Field reconnaissance of the Anti-Atlas coastline, Morocco: Fluvial and marine evidence for Late...

Global and Planetary Change 68 (2009) 297–310

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Global and Planetary Change

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Field reconnaissance of the Anti-Atlas coastline, Morocco: Fluvial and marineevidence for Late Cenozoic uplift

Rob Westaway a,⁎, Ali Aït Hssaine b, Tuncer Demir c, Anthony Beck d,1

a The Open University, Abbots Hill, Gateshead NE8 3DF, UKb Department of Geography, Ibn Zohr University, Agadir, Moroccoc Department of Geography, Harran University, 63300 Şanlıurfa, Turkeyd Department of Geography, Durham University, South Road, Durham DH1 3LE, UK

⁎ Corresponding author. Also at: IRES, Newcastle UnNE1 7RU, UK.

E-mail address: [email protected] (R. Westa1 Present address: School of Computing, University of

0921-8181/$ – see front matter © 2009 Elsevier B.V. Adoi:10.1016/j.gloplacha.2009.03.016

a b s t r a c t
a r t i c l e i n f o
Article history:Received 3 May 2008Accepted 20 February 2009Available online 2 April 2009

Keywords:MoroccoAnti-AtlasPliocenePleistoceneupliftfluvial terracesmarine terraces

The available evidence regarding the disposition and chronology of Pliocene–Pleistocene fluvial terraces,coastal rock flats, raised beaches and lacustrine sediments adjoining the Anti-Atlas coastline of Morocco hasbeen reviewed and supplemented by additional information from our own field reconnaissance. It is thussuggested that the study region has experienced uplift by ∼130 m since the Mid-Pliocene climatic optimum(∼3.1 Ma), by ∼90 m since the latest Pliocene (∼2 Ma), and by ∼45 m since the Mid-Pleistocene Revolution(∼0.9 Ma). Each of these phases of uplift correlates with a phase of global climate change knownindependently, and it is thus inferred that the observed uplift is being driven by climate through mechanismssuch as erosional isostasy and the associated induced lower-crustal flow. Numerical modelling of theobserved uplift history indicates that the mobile lower-crustal layer in the study region is ∼9 km thick, with atemperature at its base of ∼500 °C. The base of this mobile layer is inferred to be at ∼24 km depth, thedeepest crust consisting of a layer of mafic underplating that does not flow under ambient conditions. Theprincipal landform in the study region, the coastal rock platform at ∼60 m a.s.l., thus formed during asuccession of interglacial marine highstands in the late Early Pleistocene when uplift rates were low.Although control on the ages of young sediments and landforms is currently extremely limited, beingdependent on regional correlation schemes rather than on absolute dating, the study region fits the pattern,emerging worldwide, that climate change is driving the systematic growth of topographic relief evidentduring the Late Cenozoic.

© 2009 Elsevier B.V. All rights reserved.

1. Introduction

The Souss Basin in central Morocco separates the High Atlasmountain range to the north from the Anti-Atlas mountains to thesouth (Fig. 1). Previous studies of the Quaternary sediments andlandscape development in the Souss Basin and High Atlas includeWeisrock (1980), Aït Hssaine (1994) and Missenard (2006); theselocalities have also been re-examined by Aït Hssaine and Bridgland(2009-this issue). The Anti-Atlas mountains are truncated at theirwestern end by the Atlantic Ocean coastline (Figs. 1, 2). Previous workon the geomorphology of this isolated coastal region has been verylimited; the local literature consists essentially of just one publication,that by Oliva (1972), the region farther inland around Tiznit havingalso been studied by Ouammou (1994). Oliva (1972) reported raisedbeach deposits along this coastline and fluvial terraces along theregion's rivers (Fig. 2), thus providing evidence of Late Cenozoic uplift.

iversity, Newcastle-upon-Tyne

way).Leeds, Leeds LS2 9JT, UK.

ll rights reserved.

The aims of this study are to report the results of field reconnaissanceof the landscape in the region, which confirms that raised beachdeposits and fluvial terraces are indeed present, to tentatively suggesta preliminary chronology for the observed intraplate vertical crustalmotion, to model the associated vertical crustal motion using aphysics-based technique, and to thus place the evidence from thisregion in its global context, as inferred, for instance, by Bridgland andWestaway (2007).

2. Geological and geomorphological background

The Anti-Atlas mountain range is formed in Precambrian (? EarlyProterozoic) igneous and metamorphic basement (comprising schist,quartzite, granite and gabbro). These basement rocks becamedeformed in the Late Precambrian during the emplacement of theAnzi Complex (comprising rhyolite, ignimbrite, conglomerate, turbi-ditic sandstone and more granite), interpreted as a subduction–accretion complex. The Anzi Complex was in turn overlain by latestPrecambrian and Early Palaeozoic (Cambrian–Ordovician) sediments,subsequently metamorphosed (now consisting of quartzite, dolomite,schist and calcschist). Recent analyses or syntheses of the tectonic

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Fig. 1. Map of northern Morocco showing the location of the study region (around Sidi Ifni and Tiznit) and other localities discussed in the text. De facto international borders areshown approximately using thick dashed lines. Rivers with seasonal flow are shown using thin dashed lines.

298 R. Westaway et al. / Global and Planetary Change 68 (2009) 297–310

history of the basement structures and associated metasediments inthis region include Ikenne et al. (1997), Ennih and Liegeois (2001),Thomas et al. (2002), Soulaimani et al. (2003, 2006); Ahmed et al.(2005), Gasquet et al. (2005), and D'Lemos et al. (2006).

Since the Early Palaeozoic the region has experienced prolongeddenudation; except for localizedpockets of Cretaceous sediment (Fig. 2),there is no onshore outcrop aged between the Early Palaeozoic and LateCenozoic (e.g., Oliva, 1972). The passive continental margin offshore ofthewesternAnti-Atlas developed in the latest Triassic–Early Jurassic as aresult of the opening of the Central Atlantic Ocean, and was associatedwith the deposition of many kilometres of sediment (e.g., Davison,2005). The continental margin of NW Africa was conjugate to that ofeastern North America, the present study region having adjoined NovaScotia (e.g., Davison, 2005). The location of the offshore transition fromcontinental to oceanic crust has been the subject of lively debate overseveral decades; it is now generally accepted as corresponding to the S1magnetic anomaly that is ∼50–100 km offshore (see, e.g., Schminckeet al., 1998, for a discussion of this issue), roughly halfway between thecoastlines of the Africanmainland and of the inner Canary Islands, suchas Lanzarote (Fig. 1). Although basalts that erupted at this time arewidespread in the eastern Anti-Atlas and elsewhere in Morocco(including the Draa basalts and associated mafic intrusions; e.g.,Hailwood and Mitchell, 1971; see also Davison, 2005) they are notfound in the western Anti-Atlas. However (like in its much better-studied conjugate counterpart in eastern North America), the emplace-ment of mafic underplating at the base of the continental crust seems tohave been widespread during the formation of this particular passive

margin, being indicated by positive Bouguer gravity anomalies and high(N∼7 km s−1) seismic velocities (see below). As is discussed below, thecrustal properties inherited fromthis timeare shown toplayamajor rolein controlling the Late Cenozoic uplift of the region.

Numerous local names have been introduced into the literature overmany decades to represent subdivisions of the Pliocene and Quaternaryof Morocco (see, e.g., Stearns, 1978, for a synthesis). In the “classical”literature, dating back to the early and mid twentieth century (e.g.,Lecointre, 1918; Gigout, 1949; Lecointre, 1952; Choubert and Ambroggi,1953; Biberson, 1958; Biberson, 1961; Choubert, 1962; Lecointre, 1965),some of these terms have been taken tomean different spans of time bydifferent authors, creating a potential source of confusion. The mostimportant Pliocene and Quaternary sites in Morocco adjoin the city ofCasablanca in the north, providing biostratigraphic evidence andevidence of early human occupation (e.g., Raynal et al., 2001; Geraads,2006; Rhodes et al., 2006; El Graoui et al., 2007). However, the recentapplication of modern dating techniques (e.g., Raynal et al., 2001;Rhodes et al., 2006; El Graoui et al., 2007) has indicated that additionalproblems exist with the “classical” stratigraphic nomenclature. Forinstance, deposits assigned to one stratigraphic subdivisionmayoverlapin age with deposits assigned to another, and the existing subdivisionsthus cannot be uniquely and unambiguously associated with climatecycles defined by the marine oxygen isotope stage (MIS) timescale.Becauseno satisfactory stratigraphic nomenclature thus currentlyexists,the “classical”Moroccan stratigraphic stages, such as “Moghrebian”, areused in this study but with their names placed in inverted commas toindicate that they may have questionable validity.

