Shortening, structural relief and drainage evolution in ......Abstract: The Atlas, Eastern Cordillera and Pyrenees are thick-skinned thrust-fold belts formed by tectonic inversion

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Shortening, structural relief and drainage evolution in inverted

rifts: insights from the Atlas Mountains, the Eastern

Cordillera of Colombia and the PyreneesQ1

ANTONIO TEIXELL* & JULIEN BABAULT

Departament de Geologia, Universitat Autonoma de Barcelona,

08193 Bellaterra (Barcelona), Spain

*Corresponding author (e-mail: [email protected])

Abstract: The Atlas, Eastern Cordillera and Pyrenees are thick-skinned thrust-fold belts formedby tectonic inversion of rift basins in continental settings. A comparison of shortening betweenthem shows a gradation from 20–25% in the central High Atlas, to 25–30% in the Eastern Cordil-lera, and c. 40% in the Pyrenees. Accordingly, there is a structural variation from interior zoneswith low structural relief and isolated basement massifs in the first two cases, to an axial culmina-tion of stacked basement thrust sheets in the Pyrenees. This results in marked topographic and drai-nage variation: the High Atlas and Eastern Cordillera contain axial plateaus dominated bystructure-controlled longitudinal rivers and orogen flanks with slope-controlled transverse rivers,whereas the Pyrenees show a two-sided wedge profile dominated by transverse rivers. In spiteof singularities exhibited by each orogen, we propose that this spatial variation can be understoodas reflecting different degrees of evolution in mountain building. Rapidly incising, transverse riversare capturing earlier longitudinal streams of the Atlas and Eastern Cordillera, thus reducing theiraxial plateaux, which will eventually disappear into a transverse-dominated drainage. This patternof landscape evolution may be characteristic of inversion orogens as they develop from initialstages of inversion to full accretion.

The Atlas Mountains, the Eastern Cordillera ofColombia and the Pyrenees are typical examplesof thrust-fold belts formed by tectonic inversion offormer continental rifts. Rift inversion is a commonmodality of compressional deformation in intra-plate regions, in addition to other mechanisms char-acteristic of craton deformation such as lithosphericfolding or upthrusting of isolated basement mas-sifs (e.g. Ziegler et al. 1998). Inverted rifts typi-cally show a lesser amount of orogenic shorteningthan collisional mountain belts at plate-boundariesbecause they lack powerful slab-pull forces by asubducting oceanic plate, and are driven by far-fieldstresses instead.

In spite of the singularities in structural styleexhibited by each case, we aim to illustrate thatthe studied inverted rifts possess some commoncharacteristics in terms of topography and drainageevolution, and at least two of them demonstrablydiffer from self-similar, growing wedge modelscommonly envisaged for collisional belts. Hovius(1996) remarked that many actively uplifting moun-tain belts around the world have simple drainagepatterns transverse to their main structural trend.On the other hand, Babault et al. (2012) show thatin the High Atlas hinterland, the drainage is longi-tudinal and inherited from an early stage of fluvialorganization controlled by the tectonic structures

developed during upper crustal folding and thrust-ing. Amplification of regional slope perpendicularto the tectonic trend in the western High Atlas bycontinued crustal shortening and thickening trig-gered a new organization of the drainage systemtowards the regional slope, thus recording an evol-ution from longitudinal to transverse-dominateddrainage that may represent a common mecha-nism of fluvial network development in mountainbelts where lithospheric convergence progressivelyincreases the regional slopes.

In this paper, we first review and synthesize thestructural features of each of the three mountainbelts (largely based on earlier works; Figs 1–3),showing that they represent different amounts ofdeformation and cumulative shortening, and thenmake a comparison of their drainage network, whichwe propose defines a consistent pattern of drain-age gradation from longitudinal- to transverse-dominated.

The High Atlas

The Atlas Mountains of Morocco formed during theCenozoic in the interior of the African plate but inclose association with the Europe–Africa conver-gence (Mattauer et al. 1977). They consist of two

From: Nemcok, M., Mora, A. R. & Cosgrove, J. W. (eds) Thick-Skin-Dominated Orogens: From InitialInversion to Full Accretion. Geological Society, London, Special Publications, 377,http://dx.doi.org/10.1144/SP377.14 # The Geological Society of London 2013. Publishing disclaimer:www.geolsoc.org.uk/pub_ethics

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thrust-fold belts (the NE-trending Middle Atlasand the ENE-trending High Atlas), and a wide archwith modest post-Palaeozoic deformation (the Anti-Atlas; Fig. 1). The High and Middle Atlas belts werecreated by the tectonic inversion of pre-existingextensional/transtensional basins of Triassic–Jur-assic age (Mattauer et al. 1977; Beauchamp 1988;Warme 1988; Frizon de Lamotte et al. 2000; Gomezet al. 2000; Pique et al. 2002; Teixell et al. 2003,Arboleya et al. 2004), whereas the Anti-Atlas,lying originally outside of the rift basins, conformsto a large (100 km-scale or lithospheric-scale)domal fold cut by minor faults (Guimera et al.2011).

We will focus this review on the High Atlas, a50–120 km wide belt of thrusts and folds deforminga thick succession of sedimentary rocks, dominantlyJurassic in age. The structural style of compressivedeformation is mainly thick-skinned (Frizon deLamotte et al. 2000; Teixell et al. 2003), whereasin the southern part of the central High Atlas thereis a belt of south-verging detached folds and thrusts(Laville et al. 1977; Beauchamp et al. 1999; Teson2009). This is related to salt or other favourableweak layers in the area. Salt diapirs formed duringthe earlier extensional episode are occasionallyobserved (Teixell et al. 2003; Michard et al. 2011;Fig. 4).

The cross-sections of Figure 4a, b illustrate thecharacteristic structural style of the central HighAtlas. Compressional deformation is heteroge-neously distributed: narrow deformation bands con-stituted by anticlines or thrust faults are commonlyseparated by broad synclines or tabular plateaux.

