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Journal of African Earth Sciences 81 (2013) 60–81

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Journal of African Earth Sciences

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The Jurassic–Cretaceous basaltic magmatism of the Oued El-Abidsyncline (High Atlas, Morocco): Physical volcanology, geochemistryand geodynamic implications

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⇑ Corresponding author at: Geology Dept., Fac. of Sciences-Semlalia, Cadi AyyadUniversity, Prince Moulay Abdellah Boulevard, P.O. Box 2390, Marrakech, Morocco.

E-mail addresses: [email protected] (M.K. Bensalah), [email protected] (N. Youbi).

Mohamed Khalil Bensalah a,b,c, Nasrrddine Youbi a,b,c,d,⇑, João Mata b,c, José Madeira c,e, Línia Martins b,c,Hind El Hachimi a, Hervé Bertrand f, Andrea Marzoli g, Giuliano Bellieni g, Miguel Doblas h, Eric Font i,Fida Medina j, Abdelkader Mahmoudi k, El Hassane Beraâouz l, Rui Miranda b,c, Chrystèle Verati m,Angelo De Min n, Mohamed Ben Abbou o, Rachid Zayane a

a Geology Dept., Fac. of Sciences-Semlalia, Cadi Ayyad University, Prince Moulay Abdellah Boulevard, P.O. Box 2390, Marrakech, Moroccob Centro de Geologia da Universidade de Lisboa (CeGUL), FCUL, Campo Grande C6, 1749-016 Lisboa, Portugalc Universidade de Lisboa, Faculdade de Ciências, Departamento de Geologia (GeoFCUL), Campo Grande C6, 1749-016 Lisboa, Portugald National Centre for Scientific and Technical Research, Angle avenues des FAR et Allal El Fassi, Madinat Al Irfane, P.O. Box 8027, Nations Unies, 10102 Rabat, Moroccoe LATTEX, Instituto Dom Luiz – Laboratório Associado (IDL – LA), Lisboa, Portugalf Laboratoire de Géologie de Lyon, UMR-CNRS 5276, Université Lyon 1 et Ecole Normale Supérieure de Lyon, 46 Allée d’Italie, 69364 Lyon, Franceg Dipt. di Geoscienze, Univ. Padova, I-35137 Padova, Italyh Instituto de Geociencias (CSIC-UCM-Consejo Superior de Investigaciones Científicas-Universidad Complutense de Madrid), c/José Gutiérrez Abascal 2, 28006 Madrid, Spaini IDL-UL, Instituto Dom Luís, Universidade de Lisboa, Edifício C8, Campo Grande, 1749-016 Lisboa, Portugalj Laboratory GEOTEL (URAC 46), Scientific Institute, University Mohammed V-Agdal, Rabat, Moroccok Geology Dept., Fac. Sciences de Meknès, Moulay Ismail Univ., Meknès, Moroccol Geology Dept., Faculty of Sciences, Ibnou Zohr University, P.O. Box 28/S, Agadir, Moroccom UMR GéoAzur, 250 rue Albert Einstein, Bat.1, Sophia Antipolis, 06 560 Valbonne, Francen Dipartimento di Scienze della Terra (DST), Università degli Studi di Trieste, Via Weiss 8, 34127 Trieste, Italyo Geology Dept., Fac. Sciences Dhar Al Mahraz, Sidi Mohammed Ben Abdellah Univ., Fès, Morocco

a r t i c l e i n f o

Article history:Received 26 August 2011Received in revised form 17 December 2012Accepted 6 January 2013Available online 16 January 2013

Keywords:MoroccoHigh AtlasJurassic–CretaceousVolcanologyGeochemistryAlkaline-transitional magmatism

a b s t r a c t

Basaltic lava flows, dykes and sills, interbedded within red clastic continental sedimentary sequences (theso called ‘‘Couches Rouges’’) are widespread in the Oued El-Abid syncline. They represent the best can-didates to study the Jurassic–Cretaceous magmatism in the Moroccan High Atlas. The volcanic succes-sions were formed during two pulses of volcanic activity, represented by the Middle to Upper Jurassicbasaltic sequence B1 (1–4 eruptions) and the Lower Cretaceous basaltic sequence B2 (three eruptions).Whether belonging to the B1 or B2, the lava flows present morphology and internal structures typicalof inflated pahoehoe. Our geochemical data show that, at least for Jurassic magmatism, the dykes, andsills cannot be considered as strictly representing the feeders of the sampled lava flows. The Middle toUpper Jurassic pulse is moderately alkaline in character, while the Lower Cretaceous one is transitional.Crustal contamination plays a minor role in the petrogenesis of these magmas, which were generated byvariable partial melting degrees of a garnet-bearing mantle source. Magmatism location was controlledby pre-existing Hercynian fault systems reactivated during a Middle to Upper Jurassic–Cretaceous riftingevent. The associated lithospheric stretching induced melting, by adiabatic decompression, of enrichedlow-solidus infra-lithospheric domains.

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1. Introduction

Magmatism is generally active at different stages of continentalfragmentation. From the inception of continental rifting to the finaloceanic stage, when a volcanic passive continental margin devel-

ops, magma chemistry is a consequence of the dynamics of themantle, in particular of asthenosphere–lithosphere interactionsand of the possible role of mantle plumes (e.g., Saunders et al.,1992; Lassiter and DePaolo, 1997; Head and Coffin, 1997).

The Mesozoic was a period of unusually active magmatism dur-ing which ocean crust formation rate and off-ridge volcanism weregreater than at any time since (e.g., Larson, 1991; Marzoli et al.,1999; Phipps Morgan et al., 2004; Matton and Jébrak, 2009). Thefragmentation of Pangaea started in the Mesozoic with the

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M.K. Bensalah et al. / Journal of African Earth Sciences 81 (2013) 60–81 61

formation of intracontinental rifts, which later evolved until theopening of the Atlantic Ocean. Resulting from this continental frag-mentation, the eastern (African and European) Central Atlanticmargins were the locus of three geochemically distinct Mesozoicmagmatic cycles separated by time lags of �50 Ma (e.g., Martins,1991; Martins et al., 2010 and references therein). The first cycle,tholeiitic, occurred at around �200 Ma and is thought to be relatedwith the Pangaea break-up and the Large Igneous Province (LIP) ofthe Central Atlantic Magmatic Province (CAMP), as well as theassociated Triassic–Jurassic mass extinction (e.g., Marzoli et al.,1999, 2004, 2011; Knight et al., 2004; Vérati et al., 2007; Nomadeet al., 2002, 2007; Deenen et al., 2010; Font et al., 2011a,b). The vol-umetrically less important second cycle occurred close to theJurassic–Cretaceous boundary and affected essentially Morocco(e.g., Beraâouz and Bonin, 1993; Beraâouz et al., 1994; Armando,1999; Lhachmi et al., 2001; Zayane et al., 2002; Bensalah et al.,2006) and Iberia (e.g., Grange et al., 2008; Alves et al., 2010a,b).It is represented by numerous layered intrusions, basaltic lavasflows, dykes and sills with alkaline to transitional composition.Magmatic production increased during the last magmatic cycle,which took place essentially in Iberia at the end of the Cretaceous(e.g., Merle et al., 2006, 2009; Miranda et al., 2009; Grange et al.,2010). It comprised two pulses of alkaline magmatism: a firstone, between 94 and 88 Ma, related to the opening of the Bay ofBiscay and consequent rotation of Iberia, and a second one ataround 75–72 Ma, contemporaneous with the initial stages of theAlpine orogeny that led to the tectonic inversion of the Mesozoicbasins (Miranda et al., 2009). After these three Mesozoic cycles,carbonatites, lamprophyres, and related dykes were generated inMorocco during the Eocene (Bouabdli et al., 1988; Bernard-Grif-fiths et al., 1991; Wagner et al., 2003; Bouabdellah et al., 2010).

Moreover, Matton and Jébrak (2009) found that nearly half ofthe peri-Atlantic Mesozoic alkaline magmatic rocks fall withinthe 125–80 Ma interval with two major peaks: the first clusteringat about 125 Ma, and the second at about 85 Ma. The authors calledthis intense and widespread alkaline activity of the Atlantic realm‘‘the Peri-Atlantic Alkaline Province Pulse’’ (PAAP). Whether therocks studied here may be considered part of this widespread prov-ince (Fig. 1a) will be discussed.

While several Jurassic–Cretaceous layered intrusions from theHigh Atlas have been extensively studied (Hailwood and Mitchell,1971; Rolley, 1978; Martin et al., 1978; Westphal et al., 1979; Ber-aâouz and Bonin, 1993; Beraâouz et al., 1994; Armando, 1999;Lhachmi et al., 2001), a few is known on the lava flows, dykesand sills associated with the Jurassic–Cretaceous ‘‘red beds’’ Group(the so called ‘‘Couches Rouges’’) of the Oued El-Abid syncline inthe Central High-Atlas (Fig. 1b and c; Bensalah et al., 2006 and ref-erences therein). These intrusive and volcanic formations have analkaline or transitional intra-plate affinity (e.g., Beraâouz et al.,1994; Zayane et al., 2002; Bensalah et al., 2006). Based on biostrati-graphical, sedimentological, paleomagnetic and geochronologicaldata, Upper Jurassic to Cretaceous ages have been ascribed to theseformations (Hailwood and Mitchell, 1971; Bardon et al., 1973,1978; Rolley, 1978; Westphal et al., 1979; Jenny, 1985; Le Marrec,1985; Monbaron, 1988; Souhel and Canérot, 1989; Souhel, 1996;Haddoumi et al., 2002, 2010). These suggest a link with the PAAP,providing an excellent opportunity to contribute to the under-standing of the nature and origin of this peri-Atlantic province, gi-ven that volcanism generally reflects the evolution of theunderlying lithosphere and asthenosphere.

We will provide new and detailed volcanological, petrographic,mineralogical and geochemical data of the Jurassic–Cretaceousbasaltic lava flows and associated minor intrusions from the OuedEl-Abid syncline in the Central High-Atlas, by (i) describing thephysical volcanology of the volcanic successions; (ii) presentingtheir mineralogical characteristics and magmatic affinities; (iii)

discussing their petrogenesis; and (iv) proposing a geodynamicmodel for the Jurassic–Cretaceous magmatism of Morocco.

2. Geological setting

The magmatic formations studied here are located in the Cen-tral High Atlas of Morocco, belonging to the High Atlas intraconti-nental Cenozoic fold belt extending from western Morocco toTunisia (Fig. 1b). In the studied area a thick and deformed Meso-zoic–Cenozoic sedimentary cover, pervaded by abundant mag-matic rocks, rests unconformably on the Precambrian–Paleozoicbasement (Mattauer et al., 1977; El Harfi et al., 2006). It is boundedto the south by the South Atlas Fault system (separating the HighAtlas from the Anti-Atlas), and to the north by the NE-trendingMiddle Atlas.

