Cenozoic landscape development within the Central Iberian Chain, Spain

Geomorphology 44 (2002) 19 – 46 www.elsevier.com/locate/geomorph

Cenozoic landscape development within the Central Iberian Chain, Spain Antonio M. Casas-Sainz a,*, Angel L. Corte´s-Gracia a,b a Departamento de Geologı´a, Facultad de Ciencias, Universidad de Zaragoza, 50009 Zaragoza, Spain Departamento de Dida´ctica de las Ciencias Experimentales, Universidad de Zaragoza, 50009 Zaragoza, Spain

b

Received 6 November 2000; received in revised form 26 June 2001; accepted 10 July 2001

Abstract In the Central Iberian Chain (Spain), covering an area of some 7400 km2, up to seven stepped erosion surfaces are determined from aerial photographs and field observations. These planation surfaces situated between 1000 and 1600 m a.s.l. are separated by erosional scarp slopes. Most parts of the erosion surfaces are developed on Mesozoic limestones and form stepped rims around the main elevations of the Iberian Chain. It is proposed that the surfaces were formed by erosional events associated with periods of uplift related to Eocene – Miocene compressional tectonic activity. The palaeogene uplift was related with syntectonic sedimentation in internal basins with up to 3500 m of accumulated thickness. During the Neogene (Middle Miocene – Pliocene) post-tectonic period, the lowered synclinal areas formed residual, internally drained basins that were filled with non-marine deposits (mainly clays and limestones). The filling of the internally drained sedimentary basins located within the Iberian Chain onlapped onto the ancient erosional surfaces, covering the lower younger levels. The recent landscape evolution of the whole area is controlled by the capture of the internal basins and the dissection of the ancient relief by the present-day fluvial network. This process is related with the transition from internal to external drainage of the Ebro basin. D 2002 Elsevier Science B.V. All rights reserved. Keywords: Erosion surface; Neogene; Tectonics – sedimentation relationship; Iberian Chain

1. Introduction Long-term landscape evolution in continental areas is a balance between tectonic uplift, erosion and sedimentation (Merrit and Ellis, 1994 and references therein; Kooi and Beaumont, 1994; Leeder, 1997; Roessner and Strecker, 1997). Erosion surfaces separated by scarps are common features in large cratons and are usually associated to extensional tectonics and *

Corresponding author. Tel.: +34-976-762072; fax: +34-976761088. E-mail address: [email protected] (A.M. Casas-Sainz).

rifted margins (Kooi and Beaumont, 1994). Erosion or planation surfaces can be defined independently of their origin as low-relief plains cutting across varied rocks and geological structures (Baker, 1986). The problem of the origin and evolution of planation surfaces and their relationships with the tectonic evolution of cratonic areas has been extensively discussed in the literature since the early works of Davis and Penck and later of King (see, e.g. Baker, 1986; Summerfield, 1991; Gunnell, 1998; Burbank and Anderson, 2001 and references therein). Erosion surfaces may form under different climatic environments, receiving the genetic names of etchplains, peneplains

0169-555X/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved. PII: S 0 1 6 9 - 5 5 5 X ( 0 1 ) 0 0 1 2 9 - 5

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or pediplains (e.g. Fairbridge and Finkl, 1980; Summerfield, 1991). One of the main problems in deciphering the origin of planation surfaces is the precise determination of their geometry. Ancient planation surfaces can be reconstructed either from the preserved summit level (Lidmar-Bergstro¨m, 1996) or from their remnants, considering them as flat surfaces with small slope and assuming a non-eroded topography (Ollier, 1981). In this sense, quantitative analysis is a useful tool to avoid a priori interpretation of data (Baker, 1986). Dating of planation surfaces may be done by means of coeval sedimentation, taking into account their origin as (i) palaeosurfaces covered by sediments and later exhumed, or (ii) as surfaces formed coevally with sediment deposition in adjacent basins (Ollier, 1981). The examples of erosion surfaces developed under compressional tectonic regimes (orogens or intracratonic mountain belts) are few (Amato and Cinque, 1999 and references therein). In these cases, the reconstruction of ancient palaeosurfaces and their relationships with syntectonic sedimentation together with the structural evolution allow data to be obtained about the uplift history of mountain ranges (Amato and Cinque, 1999). The Iberian Chain (Spain) provides a good natural laboratory to study the relationships between tectonics, sedimentation and landscape evolution because (i) its tectonic evolution is well known, (ii) erosion surfaces are well preserved, and (iii) there is a complete sedimentary record for all the Tertiary. These conditions allow us to deal with the methodological problem of defining erosion surfaces from their preserved remnants, and to analyse the problems that arise when dating planation surfaces with correlative sedimentation. Considering the linking between planation surfaces and internally drained basins, the evolution of the Iberian Chain can be considered as an ancient analog for internally drained basins in semi-arid areas of active orogeny as in central Asia (e.g. Bridges, 1990). Erosion surfaces have long been recognised in the central –eastern part of the Iberian Chain (Birot, 1959; Sole´ Sabarı´s, 1978; Simo´n Go´mez, 1983; Pen˜a et al.,

1984; Moissenet, 1985; Gracia-Prieto et al., 1988; Gutie´rrez Elorza and Gracia, 1997). Nevertheless, no consistent model has been proposed to date to explain the complex relationships between erosion, sedimentation and tectonics during the Tertiary in this area. In this study, we analyse the long-term landscape evolution in the central sector of the Iberian Chain, characterised by the development of erosion surfaces and coeval sedimentation in intra-mountain internally drained basins. Finally, a model is presented in order to relate the geomorphological features with the tectonic evolution of the whole area during the Tertiary.

