Review Article
The South-Western Branch of the Variscan Belt: Evidence from Morocco
A. Michard
a,⁎ , A. Soulaimani
b, C. Hoepffner
c, H. Ouanaimi
d, L. Baidder
e, E.C. Rjimati
f, O. Saddiqi
ea10, rue des Jeûneurs, 75002 Paris, France
bDépartement de Géologie, Faculté des Sciences, Université Cadi Ayyad, B.P. S 20, Marrakech, Maroc
cDépartement de Géologie, Faculté des Sciences, Université Mohamed V, B.P. 1014, Rabat Agdal, Maroc
dDépartement de Géologie, ENS, Université Cadi Ayyad, BP S2400, Marrakech, Maroc
eDépartement de Géologie, Faculté des Sciences Aïn Chock, Université Hassan II, B.P. 5366 Maarif, Casablanca, Maroc
fDirection du Développement Minier, Ministère de l'Energie et des Mines, B.P. 6208, Rabat Instituts, Haut Agdal, Rabat, Maroc
a b s t r a c t a r t i c l e i n f o
Article history:
Received 19 January 2010
Received in revised form 11 May 2010 Accepted 28 May 2010
Available online 4 June 2010 Keywords:
Variscan Belt Mauritanides Morocco NW Africa Tectonics Paleozoic
This work is based on the compilation and re-evaluation of the most significant data, either personal or from the literature, concerning the Moroccan Variscides. The latter constitute the only, moderately disturbed or even undisturbed part of the South-Western Branch of the Variscan Belt, facing directly NW Gondwana. They include two orogenic segments, namely the northern Mauritanides and the Meseta Domain exposed in the Saharan and Atlas–Meseta regions respectively, and a foreland belt cropping out essentially in the Anti-Atlas.
The eastward thrust units of Saharan Morocco (Oulad Dlim) mostly originate from the West African Craton (WAC) border in an area of thin Palaeozoic sedimentation. Thin-skinned fold–thrust foreland arcs develop progressively northward (Zemmour) at the expense of the increasingly thick Palaeozoic series, whereas thick-skinned deformation characterizes the inverted proximal paleomargin in the Anti-Atlas Domain. As suggested by the Meseta and Anti-Atlas stratigraphic similarities, the Meseta Domain corresponds to a collage of moderately displaced, thinned crustal blocks from the distal Gondwana paleomargin. Variscan deformation is dominated by NW-verging thrusts, and metamorphism developed in the thickened tectonic prism in relation with crustal anatexis at depth. The Meseta–Anti-Atlas boundary is a major, ENE-trending transpressional dextral fault referred to as the South Meseta Fault (SMF). Discussing the correlations between the Variscan segments of Morocco and SW Iberia allows us to suggest that a latitudinal transform zone similar to the SMF separated these segments during the Late Palaeozoic. Subduction of the Rheic Ocean crust would have been directed SE-ward along both the Iberian and Moroccan Meseta, and NW-ward south of the SMF, i.e. along the WAC.
© 2010 Elsevier B.V. All rights reserved.
Contents
1. Introduction . . . 2
2. Anti-Atlas Domain: the thick-skinned foreland belt. . . 3
3. The Moroccan Mauritanides and their thin-skinned frontal units . . . 3
3.1. Oulad Dlim (Adrar Souttouf) transect . . . 4
3.2. The Mauritanide frontal units (Dhlou–Zemmour and westernmost Anti-Atlas) . . . 6
4. The Meseta Domain: a polyphase collage . . . 8
4.1. Sehoul Block, Rabat–Tiflet Fault Zone and Mazagan Escarpment . . . 8
4.2. Eastern Meseta . . . 10
4.3. Western Meseta . . . 10
4.3.1. Structural zones . . . 11
4.3.2. Structural framework of the Rehamna metamorphic culmination . . . 11
4.3.3. Felsic magmatism . . . 14
4.4. South Meseta Fault and Sub-Meseta Zone . . . 14
5. Interpretation and discussion . . . 14
5.1. Crustal-scale cross-section . . . 14
5.2. Pre-orogenic restoration . . . 15
⁎Corresponding author.
E-mail address:andremichard@orange.fr(A. Michard).
0040-1951/$–see front matter © 2010 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2010.05.021
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Tectonophysics
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c to
insert). The formation of the belt from the birth of the so-called Rheic Ocean to the Devonian–Carboniferous orogenic processes has been repeatedly addressed during the last decade (e.g.Matte, 2001;
Stampfli and Borel, 2002; Robardet, 2003; Simancas et al., 2005; Von Raumer and Stampfli, 2008) and was concerned ultimately by three thematic issues of international journals (Bozkurt et al., 2008; Pereira et al., 2008; Schulmann et al., 2009). Remarkably, the structure of the Moroccan segment of the belt is generally overlooked in these publications, although the wealth of data on the Morocco Hercynides or Variscides, accumulated since their early descriptions byLecointre (1926), Termier (1936) and Roch (1950), has been summarized in several works and papers until the last years (Michard, 1976; Piqué and Michard, 1989; Piqué et al., 1991; Hoepffner et al., 2005, 2006).
The recent papers by an active, Spanish-Moroccan team (Simancas et al., 2005, 2009; El Hadi et al., 2006a) are the only and valuable exceptions to this rule.
references therein), which are not considered here due to their controversial paleogeographic origin. South of the Maghrebide Belt, i.e. in the Atlas and Pre-Saharan/Saharan realms, three major Variscan domains are classically distinguished:
– theMeseta Domainwith the Western and Eastern Mesetas, the intervening Middle Atlas and most of the High Atlas to the south, except part of the Marrakech High Atlas massif (“Ouzellarh promontory”, Choubert, 1952). The Palaeozoic massifs of this domain show intense folding and thrusting, greenschist- to amphibolite-facies metamorphism and widespread syn- to late- orogenic magmatism. Tectonic polarity is dominantly to the WNW.
– theMauritanide Beltis represented by the Saharan Oulad Dlim (Adrar Souttouf) Massif, which is characterized by its nappe structure with craton-ward vergence. Variscan granites are lacking, but Variscan metamorphism is widespread.
Fig. 1.Major domains of the Variscan Belt in Morocco and adjoining countries of NW Africa in their present-day geological setting (afterMichard et al., 2008). Palaeozoic massifs (not shown here) also crop out in the Atlas Mountains and mostly belong to the Meseta Domain (seeFig. 8A).Insert: Location of the Moroccan segment in the Circum-Atlantic Variscan Belt (yellow). Orange: Caledonian Belt (fromMatte, 2001). I: Iberia; Ma: Mauritanides; Mo: Moroccan Meseta.
– theAnti-Atlas Domainforms the common, low- to very low-grade foreland belt of both others domains, being transitional to the cratonic Tindouf–Zag and Bechar basins. The Anti-Atlas belt connects with the Ougarta belt of Algeria (Haddoum et al., 2001), both forming an arcuate chain around the Saharan platform. Tectonic vergence is to the ESE, S and SW according to the western, central or eastern position in the arc, respectively.
Incorporating the Moroccan Variscides in any tentative restoration of the belt is all the more mandatory as they represent the only segment of the belt that maintains direct relationships with Gondwana. Further to the NE, the Ibero-Armorican Arc surrounds a Gondwana“promontory” (West Asturian and Cantabrian Zones) disconnected from Africa because of the Permian wrench tectonics (Matte, 2001, with ref.
therein), followed by Mesozoic–Cenozoic rifting and drifting processes.
Further to the east, the latter events resulted in disruption of the South Branch of the Variscides within the Alpine realm (Von Raumer et al., 2008; Guillot and Menot, 2009; Rossi et al., 2009).
Simancas et al. (2005, 2009)adopt an“autochthonist”view of the Variscides of northern Morocco, e.g. they assume that the Palaeozoic units of the so-called Meseta Domain were not displaced with respect to the Anti-Atlas Domain, and then with respect to Gondwana, except a moderate dextral slip along the so-called“South Atlas Fault”during the Late Carboniferous. We feel that this way of thinking is controversial.
On the other hand,Simancas et al. (2005, 2009)do not consider the SW-ward continuation of the Variscan Belt along strike, i.e. the Mauritanide Belt, notwithstanding the occurrence of large outcrops of this belt in the Saharan Province of Morocco (Lécorché et al., 1991;
Villeneuve and Cornée, 1991; Villeneuve et al., 2006). Accordingly, the aim of this paper is,first, to present an overview of the structure and evolution of the Moroccan Variscides (including their Mauritanide part) and, second, to infer a new scheme for their relationships with SW Iberia and discuss their geodynamic development in the frame of the South-Western Branch of the Variscan Belt.
2. Anti-Atlas Domain: the thick-skinned foreland belt
The structure of the Anti-Atlas Palaeozoic belt is relatively simple.