Variations in Mesozoic stratigraphy and thicknessacross many thrust faults attest to their origin assynsedimentary extensional faults. The trend ofthe main individual thrusts and folds is NE–SW,slightly oblique to the general orientation of theHigh Atlas belt (Fig. 1), and whose inheritanceargues for oblique rifting during the Jurassic(El Kochri & Chorowicz 1995; Arboleya et al.2004). Consistent with the original dip of the Meso-zoic basin-bordering faults, the present High Atlasis a doubly verging chain, in which a stronglyexhumed internal zone cannot be defined as inother orogenic belts. In fact, owing to the moderatedegree of inversion, in much of the interior of thechain, basement is at lower structural elevationthan in the peripheral forelands (Fig. 4a, b). In thecentral High Atlas there is an axial plateau-likestructure (e.g. Fig. 4b), where the topographic sur-face is at an elevation of ca. 2000 m. Hence, mostof the total orogenic shortening across the HighAtlas is concentrated in its northern and southernmargins (e.g. Bennami et al. 2001; Teson 2009) Q2.In addition to thrusting, buckle folding (not necess-arily fault-related) is also common in the HighAtlas. These folds, even if tight, may involve base-ment rocks (e.g. as observed at the basement massifSE of Midelt, Fig. 1). Section construction resultedin less line-length shortening in basement thanin the post-palaeozoic cover (Teixell et al. 2003);additional shortening mechanisms such as homo-geneous flattening or underthrusting must haveoccurred in the former. In spite of these difficulties,total orogenic shortening from the restoration ofpost-palaeozoic layers in the central High Atlas

Fig. 1. Te

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sketch map of the Moroccan High Atlas Mountains, indicating the lines of section of Figure 4a, b.

A. TEIXELL & J. BABAULT

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has been estimated as 24% for the eastern sectionof Figure 4a and 18% for the western section ofFigure 4b (Teixell et al. 2003). Total shorteningstill decreases further to the west, in spite of highertopographic elevations and wide exposure of base-ment in the western High Atlas, a fact not comple-tely understood yet.

Timing of deformation and surface uplift

The Cretaceous post-rift conditions in the Atlasdomain gave way to the initiation of compressionaldeformation in the latest Cretaceous or earliestCenozoic. Adjacent to the present-day High Atlasis a discontinuous system of foreland basins fil-led with Neogene deposits (the Souss, Ouarzazate,Haouz–Tadla and Moulouya–Missour basins;

Fig. 1). Cenozoic sediments provide timing of therift inversion process from tectonics–sedimentationrelationships, although there have been markeddiscrepancies among different authors that havefaced the subject (e.g. Laville et al. 1977; Fraissinetet al. 1988; Gorler et al. 1988; Frizon de Lamotteet al. 2000). For the main deformation episode wefollow the recent attributions by Teson & Teixell(2008) and Teson et al. (2010), which combinefield observations with magnetostratigraphy.

The earliest indications of folding are providedby the uppermost Cretaceous to Eocene inlierswithin the High Atlas or the southern thrust front.They occasionally display angular and progressiveunconformities (Laville et al. 1977; Froitzheimet al. 1988; Teson 2009), indicating that contrac-tional deformation has been initiated in this time.

Fig. 2. Tectonic

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sketch map of the Eastern Cordillera of Colombia. This map is a synthesis of Toro et al. (2004), Parraet al. (2009b) and unpublished maps by ICP-Ecopetrol. Indicated is the line of section of Figure 4c. CC, CentralCordillera; HT, Honda thrust; CT, Cambao thrust; BT, Bituima thrust; US, Usme syncline; TT, Tesalia thrust;GT, Guaicaramo thrust.

STRUCTURE AND DRAINAGE IN INVERTED RIFTS

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However, the sparse occurrences and degree ofdevelopment of the unconformities indicate thatthis deformation was very limited, consisting of iso-lated fold ridges of local extent (occasionally salt-related) growing within the post-rift basin. Thepalaeogeography of the correlative subtabular lime-stone deposits of Paleocene to mid Eocene age stilldoes not conform to a foreland basin system (Herbig& Trappe 1994; Teson & Teixell 2008). Above thelimestone, fluvial–alluvial red shales, sandstonesand microconglomerates of mid to late Eoceneage, occurring only in the southern frontal thrustbelt of the High Atlas, have been interpreted as thefirst foreland basin deposits (El Harfi et al. 1996;Frizon de Lamotte et al. 2000; Teson & Teixell2008). It must be pointed out, however, that synse-dimentary deformation structures have never beenobserved in these deposits, and that fission-trackstudies have failed to detect contemporaneous exhu-mation in the High Atlas hinterland. Still, prove-nance from an eroding High Atlas seems probable(Teson 2009); microconglomerates are composeddominantly of recycled pebbles from lower Cretac-eous red beds (i.e. approximately the youngestdeposits of the Mesozoic Atlas basins), attesting tomoderate deformation and exhumation prevailingduring the entire Eocene.

The best record for the age of main thrustingin the Atlas comes from the Ouarzazate Basin(Fig. 1) and the marginal thrust belt adjacent to it.The main infill of the basin (Ait Kandoula fm)has been recently dated as Middle to Late Mio-cene (Teson et al. 2010). This formation rests

unconformably over selected thrusts and folds,whereas it is affected by others (Fraissinet et al.1988; Teson & Teixell 2008). Hence, main thrustingmust have initiated sometime within a generalizedsedimentary hiatus below the Ait Kandoula For-mation (Oligocene to early Miocene), and has con-tinued at similar rate until the Quaternary (Sebrieret al. 2006; Arboleya et al. 2008, Pastor 2008).Apatite fission track ages of 9–25 Ma recorded inthe High Atlas of Marrakech (Missenard et al.2008; Balestrieri et al. 2009) and thermal modellingin the central High Atlas (Barbero et al. 2007) arecompatible with the Neogene age of main thrust-ing. Shortening and crustal thickening since lateEocene times, probably with relatively faster ratessince the Miocene, has led to surface uplift in theAtlas.