2. Geological and geomorphologic setting The Iberian Chain is an intracratonic mountain range located in the central –eastern part of the Iberian Peninsula (Fig. 1A). It is an elevated area (up to 2300 m a.s.l.) separated from the Pyrenees by continental Tertiary basins. To the SE, the Iberian Chain is cut by normal faults linked to the Neogene opening of the Mediterranean Sea. Its geological history as an independent geological unit began in the Mesozoic, with two stages of rifting (Triassic and Early Cretaceous, with deposition of terrestrial sediments), followed by two episodes of thermal subsidence (Jurassic and Late Cretaceous; Salas and Casas, 1993). The latter were characterised by transgressive marine platform deposition (Salas and Casas, 1993). During the Tertiary (Late Eocene – Oligocene), the whole Iberian plate was subjected to a generalised compression with the subsequent uplift of intracratonic mountain ranges (Guimera` and Alvaro, 1990; Fig. 1B). The change in stress regime from Mesozoic extension to Tertiary compression was driven by the large-scale geodynamic setting of the Iberian plate: Mesozoic rifting was linked to the opening of the Atlantic and Tethys oceans, and Tertiary compression to the convergence between Europe and Africa. Finally, the Neogene was characterised by a SSE – NNW compression, probably originated by subduction –collision in the southern

Fig. 1. (A) Geological sketch of the Central Iberian Chain with the main structures and basins surrounding the Chain. Inset shows the location of the studied area within the Iberian Peninsula. The position of cross-section of (B) is shown. (B) General cross-section of the Iberian Chain showing the envelope for the general upwarping of the chain during Tertiary times. The reference line corresponds to the top of the Upper Cretaceous considered to be at sea level until Eocene times.

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plate margin (Vegas, 1992), and ESE extension related to the opening of the western Mediterranean (Roca, 1994). 2.1. Stratigraphy and rock types The central part of the Iberian Chain contains a whole stratigraphic sequence ranging from the Cambrian to the Quaternary. The variety of rock types influences the development and preservation of planation surfaces, strongly dependent on the lithology. The Paleozoic crops out in the cores of several NW – SE anticlines (Fig. 1A and B). It comprises quartzitic sandstones and shales with limestone and dolostone levels (Ferreiro et al., 1991). The Triassic lies unconformably on the Paleozoic and consists of fluvial sandstones and shales, dolostones and gypsiferous clays. The Jurassic is a marine sequence and homogeneous throughout the studied area with a thickness between 500 and 900 m. It consists of limestones and dolostones and alternating marls and limestones. The Lower Cretaceous consists of alternating episodes of continental sandstones and shales with marine and lacustrine limestones (Soria et al., 1995). The Upper Cretaceous, up to 500 m thick, is also homogenous, with marine platform limestones and dolostones throughout the Iberian Chain (Ferreiro et al., 1991). The Iberian Chain and adjacent continental basins (Fig. 1A and B) preserve a good record of the syntectonic Tertiary sedimentation. This allows us to relate and discuss the formation of planation surfaces with basinal deposits. One of the particularities of the Tertiary basins surrounding the Iberian Chain is that they were internally drained from the Late Eocene until the Late Miocene (Anado´n and Moissenet, 1996; Villena et al., 1996; Mun˜oz-Jime´nez and Casas-Sainz, 1998; Casas et al., 2000a). This implies that local base levels for the formation of planation surfaces were not controlled by fluctuations of sea level but by changes in the sedimentation level and filling of these basins. Paleogene deposits were contemporary with compression in the Iberian Chain. The sediments filling the Paleogene basins are linked to alluvial fans with consolidated conglomerates in the proximal areas passing to sandstones, clays and lacustrine limestones and gypsum in the basin centers. The maximum thickness of the Paleogene (3500 – 5000 m) is found at the cores of the synclines and at the foreland of the

main thrusts (Mun˜oz-Jime´nez and Casas-Sainz, 1998; Casas et al., 2000a,b). The Neogene deposits reach maximum thicknesses of about 700 m, and they spread over large areas within the Iberian Chain and its adjacent basins (Anado´n and Moissenet, 1996; Alcala´ et al., 2000). Neogene sedimentation is also linked to alluvial fans with larger extension of the lacustrine facies. Plio – Quaternary deposits up to tens of meters thick can be found lying unconformably on the older units. Most of these deposits are non-cemented gravels forming alluvial fans at the piedmont of the main reliefs, and fluvial terraces associated to the Turia and Jiloca rivers. Lacustrine deposits of this age are also found in the residual internally drained areas isolated from the fluvial network (Gracia et al., 1999). 2.2. Structure The development of planation surfaces in the Iberian Chain is closely linked with its upwarping history, developed during Tertiary times. Uplift of the Iberian Ranges since the Late Cretaceous involved formation of a structural relief of more than 3000 m (Casas et al., 2000a), with erosion of the elevated areas and deposition in continental sedimentary basins (Fig. 1B). The Tertiary compressional structure of the Iberian Chain shows thrusts and folds with main NW – SE, E –W and NE –SW trends (Guimera` and Alvaro, 1990). Typical wavelengths of the folds are 5 –10 km, contrasting with the large wavelength of the general upwarping of the mountain range (Guimera` and Gonza´lez, 1998; Fig. 1B). Neogene extensional faults are also found in the easternmost part of the Iberian Chain and the Catalonian Coastal Range (Simo´n Go´mez, 1983; Guimera` and Alvaro, 1990). The structure of the studied area is defined by Paleozoic-cored anticlines with main NW – SE and N – S trends. This regional trend strongly conditioned the development and preservation of erosion surfaces. The most important of these folds is located in the Aragonian Range (Fig. 1) and forms a tectonic relief of more than 1500 m with respect to adjacent Tertiary basins. The southern limbs of the basement anticlines show average dips of 20j. Their northern limbs are usually faulted by steeply dipping thrusts that cut across the whole stratigraphic sequence. The NW – SE-trending folds west of the Jiloca depression can be

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followed below the Pliocene and Quaternary deposits, and most of them die out along the N – S line that coincides more or less with the Teruel meridian (Fig. 1A). In an overall view, the areas between anticlines and all the southeastern part of the studied area show shallow dips.

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Normal faults, with sizes ranging from a few centimeters to hundreds of meters, are widespread over the studied area. They are usually associated to the Late Neogene and Quaternary extensional movements (Capote et al., 1981; Simo´n Go´mez, 1983; Simo´n and Soriano, 1993; Corte´s and Casas, 2000).

Fig. 2. (A) Topography of the studied area. Drawn from the Carta Digital de Espan˜a (CDE), Servicio Geogra´fico del Eje´rcito (Spain). See location in Fig. 1A. (B) Cumulative curve of elevation with respect to surface within each contour line in the central Iberian Chain.