The geometry of the Central High Atlas trough is marked by anumber of ENE–WSW narrow rift basins depicting a complex his-tory with four major tectonic phases: pre-, syn- and post-Mesozoicrift phases, and Cenozoic Alpine deformation (e.g., Baudon et al.,2009 and references therein). The pre-rift phase is characterizedby compressional tectonics associated to the Hercynian Orogenyand the assembly of Pangaea. Extensional tectonics and crustalthinning occurred during an early stage of Pangaean break-up inthe Late Palaeozoic. An early syn-rift stage produced NE-SW toENE–WSW normal faults with a horst/graben geometry, possiblyreactivated from previous compressional Hercynian structures(e.g. Piqué and Laville, 1996). The main syn-rift phase (UpperPermian to Late Triassic) was characterized by elongated NE–SWto ENE–WSW rift-basins resulting from the opening of the AtlanticOcean to the west and the Tethys Sea to the north (Manspeizeret al., 1978; Manspeizer, 1988; Medina, 1995; Ellouz et al., 2003;Zühlke et al., 2004). These basins were first filled by continentalsiliciclastic sediments (Beauchamp, 1988; El Arabi et al., 2003)and then affected by huge emissions of CAMP basaltic flows (e.g.,Youbi et al., 2003; Marzoli et al., 1999, 2004; Knight et al., 2004;Nomade et al., 2002, 2007; Vérati et al., 2007; Font et al.,2011a,b). Post-rift thermal relaxation (linked to the abortion ofthe Atlasic rift) and associated sea-level rise led to the depositionof Jurassic to Cretaceous carbonate platforms. Later on, duringthe Alpine orogeny, these Triassic and Jurassic extensional struc-tures underwent inversion tectonics with reverse faulting, uplift-ing and folding of the Mesozoic sequences (Mattis, 1977;Manspeizer et al., 1978; Monbaron, 1981; Ellouz et al., 2003; ElHarfi et al., 2006).

The targets of our study correspond to the Jurassic–Cretaceousmagmatic formations that are contemporaneous with the mainpost-rift phase and, as suggested above, part of the so-called ‘‘Cre-taceous Peri-Atlantic Alkaline Pulse’’ (PAAP, Matton and Jébrak,2009). This magmatism is represented by intrusions (dykes andsills) and lava flows emplaced in siliciclastic red beds and with agesranging from Middle–Late Jurassic to Early Cretaceous (Chèvre-mont, 1975, 1977; Smith and Pozzobon, 1979; Monbaron and Just,1980; Monbaron, 1980; Laville and Harmand, 1982; Beraâouz andBonin, 1993; Beraâouz et al., 1994; Armando, 1999; Lhachmi et al.,2001; Zayane et al., 2002; Bensalah et al., 2006).

Sedimentological studies (Souhel and Canérot, 1989; Souhel,1996; Haddoumi et al., 2002, 2010) show that the red beds corre-spond to alluvial sequences. The Guettioua and the Jbel Sidal For-mations are fluvial deposits, while the Iouaridène Formationrepresents an evaporitic playa lake. The biostratigraphy of thered bed formations as revealed by a series of vertebrate, ostracodand charophyte findings has been progressively improved duringthe last decade (e.g., Haddoumi et al., 2010 and references therein),allowing to constrain the stratigraphic position of the basaltic lavaflows B1 and B2 (Fig. 2). The lower unit (Guettioua Formation) isnow assigned to the Bathonian-? Callovian; the middle formation

Fig. 1. (a) Reconstruction of Africa–South America–North America, Greenland and Europe at time of the Peri-Atlantic Alkaline Pulse (PAAP) emplacement and schematicextent of the PAAP event (Matton and Jébrak, 2009); (b) localization of the High Atlas fold belts; and (c) schematic geological map of the Central High Atlas and location of thesynclines (Haddoumi et al., 2002). The synclines (Igoudlane, Karia, Aït Attab, Ouaouizaght, Aghzif–Naour, etc.) reported in this map are grouped under the name of the OuedEl-Abid syncline. The Roman numerals I, II, and III indicate the emplacement of sections (see also Fig. 4).

62 M.K. Bensalah et al. / Journal of African Earth Sciences 81 (2013) 60–81

(Iouaridene Fm.) is Bathonian–Callovian to Kimmeridgian at itsbase, and Lower Barremian at the summit; finally the upper unit(Jbel Sidal Formation) is Late Barremian. Consequently, B1 can bebiostratigraphically assigned to the latest Bathonian to Callovianor even Kimmeridgian, while B2 is earliest Late Barremian. More-over, K–Ar ages range from Dogger to Barremian, but seem to formtwo distinct groups, 175–155 ± 5 Ma and 135–110 ± 5 Ma, respec-tively, which are consistent with the biostratigraphic dating.

The Central Atlas Jurassic–Cretaceous magmatism is repre-sented by a large variety of magmatic bodies: (i) intrusive compos-ite massifs formed by basic (troctolites, gabbros), intermediate(diorites and monzodiorites) and acid (syenites) rocks and, locally,by complex mixings of felsic to intermediate magmas; (ii) NW–SEoriented metric-scale gabbro/doleritic dykes: they intrude the Tri-

assic–Jurassic sedimentary cover of the Atlas Chain and are partic-ularly abundant in the Tagleft region, where they form a uniquestructure surrounding the Aït-Boulmane gabbroic massif(Fig. 1c); (iii) doleritic/gabbro sills intruding mid-Jurassic continen-tal red beds, limestones (Fig. 3a), and sandstones, locally connectedto feeding dykes; (iv) B1/B2 lava flows interbedded within the redbeds (e.g., Haddoumi et al., 2010 and references therein) (Fig. 3b).These basaltic lava flow successions are well exposed in the Aït At-tab syncline where they are very useful for identifying the differentred bed formations (Fig. 3b).

Previous paleomagnetic investigations of the Jurassic intrusionsfrom the Central High Atlas in southern Morocco (Hailwood andMitchell, 1971) provided a preliminary paleomagnetic pole at61�S, 71�E (A95 = 14�) that perfectly agrees with the Jurassic–

Fig. 2. Synthetic stratigraphical column of the Mesozoic deposits of the Oued El-Abid syncline (from Haddoumi et al., 2002, 2010). The ‘‘red beds’’ Group comprisethe Guettioua, Iouaridène and Jbel Sidal Formations. Az., Azilal Formation;Tilougguit, Formation; A. Tf, Aït Tafelt Formation; A. At, Aït Attab Formation.


Cretaceous global apparent polar wander path (APWP) of Torsviket al. (2008). By contrast, paleomagnetic data for the lowerCretaceous volcanics showed shallower inclinations as indicatedby a preliminary paleomagnetic pole (44�N, 251�E, N = 11, K = 24,A95 = 10�) obtained from the basaltic flows and sills of the OuedEl-Abid syncline (Bardon et al., 1973, 1978). In the latter, the pres-ence of both normal and reverse polarities may suggest ages span-ning from the upper Jurassic to the lower Cretaceous, i.e. betweenthe M25 and M0 magnetic sea-floor anomalies (Ogg et al., 2008).Westphal et al. (1979) reported concordant paleomagnetic results(pole position at 45�N, 251�E, 13 sites, K = 26, A95 = 8�) and K–Ardata corroborating an upper Jurassic to lower Cretaceous age forthe basalts and dolerites of the Oued El-Abid syncline (Westphalet al., 1979). However, these paleomagnetic and K–Ar data arepresently under discussion because of the outdated analytical pro-cedure, the reduced number of samples, the lack of field tests andthe absence of magnetic mineralogical analyses. These latter dataare crucial to check for the primary (syn-cooling) versus secondary(remagnetized) origin of the magnetic remanence isolated in therock.

3. Volcanology of the lava flows from the ‘‘red beds’’ of the OuedEl-Abid syncline

The terminology for the morphological description of the basaltflows adopted in this study is that of Self et al. (1997, 1998) and

Thordarson and Self (1998), which is based on three criteria: (i)the vesiculation pattern, which is defined by distribution, mode,shape, and size of vesicles and other degassing features; (ii) thejointing style, which refers to the arrangement and morphologyof cooling joints and columns; and (iii) the petrographic texture,which refers to crystallinity and crystal size, properties that arecontrolled by a number of parameters, such as volatile content,cooling rate, and crystal nucleation rate. In vertical section, inflatedpahoehoe flows show three zones: a basal vesicular lava crust con-taining pipe vesicles, a central dense lava core with different segre-gation structures, and an upper lava crust, displaying alternatingvesicular and massive layers.

3.1. Middle to Upper Jurassic B1 lava flows

In the Aït Attab syncline, the B1 sequence (up to 10 m thick) isformed by, at least, two lava flows cropping out continuously onthe north and west flanks of the syncline (Figs. 3c and 4).

In the Aghzif–Naour syncline, the B1 episode crops out on theNE extremity of the structure, conformably overlying a sequenceof red clays and sandstones of the Guettioua Fm. It is materializedby a succession of eleven lava flows (Fig 4). Three sedimentary lev-els are interbedded in the volcanic succession. The first volcanic se-quence is composed of two subaerial lava flows covered by apillow-breccia horizon containing lithic fragments of dolomiteand compact red siltstone, which represents a peperitic event(see Skilling et al., 2002). Those volcanic units are covered by a6 m thick sequence of red gypsiferous siltstones presenting a pyro-clastic level in the base. The sediments are overlain by a lava flow,representing a second volcanic event. On top of it, the second sed-imentary unit is a 5 m thick sequence of red gypsiferous siltstones.These sediments are covered by a sequence of six subaerial lavaflows that represent a third eruptive event. This is separated fromthe next eruptive sequence by a discontinuous dolomite layer, afew cm thick. The upper volcanic sequence is composed of a pil-low-breccia horizon containing lithic fragments of dolomite andcompact red siltstone, representing a second peperitic event, cov-ered by two subaerial lava flows.

In the Ouaouizaght syncline, on top of the Guettioua Fm., twosuccessions of lava flows can be continuously followed along theNW flank. The outcropping conditions do not allow the determina-tion of the exact number of flows in each succession. Pillow-brec-cias containing lithic fragments of dolomite occur frequently.Dolomitic layers cover the second succession of lava flows.

In the Tagleft syncline, the B1 succession is probably formed bytwo (Jbel Sgat) and three (Igharghar) lava flows intercalated in theuppermost Guettioua Fm. The exact number of flows is uncertaindue to unfavourable outcropping conditions.

The disposition of the flows in most synclines suggests that thelavas flowed towards the north and west of the basins, whichagrees with the paleo-slopes deduced from the sedimentologicalstudy of the Guettioua and Iouaridène formations (Souhel andCanérot, 1989; Souhel, 1996; Löwner, 2009).

3.2. Lower Cretaceous B2 lava flows

The second basaltic succession (B2) is intercalated in the lowerlayers of the Jbel Sidal Fm. In the Aït Attab syncline, the B2 basaltsare represented by variable sequences. On the north and westflanks of the fold the succession is composed of two lava flows,5–6 m thick, often eroded and discontinuous (Fig. 3b). On the southflank of the syncline (Aït Rhodja region) a 60–70 m thick lava se-quence extends eastwards to Souk El-Had where it is reduced tohalf of the thickness, while to the west it is reduced to a singlepyroclastic layer (ash tuff) a few meters thick (Abrarag region).