It is characterized everywhere by thick-skinned tectonics (Burkhard et al., 2006; Raddi et al., 2007), and the Precambrian basement is involved in large faulted antiforms (erosional“boutonnières”) in the axis of the domain (Fig. 2). Deformation decreases eastward and southward, i.e. getting farther from the collision boundaries of the Mauritanide and Meseta domains, respectively (Sections 3 and 4). In the western and central areas, the Palaeozoic series are detached from the basement on the latest Neoproterozoic and Lower Cambrian ductile layers, and the décollement levels that occur higher in the stratigraphic sequence (Lower Ordovician and Silurian) allow dis- harmonic folding to develop (Fig. 3). In the eastern part of the belt, the Palaeozoic series remained essentially undetached from the base- ment, and folding is directly controlled by the basement fault pattern (Fig. 4). In both the western and eastern areas, the reverse/
transpressive fault pattern seems to be basically inherited from the palaeofault system developed at the NW passive margin of Gondwana after the Pan-African orogeny, i.e. from the Late Ediacaran–Early Cambrian to the Late Devonian (Soulaimani et al., 2003; Raddi et al., 2007; Baidder et al., 2008; see also Section 5hereafter). Thus, the Anti-Atlas can be regarded as the inverted passive margin of NW Gondwana. It is worthwhile noting that this Palaeozoic paleomargin developed onto the metacratonic rim of the 2 Ga-old West African Craton affected by the Pan-African orogeny 650 Ma ago (Bousquet et al., 2008; Caby et al., 2008; Gasquet et al., 2008).
3. The Moroccan Mauritanides and their thin-skinned frontal units
In the early 60s, the occurrence of a Palaeozoic (mainly Hercynian) fold belt bordering West Africa from Senegal to northern Morocco was
Fig. 2.Structural map of the Anti-Atlas Domain, with location ofFigs. 3 and 4. A. Mel.: Agadir Melloul.
covered by the Atlantic coastal basins, and characterized by east- verging thrust sheets emplaced over the WAC border (Sougy, 1969).
Subsequent works outlined the polyphase character of the belt and documented Neoproterozoic (Pan-African) and Late Palaeozoic (Hercynian) metamorphic events (Lécorché et al., 1991; Caby and Kienast, 2009, with ref. therein). Most studies have been concentrated in the central and southern part of the belt. To the north (Fig. 5), the Oulad Dlim (Adrar Souttouf) Massif is still poorly understood in the absence of modern metamorphic-structural studies and robust isotopic datings. Several drill cores reached a crystalline basement (Mauritanide nappes?) beneath the Cretaceous–Tertiary sediments of the Boujdour (Tarfaya–Laayoune) Basin (J. Sougy, in Michard and Sougy, 1977).
3.1. Oulad Dlim (Adrar Souttouf) transect
The Oulad Dlim Massif consists of a nappe stack (Fig. 6A, B) overlying to the east the Reguibat Arch either directly (Tichla area,
granites, intruded by undeformed magmatic rocks, i.e. the Aousserd nepheline-bearing syenite and a younger, dense array of basic dykes.
U–Pb zircon ID-TIMS and Lu–Hf datings of felsic volcanics and granites in the Mauritanian Tasiast and Tijirit areas yielded Mesoarchean U–Pb zircon ages ranging between 2.96 and 2.73 Ga (Chardon, 1997; Key et al., 2008). The Aousserd syenite compares with that of Bou Naga (Mauritania) dated at ca. 680 Ma (U–Pb zircon;Blanc et al., 1992).
The autochthonous sedimentary cover begins with Upper Ordo- vician tilloids and arenites, ca. 30 m-thick, without any of the older sedimentary levels as a result of the Hirnantian glacial erosion (Destombes et al., 1969; Lécorché et al., 1991). The arenites are overlain by less than 200 m of Silurian shales and Devonian shallow- water limestones with crinoidal debris and poorly preserved brachiopods and trilobites (Fig. 6C). Further north along strike, i.e.
in the Zemmour Arc, the Devonian sequence has been chronostrati- graphically defined up to the Frasnian (Sougy, 1969). This Palaeozoic succession, which was first explained as a graben isolated in the basement (Alia, 1960; Arribas, 1968), rather represents the eastern margin of a shallow sedimentary basin (Sougy, 1962, 1969). The Devonian limestones are recrystallized and show tight, upright or E- verging folds detached from the autochthonous Ordovician deposits on the Silurian décollement level.
At the bottom of the nappe stack, the two slivers of the Quartzite Nappe consist of metaquartzites and metapelites (Amzili Tiznig Fm.;Rjimati and Zemmouri, 2002) affected by SE-verging folds and NW-dipping cleavage, and recrystallized under greens- chist-facies conditions. Quartz-mylonites with shallow-dipping, E- trending stretching lineation occur between and beneath the slivers. The protoliths are referable to the Early Neoproterozoic by comparison with the Quartzite Series of the Anti-Atlas (e.g.
Gasquet et al., 2008).
The overlying Laglat Nappe (lowest Sebkha Matallah Unit in the sense ofVilleneuve et al., 2006) consists of quartz-rich, intensely folded metapelites and meta-arenites, recrystallized under greens- chist-facies conditions. These rock-types compare with some parts of the Archean greenstone belts to the east, but they show a distinct internal structure with polyphase folds crosscut by W-dipping top-to- the-east shear bands (Plate 1, photo A). The overlying terranes include migmatites, layered leucocratic orthogneisses (Derraman), calcsili- cate marbles associated with ferruginous quartzites, garnet-bearing amphibolites and large metagabbro massifs (Entajat, Matallah, Dayet Lawda). In the latter massifs,Arribas (1968)described a wide range of rock-types (gabbro–diorite, norite and charnockite). Again, these lithologies compare with those of the adjoining Reguibat Shield, but they display a WNW-dipping foliation, which contrasts with the ESE- dipping foliation of the autochthonous domain. Cross-cutting basic dykes are observed in the allochthonous basement nappes, but they are deformed and much less frequent than in the Aousserd–Tichla area. Besides of the basement nappes which form most of the massif west of the Quartzite Nappe, the occurrence of another Neoproter- ozoic quartzite unit has been observed south-west of Imliliy (Rjimati et al., 2002b).
Fig. 4.Inverted mosaic of tilted blocks and related folding in eastern Anti-Atlas, after Raddi et al. (2007), modified. A: Top basement isobaths superimposed to the fault map of the Ougnat Massif area, with indication of the thickness of the Cambrian formations.
B: Cross-section of the Jebel Angal anticline (location in A) showing the cataclastic deformation of the basement in the fault zone. ki 1-2: Lower Cambrian; kmb: Lower–
Middle Cambrian (“Grès terminaux”); km1:“Schistes àParadoxides”); km2:“Grès du Tabanit”.
The Oulad Dlim tectonic structure shown here (Fig. 6A) must be regarded as provisional, in the absence of detail mapping of the core of the massif. Deformation is very heterogeneous, with large zones of almost non-foliated rocks contrasting with strongly recrystallized zones showing 20–60° west-dipping foliations. Such west-dipping foliations also characterize the westernmost part of the massif (Rjimati et al., 2002b). Accordingly, we confirm the interpretation bySougy (1969) and Villeneuve et al. (2006)suggesting a broadly monoclinal, NW-dipping nappe stack rooted beneath the Atlantic coastal basin (Fig. 6B) rather than the hypothesis of a synform of extremely thin nappes (Carte géol. Maroc 1:1,000,000, 1985;Lécorché et al., 1991). The thickness of the internal nappe stack is unknown, in the absence of drilling and/or geophysical profile, but could be N20 km (Fig. 6B). The frontal part of the belt would be much less thick
(a few thousands metres), should we assume a flat-and-ramp structure as illustrated in the Southern Mauritanides (Burg et al., 1993) or the Central Mauritanides (Caby and Kienast, 2009).
The oldest isotopic dates measured in the Oulad Dlim nappes are in the range 1760–1000 Ma (K–Ar WR on Matallah gabbro and gabbro– diorite, respectively;Villeneuve et al., 2006), but they are regarded as rejuvenated ages due to Pan-African and/or Variscan overprint. The Variscan event is recorded by another K–Ar WR age at 274 Ma from a strongly deformed gabbro from Dayet Lawda (Villeneuve et al., 2006).
The Variscan event is also recorded by several40Ar–39Ar dates in the range 300–325 Ma from the frontal units of the Mauritanide Belt in central and southern Mauritania (Villeneuve and Cornée, 1991, with ref. therein). Moreover, in northern Mauritania close to the southern tip of the Adrar Souttouf massif, garnet–omphacite associations Fig. 5.Sketch map of the Mauritanide thrust belt and frontal units in the Saharan Provinces of Morocco. Geological background after the Geological map of Morocco, scale 1:1,000,000 (1985). Location of wells that reached the crystalline basement of the Boujdour Basin after Sougy, inMichard and Sougy (1977). A–D: traces of cross-sections inFig. 7.
equilibrated at 550–600 °C, 13–15 kbar have been described and dated by Sm–Nd at ca. 330–340 Ma (Le Goff et al., 2001). The authors also measured from the same eclogite U–Pb zircon age at 595 Ma, interpreted as the age of the gabbroic protolith. The origin of the protolith is not clear (the Pan-African ophiolitic gabbros are older by ca.150 Ma;Gasquet et al., 2008, with ref. therein), but the Sm–Nd date strongly suggests that these units have been deeply buried in a Variscan subduction zone during the Late Devonian–Early Carbonif- erous, before being exhumed and included in the Mauritanide tectonic prism.