24 A.M. Casas-Sainz, A.L. Corte´s-Gracia / Geomorphology 44 (2002) 19–46 Fig. 3. (A) Erosion surfaces in the Central Iberian Chain mapped from the study of aerial photographs and field observations. (B) Histograms indicating the elevation of the mapped remnants of erosion surfaces. Stripped bars indicate the total area of each erosion surface. In S6, the plains of the present-day Gallocanta Lake have been considered. (C) Cumulative curve of the area of remnants of planation surfaces vs. elevation. This curve shows well-defined subhorizontal segments that correspond with the elevation of planation surfaces, in contrast with the cumulative curve of elevation considering all areas (Fig. 2B) that give no key for the definition of planation surfaces.

A.M. Casas-Sainz, A.L. Corte´s-Gracia / Geomorphology 44 (2002) 19–46 Fig. 4. Stereographic aerial photographs showing two remnants of stepped erosion surfaces in the northwestern part of the studied area (see location in Fig. 3). The lower erosion surface (to the SW) corresponds to S5 (1180 – 1200 m) and is developed upon Lower Triassic sandstones. The upper erosion surface (to the NE) is S3 (1300 – 1340 m) and is located upon marine Jurassic limestones.

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Fig. 5. Digital elevation model of the studied area obtained from the CDE (see location in Fig. 6). The larger remnants of erosion surfaces and the effects of the incision of the recent fluvial network can be seen.

Fig. 6. Topographic sections showing the stepping of erosion surfaces in the studied area. See location in Fig. 3.

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These faults are parallel to the main anticlines of the studied area and involve Pliocene and Quaternary rocks. These normal faults show heaves lower than 100 m although they can occasionally reach 300 m of vertical displacement (Corte´s and Casas, 2000). Other recent deformational structures are related to Upper Triassic clays and gypsum, which crop out in the core of diapirs (Moissenet, 1985, 1989; Anado´n et al., 1990; Alonso-Zarza and Calvo, 2000). Karstic collapses due to the solution of gypsum-based rock units (mainly the Upper Triassic and the Middle Miocene)

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are also responsible for some faults and fractures of up to several hundred meters in size (Gutie´rrez, 1995; Calvo et al., 1999). 2.3. Geomorphology The central part of the Iberian Chain shows an average elevation of 1200 m a.s.l. (Fig. 2A and B). In spite of this elevation, relief contrasts and slopes are not pronounced, and the whole area appears as a high plateau (high plains landscape) with minor incisions

Fig. 7. Geological map of the studied area showing the location of cross-sections of Fig. 8.

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Fig. 8. Geological cross-sections and topographic profiles (exaggerated 4:1) through the Neogene basins and erosion surfaces of the Central Iberian Chain.

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Fig. 8 (continued ).

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due to the recent fluvial network (Fig. 2A). Much of the morphology of the studied area corresponds to an ancient landscape inherited from Miocene times when most basins were internally drained. The transition from internally drained basins to an external drainage (mainly driven by the Ebro, Tagus and Turia Rivers; Fig. 2A) brought about the rejuvenation of the ancient landforms and sedimentation of coarser deposits (Gutie´rrez et al., 1996; Gutie´rrez Elorza and Gracia, 1997). Nevertheless, many parts of the Central Iberian Chain are not incised by the tributaries of the main rivers, and the Tertiary alluvial network is still preserved in many areas (Casas et al., 2000a). The present-day climate is continental subarid with periglacial features in the areas situated above 1000 m. Most of the region exhibits daily temperatures below 0 jC more than 90 days per year. The average annual rainfall is between 400 and 800 mm. The present-day fluvial network shows a strong structural and lithological control, with many segments of rivers aligned in NW –SE, N – S and NE – SW direction (Fig. 2A).

3. Landscape features: erosion surfaces Erosion surfaces are by far the most outstanding geomorphologic feature of the studied area. Their remnants occupy more than 1500 km2, 20% of the studied area (Fig. 3A –C). Remnants show areas of up to 50 km2 (or more than 200 km2 when considering the Gallocanta internally drained depression or the Jiloca depression). 3.1. Methods of study and correlation Remnants of erosion surfaces were mapped from aerial photographs at 1:30,000 scale (Figs. 3 and 4). Although automatic methods have also been proposed for the recognition and mapping of flat erosion surfaces (Roessner and Strecker, 1997), in our case, the remnants of erosion surfaces are not completely horizontal and they are strongly dissected. Therefore, the visual identification is necessary to recognise them (Fig. 4). The remnants of planation surfaces can be characterised as areas with flat topography developed on non-horizontal geological units, and normally dissected and separated from other remnants by the

Quaternary fluvial network (Figs. 5 and 6). Finally, the remnants were plotted on 1:50,000 topographic and geological maps and represented in cross-sections to define their variations in elevation and their relationship with the geological framework (Figs. 7 and 8). Correlation of remnants of erosion surfaces is a classical problem in geomorphology especially in tectonically active areas (Amato and Cinque, 1999). In the Iberian Chain, the criteria relying on the morphology of surfaces are difficult to apply because: (i) karstification of remnants is uneven; (ii) the degree of dissection of remnants depends more on their position with respect to the present-day fluvial network than on their relative age; (iii) tilting is difficult to be determined accurately owing to the subsequent dissection, and seems to be more dependent on the position of the remnant than on its age; and (iv) paleosols are not generally preserved, except for clayey fillings of karstic cavities. Furthermore, the development of karst and paleosols may be independent of the age of the remnants and conditioned by their elevation, orientation or subsequent soil formation processes. Nevertheless, further work is in progress to characterise the mineralogical composition of karstic fillings in different remnants of planation surfaces. Quantitative analysis, considering the different elevation of remnants, may give some clues about the age and correlation of these remnants. The histogram of elevation vs. area of remnants of planation surfaces (Fig. 3) gives several well-defined maxima separated by relative minima. Furthermore, remnants with similar heights are systematically distributed in rims surrounding the main elevations of the Iberian Chain. Therefore, we have made a first attempt of correlation on the basis of their topographic height. The presence of erosional scarps (Figs. 6 and 9) separating the remnants also favours the correlation on the basis of height. Errors in this correlation method may arise from tectonic faulting and folding of erosion surfaces. Nevertheless, as will be discussed later, these errors should be minimised when considering the wavelength of particular tectonic structures vs. the general upwarping of the range (Fig. 1B). The distribution of the remnants allows us to reconstruct the initial surfaces (Roessner and Strecker, 1997), assuming ideal non-eroded surfaces prior to fluvial incision.