However, an enigmatic feature of the Atlasregion as a whole, which has received the attentionof many geoscientists, is its high topographic ele-vation compared with the modest values of shorten-ing and crustal thickening. An inverse correlationbetween shortening and elevation along the centralHigh Atlas belt indicates that crustal thickeningdoes not fully explain the observed topography andsuggests a mantle-sourced, thermal contributionto uplift (Teixell et al. 2003), which is corroboratedby geophysical data and modelling (Seber et al.1996; Ayarza et al. 2005; Teixell et al. 2005;Zeyen et al. 2005; Missenard et al. 2006). Whilecrustal thickness is moderate (only locally reach-ing values ca. 40 km, Wigger et al. 1992, Ayarzaet al. 2005), models show a prominent lithospheric

Fig. 3. Tectonic

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sketch map of the Pyrenees (after Teixell 1996) indicating the lines of section of Figure 5. NPF, NorthPyrenean fault; GT, Gavarnie thrust; PB, Pobla and Senterada basins; SP, Sis palaeovalley.


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thinning that accounts not only for the high topog-raphy at long wavelengths but also for the poorpreservation of foreland basins (note that much ofthe High Atlas lacks peripheral foredeeps or theyare very shallow; Figs 1 & 4a, b), and for the occur-rence of Cenozoic alkaline magmatism contempo-raneous to compression (see a review in Teixellet al. 2005).

The timing of long-wavelength surface upliftrelated to the mantle structure has been indirectlyinferred from the chronology of associated mag-matic events (last 15 Ma for the later phase; Teixellet al. 2005; Missenard et al. 2006), and from strati-graphic palaeoelevation markers (1000 m in thepast 5 Ma at c. 0.2 mm a21; Babault et al. 2008).The reasons for the thinned lithosphere in a regionthat has been subject to compression during muchof the Cenozoic remain unresolved.

The Eastern Cordillera of Colombia

The Eastern Cordillera is the easternmost branch ofthe northern Andes of Colombia. It is a NNE-trending, 110–200 km wide thrust-fold belt that isseparated from the subduction/magmatic com-plexes of the Western and Central Cordilleras byan intervening depression, the Middle MagdalenaValley Basin. The Eastern Cordillera appears nowas a doubly verging thrust system formed duringthe Cenozoic by the inversion of a Mesozoic back-arc rift (Colletta et al. 1990; Cooper et al. 1995;Mora et al. 2006).

Uplift of the Eastern Cordillera was related totransmission of stresses to the South Americanplate by the accretion of arcs in the northwesternAndes, and was associated with subsidence in theadjacent foreland basins of the Middle Magdalena

Fig. 4. Structural

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cross-sections across the High Atlas and the Eastern Cordillera of Colombia. (a, b) Sections across theeastern and central High Atlas (from Teixell et al. 2003). These sections are based on field data and were modified fromthe original according to gravity modelling by Ayarza et al. 2005 (see location in Fig. 1). Although largely eroded, theCretaceous sediments probably formed a tabular body that covered the entire Atlas domain, representing post-riftconditions. (c) Simplified structural cross-section of the Eastern Cordillera of Colombia, approximately through thelatitude of Bogota (see location in Fig. 2). This section was constructed on the basis of maps, seismic profiles andstructural data provided by ICP-Ecopetrol. The deep structure of the Sabana de Bogota region is conjectural as it ispoorly imaged in the seismic profiles. The lower–upper Cretaceous boundary is taken for the sake of convenience at thetop of the Une and Hilo formations.


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Valley and the Llanos (Fig. 2). The general structureof the Cordillera has long been reported by cross-sections that illustrate diverse conceptions of itsstructural style (Campbell & Burgl 1965; Julivert1970; Colletta et al. 1990; Dengo & Covey 1993;Cooper et al. 1995; Roeder & Chamberlain 1995;Restrepo et al. 2004; Cortes et al. 2006; Mora et al.2008; Teson et al. 2013; Toro et al. 2004). There areconflicting views regarding the role of low-anglethrusting and the magnitude of associated trans-lations, from thin-skinned interpretations (Dengo &Covey 1993; Roeder & Chamberlain 1995) to thick-skinned models dominated by basement-involved,inverted extensional faults (see a review in Tesonet al. 2013). Rifting history of the Eastern Cordilleradates back to the Triassic–Jurassic for a first phase,and to the early Cretaceous for a main phase whena graben system roughly coincident with the pre-sent Cordillera was developed (Etayo et al. 1969;Cooper et al. 1995; Sarmiento-Rojas 2001, andreferences therein). Post-rift deposits accumulatedfrom the late Albian through the late Cretaceous.Triassic–Jurassic rift deposits are areally restrictedand consist mainly of terrestrial red beds, whereaswidespread lower Cretaceous sediments (up to5 km thick) are dominantly marine, including alter-nating sandstone and shale formations in the easternflank of the Cordillera, and mainly deeper-watershales in the west, with turbiditic formations atthe base. The upper CretaceousQ3 of the Sabana deBogota area still includes marine shales and sand-stones; the first terrestrial deposits indicative of anoverfilled basin appear in the late Maastrichtian–Paleocene.

The structure of the Eastern Cordillera is illus-trated by a cross-section through its central seg-ment near the latitude of Bogota (Fig. 4c). Clearstructural and geomorphic variations can be seenin this cross-section: from steep orogen flanks ofhigh relief dominated by outward-verging thrustsystems to the axial plateau of the Sabana deBogota, which is characterized by relatively openand symmetric folds with minor thrust displace-ments (see also Julivert 1970; Mora et al. 2008).

As observed in the case of the High Atlas, thrustdeformation is concentrated along the formerrift margins. The thrust belts at the flanks of theEastern Cordillera are asymmetric with regard todeformation intensity, topography and morphology,and magnitude and age of exhumation (Mora et al.2008). The eastern margin shows clear evidenceof thick-skinned deformation in the Quetamebasement massif (Figs 2 & 4c). This massif wasuplifted by a major, west-dipping thrust (TesaliaFault), formed by inversion of a formerly exten-sional fault. Extensional faults, inverted or not, arecommon around the massif (Mora et al. 2006). Infront of the Quetame massif, the Guaicaramo

thrust sheet carries a reduced Cretaceous succes-sion and the Cenozoic Medina Basin (Parra et al.2009a, 2010), and is separated from the autoch-thonous Llanos foreland basin by another thrustderived from the inversion of a graben margin.The western flank of the Eastern Cordillera showsless structural relief; basement is not exposed butprevious authors, on the basis of stratigraphic vari-ations, reported inverted extensional faults defin-ing major thrust units (Salina, Bituima, Cambaothrusts; Gomez et al. 2003; Restrepo et al. 2004;Cortes et al. 2006). Thin-skinned thrusts do occur,either in the lowermost Cretaceous successions ofthe Villeta anticlinorium (Cortes et al. 2006) orwithin the Cenozoic succession of the proximalMiddle Magdalena Valley foreland (Honda thrustsheet, Fig. 4c).