3.2. The Mauritanide frontal units (Dhlou–Zemmour and westernmost Anti-Atlas)
North of the Oulad Dlim transect, which is typified by direct thrusting of the Mauritanide metamorphic nappes over a thin, shallow marine cratonic cover (Figs. 5 and 7A), the Dhlou–Zemmour arcuate belt (Dhloa Belt inLécorché et al., 1991), partly mapped by Rjimati et al. (2002c)involves three main folded units thrust over a thin, autochthonous platform series (Fig. 7B). The stratigraphic sequence, 500 to 2000 m-thick, includes Ordovician clastics (the thickness of which increases steeply westward;Destombes et al.,
1969), Silurian slates, Devonian limestones and pelites with reef buildings (Wendt and Kaufman, 2006). Cambrian and Late Neopro- terozoic formations only occur in the most internal units. The Dhlou– Zemmour Arc can be regarded as a thin-skinned fold-and-thrust foreland belt in front of the main Mauritanide orogen. This arc developed in the transition zone from the very thin, incomplete Palaeozoic cover series of the Reguibat Shield on the Oulad Dlim transect to the thick and comprehensive series of the Tindouf Basin and Anti-Atlas domain.
After a 100 km-long hiatus in the outcrops of the deformed Palaeozoic units, the Bas Draa segment of the Anti-Atlas (Tilemsoun transect;Fig. 7C) differs from the Dhlou–Zemmour belt in three ways, i) the much greater thickness of the Palaeozoic series (nowN6000 m instead of∼2000 m); ii) the obliquity of its external boundary with respect to the cratonic domain (Fig. 5), resulting in the development of en-échelon folds along this boundary (Soulaimani et al., 1997), and iii) the involvement of the Precambrian basement in the shortening deformation, documented by the distribution of foliation and lineation structures within and around the Bas Draa inlier (Soulaimani et al., 1997). Here, the context is that of thick-skinned tectonics. We may assume that the evolution from thin-skinned to thick-skinned is linked to the increased Palaeozoic thinning of the continental crust, Fig. 6.Sketch map (A) and generalized cross-section (B) of the Oulad Dlim (Adrar Souttouf) Massif, afterVilleneuve et al. (2006), modified and completed within the framed area afterRjimati et al. (2002a,b,c). The following units (U) are defined afterVilleneuve et al. (2006): Oued Togba (OT), Sebkha Gezmayt (SG), Dayet Lawda (DL) and Sebkha Matallah (SM). In the frontal domain, the Quartzites (Tisnigaten) and Laglat units are defined in this work. Carb.: Carbonatites; Sk: sebkha; Sy: Nepheline-bearing syenite. C: Detailed cross- section of the Mauritanide front near Aousserd (Awsard), afterRjimati et al. (2002a)and personal observations. L: Laglat; T: Tisnigaten.
and then to the occurrence of more paleofaults (prone to be inverted during the Variscan collision) in the northern areas relative to the south-western ones.
North of the Draa Valley, the Ifni transect of the Anti-Atlas can be also interpreted in terms of thick-skinned, ESE-verging tectonics (Fig. 7D). The importance of shortening and thrust tectonics in the Cambrian series south-west of the Ifni inlier (Plage Blanche area) has
been stressed byBelfoul et al. (2001), Hoepffner et al. (2005) and Soulaimani and Burkhard (2008). The Plage Blanche fold–thrust belt shows inverted Middle Cambrian sequences with west shallow- dipping foliation thrust over the Cambrian folds further east. The Cambrian succession detached on its deepest, layered carbonate levels where tight minor folds with conspicuous tectonic foliation are observed both west (Plate 1, photo B) and east of the Ifni inlier Plate 1.Typical structures of the Variscan units of Morocco. A: Polyphase deformation in the Laglat Unit, Oulad Dlim Massif, Southern Provinces. B: Ductile microfolds in Lower Cambrian marbles–metasiltites from the core of a reclined fold, SW of Ifni, western Anti-Atlas. C: Large recumbent folds in the Cambrian metapelites and amphibolites, Aouli- Mibladen inlier, Eastern Meseta. D: Upright fold in the very low-grade Carboniferous turbidites, Fourhal Basin, Western Meseta. E: Kef-el-Mouneb metaconglomerates (Early–Middle Devonian, Central Rehamna). The quartzite pebbles (Ordovician?) are cigar-like stretched along the NNE fold axis direction, whereas they areflattened in the Sidi Abdallah outcrops further north (Piqué, 1973, 1975; Lagarde and Michard, 1986). F: Plagiogranite dyke crosscutting the cumulate gabbro of the Kettara mafic sill, Central Jebilet Massif. The dyke thickness is about 2 m.
(Lakhsass Plateau;Soulaimani and Bouabdelli, 2005). Deformation occurred under low-grade greenschist-facies conditions at the bottom of the Palaeozoic pile up to the Kerdous area to the east (Ruiz et al., 2008; Sebti et al., 2009). The Plage Blanche and Ifni units represent internal fold–thrust units relative to the Bas Draa and Tazeroualt– Kerdous inliers. All these units developed as frontal belts relative to the Mauritanide high-grade metamorphic axis, which acted as a tectonic buttress and is inferred to continue northward beneath the Atlantic margin.
4. The Meseta Domain: a polyphase collage
The Variscan belt of northern Morocco was sometimes referred to as“Meseta Block”(e.g.Stampfli and Borel, 2002). Indeed, this domain is far from being a single structural block (Fig. 8A). In contrast, it results from the Late Devonian–Late Carboniferous accretion of distinct sub-domains (structural zones) with different ages of folding and grades of concomitant metamorphism (Fig. 8B). These are, from
the older to the younger age of paroxysmal folding, the Sehoul Block, the Eastern Meseta and the Western Mesetasensu stricto(i.e. outside of the Sehoul Block). The Western Meseta is in turn subdivided into the Nappe Zone, Central Zone, Coastal Block and the intervening Western Meseta Shear Zone (WMSZ). The South Meseta Fault (SMF) corresponds to the collision zone of both the Western and Eastern Mesetas against the Anti-Atlas. Finally, Variscan metamorphic units mostly hidden beneath the Mesozoic–Cenozoic formations occur along the Western Meseta border, either north of the Sehoul Block or west of the Coastal Block.
4.1. Sehoul Block, Rabat–Tiflet Fault Zone and Mazagan Escarpment The outcrops of the Sehoul Block (Fig. 9A) occur north of the E- trending Rabat–Tiflet Fault Zone (RTFZ), which basically represents the uplifted basement of the Western Meseta Sidi Bettache Carbon- iferous Basin. The Sehoul Block consists of Lower–Middle Cambrian (and Ordovician?) siliciclastic phyllites and metagreywackes affected Fig. 7.Serial cross-sections of the Mauritanide Front and associated fold–thrust foreland belts in southern Morocco. Location of transects A–D: seeFig. 5. A: Aousserd transect.
B: Dhlou–Zemmour transect. C: Bas Draa–Tilemsoun transect. D: Plage Blanche–Ifni transect. The numbered arrows refer to successive senses of displacement along the listric faults (1: Latest Precambrian–Ordovician; 2: Early Carboniferous). Abbreviations as inFig. 3with Qz: Quartzite nappe; ph: phyllites; og: orthogneiss; LCr-Ng: Lower Cretaceous–Neogene.
by low-grade metamorphism and E-trending folds (Fig. 9B) and subsequently intruded by the calc-alkaline Rabat granite (Piqué, 1979; El Hassani, 1994a). The Palaeozoic sequence of the Sehoul Block ends with unconformable continental conglomerates dated from the Early Visean by their fern fossils (Danzé-Corsin, 1960; Izart and Vieslet, 1988; Fadli, 1994a,b). Based on ancient isotopic datings of the granite at ca. 430 Ma (Rb–Sr isochron; Charlot et al., 1973, recalculated byEl Hassani et al., 1991), and of the Cambrian phyllites at ca. 450 Ma (K–Ar onfine-grained phyllosilicates;El Hassani et al., 1991; El Hassani, 1994b), the Sehoul Block has been regarded for long, and still is generally regarded as an exotic“Caledonian terrane”(e.g.
Piqué and Michard, 1989; Hoepffner et al., 2005; Michard et al., 2008;
Simancas et al., 2009). Accretion of the Sehoul Block against Western Meseta was thought to be also a“Caledonian”event, based on the occurrence of unconformable Late Silurian–Early Devonian“Old Red Sandstones”-type red beds on top of granite slivers south of Tiflet (Fig. 9A, C;El Hassani, 1994a). However, new datings of zircons from the Rabat granite at 367 ± 8 Ma (U–Pb La-ICPMS) and 367 ± 3 Ma (207Pb–206Pb stepwise evaporation), and of the Tiflet granite at 605 ± 4 Ma (both methods) have been published byTahiri et al. (2010).