Fig. 3. Some field characteristics of the Jurassic–Cretaceous basaltic lava flows sequences and associated intrusive bodies of the Central High Atlas. (a) Sill intruding theJurassic limestones. Photo taken near the town of Ouzoud; (b) panoramic view of the southern flank Aït Attab syncline showing the ‘‘red beds’’ Group (Iouaridene and JbelSidal Formations) with the B2 basaltic lava flows sequence; (c) detail of the B1 basaltic lava flows sequence of the northern flank Aït Attab syncline. Only two lobes can bedistinguished; (d) injections of siltstone (Neptunian dykes), randomly dispersed within flow lobe ‘‘b’’ of the west Aït Attab section. See Figs. 1 and 5.

Fig. 4. Lithostratigraphic columns across the Jurassic–Cretaceous volcanic successions of the Oued El-Abid synclinal zone. See Fig. 1b for location of sections I, II, and III.


Detailed sections made on the volcanic pile at Aït Rhodja show thatthe B2 episode is formed by a succession of 5 to 8 flows, dependingon the section (Fig. 4).

Pillow lavas occur at the base of the first lava flow and, towardsthe west, also at the base of the second lava flow, indicating thatflows entered shallow lakes. The basal part of these flows


developed pillowed structures, while the upper part remainedabove water level. The remaining lava flows are subaerial and pres-ent an internal structure formed by a basal vesicular lava crust, acentral dense lava core with segregation structures, and an upperlava crust. Discontinuous and thin (<1 m) sediment layers are int-erbedded in the volcanic succession separating the second, thirdand fourth lava flows. These sediments mark significant time inter-vals separating three eruptions and the emplacement of the corre-sponding sequences of lava flow-units. The cores of the four lowerflow-units are injected with brownish red baked siltstone formingirregular bodies (0.4–2 m long) randomly dispersed within the lava(Fig. 3d). These structures can also be found at the lava core-uppercrust interface and in the upper crust, but their occurrence be-comes less abundant. These are interpreted as peperitic injectionsresulting from magma/sediment interaction (see Skilling et al.,2002). These occurrences indicate that, at the time of the lava flowemplacement, the basal sediments were still soft or only veryslightly consolidated and prove that the extrusions are contempo-raneous of the Iouaridène Fm. Similar structures were observed inpahoehoe lava flows of CAMP sections both in the High Atlas andsouth Portugal (e.g. Martins et al., 2008; El Hachimi et al., 2011).Analogous structures were described in the Western Deccan Volca-nic Province by Duraiswami et al. (2003). These authors suggestthat the fine grained sediments injected into the lava flow mayhave accumulated in rain-fed puddles developed in shallowdepressions on large sheet lobes and hummocky flow fields. Subse-quent emplacement of sheet lobes over these puddles producedslender, steam injection structures within the flow. Waichel et al.(2008) observed sand-lava interaction (sand diapirs and peperite-like breccia) produced by lava flowing over active aeolian dunesin the continental flood basalts of Paraná.

On the Ouaouizaght syncline, the B2 basalts form a discontinu-ous horizon, less than 4 m thick, on the south-east.

3.3. Volcanological summary

The volcanic pile of the Oued El-Abid syncline was formed dur-ing two pulses of volcanic activity, represented by the Middle toUpper Jurassic B1 and the Lower Cretaceous B2 events. On thestudied sections (Fig. 4) there is evidence for a variable numberof eruptions in each event. The products of distinct eruptions canbe separated by the deposition of fine clastic sediments, represent-ing time intervals separating the emplacement of each package oflava flow-units. The B1 volcanic pulse was built up by 1–4 erup-tions (Aït Attab and Aghzif Naour synclines, respectively), eachproducing flow fields composed of 1–6 flow units. The B2 volcanicpulse produced three eruptions composed of 1–5 pahoehoe flows.Both the B1 or B2 lava flows present internal structures typical ofinflated pahoehoe (Self et al., 1997), namely the absence of basaland top clinker, vertical internal structure characterized by threesectors differing in vesiculation, jointing and texture (lower crust,lava core et upper crust), and presence of ‘‘pipe vesicles’’. The factthat nearby sections present the same number of flows (8 in thecase of the Aït Attab syncline) may indicate that these are simplepahoehoe flows. The occurrence of pillow breccias mixed withangular fragments of sedimentary rocks (dolomite and/or siltite)indicates that some flows penetrated lakes and interacted withsemi-consolidated sediments.

4. Analytical procedures and sampling

Sampling was done on the best exposed and freshest outcropsof lava flows, sills and dykes occurring from the Aghzif–Naour, Tag-left, Aït Attab and Ouzoud areas, in order to obtain a set of samplesrepresentative from the studied formations. Petrographic examina-

tion was done to eliminate highly porphyritic samples (cumulates)or those characterized by a large proportion of secondary minerals.A total of thirty-one samples were chosen for whole-rock majorand trace-element analyses and for electron-microprobe study ofthe mineral chemistry.

Dykes associated with the Middle to Upper Jurassic B1 lavaflows (Aghzif-Naour/Tagleft area) and sills associated with theLower Cretaceous B2 lava flows (Ait Attab/Ouzoud area) will behereafter referred as B1 dykes and B2 sills, respectively.

4.1. Whole-rock analyses

The majority of samples were analyzed for both major- andtrace-elements at the Activation Laboratories Ltd., Ancaster, Ontar-io, Canada. For whole rock analysis, major oxide contents were ob-tained using Inductively Coupled Plasma-Optical EmissionSpectrometry (ICP-OES). Trace element contents were obtainedusing Inductively Coupled Plasma-Mass Spectrometry (ICP-MS),with the exception of Ba and Sr, which were analyzed by ICP-OES. Samples were mixed with a flux of lithium metaborate andlithium tetraborate and fused in an induction furnace. The moltenmelt was immediately poured into a solution of 5% nitric acid con-taining an internal standard and mixed continuously until com-pletely dissolved (�30 min). Samples were prepared andanalyzed in a batch system. Each batch contained a method re-agent blank, certified reference material and 17% replicates. Cali-bration was performed using seven prepared USGS and CANMETcertified reference materials. Duplicate measurements give an esti-mate of the total reproducibility of our analyses. The reproducibil-ity is better than: (i) 1 relative-% for major element contents (SiO2,Fe2O3, MnO, MgO and CaO); (ii) 3 relative-% for Rare Earth Ele-ments; and (iii) 2 relative-% for elements generally highly incom-patible in an oceanic context (Rb, Ba, U, Th). The accuracy of theanalyses, evaluated by analyzing international standards, is gener-ally better than 12 relative-%, with results for many elements with-in ±6% of the recommended values.

Some samples (labeled with asterix on Table 1) were analyzedat the Laboratoire de Géologie de Lyon, University Lyon 1 (France).Major and some trace elements (Sc, V, Cr, Co, Ni, Rb, Sr, Ba, Zr, Nband Y) have been analyzed using a X-ray fluorescence spectrome-ter (Phillips PW 1404). The precision is 1–2% for major elementsand 10–15% for trace elements. REE, Th, U, Hf and Ta were analyzedat the Ecole Normale Supérieure of Lyon using inductively coupledplasma mass spectrometry (ICP-MS). 200 mg of rock powder weredissolved in a mixture of 3 ml HF and 1 ml HNO3 during 48 h on ahot plate (130 �C) under a 50 bars pressure. The solutions weredried up and the residues dissolved in 25 ml HNO3 (0.5 N). Thesesolutions were diluted to 0.1 ml and analyzed with a VG Elementplasma quadripole II ICP-MS with electron multiplier. Proceduralblank analyses yielded trace elements below detection limit(<10 ppb). The standard used for all analyses was BHVO.

A comparative test conducted by analyzing the same samplewithin the different laboratories used, showed no significant inter-laboratory bias.

The major and trace element compositions of the studied sam-ples with loss on ignition (LOI) <6 wt% are listed in Table 1. Thecomplete data set of whole rocks chemical analyses is presentedas Table I in the electronic supplementary material accompanyingthis manuscript.

4.2. Electron micro-probe analyses

Electron-microprobe (EMP) analyses of olivine, pyroxene, andfeldspar were carried out with a WDS-CAMECA-CAMEBAX 799housed at Consiglio Nazionale delle Ricerche (C.N.R.) – Istituto diGeoscienze e Georisorse (I.G.G.), Padova, Italy. Operating condi-

Table 1Whole-rock geochemical analyses of the Oued El-Abid synclinal zone igneous rocks. Only samples with LOI < 6 wt% are reported. Magnesium number (Mg#) calculated as 100 �Mg2+/(Mg2+ + Fe2+), with FeO = 0.9 Fe2Ot

3. Fe2O3t = total Feexpressed as Fe2O3. Loss-on-ignition (L.O.I.) determined by weight loss after igniting sample at 1000 �C.

SampleNumber

NA-2a� NA-2b NB 3-1� NB3-2� BBKH9-4 BBKH16-1 BBKH16-2 BBKH15 TAG-2a� TAG-2b ABKH5-4 ABKH5-6� ABKH7� AT-1 AT-3 AT-5 OZ-1 OZ-2 OZ-3 OZ-5 ABKH 1 ABKH2-1� ABKH2-4

Locality Aghzif-Naour Tagleft Aît Attab OuzoudEmplacementmode

Lava flow Dyke Lava flow Sill

Group B1 B1 B2 B2

Major elements (wt%)SiO2 46.29 46.62 46.03 45.9 46.15 48.46 48.28 48.72 46.27 46.8 51.25 50.06 48.74 51.37 50.59 51.24 50.63 49.5 50.36 51.27 50.51 50.01 49.08TiO2 1.88 1.869 1.96 1.91 1.49 1.7 1.5 1.6 1.51 1.53 1.67 1.46 1.75 1.66 1.62 1.78 1.78 1.77 1.96 1.8 1.79 1.75 1.77Al2O3 14.88 14.91 14.91 14.94 15.84 16.65 16.39 15.03 14.17 14.44 15.57 14.99 13.68 15.46 15.4 15.63 15.08 15.09 15.57 15.09 14.6 14.54 14.86Fe2O3t 11.87 10.96 12.56 12.81 11.57 10.25 10.64 11.84 12.19 12.34 9.8 11.69 11.16 10.08 10.19 14.46 12.03 11.76 9.63 11.77 12.17 11.69 12.36MnO 0.18 0.195 0.21 0.17 0.13 0.1 0.12 0.17 0.11 0.11 0.11 0.15 0.15 0.11 0.12 0.05 0.16 0.14 0.11 0.16 0.15 0.16 0.15MgO 7.53 7.18 7.22 6.71 7.62 6.41 8.18 7.49 7.52 7.42 5 7.33 8.92 5.07 6.7 3.1 7.03 7.73 4.3 6.83 6.81 7.7 6.8CaO 7.18 6.97 7.62 7.95 6.36 6.43 2.38 8.14 9.03 8.65 7.73 8.11 7.58 7.75 7.18 4.99 7.89 8.06 7.15 8.06 7.95 7.73 7.59Na2O 3.34 3.29 3.01 3.11 3.56 5.34 5.42 3.38 2.41 2.53 3.43 3.3 2.95 3.29 3.55 3.39 3.7 3.48 5.04 3.58 3.68 3.79 3.6K2O 1.22 1.09 1.38 1.26 0.34 0.1 0.58 0.41 0.41 0.25 1.01 1.04 0.93 0.94 1.17 0.79 1.07 0.59 1.2 1.24 1.1 1.3 1.22P2O5 0.36 0.29 0.39 0.35 0.25 0.21 0.2 0.2 0.19 0.21 0.27 0.19 0.26 0.25 0.25 0.24 0.29 0.28 0.3 0.29 0.29 0.29 0.28L.O.I. 3.91 5.33 3.16 2.96 5.29 3.37 5.19 2.52 4.89 5.73 2.78 1.23 2.8 2.68 3.31 4.79 1.35 2.8 4.41 0.74 1.26 0.89 0.94H2O- 1.33 1.1 1.51 0.75 0.47 1.07 0.36