The interior of the Eastern Cordillera at theSabana de Bogota keeps a rather homogeneousstructural elevation as defined by fold envelopes(Fig. 4c). The Sabana preserves Palaeogene sedi-ments (mainly terrestrial), largely eroded in theorogen flanks. The deep structure of the Sabana deBogota is poorly constrained owing to a constantexposure level and poor imaging in seismic reflec-tion data. The Floresta basement massif located150 km northwards along-strike (Fig. 2) indicatesthat basement faults may exist at depth, althoughfold geometry at the Sabana suggests a detachmentin the lower part of the Cretaceous succession.Depth-to-detachment calculations tentatively indi-cate that such detachment level might be some4 km below the surface of the Sabana. A salt for-mation sourcing the Zipaquira and Nemocon dia-pirs and many other salt occurrences throughoutthe Sabana (McLaughlin 1972) is thought to haveaccumulated at the level of the earliest CretaceousMacanal or Fomeque formations (Lopez et al. 1988).This is approximately the position of our estimateddetachment level in the subsurface, and accordinglywe interpret the Sabana as a salt-detached foldbelt of the style represented in Figure 4c. That salttectonics may have played an important role in theSabana folding is also suggested by the exposedfold geometries. Systematic limb overturning inmany anticlines (e.g. Julivert 1963, 1970) mayrepresent diapir-margin deformation similar to thatdescribed in well-documented salt provinces (e.g.Giles & Lawton 2002; Rowan et al. 2003); foldsmay have originated as salt walls, later squeezedduring continuous shortening, so the salt formationhas been removed from most of them (although itsexistence at depth is suggested by the brine springsreported in Campbell & Burgl 1965 and McLaugh-lin 1972). Synclinal topographic depressions inthe Sabana de Bogota were eventually filled dur-ing the late Neogene and Quaternary with fluvio-lacustrine deposits (Tilata and Sabana formations,


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Julivert 1963, Andriessen et al. 1993; Torres et al.2005), which smoothed the relief of the plateau.

Estimates of total orogenic shortening in theEastern Cordillera have been published in severalpapers (see Toro et al. 2004 and Teson et al. 2013for a review). On one hand, dominantly thin-skinnedinterpretations have come up with magnitudes from150 to 200 km (Dengo & Covey 1993; Roeder &Chamberlain, 1995). On the other, thick-skinnedmodels arrived at shortening magnitudes of c. 70–100 km, which is about 25–30% of the originallength (Colletta et al. 1990; Toro et al. 2004;Teson et al. 2013). We favour the later values,appreciating the structural style displayed in theabove-cited publications and a preliminary estimateof line lengths in Figure 4c.


Many studies indicated a main Andean phase start-ing in the mid Miocene and continuing into recenttimes as the principal episode of mountain buildingin the Eastern Cordillera of Colombia, which wasthe locus of the more or less disrupted forelandbasin of the Central Cordillera before that time(e.g. Dengo & Covey 1993; Cooper et al. 1995;Hoorn et al. 1995; Toro et al. 2004; Bayona et al.2008; Horton et al. 2010). Increasing evidence forPalaeogene deformation indicating that the formerEastern Cordillera Basin was far from stable andwas strongly compartmentalized includes: (1) aregional mid Eocene unconformity in the MiddleMagdalena Valley and western foothills (Gomezet al. 2003; Restrepo et al. 2004; Parra et al.2012); (2) growth stratal geometries in the Palaeo-gene of the western foothills and of the Sabanade Bogota (Julivert 1963; Gomez et al. 2003,2005); and (3) flexural modelling of the Paleo-cene–Eocene formations of the Cordillera interior(Sarmiento-Rojas 2001; Bayona et al. 2008). Thesewere relatively mild deformations in the formerrift basin that did not prevent connection betweenthe Central Cordillera and the Llanos foreland,although they controlled significantly the sedimen-tary thicknesses of the Paleocene–Eocene forma-tions, their palaeogeography and palaeocurrents,which locally flowed NE parallel to the growingfolds (Gomez et al. 2005). Fold limb rotation andoverturning in the Sabana de Bogota is a persistentand long-lived feature recorded by uppermost Cre-taceous to lower Oligocene growth strata (Julivert1963; Gomez et al. 2005). Neogene main emer-gence of the major thrust faults that characterizethe Cordillera flanks attest to strong basin-margininversion that had probably commenced already inthe late Oligocene to early Miocene, as recentlyindicated by subsidence and exhumation analysis(Parra et al. 2009a, b; Mora et al. 2010) and detrital

sediment provenance. Detrital zircon geochronol-ogy studies reveal a recycling of the PalaeogeneSabana sediments into the Carbonera Formation ofthe Medina and Llanos basins during this time(Horton et al. 2010). The Eastern Cordillera devel-oped into an effective topographic barrier thatseparated the Central Cordillera from the LlanosBasin at least before mid to late Miocene times,when detrital zircons of Meso-Cenozoic age (whichindicate a Central Cordillera provenance or recy-cling of cenozoic clastic rocks of the Eastern Cor-dillera) disappear in the stratigraphic succession ofthe eastern foothills (Horton et al. 2010). The con-temporaneous conglomeratic fluxes of the Hondaand Guayabo formations into the Middle MagdalenaValley and Llanos basins may be the expression ofthese uplift events (Hoorn et al. 1995; Gomezet al. 2003).

The late Neogene fluviolacustrine deposits of theSabana de Bogota contain palaeoflora which hasbeen compared with the range of temperature toler-ance for modern taxa presumed to be their nearestliving relatives. Using this method, Van der Ham-men et al. (1973) and Hooghiemstra et al. (2006)inferred low altitudes until the late Miocene, anda rapid surface uplift of c. 1500 + 500 m between6 and 3 Ma ago, a period in which the Sabanade Bogota (c. 5000 km2) experienced sedimen-tary aggradation. However, following Gregory-Wodzicki (2000), the error margins may be signifi-cantly larger (+1500 m), which implies that theproposed palaeoelevation changes may not bereliably resolved by the nearest living relativesmethod. Therefore, we cannot discard that the East-ern Cordillera has experienced continuous thicken-ing and surface uplift since the onset of mountainbuilding.