Assuming that the K/Ar 450 Ma age previously obtained from the Cambrianfine-grained phyllosilicates results from excess argon in the poorly recrystallized detrital muscovite grains, we may now regard the Sehoul Block merely as an Eo-Variscan or Acadian block, folded and intruded by the Rabat granite during the (Late?) Devonian. This low-grade block was exhumed and accreted to the RTFZ dextral strike-slip zone (Cailleux et al., 1984) during the Visean (in the RTFZ south of Tiflet, pebbles of granite and thermally metamorphosed rocks similar to those of the Sehoul Block are lacking in the Tournaisian conglomerates, but they do appear in the Upper Visean).
Concerning the RTFZ, the new data show the occurrence of a Neoproterozoic basement as in other areas of Western Meseta (El Jadida, Central Rehamna, Zaian). However, the unconformity of the Late Silurian–Early Devonian red beds upon the Early Ordovician pelites and granitic basement witnesses the occurrence of a pre- Variscan tectonic and erosional event, which do not occurs elsewhere in the Meseta, except in the Coastal Block and Western Meseta Shear Zone (see below).
The northward sub-surface extension of the Sehoul Block is very likely beneath the western Rif foredeep (Fig. 8A) as granite and
Fig. 8.The Meseta Domain and adjacent Anti-Atlas foreland. A: Structural map of the Variscan Meseta Domain, fromMichard et al. (2008), modified afterOuanaimi and Petit (1992), Baidder et al. (2008), Soulaimani and Burkhard (2008) Michard et al. (2010), and Tahiri et al. (2010). White: Mesozoic–Cenozoic cover. B: Stratigraphic evolution and geodynamic events, afterHoepffner et al. (2005) and Michard et al. (2008), modified.
greenschist-facies phyllites have been reached by some drill cores close to Kenitra and Sidi Kacem (Wdowiarz, 1987). Further east in the Fes region, the extension of the RTFZ and its Neoproterozoic basement is suggested by the occurrence of granite and rhyolite pebbles in the Devonian conglomerates of the Immouzzer Kandar inlier (Charrière and Regnault, 1989), but the extension of the Sehoul Block itself is not documented. Finally, in the Mazagan Escarpment offshore El Jadida, mylonitic granodiorite have been drilled and yielded K/Ar age at
∼450 Ma (Kreuzer, inRuellan, 1985), suggesting some continuity with the allegedly“Caledonian”Sehoul Block (Michard et al., 2008). In the present state of knowledge, this interpretation must be discarded. In fact, the Mazagan Escarpment also yielded dredged charnockite samples dated at 900–1000 Ma (K–Ar). Therefore, even not taking into account these controversial K/Ar ages, the occurrence of such high-grade rocks contradicts any correlation with the low-grade Sehoul Block.
4.2. Eastern Meseta
This sub-domain comprises the Palaeozoic massifs of Eastern Mesetasensu stricto(i.e. the eastern Mesozoic platform of the Atlas domain), the High Atlas inliers of Mougueur and northern Tamlelt, and the eastern Tazekka Massif (Fig. 8A). West of the latter massif, the boundary of the sub-domain is marked by two vertical or steeply east- dipping faults (Fig. 10A), namely the Tazekka–Bsabis Fault Zone (Hoepffner, 1977; Hoepffner et al., 2005) and the Tizi n'Tretten Fault (Charrière and Regnault, 1989; Willefert and Charrière, 1990), which together form the Middle Meseta Fault Zone (MMFZ). Further south, the MMFZ corresponds to the basal thrust of the Khenifra nappes. The Eastern Meseta sub-domain is characterized by an Eo-Variscan folding event, postdating the Givetian–Frasnian (Marhoumi et al., 1983) and predating the Upper Visean transgression. Eo-Variscan folds are recumbent or overturned and mostly W- or SW-verging (Plate 1,
photo C), except in some areas (e.g. Beni Snassen) where eastward vergence can be observed (Hoepffner et al., 2006). The metamorphic conditions associated with this early folding vary from low-grade greenschist- (Debdou–Mekkam) to amphibolite-facies (Midelt). Phyl- lite samples from Midelt were dated at 366 ± 7 Ma (Rb/Sr whole rock, Clauer et al., 1980) and Debdou–Mekkamfine-grained mica fractions yielded poorly reliable K–Ar ages at 368 ± 8 and 372 ± 8 Ma (Huon et al., 1987). The Eastern Meseta basement was exhumed andflooded with andesitic basalts, andesites, trachytes, rhyolites and ignimbrites with calc-alkaline to shoshonitic character during the Late Visean– Namurian (Kharbouch et al., 1985). In the Tazekka Massif, volcanics includes rare basalts and abundant high-K rhyolitic ignimbrite and lavaflows probably originating from anatexis of an intermediate crust (Chalot-Prat, 1995). Monzogranite and granodiorite plutons emplaced contemporaneously at depth (Tanncherfi, 344 ± 6 Ma and 325 ± 15 Ma, Rb–Sr,Ajaji et al., 1998; Midelt, 333 ± 2 Ma and 319 ± 2, U– Pb zircon,Oukemeni et al., 1995). Part of Eastern Meseta was emergent (Tazekka) at that time, whereas other parts remained under shallow marine conditions up to Middle Westphalian (Jerada coal basin). Thus, the Eastern Meseta Eo-Variscan domain was a magmatic arc during the Visean–Westphalian. Consistently, it behaved almost rigidly during the Late Carboniferous shortening phase.
4.3. Western Meseta
The Western Meseta s.str. is characterized by folding and metamorphic events concentrated during the Late Carboniferous (Variscan events s.str.). However, Eo-Variscan (Late Tournaisian) folding is also observed in the Cambrian–Ordovician formations at the bottom of the eastern Azrou–Khenifra Basin (Zaian Massif,Fig. 10A), suggesting an initial continuity with Eastern Meseta.
}
Fig. 8(continued).
4.3.1. Structural zones
The Central Zone of Western Meseta includes uplifted massifs (mainly the Zaian and Zaer anticlinoria with Cambrian–Ordovician outcrops), and downthrown Devonian–Carboniferous basins (Azrou– Khenifra, Fourhal, Sidi Bettache, eastern Rehamna and central-eastern Jebilet). The basinal areas are characterized by turbiditic infilling associated with bimodal (gabbroic and felsic) magmatism during the Early Carboniferous. Shortening initiated as early as the Late Visean– Early Namurian (Serpukhovian) in the eastern regions, in association with the emplacement of the synsedimentary Ziar nappes, thus defining theNappe Zone(Fig. 10A). Such synsedimentary nappes can be followed southward up to the Eastern Jebilet and Ait Tamlil massifs (Huvelin, 1977; Jenny et al., 1989). In the north-eastern part of the Central Zone, the Ziar nappes are overlain by the Khenifra nappes, which can be regarded as the front of the advancing Eastern Meseta (Bouabdelli, 1994). Except in the high-grade metamorphic units of Central and Eastern Rehamna (see below), the Central zone was shortened through upright or NW-verging folds (Plate 1, photo D) associated withflat-ramp faults that developed in sequence from east to west (Fig. 10B;Ben Abbou et al., 2001). The Central Zone is partly crosscut axially by the Smaala-Oulmes dextral transpressional fault (SOF), which carries the Fourhal Basin border onto the Zaer anticlinorium. The SOF likely continues SW-ward beneath the Mesozoic–Cenozoic plateaus up to the Eastern Rehamna (see below, Section 4.3.2).
TheCoastal Block(“Môle côtier”) of Western Meseta was defined (Michard, 1967, 1976) by its weak Variscan deformation (kilometre- scale open folds) and deep level of erosion (large outcrops of Cambrian rocks) contrasting with the strongly folded Devonian and Carboniferous series further east. The Coastal Block includes not only
the Casablanca anticlinorium and the western parts of the Rehamna and Jebilet massifs, with dominant Cambrian outcrops, but also the Doukkala Basin, which corresponds to a Silurian–Devonian basin beneath the Mesozoic cover (Echarfaoui et al., 2002). This moderately deformed zone extends southward up to the westernmost High Atlas Palaeozoic Massif (Cornée, 1989; Piqué and Michard, 1989).
TheWest Meseta Shear Zone(WMSZ) separates the Coastal Block from the Central Zone (Fig. 10A). Along this zone, from the Central and Eastern Rehamna to Central Jebilet, to Jebel Tichka Massif in the High Atlas, chlorite–chloritoid ± biotite and biotite–garnet ± staurolite ± kyanite or andalusite metapelites occur close to basic and felsic intrusions, suggesting high geothermal gradients during folding (Michard, 1968a,b; Hoepffner et al., 1982). However, the latter interpretation has been challenged (Aghzer and Arenas, 1995, 1998;
Baudin et al., 2003), and then deserves additional discussion.
4.3.2. Structural framework of the Rehamna metamorphic culmination The Rehamna Massif straddles the WMSZ (Fig. 11). The Western Rehamna belongs to the Coastal Block, but the Lower Cambrian carbonates and Middle Cambrian greywackes (Schistes àParadoxides Fm.) of its eastern border are strongly folded, weakly metamorphic and intruded by the Sebt–Brykiine post-tectonic granite (Michard, 1967; Guezou and Michard, 1976). The Central Rehamna is conven- tionally bounded by the NE-trending Median and Ouled Zednes faults and underlines the axis of the WMSZ. This narrow stripe exposes a more metamorphic and less clearly dated stratigraphic sequence.