Fig. 2.Geomorphological/geologicalmapof thestudy region, simplified frompartof Fig.3 ofOliva (1972). ThePrecambrianornament includesboth the (?Early)Proterozoic basement and theLate Proterozoic “Anzi Complex” rocks. The Cretaceous conglomerate is only shown where above the level of the Pleistocene wavecut platform; this platform is bevelled into this bedrocklithology atmany localities betweenMirleft and FoumAssouka, as in Fig. 5. The roads along this coastal platform and between Gourizim and Souk-el-Arba have been omitted to avoid clutter.

299R. Westaway et al. / Global and Planetary Change 68 (2009) 297–310

As Fig. 2 indicates, the landscapeof the Anti-Atlas is heavily dissectedby fluvial incision. The highest summits in the central Anti-Atlas (theKerdous Massif) are ∼2400 m a.s.l. In the western Anti-Atlas, known asthe IfniMassif, the landscape ismore subdued, thehighest summit being∼1200 m a.s.l. As is discussed in detail by Griffiths (1972), Oliva (1972),and others, most of the Anti-Atlas region has an arid climate. Meanannual rainfall is only ∼140–150 mm at the coast (e.g., at Sidi Ifni andMirleft; Fig. 2), increasing inland to ∼400 mm at altitudes of ∼2000 m.Vegetation is thus limited (typically consisting at low altitudes only ofdrought-tolerant species such as Euphorbid cacti) and outcrop isconsequently typically well-exposed, with good preservation of manydeposits of inferred Pliocene or Pleistocene ages. Oliva (1972) reportedlacustrine deposits of inferred “Hamadian”, or (?) Early– (?) MidPliocene, age that he traced from ∼400 m a.s.l. at thewesternmargin of

the Kerdous Massif, ∼40 km inland, to ∼200 m a.s.l. west of Assersif,north of the Ifni Massif, ∼20 km inland (Figs. 2 and 3a,b). Oliva (1972)thus inferred that, since the Pliocene, the landscape has tiltedwestward,towards the Atlantic Ocean, a view consistent with the location of thehighest topography in the IfniMassif near its easternmargin (Fig. 2). Themost striking features of themodern dissected landscape of thewesternAnti-Atlas are the coastal platform, a fossil wavecut platform at a typicalaltitude of ∼60 m a.s.l., and the extensive fluvial terrace depositsflanking the principal river valleys in the region (see below).

North of the Anti-Atlas range are the much higher High Atlasmountains (the highest summit, Jebel Toubkal, reaches 4165 m). Thisrange is bounded to the south by the South Atlas Fault, a north-dippingactive oblique reverse fault that ruptured at the Earth's surface mostrecently in the magnitude 5.9 Agadir earthquake of 29 February 1960.

Fig. 3. Cross-sections illustrating the disposition of Pliocene lacustrine deposits and Pleistocene fluvial terrace deposits in the area around Tiznit (see Fig. 2 for locations). (a) and (b)NW–SE cross-section illustrating the typical coastward tilting of the (?) Early-Mid-Pliocene lacustrine deposits in the area between Asserssif and the NW margin of the KerdousMassif. Modified after Fig. 5 of Oliva (1972). S3 is a Late Cenozoic erosion surface, identified by Oliva (1972), on which these lacustrine sediments were deposited; S2 is an older,higher-altitude erosion surface. (c) West–east profile through the area east of Tiznit, modified after Fig. 9 of Oliva (1972).


The long-timescale deformation associatedwith this fault zonehas beeninvestigated in detail (e.g., Meghraoui et al., 1998;Weisrock et al., 1999).Meghraoui et al. (1998) showed that progressive reverse slip on it hascaused warping of Pleistocene interglacial marine terraces along theAtlantic Ocean coastline at the western end of the High Atlas betweenAgadir and Cape Rhir (Fig. 1). Meghraoui et al. (1998) reportedfour marine terraces in this area at maximum heights of 28, 36, 70and 210m a.s.l. The youngest and second youngest were dated, respect-ively, to the “Ouljian” (MIS 5e) and “Harounian” (MIS 7) interglacials,indicating (assuming sea-level the same as at present) an uplift rate of∼0.2 mm a−1. The third terrace was thought to be “Anfatian” or mid-Middle Pleistocene; if assigned to MIS 9 it would, likewise, indicate anuplift rate of ∼0.2 mm a−1, whereas if assigned to MIS 11 (cf. El Graouiet al., 2007) it would indicate a rate of∼0.17mma−1. The oldest terracewas thought to be “Moghrebian”, with an age of ∼2 Ma. Its time-averaged uplift rate has thus been ∼0.1 mm a−1, therefore indicatingfaster uplift in this locality in the Middle Pleistocene than in the EarlyPleistocene (i.e., suggesting an increase in uplift rates in the MiddlePleistocene). Pleistocene marine terraces are likewise evident north oftheHigh Atlas range, for instance in the Casablanca area, as noted above.Recent reinvestigations (e.g., Raynal et al., 2001; Rhodes et al., 2006; ElGraoui et al., 2007) have shown that the number of distinct marineterrace deposits in this area is much greater than was thought in the“classical” local stratigraphy. Deposits are thus now recognised atmaximum altitudes a.s.l. including the following, with inferred ages,uplift rates and “classical” nomenclature in brackets: 4 m (MIS 5e;0.03 mm a−1; “Ouljian”), 8 m (MIS 7; 0.03 mm a−1), 13 m (MIS 9;0.04mma−1), 24m(MIS11;0.03mma−1; “Anfatian”), and32m(?MIS25; 0.03mma−1; “Maarifian”). However, each of themarine terraces inthis areamaintainsnear-constant altitudes,withnoevidence ofwarpingdue to active faulting (e.g., Stearns, 1978). They thus record regionaluplift, not local effects of active faulting, the uplift rate being much lessthan in the vicinity of Agadir.

The High Atlas and Anti-Atlas mountain ranges are separated bythe ∼30 kmwide Souss Basin, which contains a stacked succession ofup to ∼1000 m of Cenozoic sediment of which up to ∼250 m isthought to be Pliocene or Pleistocene (e.g., Ambroggi, 1963; Nairnet al., 1980; Aït Hssaine, 1994; Occhietti et al., 1994; Aït Hssaine andBridgland, 2009-this issue). In addition to the active South Atlas Faultat the northern margin of this basin, there is evidence of deformationof the sediments within the basin, due to the presence of other activefaults, but the southern margin of the basin is not fault-bounded,being instead formed by the progressive onlap of the basin fill onto thenorthern margin of the Anti-Atlas (e.g., Ambroggi, 1963). Like that inthe Casablanca area, we thus regard the uplift of the Anti-Atlas asregional uplift, unrelated to active faulting.

There is little published information on rates of surface processes inMorocco. Assuming a sediment density of 2000 kg m−3, the datacompiled by Milliman and Syvitski (1992) equate to spatially averagederosion rates of ∼0.05 and ∼0.07 mm a−1, respectively, throughout thecatchments of the rivers Souss and Moulouya (Fig. 1). It can thus beinferred that erosion rates are several times greater, at least, than thesevalues in the parts of the catchmentswithin the High Atlas, with greatestrelief. Like elsewhere, erosion rates inMoroccowill alsohavevaried in thepast in response to climate change, having probably been greater duringcold climate stages when less vegetation cover is expected. Quaternarylandscape development inmany regions is known to not involve a steadystate; uplift rates typically exceed erosion rates (e.g., Bridgland andWestaway, 2007). The erosion rates estimated above thus arguablyprovide lower bounds to Quaternary uplift rates in the High Atlas region.