The record of ongoing shortening to recent timesis provided by the strong deformation of the Hondaand Guayabo formations in the most external thrustsheets of the Cordillera, also in agreement with veryyoung apatite fission track ages in the Quetamemassif (c. 3 Ma and younger) that attest strong exhu-mation in the Cordillera margins in the latestNeogene and Quaternary (Mora et al. 2008).

The Pyrenees

The Pyrenees formed in late Cretaceous to Cenozoictimes as a result of convergence between the Iberianand European continental masses, which closed arift basin located in between. The Pyrenean conti-nental crust was severely attenuated during Meso-zoic rifting, to a point where peridotites of theupper mantle were exhumed to the rift-basin flooraccording to recent interpretations (Jammes et al.2009; Lagabrielle et al. 2010), although oceanic


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crust did not form in the segment now separatingFrance from the Iberian peninsula (Fig. 3). Assuch, the central and eastern Pyrenees can be con-sidered an inverted rift much like the Atlas orEastern Cordillera, but with greater amounts ofextension and subsequent contraction.

Rifting in the Pyrenees commenced in thePermian–Triassic, with the main phase of exten-sion during the early Cretaceous, when an east–west-trending basin system, approximately spatiallycoincident with the present mountain belt, wasdefined (e.g. Puidefabregas & Souquet 1986;Verges & Garcıa-Senz 2001)Q4 . Upper Cretaceoussediments are laterally extensive and could beviewed as post-rift deposits, although extensionalfaulting was locally still under way. The Mesozoicbasins were strongly inverted during the Pyreneanorogeny, in association with rapid subsidence inadjacent foreland basins, the proximal parts ofwhich were deformed and incorporated in thethrust system. Hence, much of the southern Pyre-nees are built by deformed Cenozoic sedimen-tary rocks that provide an unique record of thetiming and kinematics of the thrust-fold belt (e.g.Seguret 1972; Labaume et al. 1985; Puigdefabregaset al. 1986, 1992; Martınez et al. 1988; Muttiet al. 1988; Verges & Munoz 1990; Burbank et al.1992; Teixell 1996; Millan et al. 2000; Teixell &Munoz 2000).

Recognition of the structure of the Pyreneesstarted approximately a century ago, and has bene-fitted from the deep seismic profiling of theECORS programme. Complete cross-sections basedon extensive field geology and inferences fromseismic profiles (along the section trace or pro-jected) include Munoz (1992), Verges et al. (1995)and Teixell (1998). The cross-sections of Figure 5illustrate the structure of the central Pyrenees,showing a north-verging thrust belt in the northernpart (the North Pyrenean Zone) with very thickCretaceous successions corresponding to the axisof the previous rift or transtensional basin. Loca-lized high-temperature metamorphism of mid Cre-taceous age attests to severe crustal thinning andhigh heat flow. The North Pyrenean belt shows athick-skinned style with thrust faults that usuallyinvolve the Hercynian basement and that havebeen formed by the inversion of Cretaceous nor-mal faults (Peybernes & Souquet 1984). Addition-ally, local detachments and diapirism sourced inTriassic salt are common (Canerot et al. 2005;Lagabrielle et al. 2010).

A southern, wider part of the Pyrenees is charac-terized by south-verging thrusting and includes theAxial Zone, a stack of basement-involved thrustsheets that forms the core of the chain, and a deta-ched imbricate fan in front of it (South PyreneanZone; Fig. 5). Hence the Pyrenees differ from the

Fig. 5. Cross-sections

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of the central Pyrenees following the trace of (a) the ECORS-Pyrenees (simplified after Munoz1992) and (b) the ECORS-Arzacq (simplified after Teixell 1998) seismic profiles. The sections are based on field datacombined with commercial and deep (ECORS) seismic profiles.


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central High Atlas and the Eastern Cordillera in theoccurrence of an axial massif with high structuralrelief (and high topographic relief as well). Themassif consists of large antiformal stack in theEastern Pyrenees and in the eastern Central Pyre-nees (Fig. 5a), whereas to the west the basementthrust sheets exhibit a lesser degree of overlap andstructural relief decreases (although a central culmi-nation can still be defined; Figs 3 & 5b). Much of theSouth Pyrenean Zone shows a major detachment inTriassic salt (and local diapirism), although thereare marked variations along-strike. Cretaceous suc-cessions are thick beneath the Tremp Basin (on topof the Montsec thrust in Fig. 5a), as the main thrustsheets there derive from the inverted basin, whereasin the Jaca Basin transect the southern Pyreneescontain only a reduced Upper Cretaceous succes-sion accumulated out of the early Cretaceous deepbasin trough. Large basement thrusts as the Gavar-nie and Guarga are thus shortcut thrusts from thebasin margin existing to the north (Fig. 5b). Allthis is because the original southern margin of therift basin was not linear, but formed salients andre-entrants, a geometry that resulted in the frequentlateral thrust ramps of the present-day Pyrenees(Fig. 3).

Total shortening for the ECORS-Pyrenees sec-tion of Figure 5a has been estimated as c. 100 km(Roure et al. 1989) or 150–165 km (Munoz 1992;Beaumont et al. 2000). For the narrower, west-central Pyrenees (Fig. 5b), shortening was estimatedas c. 80 km (Teixell 1998). In both transects thepercentage shortening across the mountain chainis probably not less than 40%.


Our present knowledge of the timing of the Pyre-nean mountain building is mainly based in theSouth Pyrenean and Axial zones, where there arelarger datasets on tectonics–sedimentation relation-ships and thermochronology. The beginning of Pyr-enean contraction was in the late Cretaceous (lateSantonian), on the basis of growth folding in thenorthernmost Boixols thrust sheet of the southernPyrenees, which derives from the inversion of alower Cretaceous fault system (Garrido-Mejıas &Rıos 1972).Q5 Late Cretaceous to early Cenozoic lim-ited contraction has been postulated for the NorthPyrenean Basin as well. Later, during the Palaeo-gene, the Pyrenees experienced a broad piggy-backsequence of imbrication where the main thrustunits were defined (e.g. Camara & Klimowitz 1985;Labaume et al. 1985; Verges & Munoz 1990;Teixell 1998; Beaumont et al. 2000). Recentaccounts for the end of the orogenic activity in thePyrenees place it in the Burdigalian (c. 18–16 Maago), on the basis of tectonics–sedimentation

relationships in the south Pyrenean thrust front andfission-track thermochronology in the Axial Zone(Millan et al. 2000; Jolivet et al. 2007).