Further east, the Eastern Rehamna includes both metamorphic units at the bottom and low- to very low-grade units above them (Cornée et al., 1982; Michard et al., 1982; Baudin et al., 2003; Razin et al., 2003).
Fig. 9.The Sehoul Block and Rabat–Tiflet Fault Zone (RTFZ) in the eponymous regions, afterEl Hassani (1994a,b), modified afterTahiri et al. (2010)who use the name of“Bou Regreg Corridor”for the RTFZ. A: Sketch map with location of B–C cross-sections. B: Cross-section within the Sehoul Block. C: Cross-section of the RTFZ. S1: metamorphic foliation; or:
Ordovician; s: Silurian; di, dm, ds: Lower, Middle, Upper Devonian; hT: Tournaisian; hV: Visean.
Controversiesfirst concern the stratigraphy and structure of the metamorphic units. The Central Rehamna northeast of the Sebt– Brykiine granite shows a Late Neoproterozoic metagranite (593 ± 8 Ma, U–Pb zircon;Baudin et al., 2003) overlain by metaconglome- rates and Lower Cambrian marbles, in turn overlain by Middle Cambrian metagreywackes. The latter are overlain locally (Kef-el- Mouneb) by reddish metaconglomerates (Plate 1, photo E), and in other places by the phyllites and quartzites of the Skhour Fm. The Kef- el-Mouneb conglomerates are ascribed to the Early–Middle Devonian, based on their homology with the Sekhira-es-Slimane conglomerates from the north-westernmost (and lowest grade) Central Rehamna, which yielded Middle Devonian fossils (Destombes et al., 1982).
Aghzer and Arenas (1995, 1998), thenBaudin et al. (2003)claimed that a tectonic contact separate the Kef-el-Mouneb conglomerates from the underlying Middle Cambrian metagreywackes. In contrast, we contend that the conglomerates represent an autochthonous, unconformable continental onlap on top of the Cambrian greywackes based on the following: i) there is no local evidence of such tectonic contact; ii) both formations show the same metamorphic grade (amphibolite-facies); and iii) the eastern Coastal Block and adjoining WMSZ are characterized by“Old Red Sandstones”deposits not only in the Central Rehamna, but also in the Western Jebilet (Huvelin, 1977;
Tahiri, 1983; Beun et al., 1986) and the Western High Atlas (Talmakent area;Cornée, 1989). In the latter area, the“ORS”deposits are clearly unconformable onto the Ordovician quartzites. Likewise, this regional unconformity correlates with that observed in the Rabat–Tiflet Fault Zone where Upper Silurian–Early Devonian red beds overlie the Tiflet granite (Section 4.1above).
The age and structural position of the Skhour Fm. is more controversial. This strongly folded series of phyllites and quartzites has been first regarded as Ordovician (Gigout, 1951), then as a metamorphic equivalent of the Famennian–Strunian formations of
the northernmost Rehamna (Michard, 1969; Baëcker et al., 1965).
Later on, the Skhour Fm. was considered as Ordovician (Destombes et al., 1982) although its lithostratigraphy differs in many aspects from that of the well dated, neighbouring Ordovician series (J. Lakhdar and Imfout synclines in the Coastal Block; Jebel Kharrou in the Eastern Rehamna). However, a Late Devonian age would account for the fact that the Skhour Fm. overlies the northern tip of the Kef-el-Mouneb conglomeratic lense (Fig. 11) without invoking a major tectonic contact at the bottom of the Skhour Fm. as assumed byAghzer and Arenas (1995, 1998) and Baudin et al. (2003).
In the (lower) metamorphic units of the Eastern Rehamna, the main stratigraphic controversy concerns the Lalla Tittaf Fm., which extends from the Benguerir area to the Oulad Ouggad“half-horst” (Fig. 11). The Lalla Tittaf Fm. consists of black schists with some sandy carbonate intercalations and bimodal metavolcanics, namely amphi- bolites, metagabbros and (scarce) acidic metatuffites. This formation has been ascribed to the Early Carboniferous (Visean–Namurian) based on its homology with the Sarhlef Schists in the Central Jebilet further south (Michard, 1969; Huvelin, 1977; Destombes et al., 1982).
However,Baudin et al. (2003) and Razin et al. (2003)proposed a
“Paleoproterozoic (?)”age for this formation, based on SHRIMP ages at 2136 ± 17 Ma obtained from 7 zircon grains over 9 extracted from one sample of the Lalla Tittaf metagabbro. Two other zircons gave ages of 2345 Ma and 335 ± 6 Ma, respectively. We feel that these results do not preclude an Early Carboniferous age for the gabbroic intrusion as the corresponding zircon grains may represent xenocrysts from the underlying continental crust. This is supported by the fact that in the nearby Jebilet massif,Dostal et al. (2005) obtained U–Pb ages at
∼2000 Ma, 700 Ma, 615–540 Ma and 328–280 Ma from zircon grains extracted from xenoliths found in Triassic lamprophyre dykes. The occurrence of a few inherited zircon cores with ages around
∼2000 Ma and 680 Ma has been also observed by Tahiri et al.
Fig. 10.Two schematic profiles across Western Meseta. A: General cross-section, afterHoepffner et al. (2006), modified. Location: seeFig. 8A. UNp: Upper Neoproterozoic; Cb:
(Lower)–Middle Cambrian; Or: Ordovician; Si–D: Silurian–Middle Devonian; UD–LC: Upper Devonian–Lower Carboniferous; C: Carboniferous; hV–N: Visean–Namurian; hW: Upper Westphalian. Red and cross signature: granitoids. S0-1: slaty cleavage, foliation (Eo-Variscan phases), mostly parallel to S0 (stratification plane); S2/S1-2: crenulation cleavage/
foliation (Variscan phases). A.K.B: Azrou–Khenifra Basin. Main faults or fault zones (SOF, etc.) as inFig. 8A. B: Sketch profile across the Nappe Zone and Fourhal Basin, afterBen Abbou et al. (2001). Circled numbers refer to the succession of the thrust faults through time, based on the age of the associated debris-flows. Notice the two décollement levels associated to the Lower Ordovician and Silurian shales, respectively.
(2010)in the Tiflet granite, emplaced at 605 ± 4 Ma (seeSection 4.1).
Moreover, the metamorphism of the Lalla Tittaf Fm. is obviously Variscan, and no previous (Eburnean) phase of metamorphism has ever been observed there. Accordingly, we contend that the Lalla Tittaf Fm. can be actually ascribed to the Early Carboniferous, and that it overlies the Devonian schists of the Ouled Hassine Fm. (metapelites, metaconglomerates and marbles) through a stratigraphic contact.
In contrast, the concept of a major thrust contact at the bottom of the Eastern Rehamna low-grade or very low-grade units (Baudin et al., 2003; Razin et al., 2003) is convincing. This contact had been recognized previously south of the Jebel Kharrou Ordovician massif (Michard, 1969, 1976; Piqué, 1982) and subsequently extrapolated to the Dalaat Ordovician, Devonian and Visean massif south of the Oulad Ouggad fault (Raïs-Assa et al., 1983). As the J. Kharrou formations are in obvious continuity with those of Koudiat el-Adam and Jorf Lahmar further west, the J. Kharrou basal thrust must persist westward until to merge with the Ouled Zednes Fault (Fig. 11). The J. Kharrou unit likely correlates NE-ward with the Zaer anticlinorium (Fig. 8A) whose Ordovician series are similar (Destombes et al., 1982; Razin et al., 2001), whereas the Dalaat Visean unit, which contains gravity-driven deposits, could correlate with the Fourhal Basin.
Two main metamorphic episodes have been recognized, synkine- matic and post-kinematic, respectively, and their isogrades mapped (Michard, 1968a,b; Hoepffner et al., 1982; Lagarde and Michard, 1986;
Baudin et al., 2003; Razin et al., 2003). The post-kinematic recrystallizations (mainly static andalusite and biotite) correspond to the contact aureoles surrounding the granite stocks or extending between three granite apexes (Ras-el-Abiod, Bir-el-Gourda and
Koudiat-er-Rmel, Fig. 11). The large Sebt–Brykiine granite clearly crosscuts the Coastal Block and Central Rehamna structures. Remark- ably, the isogrades of the early, synkinematic metamorphism also crosscut the Oulad Zednes Fault (OZF), whereas they tend to parallel the J. Kharrou thrust further east. The highest-grade, biotite–garnet– staurolite and biotite–garnet–kyanite assemblages are basically found within a domal structure straddling the OZF, and along the south border of the Oulad Ouggad tilted block. This suggests that, i) the OZF is inherited from an ESE-dipping paleofault; ii) metamorphism has been at least partly controlled by the tectonic overburden; iii) the thrust complex extended onto the Central Rehamna (beyond the OZ paleofault) during the metamorphic event. The P–T estimates for the highest-grade assemblages are close to 500–550 °C, 0.5–0.8 GPa (Hoepffner et al., 1982; Aghzer and Arenas, 1995, 1998). The minimum P estimate would suggest a burial of ca. 15 km, and therefore the T estimate would require a relatively high geotherm, close to 33–35 °C/km. Such inference is supported by the occurrence of a typical Buchan-type metamorphism with synkinematic andalu- site and staurotide in the Central Jebilet (Huvelin, 1977; Essaifiet al., 2003; Lahfid, 2005), i.e. 40 km further south in the same structural zone. Accordingly, we suggest that the early, synkinematic metamor- phism of the Rehamna and Jebilet region was controlled not only by tectonic burial, but also by the development of high T structures along the WMSZ, with irregularly dipping isotherms instead of horizontal isotherms (Baudin et al., 2003; Razin et al., 2003). These putative thermal structures could have been in relation with the incipient melting of the underlying crust, possibly caused by the slightly older, mafic magmatism (seeSection 5.4). The removal of the∼15 km thick Fig. 11.Sketch metamorphic map of Central Rehamna, afterBaudin et al. (2003), modified. The tightly spaced“isograde”boundaries (mineral assemblages in metapelites) crosscut the Ouled Zednes Fault Zone, suggesting a Buchan- (rather than Barrow-) type metamorphism during deformation. Final emplacement of granites at ca. 300 Ma postdates the main, syntectonic metamorphism.