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Fig. 9. Field view of erosion surfaces. (A) Erosion surface over Jurassic limestones and residual relief of Paleozoic rocks (Sierra Menera anticline). (B) Erosional scarp between S6 (to the right) and S7 (to the left). In the lowermost planation surface (near the bottom of the Jiloca depression), the Upper Triassic – Lower Jurassic crops out. In the uppermost (S6), the Lower Jurassic crops out, which means that the contact between the two remnants of planation surfaces is an erosional scarp. (C) Steep scarp at the eastern limit of the Jiloca depression eroding the NW – SE anticline in Jurassic limestones. The village is located on Plio – Quaternary deposits covering the bottom of the depression. The S7 planation surface is slightly higher than the bottom of the depression and corresponds to the foreground.

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3.2. Planation surfaces Normally, the mapped topographic flats correspond to the eroded folded Mesozoic beds (Figs. 4 and 7), most of them consisting of marine Jurassic and Upper Cretaceous limestones. The Lower Triassic sandstones also develop erosion surfaces (Fig. 4). The Paleozoic sandstones and shales usually form non-flat reliefs, elevated between 100 and 300 m above the surrounding remnants of erosion surfaces (Figs. 7 and 8A), which can be considered as residual reliefs (GraciaPrieto et al., 1988). Incompetent beds, such as the Upper Triassic clays or the Lower Cretaceous sands, are very prone to erosion by the recent fluvial network, and erosion surfaces are not preserved upon these units. The Tertiary deposits are also less resistant rocks and easy to erode by recent fluvial processes although there are occasionally flat surfaces well preserved on Tertiary sandstones and conglomerates (see Fig. 8, cross-section F). This lower resistance and the subhorizontal geometry of Neogene strata in most parts of the studied area make it difficult to distinguish between ancient erosion surfaces or younger mesa morphologies created by denudation of the upper clayey units. In an overall view, the age-dependent morphology correlates with the results obtained by Clayton and Shamoon (1998), who found that the age of the rock is a stronger influence than the rock type in the resistance of rock to erosion. The map of erosion surfaces (Fig. 3A) shows up to seven stepped levels separated by gentle scarps between 100 and 200 m high. The angle of scarps is usually between 15j and 25j (Fig. 9B). The highest levels (S1 and S2) are located around the Albarracı´n Massif, one of the highest Paleozoic massifs within the Iberian Chain, which separates the Mediterranean hydrological basins to the east and the Tagus Basin to the west. They occupy small areas when compared with the following four levels (S3– S6) located at elevations between 1350 and 1050 m (Fig. 3). The remnants of these four main levels define a rim surrounding the Albarracı´n Massif (Fig. 10). There is a strong structural control in the shape and in the plan view of the remnants of these erosion surfaces. In most cases, these remnants are elongated in the NW – SE direction owing to: (i) late erosion of weak units folded in NE – SW direction, (ii) the presence of residual reliefs of Paleozoic rocks with the same

orientation, and (iii) the structural control of erosion by faults or fold limbs. S6 can be continued northwards to the Gallocanta depression (Fig. 2A), ending finally in the Almaza´n Basin, filled with Neogene deposits. The lowest erosion surface (S7) with an average elevation of 1050 m can be considered to coincide partly with the flat bottom of the Jiloca depression (Fig. 9C). This 50  10-km-wide basin shows an N – S to NNW – SSE orientation (Fig. 5). Tilting of remnants of erosion surfaces can be determined from topographic profiles (Fig. 8). In an overall view, the highest (S1 and S2) and lowest (S7) levels are nearly horizontal. In contrast, intermediate levels are tilted toward the north (Almaza´n Basin, Fig. 8, cross-section A) and the east (Jiloca depressions, Fig. 8, cross-sections D and E). This geometry can be interpreted as the result of general upwarping of the Iberian Chain, with a maximum curvature in intermediate zones between Tertiary basins and the highest massifs (Fig. 1B). To the east of the Jiloca depression, there are erosion surfaces with two dominant elevations: one of them between 1125 and 1200 m (S5) and the other about 1350 m (S3, see Fig. 3A – C). This higher level, developed mainly on the Upper Cretaceous limestones, continues to the north into Tertiary basins, the Lower Miocene conglomerates being in continuity with it (Gonza´lez et al., 1998). The remnants of the 1125– 1200-m erosion surface are covered by sediments of Vallesian – Turolian age (Pen˜a et al., 1984; Gutie´rrez Elorza and Gracia, 1997). The highest erosion surface in this zone (S1) is found at the top of the Sierra del Pobo (see Fig. 3), cutting the hinge zone of the N –S anticline at an elevation of more than 1600 m. Toward the Ebro Basin, there is a decrease in elevation (down to 850 m) of the remnants of erosion surfaces (Pen˜a et al., 1984; Soriano, 1990). The relationships between erosion surfaces and Mio – Pliocene deposits can be observed at the southern margin of the Almaza´n Basin (Fig. 8, crosssection A). Exhumed remnants of surface S7 (below 1050 m) surrounded by Miocene rocks and remnants of surface S6 (1050 – 1125) are located at lower elevations than the Mio –Pliocene lacustrine deposits filling the Almaza´n Basin. The top of Tertiary sediments in the Almaza´n Basin is at 1160 m a.s.l. and can be found 5 km north of the these remnants of erosion surfaces.

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Fig. 10. Reconstruction of planation surfaces with their corresponding elevations from the remnants mapped in Fig. 3.