3. Field investigation

Our field reconnaissance has investigated the fluvial terrace de-posits, coastal platform and marine deposits in the study region.Topographic imagery, derived from Shuttle Radar Topographic


Mission (SRTM) data, has been produced for key parts of the studyregion to facilitate measurement of altitude and interpretation of thegeomorphology (Fig. 4). Our investigations of fluvial terrace depositsconcentrated on the River Massa, north of the Ifni Massif, and on theRiver Ifni (Fig. 5), within this massif, each of which has produced astaircase of six terraces according to Oliva (1972). The rock coast,bounded by a marine cliff, in lithified basement of the western Anti-Atlas stretches for ∼120 km, from Aglou-Plage (or Sidi-Moussa-d'Aglou) in the northeast to a point ∼60 km SW of Sidi Ifni in thesouthwest (Fig. 2). North of Aglou-Plage, the coastline is formed indune deposits and sediments of the Souss Basin, whereas beyond thelimit of rock coast SW of Sidi Ifni the coastline (known as the PlageBlanche; Fig. 1) is formed in Mesozoic and Cenozoic sediments of theLaayoune Basin.

The catchment of the River Ifni andmostof this rock coast are locatedwithin the former Spanish enclave of Ifni, which occupied an area of1502 km2, extending for∼80kmalong the coast, centredonSidi Ifni, and

Fig. 4. SRTM topographic imagery for key parts of the study region. The data source and propublications also discuss related issues such as the resolution and accuracy of this type of datlocations) being in Zone 29R, quadrangles LN and MN. Altitude contours are shown at interva0 m contour has been omitted to avoid cluttering the figure, as it was repeatedly crossed by thSidi Ifni area. (d) The area between Sidi-bou-Ifedail and Aglou-Plage.

for up to∼25 km inland. This territorywas ceded byMorocco to Spain in1860 (although Spanish control was only nominal before 1934) but wasreturned to Morocco in 1969. The region remains isolated, being onlyaccessible by road from the south, from Goulimine (Guelmime), acrossthe Ifni Massif, and from the northeast, from Tiznit, along the coast.Cuttings along the latter road provide the principal exposures ofPleistocene raised beach deposits in the region (see below).

3.1. Terraces of the River Massa

Oliva (1972) reported that the six terraces of the River Massadiverge upstream, away from each other and from the modern level ofthe river. In the vicinity of Aït-ou-Mribete, where the main roadbetween Agadir and Tiznit crosses the river, these terraces weredepicted as occupying a swathe of land, ∼4 kmwide, perpendicular tothe river (Fig. 2), four terraces being reported locally. On inspection inthe field, no fluvial terraces were evident in the immediate vicinity of

cedure used to create this imagery were explained by Westaway et al. (2006a,b); thesea. Co-ordinates are expressed using the UTM system, all localities depicted (see Fig. 2 forls of 5 m, those at odd multiples of 5 m, in grey, being omitted in areas of high relief. Thee undulating sea surface. (a) The Mirleft area. (b) The area around Plage Ftaissa. (c) The

Fig. 5. Long profile through the terraces of the River Ifni in the area inland of Sidi Ifni, also showing stratigraphic relationships with marine deposits as interpreted by Oliva (1972).Modified after Fig. 7 of Oliva (1972), with additional labelling to link sites to those illustrated in Fig. 4c.


the road bridge, at Universal Transverse Mercator (UTM) co-ordinates[MP 42861 06594]. However, some 500 m farther north and ∼15 mabove the modern river level, at [MP 42725 07124], more than a metrethickness of polymict fluvial gravel, coarse at the base but finingupward, was observed in a roadcut section. This deposit presumablyforms part of one of Oliva's (1972)Massa terraces, but we do not knowwhich one because he provided no height information for this locality.Another ∼1 km farther north and possibly 5 m higher, at [MP 4217808130], a ∼1 m thickness of similar fine gravel and pebbly sand wasobserved in another roadcut section, and may form part of another ofOliva's (1972) terraces; the deposit was encrusted with calcite andpartly buried beneath young dune sand, thus fitting his description ofthe older Massa terraces in the area.

About 10 km east (upstream) of these sites, the River Massa flowsthrough a narrowgorge inOrdovician quartzite at thewesternmargin ofthe KerdousMassif (Fig. 2). Farther upstream, Oliva (1972) reported theMassa terrace staircase as most clearly developed, the three highestterraces (terraces 4, 5 and 6) being locally ∼70 m, ∼90 m and ∼120 mabove present river level, formed in fluvial gravels up to ∼15 m thick.However, since the time of Oliva's (1972) study the Massa valley hasbeen dammed at the quartzite gorge (the Youssef-ben-Tachfine Dam;Fig. 2), flooding the reach farther upstreamwith the clear fluvial terracedevelopment. No fieldwork to verify these terrace deposits was thuspossible. Some25kmfarther south (upstreamof thenowfloodedpart ofthe Massa valley), Oliva (1972) reported that the flat landscape aroundTiznit is formed in the surface of lacustrine deposits of inferred“Moghrebian” age (Fig. 2), which he interpreted as meaning “LateVillafranchian”, located at ∼300 m a.s.l., two fluvial terraces (at ∼270 mand∼230m a.s.l.) being evident between this level and that of the RiverTazeroualt, a left-bank tributary of the Massa, at ∼200 m a.s.l. (Fig. 3c).Ouammou (1994) attempted radiocarbon dating of these lacustrinesediments, obtaining an indeterminate age of N30 ka; the sedimentsthus remain undated.We observed the flat landscape in the former lakebed but could not identify any of the fluvial terrace deposits reported inthis area.

3.2. Terraces of the River Ifni

Oliva (1972) depicted the terraces of the River Ifni as a long profile,extending from a point ∼9 km inland to the coastline at Sidi Ifni(Fig. 5). As described by Oliva (1972), Ifni terrace 1 forms a flat influvial silt, sand and gravel, no more than a few metres above presentriver level. Terrace 2 is formed in silty deposits with lenses of gravel,its flat surface up to ∼5 m above present river level; the deposits ofterraces 1 and 2 are uncemented. Terrace 3 forms a flat ∼8–10m abovethe present river level; its deposits, including clasts of rhyolite andgranite, are partially cemented due to calcite encrustation. Terrace 4,typically ∼30 m above present river level, is formed in a well-cemented deposit consisting of up to ∼8 m thickness of fluvial gravel,

with clasts principally of rhyolite but including some of granite.Terrace 5, of limited extent, and the more extensive terrace 6, werereported at typical altitudes of ∼60 m and ∼100 m above the modernriver. Both are formed in fluvial gravels typically ∼5–10 m thick,comprising clasts of rhyolite and highly altered clasts of Precambrianvolcanic rocks; in addition, terrace 6 was reportedly capped by acalcareous encrustation typically ∼10 cm thick. As depicted in Fig. 5,the terraces diverge upstream away from each other and away fromthe modern river. Thus, the height of terrace 6 above the modern riverincreases from ∼85 m, ∼2 km inland, to ∼145 m, ∼8 km inland,whereas that of terrace 4 increases from ∼20 m, ∼1 km inland, to∼40 m, ∼7 km inland. These fluvial terraces are best developed in thereach of the River Ifni between ∼3 and ∼8 km inland, where its valleyis relatively wide; terraces 4, 5 and 6 form particularly clear landformsin this area, as illustrated in Photo 6 of Oliva (1972).

Several of the fluvial terrace flats identified by Oliva (1972) occurwithin the area of Fig. 4c and are indeed readily identifiable within thelandscape. For instance, terrace 4 occurs at site 5 (∼60 m a.s.l.; ∼30 mabove present river level) near the confluence of the two principalaffluents of the River Ifni. Terrace 5 can be observed at site 6 (∼95 ma.s.l.; ∼60 m above present river level), about half a kilometre fartherupstream, with terrace 6 evident at site C (Lalla Myriam; ∼90m a.s.l.;∼80 m above present river level), ∼1 km inland of Sidi Ifni. About2 kmfarther upstream, aflat∼55mabove thepresent river level (∼80ma.s.l.; site 7 in Fig. 4c) may be part of terrace 5, whereas higher-levelflats nearby,∼85–90maboive the present river (∼115–120ma.s.l.; sites8 and 9 in Fig. 4c) may be part of terrace 6.