Orogenic growth of the Pyrenees has been syn-thesized by Sinclair et al. (2005) in an earlyhistory of fault inversion and frontal accretion, fol-lowed by an accelerated growth of the Axial Zoneantiformal stack by underplating. The correlationbetween the ages of the main cover imbricatethrusts, well constrained from syntectonic sedi-ments, and the ages of the basement thrusts of theAxial Zone duplex is a major challenge. Recentattempts to unravel the timing of the Axial Zoneare based on fission track thermochronology. Asouthward propagation of exhumation is recordedin the ECORS-Pyrenees transect (Fig. 5a), with:(1) rapid cooling of the northern Axial Zone andNorth Pyrenean massifs initiating at 50 Ma; (2)granitic plutons of the central Axial Zone providingages of rapid exhumation of c. 30–32 Ma; and (3)cooling ages of 20 Ma in the southernmost AxialZone (Fitzgerald et al. 1999; Sinclair et al. 2005;Metcalf et al. 2009). Closer to the transect ofFigure 5b, the major Gavarnie thrust sheet has pro-vided late Eocene–early Oligocene apatite fissiontrack ages (Jolivet et al. 2007) consistent withtectonics–sedimentation relationships observed fur-ther south in the Jaca Basin (Teixell 1996). Themain emergence of thrust systems at the Pyreneanmountain fronts started in this epoch, contempora-neous with major structural relief creation in thesouthern Axial Zone. In the early Miocene, conver-gence and surface uplift ended.

Tectonics and macroscale drainage

development

The High Atlas, the Eastern Cordillera and the Pyr-enees exhibit drainage organization features in par-allel with the structural characteristics described.In the central High Atlas and in the Bogota seg-ment of the Eastern Cordillera, shortening is con-centrated along the former rift margins, and theorogen interiors consist of less deformed plateau-like areas with a homogeneous structural relief(Figs 4 & 6a, b). On the other hand, the Pyrenees,with a greater magnitude of orogenic shortening,have developed an axial culmination with highstructural relief and also strong regional topographicslopes (Figs 5 & 6c). The differences in the amountof shortening between these three mountain beltsmay allow a spatial comparison to be used as aproxy for the analysis of drainage evolution ingrowing inversion orogens. Current climates inthese orogens encompassed semi-arid in the easternHigh Atlas to temperate and humid in the Pyreneesand tropical in the Eastern Cordillera. In these


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mountainous regions, past climates are less knownthan in their adjacent foreland basins, but if pre-sent climates held in the past, these differencesmay have played a role in shaping the topography.In what follows we show that the large-scalepatterns of drainage organization in the threemountain belts are consistent with their respectiveregional slopes, and then are most probably due totectonic shortening, crustal thickening and surfaceuplift.

The fluvial drainage of the High Atlas is char-acterized by abundant reaches that flow longitudin-ally, that is, parallel to the structural grain, in theorogen interior (Fig. 6a). Transverse rivers aremore common in the orogen flanks, near the north-ern and southern mountain fronts. Most longitudinalrivers flow along synclines in relatively soft sedi-ments (Jurassic shales), and often cut across thrustsor antiformal ridges (constituted by lower Lias-sic resistant carbonates) at structural depressionsalong fold axes. A similar pattern is observed inthe Eastern Cordillera, where longitudinal rivers invalleys of relatively low local relief flow along theSabana de Bogota plateau and the areas locatedto the north along-strike (Fig. 6b). Valleys in theSabana de Bogota also occupy synclinal depressionscontaining Cenozoic sediments between anticlinalridges constituted by resistant upper Cretaceoussandstone. The flanks of the Eastern Cordillera arecharacterized by deeply incised transverse rivers(Fig. 6b), with local relief stronger than in theHigh Atlas, consistently with the greater bulk short-ening and higher regional slope of the former. Onthe other hand, the Axial Zone of the Pyrenees ischaracterized by a well-developed transverse drai-nage (Fig. 6c).

Observations in the High Atlas suggest that thelongitudinal drainage predates the developmentof transverse rivers in the orogen flanks (Babaultet al. 2007, 2012). Remnants of low-relief longitu-dinal valleys are occasionally seen perched athigh altitude, and captures of longitudinal streamsby transverse streams are common, especiallytowards the western, more elevated High Atlas. Inthe plateau of the Eastern Cordillera, longitudinalrivers also have a protracted history, dating backto Palaeogene times. For this period, palaeogeo-graphic reconstructions indicate sediment dispersalstrongly controlled by emerging tectonic structuresin the basin (Gomez et al. 2005). Since theEastern Cordillera became a topographic barrierin Miocene times, the amounts of erosion in theplateau have been very low. By contrast, the

Cordillera flanks are characterized by very activerecent erosion, yielding young apatite fission trackages frequently younger than Miocene (Mora et al.2008). Again, stream capture of the gently slopinglongitudinal rivers in the plateau by the steep andincising transverse rivers in the flanks is common.Headward erosion in the transverse rivers and cap-tures are expressed by wind gaps and depressionsin the main drainage divides bordering the Sabanade Bogota (Struth 2011; Struth et al. 2012). Thisgeomorphologic evidence indicates that drainagedivides migrate towards the centre of the plateauand that drainage areas of transverse rivers inthe flanks increase at the same time (Struth et al.2012).