Some of the remnants of planation surfaces are covered with detrital deposits (gravels, and more commonly, red clays) up to 60 – 70 m thick although they can occasionally reach 200 – 300 m in the Sarrio´n Depression (see Fig. 1). These deposits have been considered Pliocene or Plio –Quaternary in age (Gracia-Prieto et al., 1988). This is especially evident in the lowermost erosion surface that forms the bottom of the Jiloca depression, covered by gravels with pediment morphology and deposited by alluvial fans sourced in its eastern margin (Pen˜a et al., 1984). The scarps between different erosion surfaces may also be covered with slope deposits consisting of angular pebbles of probable periglacial origin. Limestone pavements and dolines of several types, up to 500 m in diameter (Gutie´rrez and Pen˜a, 1979; Pen˜a et al., 1984), are present in the higher levels of erosion surfaces, espe-

cially in the southwesternmost part of the studied area (Albarracı´n Massif). Cavities several meters wide, with a sedimentary filling of red clays and blocks from limestones of the surrounding rock, are also present in most parts of the remnants of erosion surfaces. These features have been interpreted as surface karstic landforms. Recent studies suggest karstic corrosion processes to explain the origin of large poljes like the Gallocanta depression (Gracia et al., 1999).

4. Tertiary sedimentation The development of planation surfaces was contemporary with the uplift of the Iberian Chain and sedimentation in the sedimentary basins. The Paleogene sediments are preserved in the Almaza´n, Ebro

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Fig. 11. Relationships between remnants of planation surfaces and Neogene deposits in the northern part of the area studied (see Fig. 3 for location).

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Fig. 12. (A) Synthetic stratigraphy of the Neogene in the continental basins surrounding and within the Iberian Chain. Data from Calvo et al. (1993), Anado´n and Moissenet (1996), Villena et al. (1996) and Casas et al. (2000a). (B) Average elevation of the base of continental Neogene series in the different internally drained basins: (1) Late Aragonian limestones, (2) Vallesian limestones, (3) Turolian limestones and (4) Ruscinian limestones.

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and Aliaga basins (Fig. 1A). They are related to alluvial fans sourced in the uplifted areas of the Iberian Chain. In the Almaza´n Basin, the thickness of these deposits increases progressively in a northward direction, corresponding with the general upwarping of the mountain chain. Paleogene deposits are usually covered by the Neogene in the southern border of the Almaza´n Basin, and only limited outcrops can be seen at the limits with the Aragonian Range. Nevertheless, they can be identified in seismic reflection profiles (Casas et al., 2000a). The Almaza´n, Aliaga and Ebro basins were separated from each other during the Paleogene, but the correlation of sedimentary sequences indicates a certain parallelism in the tectonic activity recorded by the sedimentation (Villena et al., 1996; Casas et al., 2000a). In many places (Fig. 11), planation surfaces in the Iberian Chain grade to the sedimentary surface of Neogene and Quaternary deposits or are covered by them (Gutie´rrez Elorza and Gracia, 1997; Fig. 8, cross-section A). These Neogene deposits are related to the internal basins of the Iberian Chain (Almaza´n, Calatayud and Teruel basins, see Fig. 1), which show a complete sedimentary record of continental deposits from the Upper Oligocene to the Pliocene (Fig. 12A). They are relatively well dated with micromammals and magnetostratigraphy (Lo´pez-Martı´nez et al., 1987; Anado´n and Moissenet, 1996; Alcala´ et al., 2000). The maximum thickness of the Neogene deposits is about 500 m, with average sedimentation rates between 0.2 and 0.5 mm/year. The Neogene succession comprises four major stratigraphic units that can be correlated between the sedimentary basins of the study area. The boundaries between these stratigraphic units are defined on the basis of disconformities, where sudden changes in grain size and lithology are recorded. According to Calvo et al. (1993), the four discontinuities correspond to the Intra-Ramblian, Intra-Middle Turolian, Late Ruscinian – Early Villafranquian and Intra-Villafranquian. The upper disconformities are interpreted to correspond with higher humidity periods (Calvo et al., 1993). All the Neogene deposits from the Lower Miocene to the Upper Pliocene are of continental origin and were sedimented in internally drained basins (Anado´n et al., 1990; Anado´n and Moissenet, 1996). These basins (Almaza´n, Calatayud and Teruel basins) were

isolated from each other and also from the large foreland basins located at the borders of the Iberian Chain (Ebro and Tagus basins) until the Late Miocene, when the sedimentation areas became larger and some of the sedimentary thresholds were overlapped (Armenteros et al., 1989). The Lower –Middle Miocene deposits are mainly conglomerates and sandstones, with some limestones and gypsum deposits confined to the southernmost part of the Teruel Basin (Anado´n et al., 1997). From the Aragonian – Vallesian limit, limestones and clays dominate. Gypsum deposits are abundant in the Lower Miocene units of the Calatayud Basin and in the Upper Miocene units in the Teruel Basin (Anado´n et al., 1990, 1997; Anado´n and Moissenet, 1996; Alcala´ et al., 2000). Turolian and Ruscinian (Upper Miocene –Pliocene) units are characterised by widespread limestone deposition (Pa´ramo limestones) and occasionally evaporites (Alonso-Zarza and Calvo, 2000). The study of the Tertiary lacustrine sedimentary sequences preserved in the Teruel and Calatayud basins and their rodent faunas suggest semi-arid climatic conditions during the Neogene, with more humid and cooler climatic episodes during Aragonian (14 Ma), Vallesian (9.4 Ma) and Turolian (7 Ma) times (van der Meulen and Daams, 1992; Alcala´ et al., 2000; Alonso-Zarza and Calvo, 2000).

5. Discussion: a model for landscape evolution in the Iberian Chain To integrate the different data obtained, especially those referring to sedimentation in continental basins and the morphological evolution of the Iberian Chain, it is necessary to take into account the probable age for the erosion surfaces, their relationships with Paleogene and Neogene sedimentation and the denudation rate of the mountain chain during the Tertiary. In Sections 5.1 –5.3 and 6, we analyse the problem of the age of planation surface and discuss the different hypotheses proposed for their origin. 5.1. The problem of dating planation surfaces One of the main problems that arise when trying to establish an evolutionary model for the geomorphology of an area is the dating of the planation surfaces,

A.M. Casas-Sainz, A.L. Corte´s-Gracia / Geomorphology 44 (2002) 19–46

especially when there is a series of stepped surfaces occurring in the Iberian Chain (Fig. 13). In our opinion, there are three questions that must be taken into account in this subject.