Boummane and Olivier, 2007). They result from magmatic mixing between mafic magmas originating from more or less enriched mantle and anatectic magmas derived from partial melting of continental crust (Gasquet et al., 1996; Haïmeur and El Amrani, 2005; El Hadi et al., 2006a). Their isotopic ages span the Visean–Autunian period (330 to 300–290 Ma) in both the Eastern and Western Mesetas (Fig. 8B). The only exception corresponds to the Rabat granite, intrusive in the Sehoul Block and dated at 367 ± 3 Ma (U–Pb zircon;Tahiri et al., 2010).
In the Carboniferous basins of Western Meseta, the felsic intrusions mostly postdate the mafic magmatism (Plate 1, photo F), although some magmatic breccias locally suggest sub-contemporaneous em- placement during the Late Visean (Essaifiet al., 2003). The youngest intrusions (Lower Permian;Mrini et al., 1992) are coeval with trachy- andesite and rhyoliteflows emplaced in pull-apart basins scattered in the entire Meseta Domain (Saidi et al., 2002).
4.4. South Meseta Fault and Sub-Meseta Zone
These important structures of the Moroccan Variscides have been ultimately discussed by Michard et al. (2010). We summarize hereafter our main points.
The southern boundary of the Meseta Domain wasfirst recognized by Mattauer et al. (1972)in the Tizi n'Test dextral fault, then by Michard et al. (1982)in the Tineghir south-verging reverse faults.
Michard et al. (1989)andPiqué and Michard (1989)argued that the Tizi n'Test Variscan boundary has to continue till eastern Morocco as the Eastern Meseta deformation strongly contrasts with that of Eastern Anti-Atlas, and coined the term of“Atlas Palaeozoic Transform Zone (APTZ)”. West of the Ouzellarh salient, this boundary zone is narrow, involving essentially the Tizi n'Test and N'Fis faults (Ouanaimi and Petit, 1992). In contrast, the boundary zone enlarges eastward from the Ouzellarh salient to Tineghir, and to the Tamlelt Massif of Eastern High Atlas (Fig. 8A). In the latter massif,Houari and Hoepffner (2003) have shown that the main boundary occurs between the North-Tamlelt region, which shows the polyphase structure typical
Tamlelt (Houari and Hoepffner, 2003) as well as in the Tineghir area (Michard et al., 1982; Soualhine et al., 2003; Cerrina Feroni et al., 2010). The SMF itself shows the characters of a crustal-scaleflower- structure in the Tamlelt Massif (Houari and Hoepffner, 2003).
5. Interpretation and discussion 5.1. Crustal-scale cross-section
The bulk structure of the northern Moroccan Variscides (Meseta Domain) by the end of the Palaeozoic can be summarized in a schematic, crustal-scale cross-section (Fig. 12). Starting from the south, the cross-section illustrates the deeply asymmetric fan structure of the belt, with narrow zone of S-verging structures in the Anti-Atlas thin-skinned foreland belt, and a dominant NW-ward vergence in the Meseta Domain itself. The occurrence of thrusting along both boundaries of the Central Zones is outlined. The section also visualizes hypothetically a significant thickening of the crust beneath most of the Meseta orogen, consistent with the importance of the anatectic component in the felsic intrusions scattered in the belt (Chalot-Prat, 1995; Gasquet et al., 1996; El Hadi et al., 2006a). The occurrence of a large, poorly shortened block (Coastal Block) at the west border of the Meseta Domain must be emphasized. To the NW, beneath the Doukkala Mesozoic–Cenozoic basin and the Atlantic margin, the section is mostly conjectural. However, the Doukkala seismic profiles suggest the occurrence of a Late Devonian unconfor- mity on top of the shortened Cambrian–Early Frasnian series (Echarfaoui et al., 2002). Further west, we suggest that the Mazagan Plateau could be underlain by a southern extension of the Sehoul Block, whose northern outcrops have been displaced eastward along the RTFZ during the Late Carboniferous. The Sehoul Block is characterized by a Devonian event including the emplacement of the Rabat granite at 367 Ma (Tahiri et al., 2010), and probably also the greenschist-facies metamorphism and folding of the Cambrian (–Ordovician?) series. Early docking of this Eo-Variscan
Fig. 12.Schematic crustal-scale section of the northern Moroccan Variscides by the end of the Variscan orogeny. Location: seeFigs. 8A or 13A.
(Acadian) block against the Coastal Block likely occurred during the Famennian. Still further west, the last basement rocks collected at the very border of the continent are those of the Mazagan Escarpment.
They are poorly characterized by high-grade metamorphism and pre- Variscan K–Ar ages (450 Ma and 900–1000 Ma). The section (Fig. 12) shows that the Variscan Orogen developed outside of the 2 Ga-old, Eburnean crust of the WAC, on top of a crust that have been created, or at least modified by the Pan-African Orogeny. Nevertheless, 2 Ga-old zircon xenocrysts or zircon cores have been found in varied magmatic rocks from all over Meseta (Oukemeni et al., 1995; Dostal et al., 2005;
Tahiri et al., 2010; see alsoSection 4.3.2), suggesting a derivation from the Eburnean crust.
5.2. Pre-orogenic restoration
The overall structure of the Moroccan Variscides (Meseta, Anti- Atlas and northern Mauritanides) may be also shown in map view (Fig. 13A). The tentative restoration of the varied sub-domains before the Eo-Variscan events, i.e. at∼360 Ma, can be approached (Fig. 13B) based on reasonable estimates of the internal shortening that affected the elementary sub-domains, and on the available stratigraphic data (Fig. 8B).
InFig. 13B, the Oulad Dlim Mauritanide nappes are restored at roughly a hundred kilometres WNW of their present-day location, which is a bulk estimate as the frontal nappes (e.g. Quartzite nappe) are less displaced than the internal ones. The typical Mauritanide nappes develop in the regions where the WAC (Neoproterozoic–) Palaeozoic cover is extremely thin, in correspondence with the Reguibat Shield. Further north, thicker Palaeozoic series were deposited and preserved, and the frontal units of the Mauritanide Belt consist of the arcuate fold–thrust units of the Zemmour and westernmost Anti-Atlas. Accordingly, the Mauritanide structural pattern should end south of the SMF as the Anti-Atlas system itself.
North of the SMF, the location of the Western Meseta Central Zone is indicated at about 100 km WNW of its present-day location, taking into account the strong internal shortening and the importance of dextral strike-slip along the Tizi n'Test Fault (Mattauer et al., 1972;
Ouanaimi and Petit, 1992). The location of the Coastal Block is inferred at significant distance further west, assuming∼50 km thrusting and continuous shortening across the WMSZ. The Sub-Meseta Zone founds its place at several kilometres north of its present contact with Eastern Anti-Atlas, i.e. before the Late Carboniferous thrust tectonics observed at Tineghir and Tamlelt. The Eastern Meseta and the Nappe Zone of Western Meseta can be approximately located at 100–200 km east of the Central Zone. Finally, the Sehoul Block and its southern extension (Sehoul Zone) can be located along the northwest boundary of the Coastal Block as discussed above (Section 5.1). Suturing of the Sehoul Zone against the Coastal Block would have marked the earliest Variscan compressional event during the Late Devonian. It is worthy to emphasize that the“Old Red Sandstones”deposits preserved in the Coastal Block, WMSZ and RTFZ (Figs. 13B and 14; cf.Section 4.3.2) do not document a “Caledonian” compressive event, but merely the subaerial erosion of a ridge close to the eastern border of the Coastal Block, followed by its collapse and the extensional opening of N- trending hemigrabens (Cornée, 1989).