Total 99.97 98.7 99.55 99.58 98.6 99.02 98.88 99.5 99.45 100.01 98.62 100.02 99.99 98.66 100.08 100.46 101.01 101.2 100.03 100.83 100.31 100.21 98.65Mg# 55.66 56.46 53.22 50.90 56.59 55.31 60.34 55.60 54.97 54.34 50.24 55.38 61.27 49.89 56.55 29.79 53.63 56.54 46.91 53.46 52.55 56.59 52.13

Trace elements (ppm) 0.7Sc 14 17 15 17 36 33 18 18 18 21 19 18 18 16 17 20 18 19 19 19 19 18 13V 161 201 163 169 198 191 190 195 171 191 163 152 151 167 162 216 198 200 206 201 205 150 20Cr 212 160 209 211 229 152 130 200 245 180 240 203 290 140 160 260 160 170 110 160 250 226 186Co 46 31 46 48 41 31 27 35 47 33 39 45 49 27 34 37 34 32 22 32 48 43 290Ni 182 179 174 186 77 49 85 162 208 220 94 174 194 83 149 107 131 126 47 116 119 131 48Rb 16.5 17 19.2 17.1 7 5 8 5 3.1 3 8 13.7 7.2 8 12 14 17 15 17 19 20 23 1.96Sr 398 401 431 415 290 204 406 340 261 265 362 251 274 364 304 293 385 375 267 425 393 358 0.73Y 18.9 16.8 19.7 19.7 21 22 16.8 18.4 19 16.8 20.8 19.8 21.9 18.7 18.6 21.5 17 20.7 21.7 22.1 22.7 23.4 381Zr 122 113 135 126 106 80 88 82 93 86 106 89 111 102 103 116 115 116 142 114 118 123 2Nb 18.5 18.7 20.2 17.7 16 12 10.3 10.7 9.8 8.7 12.4 10.3 14.7 16.3 16.7 14.3 18.2 18 24.2 19.3 18.6 17 0.9Ta 1.15 1.34 1.26 1.1 1 0.75 0.75 0.73 0.61 0.63 0.89 0.64 0.91 1.08 1.2 1.02 1.26 1.28 1.68 1.27 1.27 1.06 66Ba 164 170 147 146 141 89 109 99 83 100 212 113 166 225 200 170 225 208 333 242 231 228 14La 17.26 16.5 20.63 11 11 10.4 9.97 11.25 9.87 13.8 10.15 14.74 14.3 13.1 14 13.9 13.9 17.4 14.3 13.8 14.42 27.7Ce 36.26 33.2 43.37 45 31 21.7 21.4 24.67 21.1 28.4 21.83 31.5 29.1 26.5 28 28.3 28.4 35.3 29.5 28.7 30.78 3.61Pr 4.36 4.09 5.1 2.69 2.79 3.11 2.71 3.58 2.73 3.82 3.61 3.28 3.55 3.53 3.57 4.34 3.68 3.66 3.77 15.3Nd 18.87 17 22.08 21 11 11.7 12.5 14.27 11.8 15.1 12.42 16.83 15.3 13.9 14.8 15.2 15.5 18 15.8 15.4 16.96 4.25Sm 4.76 4.49 5.37 3.29 3.72 3.95 3.37 4.09 3.62 4.47 4.34 3.99 4.21 4.33 4.39 4.91 4.46 4.3 4.61 1.48Eu 1.48 1.56 1.67 1.18 1.35 1.3 1.22 1.41 1.18 1.34 1.49 1.39 1.52 1.52 1.53 1.68 1.55 1.53 1.39 4.73Gd 4.4 4.75 4.7 3.75 4.26 4.06 3.86 4.62 3.89 4.37 4.79 4.24 4.86 4.76 4.83 5.26 5.05 5.02 4.63 0.78Tb 0.69 0.72 0.75 0.64 0.69 0.66 0.64 0.74 0.65 0.72 0.77 0.69 0.82 0.79 0.81 0.86 0.82 0.79 0.75 4.21Dy 4 3.67 4.29 3.49 3.74 4.06 3.55 3.9 4.05 4.33 4.02 3.77 4.26 4.27 4.36 4.64 4.47 4.36 4.58 0.78Ho 0.7 0.64 0.74 0.64 0.67 0.74 0.64 0.71 0.74 0.79 0.72 0.68 0.79 0.78 0.77 0.8 0.8 0.79 0.83 2.09Er 1.85 1.68 1.95 1.79 1.84 2.09 1.75 1.87 2.07 2.19 1.91 1.77 2.1 2.05 2.07 2.17 2.16 2.13 2.29 0.28Tm 0.24 0.23 0.25 0.25 0.26 0.27 0.25 0.25 0.27 0.28 0.26 0.24 0.29 0.29 0.28 0.31 0.3 0.28 0.3 1.7Yb 1.47 1.39 1.54 1.57 1.61 1.76 1.57 1.51 1.7 1.8 1.6 1.49 1.81 1.77 1.78 1.91 1.84 1.78 1.89 0.24Lu 0.21 0.19 0.21 0.22 0.22 0.25 0.21 0.21 0.24 0.24 0.22 0.2 0.24 0.24 0.24 0.26 0.25 0.25 0.26 0.02Hf 3.23 2.8 3.55 4 1 2.3 2.3 2.66 2.3 2.2 2.52 3.19 2.9 2.7 2.5 2.9 3 3.7 3 2.9 3.21 23.5Th 2.37 2.23 2.81 5 4 1.55 1.4 1.67 1.5 2.23 1.78 2.55 2.68 2.32 2.1 2.22 2.25 3.06 2.44 2.14 2.41 3U 0.73 0.67 0.95 2 3 0.44 0.72 0.52 0.47 0.73 0.48 0.71 0.71 0.64 0.78 0.64 0.67 1.03 0.71 0.71 0.73 23

Trace elements ratios(La/Yb)N 8.42 8.51 9.6 4.75 4.44 4.58 4.5 6.55 4.28 5.87 6.41 6.3 5.54 5.63 5.6 6.53 5.57 5.56 5.47 82.78La/Yb 11.67 11.87 13.37 6.62 6.19 6.38 6.28 9.13 5.94 8.18 8.93 8.79 7.73 7.85 7.8 9.10 7.77 7.75 7.63 8.00Y/Nb 1.02 0.89 0.97 1.11 1.31 1.83 1.63 1.71 1.93 1.93 1.67 1.92 1.48 1.147 1.11 1.5 0.93 1.15 0.89 1.14 1.22 1.37 388.77

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Ti/V

69.9

155

.74

72.0

467

.47

45.1

153

.35

47.5

849

.18

52.8

748

.08

61.6

757

.54

69.2

459

.59

60.0

249

.45

54.0

453

.26

57.1

553

.89

52.5

569

.61

57.1

4Zr

/Y6.

476.

726.

876.

435.

043.

635.

234.

454.

95.

115.

094.

535.

075.

455.

535.

396.

765.

66.

545.

155.

195.

260.

006

Ti/Y

b76

20.6

680

60.9

076

13.6

457

58.2

559

57.7

651

36.4

158

49.8

966

58.0

251

21.9

658

21.1

6219

.81

6526

.10

5902

.25

6045

.859

84.8

961

64.4

958

87.4

860

52.2

555

49.9

562

52.4

3K

/P6.

447.

146.

730

6.84

2.58

0.90

5.5

3.89

4.10

2.26

7.11

10.4

16.

87.

158.

906.

267.

014.

07.

608.

137.

218.

528.

28Tb

/Yb

0.46

0.51

0.48

0.4

0.42

0.37

0.4

0.49

0.38

0.4

0.48

0.46

0.45

0.44

0.45

0.45

0.44

0.44

0.39

17.5

4D

y/Y

b2.

72.

642.

772.

222.

322.

32.

262.

582.

372.

42.

512.

532.

352.

412.

42.

422.

422.

442.

422.

47La

/Ce

0.47

0.49

0.47

0.47

0.46

0.45

0.46

0.48

0.46

0.46

0.49

0.49

0.5

0.49

0.48

0.49

0.48

0.48

0.46

0.49

Ce/

Nd

1.92

1.95

1.96

1.85

1.71

1.72

1.78

1.88

1.75

1.87

1.9

1.9

1.89

1.86

1.83

1.96

1.86

1.86

1.81

1.81

Nd/

Sm3.

963.

784.

113.

553.

363.

613.

53.

693.

423.

763.

523.

483.

513.

513.

533.

663.

543.

583.

673.

6La

/Ta

14.9

312

.31

16.3

411

.00

14.6

613

.86

13.6

518

.37

15.6

615

.50

15.7

716

.05

13.2

410

.91

13.7

211

.03

10.8

510

.35

11.2

510

.86

13.5

713

.87

La/N

b0.

930.

881.

020.

680.

911.

000.

931.

141.

131.

110.

981.

000.

870.

780.

970.

760.

770.

710.

740.

740.

840.

94


tions included an accelerating voltage of 15 kV, a sample current of15 nA, and minimum beam diameter (1 lm). Analytical accuracyand precision estimated from repeated analyses on natural stan-dards are within ±2% for major and ±5% for minor elements.

Representative analyses of the main primary rock-forming min-erals are reported in Tables 2–4. The complete data set of mineralanalyses is presented as Tables II–IV in the electronic supplemen-tary material accompanying this manuscript.

5. Results

5.1. Petrography and mineral chemistry

The textures of the studied rocks are variable, depending on thetype of structure (dyke/sill or lava flow). The central part of dykesand sills present granular facies, sometimes with porphyroid tex-tures where poikilitic crystals are set in an ophitic matrix. At themargins of intrusions, textures are mainly microgranular to hya-lo-microlitic porphyritic. Lava flows present aphanitic to porphyrictextures (Fig. 5). The usual primary mineralogical paragenesis iscomposed by olivine, clinopyroxene, plagioclase and, as accessoryminerals, opaque oxides and apatite. In some rocks secondary min-erals such as serpentine, chlorite, calcite, uralite, quartz and seri-cite replace to various extent the primary minerals.

5.1.1. OlivineIt is omnipresent in the studied rocks as phenocryst, suggesting

that was one of the first minerals to crystallize and that it played adominant role on the fractionation processes. This mineral is lessaltered in dykes and sills than in lava flows where pristine olivinegrains are rare and some grains are totally replaced by serpentine.