As Fig. 5 illustrates, Oliva (1972) reported that the fluvial deposits ofIfni terrace 6 overlie the beach deposits that cover the inner part of thecoastal platform (at Lalla Myriam; site C in Fig. 4c), indicating therelative age of the two sets of deposits. We were unable to verify thisstratigraphic relationship; our own observations of the morphology ofthe coastal platform are discussed below. Oliva (1972) tentativelyinferred that thesemarine deposits date from the “Moghrebian” stage ofthe local stratigraphy, which he regarded as “Late Villafranchian” (i.e.,latest Pliocene), such that the overlyingfluvial depositswere interpretedas either also “Late Villafranchian” or Early Pleistocene. He alsotentatively inferred that terrace 4 is “Amirian”, meaning late EarlyPleistocene or early Middle Pleistocene.

3.3. The coastal platform

The coastal platform along the Atlantic coast of the Anti-Atlas istypically ∼300–500 m wide, at a typical altitude of ∼60 m a.s.l. It isbounded to seaward by themodernmarine cliff, which is typically∼30–60 m high, and to landward by a bluff, up to ∼150 m high, which Oliva(1972) interpreted as a fossil marine cliff— an interpretation that seemsentirely reasonable from the landscape morphology (Figs. 4–6). Oliva(1972) mapped a ∼10 km section of the coastal platform SWof Sidi Ifni,

Fig. 6. Landscape photographs. (a) View NE across the coastal platform from [MN 03617 80019], between Gourizim and Sidi-bou-Ifedail. (b) View NW down the valley of the RiverTiouinite from [MN 02959 78200] to the coast at Gourizim. Note the skyline formed by the distal part of the coastal platform and the much lower and narrower rock flat which weestimate is ∼5 m above present-day sea-level. (c) View SW from [MN 0084173671] (site D in Fig. 4a), between Gourizim and Mirleft, showing the coastal platform and a higher flat(site 1 in Fig. 4a) at ∼120m a.s.l. (d) View SW from [LN 93214 59995] near Sidi Bourja, looking towards Sidi Ifni, which is visible on the skyline formed by the coastal platform, ∼10 kmaway. The viewpoint is ∼110 m a.s.l., on a pediment of slope deposits, which slopes gently downward to the level of the coastal platform.



around Sidi Ouarsik (which we have not visited), as an erosionalwavecut platform, butmapped the rest of it as cappedby beachdeposits,except in the north where it is capped by dunes (Fig. 2; see also, e.g.,Weisrock et al., 2002).We found, on the contrary, that exposure of beachdeposits is very limited (localities with beach deposits are discussedbelow),most of the platform surface being covered by angular talus thathas evidently been emplaced by slope processes. Oliva (1972)interpreted the characteristic seaward slope of the platform surface asindicating the seaward slope of a single beach deposit. Our impression,on the contrary, is that this slope simply reflects the disposition of thepediment of slope deposits that covers the true platform surface, whichis typically visible only in roadcuts and gullies.

Some 10 kmNE of Mirleft and ∼5 kmNE of Gourizim, circa [MN 065825], Oliva (1972) reported an outlier of hard fossiliferous Cretaceousmarine limestone (identified by Bourcart, 1939). It reportedly occupiesan area of ∼1 km2, its top and base being at ∼260 m and ∼200 m a.s.l.Oliva (1972) thus inferred the presence of an erosion surface at ∼200 ma.s.l., marking a low-relief land surface of Early Cretaceous age, onwhichthesemarine sedimentswere deposited. He also inferred that outliers ofterrestrial sediments, at sites farther inland, are also Cretaceous, andthus that the Cretaceous coastline lay not far inland of, butmaybe 200mabove, the present-day coastline. The SRTM imagery (Fig. 4) revealsmany flats in the landscape landward of the coastal platform, consistentwith this interpretation.

The present study mainly concerns the central ∼35 km of rockcoast, between Gourizim and Sidi Ifni (Fig. 2), access to which isprovided by the road from Tiznit. Around Gourizim (Fig. 6a) thewidthof the coastal platform varies from ∼300 m to ∼1 km; it is continuousexcept where incised by the valley of the River Touinite (Fig. 6b).About 4 km SW of Gourizim, where the platform is ∼300 m wide, ahigher rock flat, ∼150mwide and ∼120m a.s.l., is also evident (Fig. 6c,locality 1 in Fig. 4a). About 1.2 km farther NE, a similar flat (2 inFig. 4a) is also evident at ∼130 m a.s.l. Above this level, manyconcordant hilltop flats rise landward from ∼180 to ∼200 m a.s.l. overa distance of several kilometres (Fig. 4a).

About 2 km farther SW, the coastal platform is again interrupted bythe valley of the River Salogmad, which reaches the sea at Plage ImiNtarga just north of Mirleft (Fig. 4a). As Fig. 4a indicates, this riverreaches the coastal platform through a narrow gorge, up to ∼150 mdeep. Some ∼3–4 km from the coast, where the river is between∼65 m and ∼80 m a.s.l., its modern narrow gorge flanked by flats.These include sites 3 and 4 (∼30 m above the present river,respectively ∼95 m and ∼110 m a.s.l.; Fig. 4a) and site 5 (∼55 mabove the present river; ∼120 m a.s.l.; Fig. 4a). However, we have notvisited these localities, and so cannot say whether they are deposi-tional fluvial terraces or strath terraces. South of Mirleft, the coastalplatform widens to N1 km, and remains this wide for ∼9 km fartherSSW to Plage Ftaissa (Fig. 4b). Landward of Plage Ftaissa, hilltop flatscan be grouped into two ranges of height, around 110–130 m a.s.l. and∼185–220 m a.s.l. (Fig. 4b). The former grouping includes (Fig. 4b)sites 1–4 at ∼125–130 m a.s.l. and sites 5–8 circa 115 m a.s.l.

Similar landforms persist SW to Sidi Ifni (Figs. 6d, 4c). At Sidi Ifni,the coastal platform is ∼1.0–1.3 km wide (Fig. 4c) and seems to beresolvable into at least two facets; a higher one at ∼70m a.s.l., clearestNE of the town, and a lower one, at ∼45 m a.s.l., clearest SW of thetown (being ∼2000 m long and up to ∼300 m wide; most of it isoccupied by the disused Ifni airport, sites 1–2 in Fig. 4c) and alsoevident at the same level north of the town, at site 3 in Fig. 4c. As isillustrated in Fig. 5 and as was discussed above, Oliva (1972) reportedthat at Lalla Myriam (C in Fig. 4c), the inner edge of the coastalplatform is capped by beach deposits that are overlain by deposits ofterrace 6 of the River Ifni, these sediments being ∼90m a.s.l. However,we were unable to gain access to this locality and so cannot verify thisstratigraphic relationship.

In the north, between Sidi-bou-Ifedail and Aglou-Plage, the coastalplatform is relatively narrow, typically ∼200–400 mwide (Fig. 4d); it

also seems to be typically somewhat lower than farther south, being∼50 m rather than ∼60 m or more a.s.l. As Oliva (1972) noted, unlikefarther south the coastal platform in this area is typically covered bydune deposits, some of which are calcareously cemented or capped bycalcareous deposits (thus suggesting that the dunes are of consider-able antiquity). Erosion of the dune deposits means that the platformis typically bounded to seaward by a slope, rather than by the verticalcliffs observed farther south. However, it is bounded to landward by aprominent cliff beyond which the land surface rises above 200 m a.s.l.(Fig. 4d). We identified one locality (site 1 in Fig. 4d) with evidence ofa ∼200 mwide flat at ∼130 m a.s.l. and another (site 2 in Fig. 4d) witha ∼300 mwide flat at ∼195 m a.s.l. However, as is indicated in Fig. 4d,most of the land surface inland of the coastal platform is higher,forming a dissected landscape at ∼250 m a.s.l. or thereabouts.