5.3. Moroccan Meseta: a Lower–Middle Palaeozoic passive margin domain The passive margin of north-western Gondwana began to form shortly after the Pan-African orogeny, during the Late Ediacaran–Early Cambrian (Piqué et al., 1999; Soulaimani et al., 2003; El Archi et al., 2004; Landing, 2005). Early Cambrian extension is recorded in the Anti-Atlas domain by paleofaults associated with conspicuous slumpings (Plate 2, photo A). Calc-alkaline to alkali-tholeiitic basalts are widely exposed both in the northern border of the Anti-Atlas and in the Meseta Domain (Fig. 13B andPlate 2, photo B;Ouali et al., 2001,
2003; El Hadi et al., 2006b; Raddi et al., 2007; Ezzouhairi et al., 2008;
Aarab, 2009). The trend of the Variscan structures is often inherited from the Cambrian paleogeography. For example, the WMSZ parallels the“Western Meseta Cambrian Graben”characterized byN6000 m- thick Middle Cambrian greywackes (Bernardin et al., 1988; Piqué et al., 1995; El Attari et al., 1997). The Sehoul Block itself shows similar Cambrian deposits with some basaltic intercalations, and would represent a deep-sea fan complex of well-sorted sand and silt with Gondwanan derivation (Schenk, 1997).
Fragmentation of the north-western Gondwana platform went on during the Ordovician. In the eastern Anti-Atlas, chaotic breccias are observed (Plate 2, photo C) in relation with a major, NW-trending paleofault (Baidder et al., 2008). Slope facies are also documented on the Anti-Atlas northern paleomargin south of Tineghir (Malusá and Schiavo, 2007) or north of the Ouzellarh salient (Plate 2, photo D). In the Meseta Domain, contrasted sedimentary facies occur during the Ordovician. In the Coastal Block and part of Central Meseta, they are similar to those of the Anti-Atlas, i.e. rich in quartzitic sandstones sourced in the central Sahara, but elsewhere they are richer in pelites and claystones (Destombes et al., 1985). It is worthwhile emphasizing the occurrence of thick pillow lavas and gabbroic sills intercalated within the Lower Ordovician of the Rabat–Tiflet Zone (Piqué, 1982;
Tahiri and El Hassani, 1994). These mafic rocks show poorly differenciated calc-alkaline affinities and can be referred either to rift or subduction settings (El Hassani, 1994a,b). They could mark a lost oceanic hiatus (Hurley et al., 1974) southeast of the Sehoul Block (Fig. 13B).
After the global regression of the Late Ordovician, extension is documented again during the Silurian by scattered alkaline basalt outpours (Oulad Abbou and Rabat area;El Kamel et al., 1998). The coeval sedimentation is rather monotonous (shale deposits linked to melting of the Saharan icecap) except in the Coastal Block-RTFZ regions where red beds are observed. The subsequent, Devonian facies testify for a strong internal differentiation of the Meseta Domain. The Lower–Middle Devonian reefal platform series of the Coastal Block, RTFZ and NW Central Meseta closely compare with those of the Anti- Atlas platforms (Hollard, 1967; Tahiri and Hoepffner, 1988; Piqué et al., 1991; Zahraoui, 1994). On the other hand, basinal facies with shales, platy limestones, radiolarian cherts and turbidites developed from the Lochkovian to Famennian in most of the Western and Eastern Mesetas (Hollard, 1967; Marhoumi et al., 1983; Bouabdelli et al., 1989; Jenny et al., 1989; Razin et al., 2001). The activity of the paleofaults between the platform and basin areas increased during the Famennian as shown by unconformities and/or chaotic conglom- erates (e.g. west and north borders of Sidi Bettache Basin,Figs. 8A and 9;Piqué, 1984; Zahraoui, 1991; Fadli, 1994a,b). In the Sidi Bettache Basin, Upper Devonian–Lower Tournaisian deposits are accompanied locally by effusive mafic intercalations of alkaline and tholeiitic- transitional affinities (Kharbouch et al., 1985; Kharbouch, 1994). This extensional episode has its exact counterpart in the Anti-Atlas domain where platforms (e.g. N-Tafilalt) and basins (Maider, SE-Tafilalt) were individualized during the Givetian–Famennian (Plate 2, photo E;
Wendt, 1985; Baidder et al., 2008). These observations are contrary to the reconstructions based on paleomagnetic data (e.g.Stampfli and Borel, 2002), which suggest a 3000–4000 kilometre-wide ocean between the Moroccan Meseta and Anti-Atlas during the Devonian, but in accordance with reconstructions based on geological and paleobiological date (e.g.Walliser et al., 1995; Fortey and Cocks, 2003;
Robardet, 2003; Rücklin, 2009).
To summarize, comparing the Cambrian–Devonian stratigraphy of the Meseta Domain with that of the Anti-Atlas yields evidence of a derivation of the former from a wide passive margin domain, which always remained close to the continent. Consistently, the rare Precambrian outcrops that occur in the Meseta Domain (El Jadida, Tiflet, Zaian anticlinorium,Figs. 8A and 10A; Central Rehamna,Fig. 11) consist of Late Ediacaran (meta-) rhyolites comparable with those of
the Anti-Atlas. U–Pb dating of zircons extracted from xenoliths in the Triassic dykes of the Jebilet Massif yielded∼2000 Ma, 700 ma, 615– 540 Ma and 328–280 Ma ages (Dostal et al., 2005), suggesting the
occurrence of a Pan-African basement with Eburnean components. As noticed above (Section 5.1), the occurrence of such Eburnean (Gondwanan) components has been inferred at several places Fig. 13.Present structure of the Moroccan Variscides in map view (A) and tentative restoration at ca. 360 Ma, before the Variscan events (B).
beneath the Meseta Variscan Belt. In contrast, if the 900–1000 K–Ar ages of the charnockite from the Mazagan Escarpment (Kreuzer, in Ruellan, 1985) are correct, and as the Mesoproterozoic Grenville Event is lacking in NW Africa (Gasquet et al., 2008), the“undefined unit” beneath the escarpment (Fig. 12) would originate from Laurussia (either Laurentia or Avalon, if the latter has a Laurentian basement as claimed byDostal et al., 2005).
5.4. Carboniferous geodynamics
During the Tournaisian, shortening occurred in the eastern part of Western Meseta and, probably, in the Eastern Meseta although shortening could have begun there during the Late Famennian according to the isotopic ages (366± 7 Ma and 368 ± 8–372 ± 8 Ma;
seeSection 4.2andFig. 8B). In contrast, extension dominated at that time in Western Meseta, being recorded by thick turbiditic sedimen-
tation and mafic volcanism. In the Fourhal Basin, pillow basalts and gabbros sills show a transitional to tholeiitic affinity (Kharbouch, 1994; Remmal et al., 1997) and local calc-alkaline affinities (Roddaz et al., 2002; Driouch et al., 2010), whereas in the Sidi Bettache Basin, they display alkaline-tholeiitic to transitional characters (Kharbouch et al., 1985), and tholeiitic affinities in the Jebilet Basin (Bordonaro et al., 1979;
Bordonaro, 1983; Aarab and Beauchamp, 1987; Essaifi et al., 2004;
Essaifiand Hibti, 2008) and in the Northern Rehamna (Hoepffner, 1982, Remmal et al., 1997). These volcanics emplaced mostly during the Visean (Huvelin, 1977; Marcoux et al., 2008; Moreno et al., 2008), prior to the post-Serpukhovian folding event (Plate 2, photo F). Felsic magmatism followed closely the mafic one in Western Meseta (Plate 1, photo F) and developed almost simultaneously in both the Western and Eastern Meseta from 330 Ma onward (Fig. 8B).
In an earlier conception of the Meseta Variscides as an intraconti- nental orogen (Fig. 14A), Piqué and Michard (1989) assumed that Fig. 14.Interpretations of the Carboniferous tectonic/magmatic setting of the Moroccan Meseta in terms of plate tectonics. A: afterPiqué and Michard (1989). B: afterRoddaz et al. (2002). C: This work. The east-dipping subduction accounts for the Late Devonian–Early Carboniferous magmatism and corresponds to the consumption of the Rheic Ocean, as admitted in the SW Iberian transect (seeFig. 15).
Western Meseta has been affected by pull-apart tectonics during the westward thrusting of Eastern Meseta (see alsoBouabdelli and Piqué, 1996). In contrast,Kharbouch et al. (1985), Boulin et al. (1988) and Roddaz et al. (2002)proposed that a west-dipping subduction would have operated beneath the Meseta Domain, with its hypothetic suture concealed beneath the High Plateaus (Fig. 14B). In their general discussion of the SE Variscan Belt evolution,Von Raumer et al. (2008) assume a similar setting for the Moroccan Meseta transect, combining
west-dipping subduction and dextral strike-slip. However, in line with Simancas et al. (2005, 2009), we observe that along the SMF, which takes the place of the putative suture to the SW, ophiolitic remnants are lacking. Accordingly, we favour another geodynamic setting (Fig. 14C) involving an east-dipping subduction whose former trough should have been located offshore the present Meseta. The origin of the subducted crust is discussed in the following section. Notice that the dip of the subduction zone was likely opposite in the Mauritanide transect (Caby Plate 2.Pre-orogenic extensional structures in the Moroccan Variscides. A: Disruption of siltite layers in a slumped carbonate bed, Lower Cambrian, east border of the Zenaga inlier, central Anti-Atlas. B: Tholeiitic pillow lava, Bou Acila, Zaian region (eastern Central Massif of Western Meseta). C: Imzioui chaotic conglomerates, Upper Ordovician (or5), Eastern Anti-Atlas (Destombes et al., 1985); light-colored blocks are Lower Cambrian limestones, darker ones are Lower Ordovician Fe-oolithic sandstones. D: Slumped beds in the Upper Ordovician of the Tizi n'Tichka Pass, north-eastern side of the Ouzellarh salient, Marrakech High Atlas. E: Upper part of an E-trending Devonian paleofault sealed by Famennian limestones, J. Amelane, Tafilalt, Eastern Anti-Atlas (Baidder et al., 2008). F: Folded sill of layered gabbro (lgb) on top of Visean–Serpukhovian turbidites (hV–S), Kettara, Central Jebilet (Essaifiet al., 2004).
and Kienast, 2009), in accordance with the opposite vergence of the belt.