Phenocrysts and micro-phenocrysts are usually idiomorphic(polyhedral or granular olivines; Donaldson, 1976) as is typicalfrom slow crystallizing olivines, but some of them show corrosiongulfs evidencing partial resorption by magma.

The composition of the olivines ranges from Fo54 to Fo81 (Ta-ble 2). Their high CaO contents (>0.18 wt%) reflects low pressuresof crystallization, typical of crustal conditions (e.g. Köhler and Brey,1990; Hirschmann and Ghiorso, 1994).

One sample (BBKH15) yieds olivine compositions which areclose to equilibrium with the host whole-rock. For these olivinesa crystallization temperature of ca. 1220 �C was calculated, afterPutirka (2008), considering a pressure of 2–8 kbar (compatiblewith P–T estimates calculated for the clinopyroxenes, seebelow).

5.1.2. ClinopyroxeneClinopyroxene is a ubiquitous mineral in the studied rocks

occurring either as a matrix component or as phenocrysts. In somedykes and sills, it can present poecilitic tendency. In the more al-tered rocks clinopyroxene is partially replaced by calcite, epidote,oxides and chlorite.

The compositions are different in Jurassic (B1) compared to Cre-taceous (B2) rocks. The B1 lava flows are dominated by the pres-ence of diopside (Wo43.18–47.93En32.76–40.13 Fs14.12–23.31), while theB2 lava flows and sills have only augite (Wo37.33–44.51

En38.76–46.26Fs11.85–20.06). The B1 dykes, usually considered coevalfrom B1 lava flows, possess diopside and augite (see Table 3 andFig. 6a). Moreover for a given MgO concentration, clinopyroxenesfrom the B1 lava flows and dykes are characterized by lower SiO2

and higher TiO2 (1.15% 6 TiO2 6 3.06%; average = 2.28%) contentsthan those from the B2 lava flows and sills (0.70% 6 TiO2 6 1.87%;average = 1.27%) (Fig. 6b). This is in agreement with the more SiO2-undersaturated chemistry of Jurassic alkaline magmas as

Table 2Representative analyses of olivine from the Oued El-Abid synclinal zone igneous rocks.

Sample BBKH15 ABKH7 ABKH2-1

Crystal size Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst PhenocrystLocation in crystal Core Core Rim Core Core Core Rim Rim Core Core Rim RimNumber analysis 1 2 3 6 12 13 14 15 65 66 67 68

Locality Tagleft Aît Attab OuzoudEmplacement mode Dyke Lava flow Sill

Group B1 B2 B2

SiO2 38.76 38.69 38.84 39.15 38.35 39.52 38.92 32.94 35.13 35.96 36.23 36.13TiO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.05 0.03 0.09 0.03Al2O3 0.04 0.06 0.05 0.00 0.06 0.04 0.05 16.58 0.01 0.01 0.00 0.57FeO 18.95 19.22 19.09 18.76 20.15 17.78 24.26 18.80 37.85 36.02 32.90 32.76MnO 0.15 0.21 0.21 0.28 0.18 0.19 0.22 0.24 0.58 0.54 0.43 0.49MgO 42.52 42.99 42.68 40.39 40.89 43.22 37.90 32.46 25.34 27.16 29.34 28.98CaO 0.23 0.19 0.20 0.24 0.21 0.26 0.21 0.22 0.30 0.24 0.26 0.27Na2O 0.02 0.01 0.00 0.00 0.04 0.01 0.01 0.05 0.01 0.01 0.05 0.00Cr2O3 0.04 0.05 0.06 0.07 0.07 0.04 0.00 0.03 0.02 0.00 0.02 0.00

Total 100.70 101.43 101.14 98.91 99.95 101.06 101.55 101.32 99.30 99.99 99.33 99.24

Si 0.986 0.979 0.984 1.012 0.989 0.995 1.004 0.825 0.999 1.003 1.002 0.999Ti 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.001 0.001 0.002 0.001Al 0.001 0.002 0.002 0.000 0.002 0.001 0.001 0.489 0.000 0.000 0.000 0.019Fe2+ 0.403 0.407 0.405 0.405 0.434 0.374 0.523 0.394 0.900 0.841 0.761 0.758Mn 0.003 0.005 0.004 0.006 0.004 0.004 0.005 0.005 0.014 0.013 0.010 0.011Mg 1.613 1.621 1.613 1.556 1.572 1.622 1.457 1.211 1.074 1.130 1.210 1.195Ca 0.006 0.005 0.005 0.007 0.006 0.007 0.006 0.006 0.009 0.007 0.008 0.008Na 0.001 0.000 0.000 0.000 0.002 0.001 0.000 0.003 0.001 0.001 0.003 0.000Cr 0.001 0.001 0.001 0.001 0.001 0.001 0.000 0.001 0.001 0.000 0.001 0.000

Total 3.014 3.020 3.014 2.988 3.010 3.005 2.996 2.932 3.000 2.996 2.997 2.991% Fo 80.00 79.95 79.94 79.33 78.34 81.25 73.58 75.47 54.41 57.34 61.38 61.19

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Table 3Representative analyses of clinopyroxenes from the Oued El-Abid synclinal zone igneous rocks. Total Fe+Mn were considered for the calculation of the FS component.

Sample NA2a BBKH15 ABKH7 ABKH2-1

Crystal size Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst Phenocryst PhenocrystLocation in crystal Core Core Core Core Core Core Core Rim Rim Rim Core CoreNumber analysis 1 2 3 34 35 36 59 60 61 119 120 121

Locality Aghzif-Naour Tagleft Aît Attab OuzoudEmplacement mode Lava flow Dyke Lava flow Sill

Group B1 B1 B2 B2

SiO2 48.97 49.33 49.21 47.42 47.97 47.64 52.01 49.61 50.68 51.26 51.74 51.37TiO2 2.24 2.20 2.22 2.97 2.48 2.48 0.97 1.30 1.33 1.28 0.89 1.05Al2O3 3.63 3.44 3.28 4.05 3.89 3.45 1.36 3.20 2.36 2.12 1.66 1.32Cr2O3 0.01 0.11 0.00 0.04 0.00 0.06 0.11 0.09 0.22 0.10 0.30 0.01FeO 10.41 10.06 9.67 13.14 12.56 12.01 7.51 7.79 7.98 9.22 8.00 10.46MnO 0.17 0.14 0.24 0.29 0.16 0.18 0.20 0.14 0.16 0.13 0.25 0.33MgO 13.55 13.69 13.72 12.29 12.37 12.01 16.66 15.61 15.64 14.80 15.31 14.32CaO 22.04 22.23 22.38 20.42 20.33 21.28 21.47 21.01 21.15 20.66 20.84 20.09Na2O 0.46 0.44 0.49 0.57 0.49 0.52 0.34 0.39 0.36 0.37 0.39 0.30

Total 101.49 101.65 101.21 101.20 100.25 99.62 100.64 99.14 99.89 99.95 99.36 99.26

FeO 7.72 7.73 7.43 10.59 9.58 9.68 6.65 4.80 6.42 7.75 6.64 9.35Fe2O3 2.37 2.01 1.95 1.82 2.38 1.73 0.96 3.33 1.74 1.14 1.14 0.59

Si 1.813 1.822 1.822 1.785 1.818 1.817 1.902 1.844 1.875 1.913 1.932 1.943Ti 0.062 0.061 0.062 0.084 0.071 0.071 0.027 0.036 0.037 0.036 0.025 0.030Al 0.158 0.150 0.143 0.180 0.174 0.155 0.059 0.140 0.103 0.093 0.073 0.059Cr 0.000 0.003 0.000 0.001 0.000 0.002 0.003 0.003 0.006 0.003 0.009 0.000Fe3+ 0.066 0.056 0.054 0.052 0.068 0.050 0.026 0.093 0.048 0.032 0.032 0.017Fe2+ 0.239 0.239 0.230 0.333 0.304 0.309 0.203 0.149 0.198 0.242 0.207 0.296Mn 0.005 0.004 0.007 0.009 0.005 0.006 0.006 0.004 0.005 0.004 0.008 0.010Mg 0.748 0.754 0.758 0.690 0.699 0.683 0.908 0.865 0.863 0.823 0.852 0.807Ca 0.874 0.880 0.888 0.824 0.826 0.870 0.841 0.837 0.838 0.826 0.834 0.814Na 0.033 0.032 0.035 0.041 0.036 0.038 0.024 0.028 0.026 0.027 0.028 0.022

Total 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000

AlIV 0.159 0.150 0.144 0.181 0.174 0.156 0.059 0.141 0.103 0.087 0.068 0.057AlVI 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.007 0.005 0.002

Wo 45.24 45.53 45.84 43.18 43.43 45.37 42.37 42.95 42.92 42.86 43.13 41.87En 38.70 39.00 39.10 36.16 36.76 35.63 45.75 44.39 44.17 42.71 44.07 41.52Fs 16.06 15.47 15.06 20.66 19.81 19.00 11.88 12.66 12.91 14.43 12.79 16.61

Mg# 75.79 75.94 76.70 67.42 69.70 68.86 81.71 85.29 81.30 77.28 80.43 73.19

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Table 4Representative analyses of feldspars from the Oued El-Abid synclinal zone igneous rocks.

Sample NA2a BBKH15 ABKH7 ABKH2-1

Crystal size Phenocryst Phenocryst Phenocryst Microphenocryst Phenocryst Microphenocryst Microphenocryst Microphenocryst Microphenocryst Microphenocryst Phenocryst PhenocrystLocation incrystal

Core Core Core Core Core Core Core Rim Rim Core Core Rim

Number analysis 1 2 3 18 19 20 27 28 29 45 46 47Locality Aghzif-

NaourAghzif-Naour

Aghzif-Naour

Tagleft Tagleft Tagleft Aît Attab Aît Attab Aît Attab Ouzoud Ouzoud Ouzoud

Emplacementmode

Lava flow Dyke Lava flow Sill

Group B1 B1 B2 B2

SiO2 53.26 64.95 62.50 57.27 51.62 52.49 57.72 65.56 62.47 54.88 51.87 57.60TiO2 0.10 0.21 0.12 0.09 0.07 0.11 0.17 0.11 0.14 0.09 0.09 0.10Al2O3 29.20 20.35 23.30 27.25 29.45 29.09 25.46 20.35 22.31 27.82 29.42 25.11FeO 0.71 0.45 0.44 0.72 0.56 0.57 0.61 0.61 0.48 0.67 0.70 0.69MnO 0.02 0.03 0.01 0.04 0.01 0.02 0.01 0.01 0.00 0.01 0.02 0.00MgO 0.07 0.26 0.00 0.09 0.15 0.11 0.19 0.01 0.05 0.06 0.04 0.03CaO 12.53 1.69 4.79 9.12 13.40 12.73 8.54 1.98 4.51 10.98 12.95 7.83Na2O 4.30 7.28 7.85 6.13 4.01 4.36 6.14 8.46 7.97 5.25 4.19 6.64K2O 0.30 4.62 1.44 0.32 0.15 0.13 0.52 2.66 1.36 0.36 0.25 0.65

Total 100.47 99.84 100.45 101.04 99.42 99.62 99.36 99.76 99.29 100.12 99.53 98.64