This coastal platform was recognised by Pommerie et al. (1951)and tentatively interpreted as of “Calabrian” (“Messaoudian” in theregional stratigraphy; i.e., Early Pleistocene) age. Farther southwest inthe Laayoune Basin, two higher-level land surfaces are recognised inthe landscape (e.g., Alaoui Mdaghri et al., 1994), evidently markingepisodes of marine planation of the sedimentary succession withinthis basin. The higher of these, known locally as the Hamadian surface,was reported by Alaoui Mdaghri et al. (1994) at ∼200 m a.s.l. in thevicinity of Tan-Tan and was traced NE by them at a near-constantaltitude to the NE limit of the Laayoune Basin. This surface, typicallyseveral tens of km wide and again typically ∼200 m a.s.l., was alsoreported in the Tarfaya area by Brebion and Ortlieb (1976). The lowersurface, known locally as the Moghrebian surface, was reported byAlaoui Mdaghri et al. (1994) at ∼100 m in the vicinity of Tan-Tan andwas traced by them SW, decreasing in altitude; they noted it at ∼20–40 m a.s.l. between Sidi Lemsid and Tarfaya (Fig. 1). Brebion andOrtlieb (1976) also noted this ∼40 m a.s.l. coastal flat, of inferred“Moghrebian” age, in the Tarfaya area and contrasted it with afossiliferous deposit of “Messaoudian” age, also∼40m a.s.l., farther NEat Sidi-bou-Maleh near Tan-Tan, where the “Moghrebian” surface isreportedly ∼100 m a.s.l. As already noted, the “Hamadian” stage of theregional stratigraphy is generally thought to represent the Early toMid Pliocene (∼4–3 Ma), whereas the “Moghrebian” is generallythought to represent the Late (or latest) Pliocene, thus broadlycontemporaneous with the “Villafranchian” of Europe.

3.4. Pleistocene marine sediments

As is indicated in Fig. 6c and d, most of the coastal platform iscovered by dune sand and by angular slope deposits. The thickness ofthese deposits, overlying the bedrock exposed along the moderncliffline, can reach many metres. The road from Tiznit to Sidi Ifni is notheavily engineered and typically follows the surface of these aeolianand colluvial deposits. However, a few cuttings reach deep enough toexpose sections through beach deposits beneath the aeolian andcolluvial cover.

One such section is inlandof Plage ImiNtarganearMirleft (Figs. 7a, 8a;site A in Fig. 4a). About 0.5 m of exposed weathered basement iserosionally truncated andoverlainby∼0.5mofbeachgravel, then∼0.5mof beach sand, then ∼3 m of dune sand, then colluvial gravel. The beachgravel consists of polymict rounded clasts, representative of local bedrocklithologies, up to ∼50 cm in size but fining upward, between whichoccasional gastropod fragments are evident. The overlying highlycemented beach sand contains pebbles of up to ∼5 cm size along withabundant shell fragments of bivalvedmolluscs. Thedune sand, alsohighlycemented, also contains many small shell fragments. From the site co-ordinates and the SRTM topographic imagery (Fig. 4a), we estimate theheight of this beach deposit as ∼45 m a.s.l.

Another such section is farther southwest at Plage Ftaissa (Figs. 7b, 8b;site B in Fig. 4b). About2mof reddish slopedeposits composedof angularclasts are overlain by almost 3 m of yellow beach deposits, then ∼2 m ofcemented calcareous silt with fine gravel interbeds, then ∼1 m more of

Fig. 7. Photographs of raised beach sections. (a) at Plage Imi Ntarga, near Mirleft (site A in Fig. 4a), showingweathered basement overlain by beach gravel, then beach sand, then dunesand (height of section illustrated ∼2m). (b) at Plage Ftaissa (site B in Fig. 4b), showing angular alluvial fan deposits overlain by beach gravel then beach sand thenmore beach gravel(height of section illustrated ∼3 m).


colluvium. The beach deposits consist of ∼1 m of beach gravel, which isvery coarse at its base but fines upward, then∼60 cmof beach sandwithbivalve fragments, then an additional ∼1 m of beach gravel, which againfines upward, then ∼15–20 cm more beach sand. From the site co-ordinates and the SRTM topographic imagery (Fig. 4b), we estimate theheight of this beach deposit as ∼55 m a.s.l.

Neither of the above-mentioned sections in raised beach depositswere noted by Oliva (1972); this is not surprising, as the road fromTiznit to Sidi Ifni did not exist at the time of his work. Several other

Fig. 8. Summary stratigraphic logs for

sections through apparently similar beach deposits were brieflydescribed by Oliva (1972). However, no detailed location informationwas provided, and we have been unable to locate any of these sites. AtSidi Ifni he described beach sand and gravel at ∼40m a.s.l., overlain bya calcareous crust, palaeosol and slope deposits. At Sidi Bourja (Fig. 2),∼10 kmNNE of Sidi Ifni (circa [LN 930 600]) he described beach graveland sand with mollusc fossils at ∼30 m a.s.l., overlain by ∼25 m ofslope deposits. Oliva (1972) also noted that somewhere north ofMirleft a wavecut platform at ∼20 m a.s.l., cut into Precambrian

the localities illustrated in Fig. 7.


rhyolite, is overlain by ∼3 m of coarse shelly beach sand, then ∼7 m ofdune sand. He also mentioned that numerous small pockets of beachdeposits, comprising sand, shell fragments, and gravel, were observedalong the coastline between ∼5 m and ∼15 m a.s.l., and inferred thatthese represent recent marine highstands, namely the Ouljian(Eemian; MIS 5e) and Mellahian (Flandrian; MIS 1) of Morocco. Wenoted various sites with rock flats cut into basement at ∼5 m a.s.l.(such as that in Fig. 6b), which may well be wavecut platforms fromthe last interglacial.

South of the Ifni Airport flat, circa [LN 84897 48326] (4 in Fig. 4c),we noted the exposure in a gulley, consisting of what appeared to beanother beach section overlain by ∼20 m of colluvium. We were notable to reach this site to document it in detail. However, from its co-ordinates and the SRTM topographic imagery (Fig. 4c), we estimatethe height of the viewpoint as ∼45 m a.s.l., and thus estimate that thebeach deposit is ∼25 m a.s.l.

Deposits of calcareous silt similar to those capping the beachsection at Plage Ftaissa were also observed elsewhere, notably in thearea north of Sidi-bou-Ifedail, circa [MN 10908 87920]. Here anestimated ∼30 cm of calcareous “crust” capped a ∼1.5 m exposure ofpolymict gravel with rounded clasts (interpreted as another beachgravel) at an estimated height of ∼40 m a.s.l. A possible modernanalogue of the low-energy environment where such sedimentsmight have accumulated in the past is indicated at Gourizim, wherethe River Touinite flows into a pool on the landward side of themodern beach ridge (Fig. 6b).

Oliva (1972) regarded all the beach deposits between ∼40 and∼90 m a.s.l. as contemporaneous, forming a single raised beach of“Moghrebian” age, as indicated in Fig. 5. We regard this as improbable,because such a great range in altitude of a single beach deposit seemsunlikely (its distal part would have been ∼50 m below sea-level whenits proximal part was at sea-level), because the individual beachdeposits that we have examined (discussed above) are quite localized,occur at different altitudes, and have different stratigraphy. We thusinfer that a range of beach deposits of different ages are present. Wehave not attempted to test this hypothesis by examination of thebiostratigraphy; however, the issue is reminiscent of the situation atCasablanca, where a limited number of “classical” raised beaches haverecently been subdivided into a greater number of distinct deposits(e.g., Raynal et al., 2001). Westaway and Bridgland (2007) discussed asimilar example, resolving a sequence of raised beaches in an area ofsouthern Italy where others had previously interpreted a single raisedbeach that has been offset by normal faulting.