The timing relationships described above arguefor a process of drainage reorganization, fromlongitudinal-dominated to transverse-dominated,occurring alike in these two orogens of moderatetectonic inversion. Absolute chronologies are notavailable for the Atlas case, although low-relief,structure-controlled drainage can be envisaged atleast for the late Eocene times, when earliest indi-cations of growth of structural highs go in parallelwith little exhumation (no fission track coolingrecords) and with the unroofing of only the highestlevels of the preorogenic sedimentary cover(Teson 2009; Teson et al. 2010). Babault et al.(2012) suggest that the increased tectonic thicken-ing and surface uplift in later times incrementedpotential energy on both sides of the deformedwestern High Atlas, and enhanced the fluvialerosion in short transverse rivers (Fig. 7). Recordof the progressive development of the transversenetwork might be first indicated by the midMiocene Ait Kandoula alluvial fans in the Ouarza-zate Basin (El Harfi et al. 1996; Teson 2009), andby contemporaneous fission track cooling ages inthe western High Atlas (Missenard et al. 2008;Balestrieri et al. 2009).

In the Eastern Cordillera, growth strata inPalaeogene sediments along the detachment foldsof the Bogota area, together with palaeocurrents,indicate that longitudinal drainage dominatedat least the present-day interior of the Cordilleraduring the early stages of inversion (Paleocene toearly Oligocene; Gomez et al. 2005). Pliocene andQuaternary fluviolacustrine aggradation in theSabana de Bogota (Andriessen et al. 1993; Torreset al. 2005; Hooghiemstra et al. 2006) argue for alow-relief uplifting plateau, possibly internallydrained or with limited connectivity (as observedtoday; Fig. 6c). As in the High Atlas (Babault

Fig. 6. Oblique views showing shaded topography and drainage network of (a) the eastern and central High Atlas,(b) the central Eastern Cordillera and (c) the central Pyrenees. These images are based on the digital elevation modelSRTM90 draped onto Google Earth. Thick grey lines indicate main tectonic boundaries.


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et al. 2012), it can be envisaged that the inversion ofthe Cordillera during the Neogene increased themean elevation and, in its flanks, the regionalslopes, favouring the development of an activelyincising transverse drainage. In addition, the mainmechanism of transverse river development wasprobably headward erosion, as evidenced by streamcaptures (Struth et al. 2012). As the plateau area isreduced by piracy and the transverse rivers increasetheir catchments, an increase in terrigenous influxinto the foreland basins is expected. This processmay have influenced the coarse accumulations ofthe late middle Miocene Honda and Guayabo for-mations in both sides of the Eastern Cordillera.

Evidence for a previous longitudinal drainagehas not been reported in the area now occupied bythe Axial Zone in the Pyrenees. Certainly, andbefore grading to transverse from mid Eocenetimes, longitudinal sediment routing has extensively

been reported for the south Pyrenean foreland basinand thrust-top basins, especially in their early stages(e.g. Nijman & Nio 1975; Marzo et al. 1988; Barno-las et al. 1991; Puigdefabregas et al. 1992; Whitch-urch et al. 2011 and references therein) Q6. This is notsurprising as longitudinal rivers have been else-where described as characteristic of external partsof growing mountain belts (e.g. Oberlander 1985;Koons 1995; Burbank 1992; Jackson et al. 1996,Humphrey & Konrad 2000; van der Beck et al.2002; Ramsey et al. 2008). However, for compari-son with what is observed in the High Atlas andEastern Cordillera, what is relevant here is to deter-mine whether a longitudinal drainage ever existed inthe growing Pyrenean orogen, that is, the hinterland(the Pyrenean Axial Zone). This possibility hasnever been addressed before.

As mentioned above, compelling evidence for anearly longitudinal drainage in the Axial Zone is not

Fig. 7. Two-stage

Colour

online/

colour

hardcopy

sketch model for the evolution of the drainage network in the Atlas Mountains of Morocco, modifiedafter Babault et al. (2012), which can be of general application to inversion orogens which experience progressiveincrease of tectonic thickening and regional topographic slope.


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currently available, but a few provenance obser-vations in the southern Pyrenees lead us to contendthat it cannot be ruled out. A detrital zircon U–Pbgeochronological study (Whitchurch et al. 2011)shows that the upper Cretaceous and Paleocenesediments in the Tremp Basin (Fig. 6) contain zir-cons with Cadomian crystallization ages. Plutonsformed during the Cadomian orogeny are onlypresent in the Eastern Pyrenees and in the northernflank. Whitchurch et al. (2011) interpret these detri-tal ages as evidence for an Eastern-Pyrenean prov-enance. Since palaeocurrents are to the west inthese strata, the authors implicitly assume that sedi-ment was transported by transverse rivers from thehinterland of the Eastern Pyrenees to the southernforedeep, and from the eastern foredeep to thecentral foredeep, parallel to the strike of theorogen. Alternatively, we cannot discard the possi-bility that part of the longitudinal transport fromthe Eastern to the Central Pyrenees occurred in thehinterland (the growing mountain belt) and waseventually transported by a transverse river in theeastern part of the Tremp Basin. The stacking ofbasement units in the Axial Zone resulted in the can-nibalization, in the Central Pyrenees, of the northern(proximal) margin of the late Cretaceous to Paleo-cene foredeep before 40 Ma (Beamud et al. 2011),and consequently neither of these two sedimentrouting possibilities can be demonstrated with thecurrent knowledge. The same reasoning could beapplied to younger deposits preserved in the wedge-top basins of the Central Pyrenees, the Pobla andSenterada Basins at the southern margin of theAxial Zone, where north-derived, Oligocene-ageproximal fan sediments (c. 30–25 Ma, Beamudet al. 2011) also show a clear Cadomian detritalzircon signal.

At present, the Noguera Pallaresa River and itstributaries drain the Axial Zone (Fig. 6), incisingthe Pobla and Senterada basins. However, thereare no outcrops of Cadomian rocks in their catch-ments, which means that the Oligocene catch-ment of the palaeo-Noguera Pallaresa River had agreater extent or that higher structural levels thatsupplied the Pobla and Senterada basins containedCadomian sources that are now completely eroded.Oligocene conglomerates of the Sis palaeovalley(Beamud et al. 2011), located only c. 20 km west,do not show a Cadomian signal, indicating that thehigher structural levels of the central PyreneanAxial Zone, now eroded, were lacking Cadomianplutons or earlier sedimentary deposits (Palaeozoicor Mesozoic) with that signal. Consequently, wemay suggest that the upstream parts of the palaeo-Noguera Pallaresa River, previously draining theEastern Pyrenees, where the Cadomian plutons arelocated, have been captured after the deposition ofthe sediments stored in the Pobla and Senterada

basins. On the other hand, it is considered unlikelythat the palaeo-Noguera Pallaresa River was drain-ing the North Pyrenean zone, north of the NorthPyrenean Fault, during the Oligocene (Beamudet al. 2011).