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(1) According to the general evolution of the landscape, when several stepped surfaces are present and separated by erosional scarps, the lower surfaces are younger than the higher. This rule can only be

Fig. 13. Conceptual sketch showing the two end-member models for the origin of planation surfaces and the difficulty for dating erosion surfaces (only in the first case). (A) Stepped planation surfaces formed by general upwarping of the mountain chain. (B) Scarp formation triggered by faulting. In both cases, the exact age of formation of each planation surface is different for each segment and only an approximate or mean age can be assigned.

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discarded in cases of large faults limiting tectonic units (Fig. 13B). In this case, later erosion can exhume older lower planation surfaces (Gracia-Prieto et al., 1988). (2) When a planation surface grades to flat-lying sediments, its age may be either contemporary with the strata located at the same level (erosion and sedimentation at the same time; Fig. 13A, stage 3) or older than these strata (filling of an ancient basin and covering of the old surface; Fig. 13A, stage 6). (3) The age of each planation surface may not be homogeneous. That is, the creation of an erosion surface and associated scarp retreat (in pediplains, as is the case of the Iberian Chain; Gracia-Prieto et al., 1988; Gutie´rrez Elorza and Gracia, 1997) may continue even when the next erosion surface at a lower level is forming (Fig. 13A, stages 4 – 6). In this case, each stage may comprise segments developed at different ages, and the average age that should be considered is the age of formation of the erosional scarp with the previous planation surface (Twidale and Bourne, 1998). 5.2. Retreating scarps vs. deformation of erosion surfaces Scarps between erosion surfaces in the Iberian Chain have been since long interpreted as the result of deformation by normal faulting of a single erosion surface (Fundamental Erosion Surface of the Iberian Chain; Sole´ Sabarı´s, 1978; Pen˜a et al., 1984; Simo´n Go´mez, 1983; Gutie´rrez Elorza and Gracia, 1997). Many of these escarpments would then correspond to exposed fault surfaces or fault scarps. According to these authors, this erosion surface developed at the end of the Miocene and Pliocene, and deformed by the Pliocene –Quaternary extensional tectonics including doming and rifting. These extensional processes would be the responsible for the present-day elevation of the Central Iberian Chain. The geometry of the different flats would then be the result of normal faulting of a single surface, and therefore, an indicator of post-Pliocene deformation with elevations of more than 2000 m a.s.l. (Simo´n Go´mez, 1983). Nevertheless, Gracia-Prieto et al. (1988) and Gutie´rrez Elorza and Gracia (1997) admit the existence of several planation surfaces separated by erosional scarps in the Gallocanta area although considering a main

erosion surface (their S2), coinciding with the deformed Fundamental Erosion Surface, folded and faulted after the Miocene. In this case, faulting at the edge of Mio –Pliocene basins would be responsible for lowering the local base level for forming some erosional scarps (Fig. 13B). Nevertheless, in our opinion, there are some data that invalidate the model of deformation of a unique surface. . The age of the ‘‘Fundamental Erosion Surface’’ cannot be constrained within a determined time interval (also suggested by Moissenet, 1985; Gutie´rrez et al., 1996; Gracia et al., 1999). The fragments of erosion surfaces mapped in this work (most of them characterised previously as a single erosion surface) correlate with deposits of different age in different places: Lower Miocene, Middle Miocene, Upper Miocene, and Pliocene (Fig. 11). . In a general overview, in the eastern margin of the Albarracı´n Massif, the lower elevation of the erosion surfaces corresponds with lower erosion levels of the Mesozoic sedimentary cover (Fig. 8, cross-sections D and E; Fig. 9B). This can be seen in the map of eroded thickness, considering the overall envelope of anticlines and synclines, and disregarding the erosion linked to the recent fluvial network (Fig. 14, see also Guimera` and Gonza´lez, 1998). This fact also invalidates the hypothesis that the Iberian Chain was strongly eroded and subsequently uplifted in post-Pliocene times by doming and rifting (Simo´n Go´mez, 1983): the rocks cropping out in large parts of the Iberian Chain are Upper Cretaceous in age. Erosion was only important at the cores of the main anticlines (Gonza´lez et al., 1998; Guimera` and Gonza´lez, 1998). . In many places within the studied area, there is no evidence that erosion surfaces have been faulted or dislocated by tectonic structures, either compressional or extensional, and there are no geological faults coinciding with the scarps or neighbouring zones (Corte´s and Casas, 2000). In some cases, the displacement of the fault is contrary to the stepping of erosion surfaces, or faults are completely levelled by planation surfaces (Fig. 8, cross-sections A and D). Nevertheless, we do not discard that locally remnants of erosion surfaces at different elevations may be the result of fault displacement (normal or reverse) subsequent to their formation. This would give smaller scale irregularities within the general pattern of stepped planation surfaces.

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Fig. 14. Contour map showing the eroded thickness (in kilometers) of the sedimentary sequences in the central Iberian Chain. Contours were drawn from a network of square cells 5  5 km wide to avoid short wavelength variations associated to local structures. The eroded thickness was calculated in each point considering the thickness of the stratigraphic log in each area comprised between the outcropping unit and the top of the Upper Cretaceous.

. Finally, the remnants of erosion surfaces can be arranged in groups with similar elevations (Figs. 3B and 10), and these groups form stepped belts with rim geometry surrounding the main massifs. There is not a dominant elevation but four groups of remnants at average elevations of 1100, 1200, 1300 and 400 m occupying similar areas across the chain. It is not clear which one may correspond to the Fundamental Ero-

sion Surface. This supports the hypothesis of an erosional origin for stepped surfaces because it does not seem probable that after dislocation of a single surface by different faults and monoclines, its different segments remain within fixed elevation limits (see also Gutie´rrez Elorza and Gracia, 1997; Gracia et al., 1999). Evidently, this applies to parts of the chain that have evolved in connection with the same basin, as is