4. Uplift history and its relationship to crustal properties

We now attempt to model the post-Middle Pliocene uplift historyof the study region using a technique, developed byWestaway (2001),which calculates the isostatic response to surface processes takinginto account the flow in the mobile lower-crustal layer that is inducedby these processes. Previous experience indicates that to constrain themodelling of vertical crustal motions in this way requires knowledgeof crustal properties, specifically the thickness of themobile layer. Thiscan be estimated using a combination of data on surface heat flow,Moho depth, and gravity anomalies. The surface heat flow indicatesthe depth of the top of the mobile layer, which is taken as the 350 °Cisotherm (after Sibson, 1983). Gravity data can indicate whether alayer of mafic underplating is present at the base of the crust. If so, thebase of the mobile layer is taken as the top of the underplating; if not,it is taken as the Moho.

4.1. Crustal properties

The principal publication on gravity data from the study region isby Hildenbrand et al. (1988); Bakkali and Mourabit (2006) providedan updated analysis. According to these sources, the Bouguer gravity

anomaly is ∼+40 mgal in the Ifni Massif and is similar in theCasablanca area. It decreases into the central Anti-Atlas, where it is∼−80 mgal, and is even lower, ∼−150 mgal, beneath the highestparts of the Atlas Mountains. These variations are consistent withisostatic compensation of the high topography by the development oflower-crustal “roots”. However, the Bouguer gravity anomaly remainslow, ∼−80 mgal, in many localities SW and SE of the central Anti-Atlas, localities where the topography is no greater than, and in mostcases less than, that in the Ifni Massif. The Bouguer gravity “high”evident in the Ifni Massif indeed follows the Atlantic coastline,typically persisting no more than ∼40–80 km inland. Such variationsare not consistent with Airy isostatic compensation of the topography;they require a higher mean density of the Earth in the coastal regionwith the positive gravity anomaly. Using the standard slab formula,the observed ∼120mgal difference in Bouguer gravity anomaly can beexplained by a ∼6 km thickness of mafic underplating, with a densitycontrast of 500 kg m−3, at the base of the crust in this area.

Several syntheses of geothermal data from Morocco have beenpublished (e.g., Rimi, 1990, 1999, 2000; Zarhloule, 2004). Mostgeothermal data have been obtained from boreholes in sedimentarybasins; data are thus sparse or absent in basement highs such as theIfni Massif, where the thermal state of the crust can be derived only byinterpolation of data between surrounding areas. The analysis by Rimi(2000) thus indicates that the surface heat flow in the Ifni Massif is∼75 mW m−2, corresponding to a near-surface geothermal gradientof ∼30 °C km−1 for a thermal conductivity, k, of 2.5 W m−1 °C−1.Zarhloule (2004) deduced a more complex variation, with the near-surface geothermal gradient increasing SW along the coast from∼24 °C km−1 at the northern end of the Ifni Massif to ∼32 °C km−1

some 200 km SW of Sidi Ifni, the value being ∼27 °C km−1 in the SidiIfni area, corresponding to a surface heat flow of ∼68 mW m−2.

We are not aware of any direct measurement of crustal thicknessby seismic profiling in the Ifni Massif, although Rimi (1999) suggestedthat 32 km is a typical representative value for the Anti-Atlas. Aseismic refraction profile ∼200 km south of Sidi Ifni (discussed byRimi, 1999) indicates a crustal thickness of ∼30 km in that area. Theanalysis of this profile by Rimi (1999) indicates a P-wave velocity ofN7 km s−1 in the deepest ∼5 km of crust, suggesting a ∼5 kmthickness of mafic underplating, roughly consistent with the earliercalculations using gravity data. Rimi (1999) also inferred that theradioactive heat production in the area traversed by this seismicprofile decreases downward from ∼1.5 µW m−3 at ∼10 km depth tozero at the Moho, although this deduction was based on use of anempirical equation linking seismic velocity and radioactive heatproduction, rather than on direct measurement.

Regarding the overall thermal state of the crust, Rimi (1999) adopteda surface temperature To of 20 °C and estimated that the typicaltemperature at the 32 km deep Moho beneath the Anti-Atlas is only368 °C. These calculations assumed a surface heatflowof 75mWm−2 ofwhich qo=40 mW m−2 was derived from the mantle. It is not clearwhy, with this relatively high heat flow, such a low Moho temperaturewas deduced, and it thus appears that the calculations are incorrect (forinstance, they include a contributionqr from radioactiveheatproductionin the crust of 30mWm−2 to the surface heatflow qs, when it should be35 mWm−2 to maintain consistency with the other data).

Using simple theory, which assumes radioactive heat production at auniform rate Y in the uppermost D kilometres of the crust (see, e.g.,Seyrek et al., 2008, for the associated equations), one can makealternative estimates of the temperature at depth. Thus, if qo=40 mWm−2 and qr=35 mW m−2 (the values adopted for the Anti-Atlas byRimi, 1999), consistent with Y=2.5 μW m−3 and D=14 km, and ifTo=20 °C and k=2.5 Wm−1 °C−1, one may estimate the temperatureat 14 km depth as 342 °C. The depth of the 350 °C geotherm is thus14.5 km, that at the top of the mafic underplating (taken as a depth of24 km) is 502 °C and the Moho temperature (at 30 km depth) is 598 °C.Conversely, if qs=68 mWm−2, as deduced above from the analysis by


Zarhloule (2004), then with qo=40 mW m−2 and qr=28 mW m−2,consistent with Y=2.0 μW m−3 and D=14 km, one may estimate thetemperature at 14 kmdepth as 322 °C. The depth of the 350 °C geothermis thus 15.7 km, the temperature at the top of the mafic underplating is482 °C and theMoho temperature is 578 °C. The thickness of themobilelower-crustal layer is thus estimated as 9.5 km for the first set ofparameter values and 8.3 km for the second.

It is noted inpassing thatmaficmagmatismhas been reported in theAnti-Atlas region in the Late Precambrian (e.g., Hafid et al.,1999; Barakatet al., 2002) as well as during the rifting of the central Atlantic Ocean.Some of the mafic underplating at the base of the crust may thus relateto this Precambrianmagmatism. However, if so, the above arguments inrelation to present-day crustal conditions are unaffected.

4.2. Uplift modelling

We now model the uplift history of the study region, using thetechnique of Westaway (2001; Westaway et al., 2002) that has beenapplied to many other fluvial sequences worldwide. This technique isbased on the assumption that the observed surface uplift is the isostaticconsequence of net inflow of lower-crust to beneath the study region,induced by surface processes. Detailed explanations have beenprovidedelsewhere (e.g., Westaway, 2001, 2002a,b; Westaway et al., 2002;Westaway, 2004) and so are not repeated here. Specifically, thistechnique assumes that the observed uplift is the net isostatic responseto repeated cycles of loading and unloading of the crust, for instance byice-sheets or fluctuations in eustatic sea-level. In general, increases inrates of erosion caused by global climate change can also induce netinward lower-crustal flow (e.g., Westaway, 2002c). Both mechanismswill in general occur together and their individual contributions can bedifficult to resolve (Westaway, 2002a). However, as has been notedbefore (e.g., Westaway et al., 2004, 2006a), by analogy with localities inthe Mediterranean region that have been studied previously it is likelythat the observed uplift is primarily the isostatic response to increasedrates of erosion. Nonetheless, tests (e.g., Westaway, 2002a; Westawayet al., 2004, 2006a) have established that very similar uplift responsescan result from both mechanisms whereby surface processes arecoupled to induced lower-crustal flow (i.e., from cyclic surface loadingor from increased rates of erosion). The chosen technique is simpler toimplement, requires fewer model parameters to be specified, and canhandle multiple phases of lower-crustal flow forcing (LCFF), and is thusused here as an approximation.

Recent studies (e.g., Bridgland and Westaway, 2007; Demir et al.,2007, 2008; Westaway et al., 2008) have established that in manyregions phases of LCFF occur at ∼3.1, ∼2.0 and ∼0.9 Ma. These relate towell-documented global events when climate deterioration can beexpected to have induced erosion, at the end of the Mid-Plioceneclimatic optimum, in the latest Pliocene (the Tiglian C of NW Europe),and at the end of the Early Pleistocene (the “Mid-PleistoceneRevolution” or MPR). Wewill investigate via uplift modelling whetherthe same pattern is also characteristic of the Anti-Atlas coastline.