Si 2.410 2.906 2.772 2.552 2.367 2.397 2.610 2.917 2.800 2.483 2.376 2.624Ti 0.003 0.007 0.004 0.003 0.003 0.004 0.006 0.004 0.005 0.003 0.003 0.003Al 1.558 1.073 1.218 1.431 1.592 1.566 1.357 1.067 1.179 1.484 1.588 1.348Fe2+ 0.027 0.017 0.016 0.027 0.021 0.022 0.023 0.023 0.018 0.025 0.027 0.026Mn 0.001 0.001 0.000 0.002 0.000 0.001 0.000 0.000 0.000 0.000 0.001 0.000Mg 0.004 0.017 0.000 0.006 0.010 0.007 0.013 0.001 0.003 0.004 0.003 0.002Ca 0.608 0.081 0.228 0.436 0.658 0.623 0.414 0.094 0.216 0.532 0.636 0.382Na 0.377 0.632 0.675 0.529 0.356 0.386 0.538 0.730 0.693 0.460 0.373 0.586K 0.017 0.264 0.082 0.018 0.009 0.008 0.030 0.151 0.078 0.021 0.015 0.038

Total 5.005 4.998 4.994 5.004 5.017 5.013 4.991 4.987 4.991 5.013 5.020 5.010

%An 60.65 8.29 23.13 44.29 64.32 61.29 42.12 9.68 21.93 52.51 62.15 37.96%Ab 37.64 64.70 68.58 53.85 34.81 37.95 54.80 74.84 70.20 45.42 36.43 58.25%Or 1.70 27.01 8.29 1.86 0.86 0.76 3.08 15.49 7.87 2.07 1.43 3.78

70M

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Fig. 5. Thin section photomicrographs of the Oued El-Abid synclinal zone igneous rocks showing some of the significant textures. (a) Poecilitic phenocryst of clinopyroxene(Cpx) including olivine (Ol) crystals and a plagioclase (Pl)-clinopyroxene intergrown groundmass; sample NB3-2 (B1 lava flow-Aghzif–Naour), crossed nicols, field of view8 mm; (b) polyhedral or granular olivine crystals partially pseudomorphosed by secondary minerals in sample TAG 2a (Tagleft Dyke), parallel nicols, field of view 10 mm; (c)porphyric texture with poecilitic phenocryst of clinopyroxene and fresh olivine crystals in sample ABKH 5-6 (B2 lava flow Aït Attab), crossed nicols, field of view 8 mm; (d)microgranular to porphyric texture in sample ABKH 2-4 (Ouzoud sill), crossed nicols; field of view 8 mm.


compared with the transitional affinities of the studied Cretaceousrocks (see Section 6.1).

The crystallization temperature and pressure of the analyzedclinopyroxene cores have been calculated after Putirka (2008), con-sidering the whole-rock composition as representative for themagma from which the pyroxenes crystallized and assuming thatthese pyroxenes are early liquidus minerals. Equilibrium criteria,suggested by Putirka (2008) are considered. Generally, theclinopyroxenes yield moderate pressure and temperature valuessuggesting crystallization in mid to upper crustal magma cham-bers (8–2 kbar at ca. 1100–1200 �C). Only sample NA2A of B1 lavaflows yields relatively high crystallization pressure, indicative of adeep crustal onset of crystallization (ca. 10 kbar, Fig. 6c).

5.1.3. FeldsparsFelsdspars phenocrysts are also ubiquitous, ranging in composi-

tion from labradorite to oligoclase (An64 to An20). The calciumrichest compositions are from the phenocryt cores. Phenocrystrims and matrix crystals are sometimes alkali felsdspars (up toOr43). See Table 4 and Fig. 7.

5.2. Whole rock geochemistry

The studied rocks were variably altered during hydrothermaland/or supergene processes (see also Beraâouz et al., 1994; Bensa-lah et al., 2006), as shown by the mineralogy (see Section 5.1) andby the large range of loss on ignition (LOI) up to 9.91 wt%. In orderto avoid the effects of such secondary processes, we will focus thisstudy on samples with LOI < 6 wt% with analyses recalculated to100% on an anhydrous basis (Table 1).

The Middle to Upper Jurassic (B1) and the Lower Cretaceous(B2) magmatic rocks are usually basic, presenting SiO2 contentsfrom 48.26 to 53.56 wt%. For a given silica content the B1 lava

flows and dykes tend to be more alkali-rich than the B2 group(Fig. 8). They span a large variation in Mg# [defined as100 �Mg2+/(Mg2+ + Fe2+), with FeO = 0.9 Fe2Ot

3], from poorly dif-ferentiated (Mg# = 61; Ni = 194 ppm; Cr = 290 ppm) to highlyevolved (Mg# < 30; Ni < 20 ppm; Cr < 100 ppm) compositions. Nosamples represent primary mantle-derived magmas.

Rocks are moderately enriched in the more incompatible ele-ments as depicted by the fractionation between Light Rare EarthElements (LREEs) and Heavy Rare Earth Elements (HREEs) [(La/Yb)n = 4.26–9.60, where n stands for chondrite normalized valuesaccording to Nakamura, 1974]. If we consider rocks withMg# > 50, in order to minimize the effects of magma evolution,the B1 lava flows present LREE/HREE ratios [(La/Yb)n = 8.38–9.60]notably higher than the other studied rocks [(La/Yb)n for B1dykes = 4.44–5.77; for B2 lavas = 4.26–6.56; for B2 sills = 5.47–5.64].

The general geochemical characteristics for all incompatibleelements (including high field strength elements, HFSE, and largeion lithophile elements, LILE) are summarized in Fig. 9. These pat-terns show the moderate enrichment in incompatible elementsand the presence of very small Nb negative anomalies, not accom-panied by negative anomalies for Ti and P (see Section 6.2.1 for adiscussion).

6. Discussion

6.1. Magmatic affinities

On the total alkali-silica (TAS) diagram (Fig. 8), the analyzedsamples (with LOI < 6 wt%) fall on both sides of the dividing linebetween the fields of alkaline and subalkaline series (Irvine andBaragar, 1971). The B1 lava flows show alkaline affinity while theB2 lava flows cluster in the subalkaline field. The compositions of

Fig. 6. (a) Clinopyroxene composition of the Oued El-Abid synclinal zone igneousrocks plotted in the Ca–Mg–Fe triangle, classification of pyroxenes according toMorimoto et al. (1988): 1, diopside; 2, hedenbergite; 3, augite; 4, pigeonite; 5,clinoenstatite; and 6, clinoferrosilite; (b) TiO2 vs. MgO diagram for clinopyroxenesof the Oued El-Abid synclinal zone igneous rocks. (c) Pression–Temperature (P–T)diagram of Putirka (2008). Clinopyroxene crystallization pressure and temperaturehave been calculated for those crystals that are close to equilibrium with the hostwhole-rock. This is not the case for the few analyzed Tagleft dykes clinopyroxenes.Symbols: open triangle, B1 lava flows; filled triangle, Tagleft dykes; open square, B2lava flows; filled square, Ouzoud sills.

Fig. 7. Feldspar composition of the Oued El-Abid synclinal zone igneous rocksplotted in an An–Ab–Or diagram. 1, anorthite; 2, bytownite; 3, labradorite; 4,andesine; 5, oligoclase; 6, albite; 7, anorthoclase; and 8, sanidine. Symbols as inFig. 6.

Fig. 8. Total alkali vs. silica (TAS) diagram for the chemical classification andnomenclature of the studied rocks (after Le Bas et al., 1986). Subalkaline vs. alkalineseries boundary (dotted lines) is from Irvine and Baragar (1971). Rocks plotting intothe field trachybasalt are named hawaiite according to Le Bas et al. (1986). Symbols:open triangle, B1 lava flows; filled triangle, Tagleft Dykes; open square, B2 lavaflows; filled square, Ouzoud sills.


B1 dykes and B2 sills are more variable and straddle the dividingline between alkaline and sub-alkaline fields.

Because the whole rock major element compositions may havebeen partially disturbed by alteration processes, the clinopyroxenecomposition may be an alternative indicator of the magmatic affin-ity (Le Bas, 1962; Leterrier et al., 1982). Indeed clinopyroxene com-position tends to mirror the composition of the magmas fromwhich they crystallise (e.g. Thy and Lofgren, 1992). Taking this intoaccount, the differences depicted on Fig. 10a support the inferencesmade about magmatic affinities based on the TAS diagram (Fig. 8).In fact, despite some overlapping, pyroxenes from Jurassic rocksare usually characterized by nepheline normative compositions,while many of the pyroxenes from Cretaceous rocks are hyperstenenormative.

The lower SiO2 and higher TiO2 contents of the pyroxenes fromthe B1 relative to B2 lava flows (Fig. 10, Table 3), suggest a largerincorporation of non-quadrilateral components such as CaTiAl2O6

(e.g. Duda and Schmincke, 1985) in the former group, which is fa-voured by the lower silica activity of the alkaline magmas, com-pared to their subalkaline counterparts. The incorporation of theCaTiAl2O6 component in the pyroxene probably reflects. the fol-lowing paired substitution:

Ti4þ þ 2AlIV ¼ ðMgþ Fe2þÞ þ 2Si4þ

Considering that the studied rocks are somewhat weathered,the relatively immobile high field strength incompatible elements,such as Nb and Y, may be used to define the geochemical affinity ofthe studied rocks, alternatively to major elements. Nb is a highlyincompatible element during partial melting and the crystalliza-tion of most silicate minerals, whereas Y is compatible with garnetand, as such, its concentration is highly dependent on the amountof residual garnet in the source, i.e. depends on the source depth. Inaddition, Nb concentration in primary magmas reflects the degreeof enrichment/depletion of mantle source and the extent of partialmelting. Thus, Pearce and Cann (1973) considered as alkaline,

Fig. 9. Primitive mantle-normalized trace element ‘spidergram’ for the Oued El-Abid synclinal zone igneous rocks. Primitive mantle normalizing values are from Sun andMcDonough (1989) and the values of the Lower Continental Crust (LCC, open diamond), and Upper Continental Crust (UCC, filled diamond) are from Rudnick and Gao (2003).Symbols as in Fig. 8.

Fig. 10. Clinopyroxene composition of the Oued El-Abid synclinal zone igneous rocks plotted in the (a) normative (CIPW) Diopside (Di) vs. Hyperstene (Hy) and Nepheline(Ne) diagram of Bellieni et al. (1984); (b) Ti vs. Ca + Na diagram of Leterrier et al. (1982); (c) the Ti + Cr vs. Ca diagram of Leterrier et al. (1982); and (d) SiO2 vs. Al2O3 diagramof Le Bas (1962). Symbols as in Fig. 6.