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the case in this study. The southwestern part of the Iberian Chain that is related to the Tagus basin (Fig. 10) probably shows different elevations for planation surfaces and related deposits. These lines of evidence support the hypothesis that the erosion surfaces formed sequentially, stepping from higher (older) levels to lower (younger) levels. The most recent pre-Quaternary planation surface is then located at an elevation of about 1050 m, and its formation would be favoured by the existence of a non-resistant stratigraphic unit corresponding to the Upper Triassic clays and gypsum (Gracia et al., 1999). The older age of the higher topographic flats is also consistent with the more intense karstification of the higher erosion surfaces. Nevertheless, no direct correlation can be established between those variables because of the climatic (humidity and temperature) control of these karstic features and their variation with topographic elevation. 5.3. Tectonics –erosion – sedimentation relationships From the relationships between erosion surfaces and sedimentation in the internal basins of the Iberian Chain, two different models can be established relating the age of erosion and sedimentation (Fig. 13; see also Ollier, 1981): . The development of each erosion surface was coeval with the stratigraphic levels located at the same topographic elevation, which are found onlapping onto the planation surface. In this case, the age for the end of the erosion surface would coincide with the age of the deposits partially covering its surface (Fig. 13A). . There is no relationship between the age of each erosion surface and the age of sediments covering them (Fig. 13A). All the scarps formed before the Neogene sedimentation extended toward the topographically elevated areas of the Iberian Chain. Sediments covered the previously formed flats and scarps, onlapping from the centers of the Almaza´n, Teruel and Calatayud basins toward their margins (Mele´ndez et al., 1979). From the results presented in this work, the second hypothesis is better supported by the data (Figs. 13B and 15) since topographic flats at higher levels within the Iberian Chain are covered, or laterally related, with younger deposits. Mio – Pliocene lacustrine limestones

in the Almaza´n Basin are at 1160 m a.s.l.). This means that the base level increased in elevation during the Neogene, and sedimentation extended to larger areas in more recent times (onlap of sediments in basin margins, Casas et al., 2000a). This implies that if erosional flats were located at approximately the same elevation throughout large parts of the basin margins, the accommodation space for sediments may have experienced dramatic increases in the change from filling a basin limited by scarps to the overlying topographic flat. This change, considering a constant discharge of sediments and no climatic change, would bring about the expansion of lacustrine areas and widespread limestone sedimentation. This factor should be considered since its sedimentary signal may interfere with changes in sedimentation due to climatic changes. The last Mesozoic marine sediments in the central part of the Iberian Chain are Late Cretaceous in age. Sedimentation during the Paleocene was of continental origin, still related to marine basins to the north and east. The general uplift of the whole area began during the Late Eocene. The Iberian Chain became the source area for the surrounding sedimentary continental basins (Fig. 15), and compressional structures developed contemporaneously with sedimentation (Guimera` and Alvaro, 1990; Casas et al., 2000a). The main stage of compressional mountain building ended in the Late Oligocene (Alvaro et al., 1978; Guimera` and Alvaro, 1990; Villena et al., 1996). The central part of the Iberian Chain shows a wide negative Bouguer anomaly of up to 110 mgal centered west of Teruel (Salas and Casas, 1993; Mezcua et al., 1996). This indicates maximum crustal thickness (up to 43 km) in this area, gradually diminishing toward the north and south, and more abruptly toward the east (Salas and Casas, 1993). Starting from a continental crust thinned during the Mesozoic, crustal thickening was achieved during the Late Eocene– Oligocene compression. This thickening was related to the formation of thrusts with northward displacements of more than 10 km (Guimera` and Alvaro, 1990; Casas et al., 2000a,b) in the northern border of the Iberian Chain. From this period onwards, very few structures formed and the remaining basins within the Iberian Chain were filled without a clear tectonic control (Casas et al., 2000b). The end of important tectonic movements since the Early Miocene within the Iberian Chain is also supported

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Fig. 15. Conceptual sketch showing the proposed model for the formation and covering of planation surfaces in the Central Iberian Chain.

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by the similar elevation shown by stratigraphic levels in the different Neogene basins (Fig. 12B). The change from tectonically driven subsidence to the atectonic filling of basins is accompanied by a slowing in the sedimentation rate within the continental basins from the Early Miocene (Fig. 16). Subsidence in the basins surrounding the Iberian Chain must have continued until at least the end of the Early Oligocene since sediments of this age are presently found at elevations below sea level in the Almaza´n and Ebro basins (Casas et al., 2000a,b). The late filling of the internal basins (Early Miocene– Pliocene) must have taken place with a gradual increase in the elevation of base level as the basins were filled. In these cases, isostatic sinking due to the sedimentary charge in internal basins (see for example, Sclater and Christie, 1980; Casas et al., 2000b) would not be enough to compensate for the increase in elevation due to basin overfilling. In spite of the large number of palaeontological sites (see for example, Lo´pezMartı´nez et al., 1987; Alcala´ et al., 2000), no palaeoelevation data in the Central Iberian Chain are available to date. The gradual change toward more

humid and cooler climates from the Early to the Late Miocene (van der Meulen and Daams, 1992; Calvo et al., 1993; Alonso-Zarza and Calvo, 2000) could be caused partly by this increase in elevation.

6. Long-term landscape evolution in the Iberian Chain From the considerations above, the exposed Neogene landscape of the Central Iberian Chain can be interpreted as a sequence of stepped erosion surfaces surrounding areas uplifted by compression and crustal thickening (Fig. 15). The general upwarping defines a long wavelength curve descending toward the sedimentary basins located in its margins. In other cases of erosion surfaces, their dissection was controlled by the tectonic activity (Peulvast et al., 1996). In the case studied, the late dissection of erosion surfaces was driven by the change from an internal drainage (until the Late Miocene) to the recent external drainage (Fig. 15). Except for small endorheic areas as the Gallocanta lake, the present-day drainage is open to the

Fig. 16. Thickness of sediments (without considering the effect of compaction) with respect to time in the depocenters of Tertiary continental basins related with the Iberian Chain. The Ebro and Tagus basins are located at the northern and western margins of the chain, respectively. The Almaza´n, Montalba´n, Calatayud and Teruel basins are located within the Chain (see location of basins in Fig. 1).