The parameters to be specified in the modelling are the magnitudeΔTe and start time to of each phase of LCFF, the geothermal gradient uand thermal diffusivity κ of the lower-crust, and Wi, a measure of thethickness of this mobile layer. As in previous studies, it is assumed thatthe global sea-level was the same as at present during all previousinterglacials from which raised beaches have been preserved; thus,the height of each raised beach provides a direct proxy for the upliftsince the raised beach formed. Use of river terraces (such as those ofthe River Ifni in the present study region; Fig. 5) as proxies for uplift isin principle more problematic. As previously discussed (e.g., byWestaway et al., 2002), river terraces can serve as proxies for upliftprovided they remain subparallel with uniform gradients andprovided no significant changes in the downstream length of theriver channel have occurred. The first of these conditions is necessaryto demonstrate that equivalent quasi-equilibrium downstream grad-

ient profiles develop during successive climate cycles when thefluvial terrace deposits accumulate, thus demonstrating that similarhydrological and sediment-transport regimes repeat during thesedifferent climate cycles. If there were no uplift, fluvial deposits wouldthus repeatedly aggrade, during different climate cycles, at the samelevels in a river valley; their differences in altitude thus indicate theuplift on the intervening timescale. The second condition is necessary,because downstream channel lengthening or shortening will respect-ively cause fluvial incision to underestimate or overestimate uplift,given the requirement of the river tomaintain a downstream gradient.The course of the River Ifni has lengthened by no more than ∼1.5 kmsince the coastal platform began to develop (Fig. 5). However, sincethe downstream gradient of this river is quite high (∼8 m km−1 nearthe modern coastline; Fig. 5), a significant correction (up to 12 m) isneeded to convert the ∼78 m of incision since the aggradation of theoldest fluvial terrace deposits (of terrace 6) to uplift.

Data used for uplift modelling include, first, the marine deposits at∼5,∼15,∼30,∼40,∼45,∼55 and∼85ma.s.l. at sites (discussed above)within or banked against the coastal platform. Second are the fluvialterraces of the River Ifni. Taking account of the gentle downstreamconvergence of the terraces of this river (Fig. 5), terraces 1–4 areconsidered indicative of 2, 5, 10 and 20m of uplift. Terrace 5 is difficultto project downstream and the uplift since its deposits aggradedcannot be estimated precisely; it is thus inferred to be somewherebetween 40 and 55 m. As noted above, due to uncertainty in thedownstream channel-lengthening correction, terrace 6 may indicatebetween 78 and 90mof uplift. The final piece of data used is the heightof the 130m a.s.l. flat, which (as noted above) is inferred to be an olderwavecut platform, above the level of the main coastal platform.

The 130 mwavecut flat is tentatively inferred to date from the MidPliocene (N3 Ma). The 90 m marine deposits are assumed, after Oliva(1972) to be “Moghrebian”, with a nominal age of ∼2Ma. Ifni terrace 6is thus inferred to be slightly younger than this, being assigned anominal age of MIS 68 or ∼1870 ka, thus marking a timewhen (due tocontemporaneous climate instability) significant terrace deposits areobserved in many other rivers (e.g., Bridgland and Westaway, 2007).The three youngest fluvial terraces are inferred to date from MIS 2, 6and 8, and the ∼5 m raised beach deposits from MIS 5e.

As noted above, we set out to fit the data using three phases ofLCFF, starting at 3.1, 2.0 and 0.9 Ma. The fit illustrated in Fig. 9 can beseen to be consistent with the assumed age constraints, noted above.It suggests that Ifni terraces 4 and 5 date, respectively, fromMIS 12 and22, and that the ∼15m,∼30m,∼40m, and∼45m beach deposits datefrom MIS 9, 15, 19 or 21, and 25, respectively. The ∼55 m raised beachat Plage Ftaissa can thus be inferred to date from one of the earlyMiddle Pleistocene interglacials, possibly MIS 41 circa 1300 ka. The∼25 m raised beach tentatively identified south of Sidi Ifni (site 1 inFig. 4c) can thus be inferred to date from MIS 13.

The solution has assumed a value for Wi of 8.0 km. As discussed byWestaway (1998) andWestaway et al. (2002),Wi is approximately 9/10of the thickness of the mobile lower-crustal layer; Wi 8.0 km is thusconsistent with a mobile layer ∼8.9 km thick. Earlier calculations takingaccount of the surface heat flow and presence of mafic underplating atthe base of the crust indicated that themobile layer in the study region isbetween 8.3 and 9.5 km thick. Themodelling is thus consistentwith theindependent evidence pertaining to the physical and thermal state ofthe crust in this region.

5. Discussion

Like many other localities, discussed by Bridgland and Westaway(2007), the study region is consistent with the view that uplift is beingdriven by climate through the mechanism of erosional isostasy andthe associated induced lower-crustal flow. The data from the studyregion is indeed consistent with timings of globally-recognised phasesof LCFF that correlate with times of global climate change (Fig. 9). The

Fig. 9. A modelling solution for the uplift history of the Anti-Atlas coastline. Altitudes offluvial and marine terraces have been measured using SRTM data or taken from theliterature, as described in the main text. Marine data are labelled using circles, fluvial datausing squares; for both, the symbol marking the preferred age is shaded and symbolsrepresenting other possible ages are unshaded. Model prediction uses the technique ofWestaway et al. (2002), with the following parameter values:Wi=8 km, u=16 °C km−1,κ=1.2 mm2 s−1, to1=18 Ma, ΔTe1=−1.6 °C, to2=3.1 Ma, ΔTe2=−1.7 °C, to3=2.0 Maand ΔTe3=−1.2 °C, to4=0.9 Ma and ΔTe4=−1.2 °C. This solution predicts 42 m of upliftsince 875 ka (MIS22), 44msince 950ka (MIS 25), 88msince 2Ma and130msince 3.1Ma.Themaximumpredicteduplift ratewas 0.059mma−1 around 2825 ka during thephase ofLCFF starting at 3.1 Ma, 0.060mm a−1 around 1725 ka during the phase starting at 2.0 Ma,and 0.063 mm a−1 around 625 ka during phase starting at 0.9 Ma. (a) Uplift history.(b) Enlargement of the most recent part of (a). (c) History of variation of uplift rate.


uplift modelling thus provides a natural explanation, in terms ofclimate change, for the principal Quaternary landforms in the studyregion, the fluvial terraces and coastal platform. The coastal platformis inferred to have begun to develop in the late Early Pleistocenewhenuplift rates were relatively low, and became “abandoned” by the sea inthe early Middle Pleistocene following the increase in uplift rate at theMPR. The raised beaches below the level of this platform representMiddle Pleistocene interglacials; the four lowest terraces of the RiverIfni mark Middle and Late Pleistocene cold stages. Terrace 5 of the Ifniis interpreted as marking MIS 22, at the time of the MPR, the firstPleistocene cold stage during which the global cooling and extent ofglaciation were comparable to what subsequently occurred in theprincipal Middle Pleistocene cold stages. Terrace 6 of the Ifni is

interpreted as marking one of the cold stages during the Olduvaisubchron (the Tiglian C stage of northern Europe), when similarconditions developed during the era, before the MPR, when the globalclimate was dominated by ∼40 ka Milankovitch periodicity. Terracesof this age are known in many other rivers worldwide (e.g., Bridglandand Westaway, 2007), which in many river systems is the age of theoldest terrace deposits that are preserved. Rose et al. (2002) haveindeed suggested that in Britain this effect has arisen because theclimate deterioration in the Tiglian C and associated reduction invegetation led to the first significant influxes of gravel into riversystems; the same may be true in the present study region. The∼130 m rock flat is interpreted as marking relatively stable conditionsduring the Mid-Pliocene climatic optimum, before the Late Pliocenedeterioration in climate triggered faster uplift, possibly as a result ofincreased rates of erosion accompanying reduced vegetation cover.The suggested pattern is indeed very similar to other instances,worldwide, recognised by Bridgland and Westaway (2007). However,it must be emphasized that currently no dating evidence exists for anyof the landforms in the study region; the only age control is providedby the various tentative age assignments suggested by Oliva (1972)and by the fit of data to the uplift-modelling solution.