Fig. 11. Tectonic discriminant diagrams for the Oued El-Abid synclinal zoneigneous rocks. (a) Ti–Zr–Y (Pearce and Cann, 1973), (b) V vs. Ti (Shervais, 1982), and(c) Zr/Y vs. Zr (Pearce and Norry, 1979). VAB, Volcanic Arc Basalts; MORB, Mid-Ocean Ridge Basalts); WPB, Within Plate Basalts).; Alk., Alkaline Basalts. CFB,Continental Flood Basalt. IAT, Island Arc tholeiite Symbols as in Fig. 8.


transitional and tholeiitic, rocks with Y/Nb < 1, from 1 to 2, andhigher than 2, respectively. Moreover, considering the significantdifference of the partition coefficients of Y (DY = 0.9 ± 0.17) andNb (DNb = 0.005 ± 0.001) between clinopyroxene and basic mag-mas (see Green et al., 1989), the Y/Nb ratio is expected to slightlyincrease by accumulation of clinopyroxene, a mineral phase abun-dant in the studied rocks. Therefore, the Y/Nb ratios measured onthe B1 (1.00 ± 0.04; 2r) and B2 (1.37 ± 0.15) lava flows are proba-bly maximum values. On the basis of all the presented mineral andwhole rock compositions, we consider B2 lavas of transitional char-acter and B1 ones as moderately alkaline.

As shown by Fig. 10, clinopyroxenes from Lower Cretaceousrocks (B2 lava flows and sills) are characterized by low Ti contentsas compared with those occurring in Jurassic B1 lava flows anddykes (see also Fig. 6b). However, it should be emphasized that,

notwithstanding the relatively low Ti contents, these clinopyrox-enes are Ti (and Cr)-rich enough to discriminate them from theclinopyroxenes usually crystallizing from magmas generated in su-pra-subduction zone environments (e.g. Beccaluva et al., 1989;Mollard et al., 1983). The anorogenic character of the studied rocksis also evidenced from whole-rock compositions, as shown by theuse of the discriminant diagrams on Fig. 11, where the high Ti/V,and Zr/Y ratios, as well the relatively high TiO2 contents (see Ta-ble 1), are clearly distinct from those usually characterizing arc-re-lated magmas.

6.2. Petrogenesis

There has been considerable discussion about the relative roleof asthenosphere, lithospheric mantle, and crust in the generationof continental basalts (e.g. Wilson, 1989; Coish and Sinton, 1992;Wilson, 1993; Dorais et al., 2005; Deckart et al., 2005). In particu-lar, most controversy has arisen on how continental basalts acquiretheir high concentrations of incompatible elements. Are theyasthenosphere (depleted or enriched) melts subsequently contam-inated by continental crust? Are they formed simply by relativelysmall degrees of melting of asthenospheric mantle? Does the litho-spheric mantle play a role in their genesis, either as source forbasaltic magmas, or as a contaminant of asthenospheric melts asthey pass upward ?

For the studied rocks, these questions will be addressed below.

6.2.1. Crustal contaminationThe comparison of the Moroccan Jurassic–Cretaceous rocks

with the mean composition of different portions of continentalcrust, enable us to assess the importance of crustal contamination.The compositions of the lower and upper continental crust are sig-nificantly different (e.g. Rudnick and Gao, 2003). However, theyshare some common characteristics such as negative anomaliesof Ti, Nb and Ta, which reflect the fact that continental crust mainlyforms at expenses of magmatic extraction at supra-subductionzone settings.

Some of the here studied rocks present small negative Nbanomalies (Fig. 9) that might reflect either source characteristicsor evidence of crustal contamination. This process would also beable to explain somewhat high Th/Yb ratios for a given Nb/Yb (orTa/Yb) ratio of the studied rocks as compared with subduction-unrelated oceanic basalts (cf. Pearce, 1982, 2008). However, thesmall negative Nb anomalies are not accompanied by negative Tiand P anomalies, which are typical of magmas contaminated bythe continental crust (e.g. Thompson et al., 1982; Dorais et al.,2005), suggesting that crustal contamination was not verysignificant.

This percept is also endorsed by Ti/Yb ratios, a parameter used,for example by Coish and Sinton (1992), to evaluate the possibleoccurrence of crustal contamination. Indeed, the studied rockspresent Ti/Yb ratios (5122–8061) clearly above those reported forthe continental crust (upper crust = 1957; lower crust = 3277; Rud-nick and Gao, 2003).

Also the K/P ratio has been considered as criterion to distinguishcontaminated from non-contaminated continental basalts giventhat K, a LILE, is enriched in the continental crust as opposed toP, a HFSE (see Fig. 9). If we take the extreme compositions of thedifferent types of within-plate oceanic island basalts as represen-tating the range of anorogenic mantle-derived magmas(DMM = 13.31; HIMU = 2.90; EM1 = 4.53; EM2 = 8.80; Jacksonand Dasgupta, 2008) and the mean compositions of the continentalcrust (upper crust = 35.51; lower crust = 11.60; Rudnick and Gao,2003) we conclude that crustal contamination was not an impor-tant player in the petrogenesis of the studied rocks. Indeed, withthe exception of one outlier, these samples are characterized by

Fig. 12. Variation of La/Yb and Tb/Yb in the Oued El-Abid synclinal zone igneousrocks compared with a fractional melting grid using primitive mantle ratios as thesource (shown as a large white circle; Sun and McDonough, 1989). The modelassumes a garnet mode in the source varying from 0–8% (with olivine at 55%,orthopyroxene at 25% and clinopyroxene at >12% entering the melt phase in theproportions 20:20:40:20% respectively). Melt increments are shown as subverticallines in 1% intervals between 1% and 7%. Partition coefficients employed are fromthe compilation of McKenzie and O’Nions (1991), which are in good agreement withcompilations from other sources (e.g. Halliday et al., 1995). After George and Rogers(2002). Symbols as in Fig. 8.


K/P ratios up to 8.90, which is close to that reported for non-oro-genic oceanic basalts.

In conclusion, if some fingerprints of crustal contamination can-not be completely ruled out, the elemental concentrations and ra-tios of the large majority of samples do not appear to have beenmodified beyond the point where they can be used as tracers ofmelting processes and mantle source compositions.

6.2.2. Magma genesis processesIf we only consider samples with Mg# P 50, in order to avoid

significant effects of crystal fractionation, a large (La/Yb)N rangeis depicted by the studied samples (4.26–9.60). Moreover, theirDy/Yb ratios (2.22–2.77), significantly higher than those usually re-ported for N-MORB (1.49; Sun and Mcdonough, 1989), are a clearindication that these rocks preserve the fingerprint of their paren-tal magmas being in equilibrium with residual garnet, i.e. at depthsin excess of 66 km, if small amounts (<0.3 wt.%) of water werepresent (Green, 1973), or in excess of 99 km in anhydrous condi-tions (e.g. Walter and Presnall, 1994).

The REE fractionation can also be emphasized by the diagramTb/Yb vs. La/Yb on which is superimposed a grid (after Georgeand Rogers, 2002), showing the effects of the variation of partialmelting degree of peridotitic sources where the amount of modalresidual garnet varies between 0% and 8% (Fig. 12). We emphasizethat Tb/Yb and La/Yb ratios are much more affected by meltingprocesses than crystal fractionation processes dominated by re-moval of olivine ± clinopyroxene, thus allowing the use of this dia-gram for rocks representatives of moderately evolved magmas.

Table 5Modeling of the effect of partial melting degrees on the La/Ce and La/Sm ratios of the Ouecoefficients for the residual paragenesis; C0 = source composition. For calculations it wacomposed by 55% olivine + 25% orthopyroxene + 16% clinopyroxene + 4% garnet. Mineral-mrefers to the relative percent difference between concentrations or ratio values calculated fobetween concentrations or ratio values calculated for 4% and 10% of partial melting degre

F 0.02 0.04D C0

La 0.00976 0.708 23.635 14.1Ce 0.01755 1.833 48.363 31.4Sm 0.05350 0.444 5.954 4.6La/Ce 0.386 0.489 0.4La/Sm 1.595 3.970 3.0

From this diagram we can infer that the amount of residual gar-net would be approximately 1–3%, or slightly higher if instead theprimitive mantle source composition a more enriched and REE-fractionated source was melting. As portrayed by Fig. 12, REEfractionation of the studied samples also emphasizes a significantvariation in the extent of melting (F) in order to explain the largevariation in the La/Yb ratios (5.94–13.37). If, for the studied rocks,we assume a homogeneous mantle source of primitive composi-tion, F would have varied between ca. 0.02 and 0.05 Obviously,these values are merely indicative, as F depends on the assumedelemental source composition, for which we have no way to fur-ther constrain, and considering furthermore that our samples arenot (near-) primary mantle melts. In conclusion, if we take into ac-count all these incertitudes and the source is considered homoge-neous, the above mentioned F values indicate a large variation(�150% relative) on the extent of partial melting to explain thechemical variability of the studied magmatism.

The homogeneity of the source can be assessed using elementalratios involving MREE and LREE, considering that ratios like as La/Sm or Ce/Nd would not be significantly altered by magmatic evo-lution processes dominated by olivine fractionation. The observedvariations for the less evolved rocks (Mg# P 50 ppm) is 43% for La/Sm (i.e. from 2.68 to 3.84) and 6% for La/Ce (i.e.from 0.47 to 0.49).Using the set of partition coefficients of McKenzie and O’Nions(1991) and taking into account the maximum amount of residualgarnet (4%) inferred from Fig. 12, the range in La/Sm and La/Ce ra-tios, imposed by a variation of F values of 150% (see above andFig. 12) would be respectively of 41% and 11% (considering a rangeof F between 0.02 and 0.05) and of 33% and 8% (for F ranging from0.04 and 0.1; see Table 5). These values are similar to variations de-picted by lavas, suggesting that variations in partial melting degreeexplain most of the observed variability on those significant ele-mental ratios (Table 5) and that the source of the Jurassic–Creta-ceous High-Atlas magmatism was reasonably homogeneous.

As discussed above, B1 lava flows and Aghzif–Naour/Tagleftdykes were emplaced during Middle to Upper Jurassic while theB2 lava flows and sills of Ait Attab/Ouzoud area intruded duringthe Early Cretaceous. This could lead to the interpretation that hyp-abyssal rocks could be cogenetic to (and feeding) the studied lavaflows. However, incompatible trace element ratios clearly showthat, at least for the Jurassic event, this is not the case, as, consid-ering the least evolved rocks (Mg# P 50), there are significant dif-ferences between lava flows and dykes (see Table 1) impeding thatthey can be considered strictly comagmatic. Moreover, consideringthe very significant differences on the La/Yb ratios (Table 1;Fig. 12), dykes point to higher degree of melting compared to thelava flows, and therefore cannot be considered as the feeder ofthe latter.

6.2.3. Source characteristicsA question always emerging from the study of continental bas-

alts is whether magma sources were located in the sub-continentallithospheric mantle (SCLM) or in the asthenosphere.

d El-Abid synclinal zone igneous rocks. F: partial melting degree; D: global partitions assumed the occurrence of batch melting and that the residual paragenesis was

agma partition coefficients were taken from McKenzie and O’Nions (1991). D% 2–5r 2% and 5% of partial melting degree. D% 4–10 refers to the relative percent differencee.