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Mediterranean Sea (by the Ebro River and its tributaries) and the Atlantic Ocean (Almaza´n Basin, captured by the Duero River). The regional base level has been dramatically lowered from an elevation of more than 1000 m (final base level in internal basins) to the present-day sea level. Nevertheless, erosion surfaces were preserved for more than 20 Ma, probably favoured by the structural barriers created by scarps in hard Mesozoic limestones (Gunnell, 1998). Large-scale erosion surfaces have been recognised in cratonic areas. They are usually associated with long-term landscape evolution (Fairbridge and Finkl, 1980; Kooi and Beaumont, 1994; Peulvast et al., 1996; Lidmar-Bergstro¨ m, 1996; Roessner and Strecker, 1997; Taylor and Howard, 1998; Gunnell, 1998; Johansson, 1999). The development of erosion surfaces in the Central Iberian Chain was a smaller scale process, involving stepped planation surfaces with amplitudes of several kilometers each. The formation of planation surfaces during the Tertiary was probably favoured by semi-arid climatic conditions (Moissenet, 1985), and the planation surfaces can be considered as pediplains (Gracia-Prieto et al., 1988; Gutie´rrez Elorza and Gracia, 1997). Nevertheless, periods with higher humidity would be responsible for etching and formation of karstic fillings and red clays over planation surfaces. The time of evolution for the retreating scarps is relatively short (Fig. 15), compared with large scarps associated with rifted margins (Kooi and Beaumont, 1994) that agrees with the elongated shape of flat areas (with an average width of about 10 km) between scarps. The small width of scarp zones is also consistent with their development under an arid climate according to models of scarp evolution presented by Kooi and Beaumont (1994). The formation and retreat of erosional scarps between successive levels of erosion surfaces in the Central Iberian Chain during the Oligocene and the Miocene could be related either to climatic changes or tectonics. In contrast with other examples of stepped surfaces (e.g. Ollier, 1981), changes in sea level cannot be invoked as the cause for the stepped erosion surfaces since the basins surrounding the Iberian Chain were continental, internally drained from the Late Eocene until the Pliocene (Anado´n and Moissenet, 1996). Therefore, they would not feel the effects of base level lowering from sea level change. In our

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opinion, in the Iberian Chain, the formation of erosional scarps must have been a tectonically driven process caused by: (i) the uplift of the Iberian Chain due to compressional structures and crustal thickening and (ii) subsidence due to tectonic load and sedimentary charge in adjacent basins. These two processes resulted in large-scale upwarping of the whole area (see Fig. 1B). The formation of planation surfaces took place during active tectonic uplift, and no tectonic quiescent period was necessary for them to develop (Kennedy, 1962; Amato and Cinque, 1999). The period of formation of the erosion surface sequence must have been of about 15 –20 Ma. Denudation can be accurately defined owing to the good stratigraphic control of Mesozoic sequences (Fig. 14). From volumetric calculations, it can be estimated that the average eroded thickness in the studied area is 0.58 km. Contours of eroded thickness (Fig. 14) show a strong structural control, elongated in NW – SE direction. Erosion shows maxima (1400 m) in the crests of Paleozoic-cored anticlines and minima in the synclines. Considering the age of upwarping in the Iberian Chain, denudation rates give relatively low rates during the Tertiary (0.02 mm/year, below the global average). These values are consistent with old orogens with mean elevations of about 1000 m (see Ollier, 1981; Einsele, 1992; Pinet and Souriau, 1988; Burbank and Anderson, 2001). Consistently, Casas et al. (2000a) showed that although during the Paleogene, this part of the Iberian Chain was uplifted, the main source areas of the Almaza´n Basin were located farther north. Nevertheless, the composition of pebbles in conglomerates of the southern part of the Almaza´n Basin and the southern provenance of alluvial fans indicate that the Central Iberian Chain was probably a weak, but constant, source area during the Tertiary (Casas et al., 2000a). Several cycles of increasing/decreasing of the [tectonic uplift/sediment supply] relationship have been recognised in the Paleogene sedimentary record of the basins surrounding the Iberian Chain (Villena et al., 1996; Mun˜oz-Jime´nez and Casas-Sainz, 1998; Casas et al., 2000a,b). The origin of the sequence of erosion surfaces could be related to episodic uplift. Nevertheless, it is difficult to establish a direct relationship between these episodes and the formation of scarps since other variables may be involved (Gillchrist and Summerfield, 1991).

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The strongest change in the sedimentary conditions in the internal basins of the Iberian Chain was caused by the change from an internal to an external drainage as occurred in all the large internal basins within the Iberian Peninsula (Ebro, Duero and Tagus). This change was diachronous (Gutie´rrez et al., 1996). The capture of the Tertiary internal basins has attained different degrees in different basins that can account for the different development of stepped Quaternary terraces and also for the different average elevation of the Duero, Tagus and Ebro basins. The change in drainage conditions brought about the lowering of the regional base level, and therefore, the incision on the ancient surfaces and scarps and the formation of alluvial fans on scarps related to large drainage areas. Although slowed by structural barriers created by the scarps, denudation rates abruptly increased after the capture of the whole drainage system during the Quaternary.

7. Conclusions In the Central Iberian Chain (Spain), several stepped erosion surfaces are determined from the study of aerial photography and field observations. Most of the remnants of erosion surfaces are developed on Mesozoic limestones, and form a rim around high massifs within the Iberian Chain. They are distributed at average elevations of 1600 (S1), 1450 (S2), 1350 (S3), 1280 (S4), 1160 (S5), 1050 (S6) and 1000 (S7) m a.s.l. They are separated by erosional scarps and show karstic features on their surfaces. The lower levels (S6 and S7) are covered by Neogene sediments corresponding to the internal basins of the Iberian Chain that indicates that they were formed before the end of the filling stage of these basins. Tilting of the intermediate levels can be related to the curvature of the general upwarping of the compressional mountain chain. From their relationships with sedimentation, it can be stated that erosion surfaces formed during the Oligocene and Early Miocene, contemporary with the uplift of the whole area and crustal thickening caused by compression in the inner part of the Iberian plate. The covering of the planation surfaces by horizontal sediments took place during a stage of low tectonic activity and progressive filling of the internal basins, inducing a rise in the lacustrine base levels. From the Pliocene, the internal basins were captured by the

externally drained fluvial network, causing the rejuvenation and incision of the ancient surfaces.

Acknowledgements This work was supported by Project PB97-0997 of the Direccio´n General de Ensen˜anza Superior (Spain). The authors are grateful to Martin Stokes for his careful revision of the manuscript.

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