The change of dip of the subduction zone would have been located along the westward extension of the SMF.
Low-angle extensional faults have been recognized by place, for example in the Western High Atlas (Cornée, 1989), in the Central Rehamna (Baudin et al., 2003; see above,Section 4.3.2) and in the Zaer anticlinorium (Chèvremont et al., 2001; Razin et al., 2001). These (rare) subtractive structures would record the late-orogenic isostatic re-equilibration of the belt.
5.5. Morocco–SW Iberia relationships in the Variscan framework The Variscan segment exposed in south-western Spain and southern Portugal is the closest part of the European Variscides respective to the Moroccan Meseta segment. Both segments belong to the“south-western”branch of the Variscan Belt (Fig. 15A) accreted either directly to north-western Gondwana (Morocco) or to the Cantabrian–South Sardinia salient of Gondwana that forms the core of the Ibero-Armorican Arc (Matte, 2001; Simancas et al., 2005; Rossi
Fig. 15.A: Tentative restoration of the Moroccan and Iberian Variscides in the Peri-Atlantic Variscan framework. Source map fromSimancas et al. (2005), modified as follows: i) the SMF is defined as a major Palaezoic intracontinental transform fault, distinct from the SAF; ii) the Mauritanides is a major Variscan segment distinct from the Anti-Atlas; iii) the Sehoul Block is no longer a“Caledonian”terrane, and iv) the Rheic suture is located offshore Morocco. The Variscan chips included in the Rif-Betic Alpine Belt (MG: Malaguides; KA:
Kabylias) might have been detached from Eastern Meseta (Hoepffner et al., 2005) or from south-eastern Iberia (Simancas et al., 2005). CA: Collector Anomaly; CZ: Cantabrian Zone;
EMM: Western Moroccan Meseta; LS: Lizard; MN: Montagne Noire; OMZ: Ossa-Morena Zone; SAF: South Atlas Fault; SIF: South Iberian Fault; SMF: South Meseta Fault; SPZ: South Portuguese Zone; S2i: Intracontinental“suture”equivalent to S2; WAL: West Asturian–Leonese Zone; WMM: Western Moroccan Meseta; WMR: Western Meseta Rift (Cambrian).
B: Schematic palaeogeography of the SW Iberia–Morocco region during the Late Palaeozoic. The Variscan segments of SW Iberia, Moroccan Meseta and Anti-Atlas–Mauritanides are separated from each other by a system of intracontinental transform zones. The NE-trending oceanic sutures that occur in SW Iberia are replaced by intracontinental sutures in the Moroccan Meseta segment.
based on dating of allegedly MORB metabasites of the BAOC at 355– 340 Ma (Azor et al., 2008), but their conclusions are controversial (Pin et al., 2008; Pin and Rodriguez, 2009). In contrast in Morocco, only one suture would hypothetically occur between the Sehoul Block and the remainder of Western Meseta, outlined by Lower Ordovician pillow lavas and deeply modified by dextral strike-slip displacements.
Simancas et al. (2009), and Tahiri et al. (2010)consider the Sehoul Block as a“Caledonian”terrane, which would derive (together with the unknown basement of the SPZ) from a salient of the SE border of Avalonia. Following these authors, the OMZ would come to an end somewhere north of Morocco, and the Moroccan Meseta would correlate with the Iberian east of the OMZ, i.e. with the Central Iberian, Asturian–Leonese, and Cantabrian zones (CIZ, ALZ, CZ).
We feel that the above proposals can be improved on the following aspects:
i) The Sehoul Block should be no longer defined as a“Caledonian terrane”as the allegedly Silurian Rabat granite is now dated at 367 ± 3 Ma (Tahiri et al., 2010), and as the metamorphism of the country-rocks is virtually undated (only one K/Ar date at 450 Ma besides 5 dates at 350–330 Ma on poorly recrystallized phyllites). We suggest to rather considering the Sehoul Block as an Eo-Variscan or Acadian terrane, affected by Late Devonian– Early Carboniferous events during its emplacement against the Rabat–Tiflet Zone.
ii) Accordingly, the Sehoul Block would not be correlated with Avalon, but rather to Meguma as previously proposed by Schenk (1997). U–Pb dating of the detrital zircon grains showed that the Cambrian turbidites of the Meguma terrane contains 566–675 Ma and ∼2000 Ma-old volcanic–plutonic components and no intermediate ages, consistent with a Gondwanan derivation (Krogh and Keppie, 1990). Acadian deformation took place in the Meguma terrane before and during the emplacement of granites (Benn et al., 1997), which are dated at 376–370 Ma (Schenk, 1997). The Meguma terrane, which is located at the SE border of Avalon, shows an Alleghenian (Hercynian) overprint as the Sehoul Block itself.
iii) Early Devonian reconstructions (e.g. Robardet, 2003) place the Meguma Zone on the southern margin of the rapidly closing Rheic Ocean. Therefore, the suture between the Sehoul Block and Western Meseta may derive from a minor (and still hypothetic) oceanic hiatus with respect to the classical Rheic suture supposed to extend between Avalon and Meguma (Cocks et al., 1997). Indeed, a complex network of seaways was seemingly created by splitting of the NW Gondwana margin during the Cambrian–Ordovician extension (Pin and Marini, 1993; Crowley et al., 2000; Faure et al., 2009).
In our proposal (Fig. 15B), we take into account the subdivisions of the Moroccan Meseta summarized inFig. 13, and suggest that they extended between two latitudinal transform zones, the SMF and the South Iberia Fault (SIF), respectively. The CIZ–ALZ system would not have equivalents south of the SIF, whereas the CZ would correspond
6. Conclusions and perspectives
The southern branch of the Peri-Atlantic Variscan Belt is widely exposed in Morocco, and a wealth of stratigraphic, petrologic and structural works are available for any tentative restoration of its Palaeozoic evolution.
Two contrasting segments constitute the Moroccan Variscides, i.e.
the Mauritanide and Meseta segments, separated by a transform zone, here above referred to as the South Meseta Fault (SMF). The SMF parallels the WAC northern boundary and the E-trending part of the Pan-African belt. This major lineament was frequently reactivated during the Mesozoic–Cenozoic along the partially inherited South Atlas Fault (SAF).
The Mauritanide crystalline thrusts extend in front of the central part of the WAC, namely the Reguibat Shield where the Palaeozoic cover is extremely thin and lacunar. Thin-skinned frontal units develop northward as soon as the Palaeozoic cover thickens (Dhlou–Zemmour arcuate belt). Further north, the westernmost Anti-Atlas forms a thick-skinned foreland fold–thrust belt derived from the proximal Gondwana paleomargin.
The Meseta Domain developed mostly on distal blocks of the Gondwana margin, from east to west, the Eastern Meseta and the Central Zone and Coastal Block of Western Meseta, separated from Gondwana and from each other by narrow zones of thinned crust.
Crustal extension occurred repeatedly during the Cambrian, Ordovi- cian, Devonian and Early Carboniferous. The varied sub-domains were shortened mainly during the Early (Eastern Meseta) and Late Carboniferous, and intruded by 330–290 Ma-old granites.
However, the Meseta Domain also includes an allochthonous Eo- Variscan or Acadian terrane, namely the Sehoul Block, includingca.
367 Ma-old granite. This block was possibly part of a larger Meguma– Sehoul terrane, located south of the main Rheic Ocean. The Moroccan and Iberian segments of the Variscan Belt were separated by a transform zone, broadly parallel to the SMF and here labelled the South Iberian Fault (SIF). This lineament would have been reactivated as a major, Permian dextral strike-slip fault, and later on as the Azores–Gibraltar transform zone. Part of the Iberian structural zones would not have equivalents south of the SIF (in particular, the Central Iberian Zone), whereas others would have suffered less extensional tectonics (in particular, the sutured oceanic domains on both sides of the Ossa- Morena Zone would be replaced on both sides of Eastern Meseta by intracontinental sutures derived from thinned crust zones). Consistent- ly, shortening of the Moroccan Meseta during the Variscan collision was much lesser than that of SW Iberia (compareFig. 15A and B).
We are conscious of the controversial character of our proposals.
We also recognize that some stratigraphic, structural or petrologic issues are still poorly explored. One of the most exciting at present is the occurrence of Cambrian breccias including fragments of sub- continental lherzolite in the Ounein area (Fig. 2), close to the Tizi n'Test Fault (Pouclet et al., 2007; Aarab, 2009). They could indicate that the Cambrian extension was strong enough to exhume the