The significance of the present work is thus mainly to highlight thepotential of the study region for more detailed analysis in future, duemainly to the excellent preservation of fluvial and marine terracedeposits, the former due to the arid climate and the latter due to burialbeneath slope deposits. First, heights of marine and fluvial terracescould be surveyed, for instance using differential GPS (cf. Demir et al.,2008; Westaway et al., 2008), to supersede the present set of heightdata based on estimation by eye and on SRTM data. Second,independent age control is required. In principle this could comefrom examination of the biostratigraphy of the raised beach deposits(cf. Stearns, 1978) or possibly from amino acid dating. The warmclimate would preclude use of conventional amino acid dating basedon racemization of isoleucine; however, the technology has recentlybeen extended to cover analysis of more slowly racemizing aminoacids such as valine (e.g., Penkman et al., 2007); in principle it couldthus be applied to the older marine deposits in the region.

The ∼45 m of uplift estimated at Sidi Ifni since MIS 25 exceeds the32 m estimated on the same timescale at Casablanca (e.g., El Graouiet al., 2007). The relatively slow uplift at Casablanca, like at Sidi Ifni, isconsistent with a relatively thin lower-crustal layer due to the presenceof mafic underplating in the area (as indicated by the Bouguer gravityanomaly; see above). The significant variations in uplift rates that resultfrom the different phases of LCFF (Fig. 9) mean that one cannot assumethat the age of any marine deposit in this area is proportional to itsheight, as others (e.g., Stearns, 1978) have done in order to estimate theages of Quaternarymarine deposits in Morocco. Any age estimate madeon this basis should thus be treated with caution.

If the evidence, discussed earlier (Fig. 3), presented by Oliva (1972)in relation to the disposition of Pliocene lacustrine deposits, has beencorrectly interpreted, then the uplift has been much greater in inlandparts of theAnti-Atlas region thanalong the coast. The apparent absenceof mafic underplating at the base of the crust in these inland regions(given the absence of the positive gravity anomaly that would indicatesuch underplating, as discussed earlier) would facilitate more rapidlower-crustalflow in response to a givenmagnitude of surfaceprocessesand, thus, faster uplift. It is difficult to envisage that the Pliocenelacustrine deposits in the study region were laid down at a significantheight above sea-level, there being no land to the northwest wherebythe lake basins couldhavebeen enclosed. It thus seemsprobable that thelake basins were close to sea-level, such that their present-day altitudesprovide rough indications of amounts of subsequent uplift. Assumingalso that the ages of the lacustrine deposits have been correctlyestimated by Oliva (1972), one may thus deduce that there has been∼400 m of uplift since the Mid Pliocene at the western margin of theKerdous Massif, decreasing to ∼200 m at Asserssif and ∼130 m at the


coast, and that there has been ∼300 m of uplift since the latest Pliocenein the Tiznit area, decreasing to ∼90 m at the coast.

To sustain the estimated amounts of uplift requires significant inflowof lower-crustal material beneath the region. The most likely source ofthis material is from the continental crust beneath the offshorecontinental shelf. One may thus envisage that loading of this offshorecrust bywater and sediment has “squeezed” some of the original lower-crustal layer out to beneath the onshore region, thus causing crustalthinning in the offshore area and crustal thickening in the Anti-Atlas, asis indicated by the observed uplift. The erosion occurring in the Anti-Atlas, in part due to the increasing relief caused by the uplift, results inthe flux of sediment into the offshore depocentre, and thus contributesto sustaining the process,which evidently resembles the typeof coupledsystem envisaged elsewhere by Westaway (2002c, 2007). In principle,such systems can be modelled fully quantitatively, for instance byassuming that they represent closed systems for crustal mass (see, e.g.,Westaway et al., 2006a; Seyrek et al., 2008). However, suchmodelling isnot yet possible for the Anti-Atlas, because the uplift data are so limitedand because no data exists either on amounts of Quaternary subsidencein the offshore continental shelf, which could indicate the crustalthinning that has occurred there and thus contribute to estimationof thecrustal mass flux. As already noted, the modelling that is presented inFig. 9 is scaled to values of the forcing parameter ΔTe, no attempt beingmade to relate the magnitude of this parameter to rates of surfaceprocesses in the study region.

By analogy with the Anti-Atlas, it is possible that much of thetopography of the High Atlas has likewise developed as a consequence ofclimate-induced erosion and the associated induced lower-crustal flow.De Sitter (1952) indeed estimated that ∼2000 m of the modern relief inthis region, as well as ∼1000 m of the relief in the central Anti-Atlas, hasdeveloped since the Early-Mid Pliocene. The High Atlas mountains aregenerally considered to liewithin theboundary zonebetween theAfricanand Eurasian plates, which locally accommodates a component of crustalshortening that can in principle cause crustal thickening and uplift.However, it is evident that the relatively slow rate of crustal shorteningcaused by the plate motions (no more than a few millimetres per yearoverall, across a zone hundreds of kilometres wide) cannot possiblyaccount for the scale of development of topography in the region. Thesame argument has also been put forward by Bridgland and Westaway(2007) in relation to thedevelopmentof topography in the∼3500mhighSierra Nevadamountain range in southern Spain (Fig.1). Missenard et al.(2006) have recently proposed, as an alternative, that the Late Cenozoicdevelopment of topography in the High Atlas and Anti-Atlas mountainsas been caused by a mantle plume. Mantle plumes have indeed pre-viously been suggested as the cause of uplift in other continental interiorregions, such as Britain and northwest Europe. However, this idea hasbeen superseded following the realization that climatic forcingprovides abetter explanation of the evidence, since it can account for variations inuplift rates that correlatewithknown timesof global climate change (e.g.,Westaway, 2006a; Bridgland andWestaway, 2007;Westaway, 2009-thisissue). Rather than being a consequence of plate motions, as is usuallythought, the active faulting evident in and around the High Atlas (e.g., atAgadir; see above) may itself be a consequence of the effects on theregional stressfieldof the induced lower-crustalflowthat is inferred tobeoccurring beneath the region, in accordance with a general explanationset out byWestaway (2006b). For instance, the componentof reverse slipon the South Atlas Fault may well occur in order to accommodate theperturbation to the stress field caused by the component of northwardlower-crustal flow, induced by the erosion in the High Atlas anddeposition in the adjoining Souss Basin, and by the resulting redistrib-ution of crustal mass and associated development of topography.

6. Conclusions

The available evidence regarding the disposition and chronology ofPliocene–Pleistocene fluvial terraces, coastal rock flats, raised beaches

and lacustrine sediments adjoining the Anti-Atlas coastline of Moroccohas been reviewed and supplemented by additional information fromour own field reconnaissance. It is thus suggested that the study regionhas experienced uplift by ∼130 m since the Mid-Pliocene climaticoptimum (∼3.1Ma), by∼90m since the latest Pliocene (∼2Ma), and by∼45msince theMPR(∼0.9Ma). Each of these phases of uplift correlateswith a phase of global climate change known independently, and it isthus inferred that the observed uplift is being driven by climate throughmechanisms such as erosional isostasy and the associated inducedlower-crustal flow. Numerical modelling of the observed uplift historyindicates that themobile lower-crustal layer in the study region is∼9kmthick, with a temperature at its base of ∼500 °C. The base of this mobilelayer is inferred to be at ∼24 km depth, the deepest crust consisting of alayerofmaficunderplating thatdoesnotflowunderambient conditions.The principal landform in the study region, the coastal rock platform at∼60 m a.s.l., thus formed during a succession of interglacial marinehighstands in the late Early Pleistocene when uplift rates were low.Although control on the ages of young sediments and landforms iscurrently extremely limited, being dependent on regional correlationschemes rather thanonabsolute dating, the study regionfits thepattern,emerging worldwide, that climate change is driving the systematicgrowth of topographic relief evident during the Late Cenozoic.

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