In summary, transverse rivers have beendescribed in the internal part of most of the Ceno-zoic orogens (Hovius 1996). Structural evidencecoupled with indications of drainage reorganizationin the western High Atlas (Babault et al. 2012) andthe Eastern Cordillera of Colombia (Struth et al.2012) suggest that these inversion orogens evolvedfrom early stages of orogenic growth dominated bystructure-controlled longitudinal rivers in its innerparts, and as deformation was accumulated and theregional slope increased, the drainage evolved intoa regional slope-controlled, transverse network(Fig. 7). The case of the Pyrenees may be consistentwith this pattern of evolution, as this orogen, in amore advanced stage of tectonic accretion, showsa double-wedge topographic pattern and a drainagenetwork almost completely dominated by transverserivers, as observed in most of the Cenozoic linearmountain belts considered by Hovius (1996).

Conclusions

The High Atlas, the Eastern Cordillera of Colombiaand the Pyrenees share in common their originfrom rift basins, where contractional deformationof continental crust, even if severely attenuated inthe rifting stage, was not driven by oceanic slabpull forces. A prominent deformation style isthick-skinned thrust faulting, either derived fromthe reactivation of early extensional faults, ornewly formed as large-scale shortcuts in their foot-walls. However, while thick-skinned thrusting iseffectively a primary mechanism in the invertedrifts described, weak levels in the syn-rift sedimen-tary infill (salt, overpressured rocks) have enableddecoupling between basement and cover, and theformation of localized detachment folds or imbri-cate thrust systems.

Variations in structural geometry between thesections selected reflect different magnitudes oftotal shortening accumulated, increasing from theHigh Atlas (20–25%) to the Eastern Cordillera(25–30%) and the Pyrenees (c. 40%). The HighAtlas and Eastern Cordillera show shorteninglargely concentrated at the outer flanks, which coin-cide with the former rift basin margins. Greaterthrust translations in the Eastern Cordillera resultin high-relief basement massifs in the easternflank, whereas much of the central and easternHigh Atlas lack basement culminations (basementis at a rather homogeneous elevation), with a fewexceptions (Fig. 1). In both orogens the axial


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regions are high-elevation plateaus with compara-tively less deformation and structural relief. Accord-ingly, much of the exhumation is concentrated in theorogen flanks, especially in the case of the EasternCordillera. In contrast, the Pyrenees can be viewedas more mature inversion orogen where a centralhigh-relief culmination of stacked basement thrustsheets occupies the core. This central core (theso-called Axial Zone), which transfers thrust displa-cement to the mountain fronts via decollementlevels in the sedimentary cover, concentrated thegreatest exhumation in the Pyrenees.

Convergence in the Pyrenees ceased in the earlyMiocene, whereas it is still active in the Eastern Cor-dillera and the Atlas, which experienced majorperiods of mountain building during the Neogene.In spite of this difference, timing relationshipsrevealed by tectonics–sedimentation relationshipsand thermochronology indicate a common patternof evolution. This pattern consists in an early stageof shallow and distributed deformation of the riftbasin fill (folding and localized thrusting, with lowstructural relief), followed by strong inversion androck uplift concentrated at the basin margins inlater stages, with large basement-involved faultscausing higher structural relief and high topo-graphic elevations. This is particularly clear in thecase of the Eastern Cordillera, while it seemslikely in the east-central High Atlas as well. In thewest-central Pyrenees, an episode of major rockuplift by shortcut thrusting at the southern basinmargin (beginning in the late Eocene) can also bedifferentiated. We conclude in a mechanism of riftinversion by which initially rift interiors deformin a distributed way, and evolve to a point wherestrong deformation is concentrated in weakerfaults at the rift margins (inverted or newly formedaccording to the specific mechanical conditions, e.g.Hilley et al. 2005), giving way to the episode whenthe prominent mountain topography is acquired.

Observations from the western High Atlas(Babault et al. 2012) showed that, as this invertedrift evolved from the initial deformation stagesto the full accretion, a parallel evolution of thefluvial drainage network was triggered. The axial,plateau-like regions existing in the High Atlas andEastern Cordillera are dominated by old longitudi-nal valleys, whereas the orogen flanks are com-monly characterized by strongly incised transverserivers. Transverse rivers are dynamically enlargingtheir catchments at the expense of longitudinalrivers, by mechanisms of headward extension andstream capture (Babault et al. 2012; Struth et al.2012). The Pyrenees show a two-sided wedgeprofile with a high-relief Axial Zone and transverserivers dominating in both sides of the main drainagedivide. Applying the longitudinal-to-transversescheme, deduced from the High Atlas and the

Eastern Cordillera, to the hinterland of the Pyrenees(Axial Zone) gives an alternative explanation forunexpected provenance features, which would beindicative of the former existence of longitudinalriver reaches. Hence, spatial variations betweenthe different orogens might be taken as a hint to tem-poral variations, arguing for a common patternwhere in the early, mild stages of inversion a longi-tudinal fluvial network controlled by local structuresis established, and as shortening and mean elevationare accumulated, the system reorganizes into aregional slope-controlled network of prevailingtransverse rivers.

The contents of this paper benefited greatly from sharinginformation and ideas with J. Van Den Driessche,M. L. Arboleya, E. Teson and L. Struth. Incisive commentsby two anonymous referees and by J. C. Ramırez substan-tially helped to improve the manuscript. A.T. acknowl-edges A. Mora and the Instituto Colombiano delPetroleo-Ecopetrol for providing maps and additionalinformation which formed the basis for section construc-tion of Figure 4c. This section was constructed with thesoftware Move, provided by Midland Valley through theASI. Financial support was provided by Spanish projectsCGL2010-15416, Consolider-Ingenio CDS2006-00041(TopoIberia) and CGL2007-66431-CO2-01 (TopoMed).Finally, we would like to thank A. Mora and M. Nemcokfor organizing stimulating discussions at the Baricharameeting and for encouragement to prepare this reviewpaper.

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