistral oblique extension in the Bay of Biscay Rift prior to mid-Aptian crustal separation between Iberia and Europe. Rifting persisted, however, in the Iberian Basin during the early phases of sea-floor spreading in the Bay of Biscay and terminated only in rnid-Albian times. During the Late Cretaceous post-rift cycle 2, Iberia acted as an independent plate. However, starting in Cenomanian times, its northem margin was gradually converted to an active margin in conjunction with the early phases of the Pyrenean orogen y. As a result of the Alpine orogen y a compres sive tectonic re gime controlled the Tethyan Margin of the lberian plate (VERA, 2001). Intraplate compressional stresses building up within cratonic Iberian controlled the Eocene and Oligocene inversion of the lberian Basin · and the development of the lberian and Catalan coastal ~hains. Many structures of the lberian chain involve compressionally and transpressionally reactivated Mesozoic Basin bounding normal faults. The kinematics of their reactivation depended on their strike with respect to the orientation of the regional compressional stress field. The structural style of the lberian chain is dominated by steep reverse and thrust faults originating in the basement, and, particularly in its south-eastem extemal parts, by thin-skinned thrust faults (Figs 2-4). Basement-involving thrust faults appear to sole out at mid-crustal levels at depths of 7 to 11 km: Gravity data indicate the presence of a crustal keel beneath the axis of the lberian chain where the Moho discontinuity is located at a depth of 40 to 43 km. This is compatible with the reinterpretation of available refraction-seismic data. Beneath the adjacent lberian meseta the crust/mantle boundary rises to 32.5 km depth. The indicated crustal keel suggests that inversion of the lberian Basin involved the entire crust, the lower part of which was probably deformed by ductile pure shear (GUIMERA et al., 1996). There are insufficient data to determine whether or not the rheologically strong upper part of the mantlelithosphere is also involved in the deep structure of the lberian chain. However, it is unlikely that Iberia was detached as a whole from its mantle lithosphere during the Palaeogene phase of intraplate compression. Therefore, imbrications of the Moho, such as deep-seated pop-up structures, may be expected (SALAS & CASAS, 1993). During the Palaeogene inversion of the lberian chain, crustal shortening probably ranged between 55 km in its north-westem part and possibly as muchas 75±12 km in its south-eastem part. The axis of the lberian chain coincides with a deep seated positive magnetic anomaly. The cause of this anomaly is unknown but could either be related to anancient suture between crustal terranes orto Mesozoic magmatic bodies that intruded into the crust during the Triassic rifting cycle 1 and the Jurassic post-rift phase l. Cumulative Mesozoic crustal extension across the Iberian Basin is difficult to estímate but may have been as much as 45 km in its south-eastem and sorne 33 km in its north-westem parts. In the areá of the Early Cretaceous Cameros Basin, extension may have been confined to upper crustallevel, as suggested by the geometry of this extensional-ramp basin, with extension rooting at deeper lithospheric levels to the south (GUIMERA et al., 1995). This would imply, at least for the rifting cycle 2, a simple-shear mode of extension for this area. However, in the south-eastem parts of the lberian Rift, the presence of Triassic syn-rift volcanic activity suggests that extension affected the entire lithosphere, at least during rifting cycle l. The Late Jurassic to Early Cretaceous rifting phase, which controlled the rapid subsidence of the Maestrat, Cameros, South lberian and Columbrets basins, was not associated with · volcanic activity. The evolution of the Late Jurassic-Early Cretaceous lberian Rift System was paralleled by the development of the North Atlantic Margin basins (TANKARD & BALKWILL, 1989), the Aquitaine Basin (BRUNET, 1984; LE VOT et al., 1996), the South Pyrenean Basin (BERÁSTEGUI et al., 1990), and the Mesozoic part ofthe Ebro Basin (DESEGAULX & MORETTI, 1988). Further crustal-scale geophysical data are required to resolve the deep structure of the lberian chain and to better understand the kinematics of lithospheric deformation during the Palaeogene phase of intraplate compression which had affected large parts of cratonic Iberia.

182

RAMON SALAS ET AL.

ACKNOWLEDGEMENTS We are grateful to our colleague K. BITZER for bis belpful comments and suggestions. We wisb to tbank W. CA VAZZA for bis critica! comments and suggestions and also P. ZIEGLER for bis kindness in improving the final version of tbe manuscript. We are indebted to G. VON KNORRING for bis technical assistance. Tbis work was funded by the project: PB95-1142-C02 (Direción General de Enseñanza Superior e Investigación Científica). Additional financia! support was provided by the Comissionat per Universitats i Recerca de la Generalitat de Catalunya (1998 SGR 00034).

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