0.05 0.1 D% D%2–5 4–10

18 11.751 6.394 101 12167 26.787 15.364 80.5 10542 4.182 2.795 42.4 66.149 0.439 0.416 11.4 7.841 2.810 2.287 41.3 32.9

Fig. 13. The Oued El-Abid synclinal zone igneous rocks plotted in the (a) Nb/Th vs.Zr/Nb, and (b) Zr/Y vs. Nb/Y diagrams (after Condie, 2003, 2005). EN, enrichedmantle ; DM, depleted mantle (DM); DEP, deep-depleted mantle; REC, recycledmantle field; OIB, oceanic island basalt field (OIB); PM, primitive mantle (PM);HIMU, high U/Pb mantle source; EM 1 and EM 2, enriched mantle sources. Vectorsshow the effects of subduction (SUB) and batch melting (F). N-MORB, normal mid-ocean ridge basalt; UC, upper crust. Symbols as in Fig. 8.


The composition of the sub-lithospheric mantle, away from theinfluence of recent subduction zones, is relatively well constrainedby the study of oceanic basalts. Its elemental and isotopic variabil-ity has been explained in terms of the existence of several mantlecomponents (DMM, HIMU, EM1, EM2, FOZO) which are thought torepresent mantle depleted by the extraction of continental crust(DMM) or ‘‘enriched’’ by ancient recycled materials (HIMU, EM1,EM2) (e.g. Hofmann, 2003 and references therein). FOZO has beenexplained either as relatively primitive mantle material (Hart et al.,1992) or as resulting from the mixture within the mantle of theabove mentioned recycled materials (Stracke et al., 2005). SCLMdomains can be very old (>1 Ga) and highly heterogeneous giventhat their rigidity does not allow the existence of convective mo-tion capable to induce chemical homogenization. All these makethe exercise of assigning a determined magma composition to alithospheric or an astenospheric source, highly difficult.

Nevertheless, some authors (e.g. Coish and Sinton, 1992 and ref-erences therein) have used, in absence of isotopes, La/Ta and La/Nbratios to distinguish lithospheric (La/Ta > 22; La/Nb > 1.5) fromasthenospheric sources which would be characterized by La/Ta < 22; La/Nb < 1.5. Plume sources dominated by HIMU or EM-type components are expected to have also La/Ta clearly belowthe lithospheric values (see for example Stracke et al., 2003). Thestudied rocks are characterized by ratios (La/Ta = 10.80–16.94;La/Nb = 0.68–1.14) which, according to such criteria, point to asublithospheric (s.l.) origin. This is apparently also supported byother HFSE ratios (see Fig. 13 diagram, based on Condie, 2003).The similarity between the studied rocks and the mantle compo-nents thought to represent ancient recycled enriched materials(EM components), and its difference relatively to the depletedmantle component (DMM), which is mainly sampled at Mid OceanRidges, may suggest a mantle plume as a constrainer of the compo-sition of the studied rocks. However, given the absence of geolog-ical evidences (e.g. uplift; Saunders et al., 1992) for the presence ofa mantle plume in the region at Middle–Upper Jurassic–Cretaceoustimes and even during the previous CAMP genesis (e.g., Marzoliet al., 2004; Martins et al., 2008), such signatures are most proba-bly inherited from lithosphere. In fact it cannot be excluded thatthe deeper portions of the lithosphere, which may melt preferen-tially, do not have the Nb depletion typical of the lithospheric man-tle modified by supra-subduction processes, while EM-likecompositions of the High Atlas Jurassic–Cretaceous magmatismmay be attributed to enriched components entrained within thelithosphere, for example as a consequence of metasomatism (cf. Pi-let et al., 2008).

6.3. Towards a geodynamic model for the Jurassic–Cretaceousmagmatism of Morocco

This section aims to integrate the geological, volcanological,geochemical and mineral-chemistry data of the region with its tec-tonics, in order to establish a coherent geodynamic model for theMiddle-Jurassic to Lower Cretaceous magmatism of the MoroccanHigh-Atlas. We will start by briefly describing the geological con-text of other occurrences of this type of magmatism in additionalregions of Morocco, Iberia, the Americas, etc.

Middle–Late Jurassic to Early Cretaceous intrusive complexes,dyke swarms, and basaltic lava flows were emplaced in the HighAtlas, the Jbilets and possibly the Middle Atlas. This is documentedby some 40Ar/39Ar plateau-ages (Armando, 1999), by several K/Arages, by paleomagnetic (e.g, Westphal et al., 1979) and biostrati-graphic studies (e.g., Haddoumi et al., 2010). K–Ar ages range fromthe Dogger to the Barremian and might be separated in two differ-ent groups: 175–155 ± 5 Ma and 135–110 ± 5 Ma, respectively.40Ar/39Ar ages obtained on biotite separates from the troctolitesand olivine gabbros from the Tirrhist massif (Jebel Hayim, High At-

las) produced isochrones and plateau ages in two distinct intervals:151 ± 1 Ma to 147 ± 1 Ma (isochrons) and 151.3 ± 0.5 Ma to145.0 ± 0.5 Ma (plateau ages). The most voluminous Jurassic–Cre-taceous rocks occur in the Central High Atlas as bimodal basic/ultrabasic (troctolites to gabbros) and intermediate/silicic (diorites,monzodiorites and syenites) intrusions or, more rarely, as basalticlava flows (e.g., Argana basin, Jbilets and, possibly, Eastern High At-las; Huvelin, 1977; Ferrandini et al., 1991; Amrhar et al., 1997; ElKochri and Chorowicz, 1996; Aït Chayeb et al., 1998; and Haddo-umi et al., 2008). In the Middle Atlas, the recognition of Jurassic–Cretaceous magmatic bodies is problematical due to the complextectonic (Fedan, 1989; Mahmoudi and Bertrand, 2007) and is prob-ably limited to small intrusive bodies and dykes (Mahmoudi et al.,2007). In general, these Jurassic–Cretaceous rocks have a bimodalchemical distribution and are characterized by moderately alkalineto sub-alkaline affinities, with typical compositions of within-plateand rift-related magmatic suites (e.g., Bensalah et al., 2006 and ref-erences therein).

The Jurassic–Cretaceous magmatism is well documented in theWestern Iberian margin in Portugal (147–141 Ma; Grange et al.,2008; Alves et al., 2010a). Magmatic rocks form two alignmentsbetween the Soure and Rio Maior latitudes in the Lusitanian basin.They are transitional (moderately alkaline to sub-alkaline) andcompletely distinct from the magmatic events of the Triassic–Jurassic tholeiitic CAMP and the Upper Cretaceous alkaline rocks.The chemical variability is accounted by different degrees of

Fig. 14. Map of the ‘‘Moroccan Hot Line’’ showing the relationships with the Late Triassic–Early Jurassic rift systems and the Upper Jurassic–Lower Cretaceous West MoroccanArch uplift (after Missenard, 2006 and Missenard et al., 2006; completed by Frizon de Lamotte et al., 2009), and the repartition of the Jurassic–Cretaceous (Huvelin, 1977;Beraâouz and Bonin, 1993; Beraâouz et al., 1994; Amrhar et al., 1997; Armando, 1999; Aït Chayeb et al., 1998; Lhachmi et al., 2001; Zayane et al., 2002; Bensalah et al., 2006)and Tertiary-Quaternary magmatism (Maury et al., 2000; Duggen et al., 2009; El Azzouzi et al., 2010). Explanations in the text.


mantle partial melting, fractional crystallization and hydrothermalalteration (Alves et al., 2010b).

In Eastern North America, the Late Jurassic to Early Cretaceousalkaline activity is widespread from Florida to Newfoundland (Hur-tubise et al., 1987 and references therein). Early Cretaceous alka-line activity is also widespread in South America (e.g. Barraganet al., 2005; Lustrino et al., 2005; Gibson et al., 2006): basaltic tuffcones of the Oriente Basin (Ecuador), Valle Chico igneous complex(eastern Uruguay), and Parana Continental Flood Basalts (Para-guay–Brazil).

Emplacement of the Jurassic–Cretaceous magmas occurredafter the earliest Jurassic break-up of Pangaea and the formationof the Central Atlantic ocean, accompanied by generalized exten-sion and rifting in the Maghrebian sector of the Alpine Tethys filleddominantly by carbonate rocks (Warme, 1988). The Jurassic–Cretaceous magmatism of Morocco mainly occurs along twoalignments: the NE–SW-oriented Argana–Jbilets axis and theENE–WSW-trending Argana–Central High Atlas axis (Fig. 14). Wecan reasonably assume that magmatism was controlled by pre-existing Hercynian fault-systems reactivated as normal faultsduring the Triassic and Jurassic/Cretaceous passive rifting stage.

Several models have been envisaged to explain the genesis ofthe Moroccan Jurassic–Cretaceous magmatism. According to theclassical interpretation of Laville and Piqué (1992), Jurassic mag-matism is related to major deformation event producing transpres-sional folding and cleavage, with subsequent erosion of theorogenic weld, leading to the exhumation of the basal plutonicrocks. However, in this case, plutonic rocks and tectonic foldsshould be unconformably overlaid by the latest Jurassic–Creta-ceous red beds. This interpretation remains a matter of debate(see Laville, 2002, and Gomez et al., 2002). In fact, recent thermo-chronologic data suggest that plutonic rocks were still located atdepth in the 90–80 Ma time lapse (Barbero et al., 2007). It thus

seems more convincing to link the High Atlas Jurassic–Cretaceousmagmatism to the continuation of the previous extensional re-gime, an interpretation consistent with the petrologic and geo-chemical features described for our magmatic rocks. In thisperpective some small-scale ductile deformations found aroundthe Jurassic intrusions (Laville et al., 1994) can be explained interms of igneous emplacement mechanisms (e.g., ‘‘diapiric folding’’of Schaer and Persoz, 1976).

Alternatively, we could suggest that the Jurassic–Cretaceousmagmatism is associated to the uprising of some kind of ‘‘man-tle-plume’’ (sensu lato): either a vertically ascending conduit (theclassical plume concept), or a laterally migrating anomalouslyhot sublithospheric channel (sensu Oyarzun et al., 1997). Thisinterpretation would be consistent with the presumed latest Juras-sic exhumation of Central Morocco termed the West MoroccanArch (WMA) (e.g., Teixell et al., 2005; Zeyen et al., 2005; Missenard,2006; Missenard et al., 2006; Barbero et al., 2007, 2011; Malusaet al., 2007; Ghorbal et al., 2008; Frizon de Lamotte et al., 2009;Saddiqi et al., 2009). However, we are unable to certify a mantle-plume geochemical signature for our rocks, while the small volumeof the here studied magmas and the probably low melting degreesdo not necessarily require a mantle-plume (i.e., anomalously hotmelting conditions). An anomalously hot (plume) mantle is alsonot supported by the relatively low crystallization temperaturescalculated for olivine and clinopyroxene phenocrysts (ca. 1200 �C).

While it is unlikely that a plume-related magmatism may belocalized in the same area for over 140 Ma (also given the small,but not insignificant, movement of the African plate), it may besuggested that zones of the African plate thinned in rift zonesmay be the preferential site of magmatic activity since the Triassic(CAMP), to the Jurassic (the here studied magmatism) to the Pres-ent (the Middle Atlas alkaline magmatism Maury et al., 2000; Dug-gen et al., 2009; El Azzouzi et al., 2010). The lithosphere stretching


during these rift events can induce adiabatic melting of either theasthenosphere or the lithosphere. The probability of lithospheremelting would be enhanced in the case of low-solidus enriched do-mains generated by metasomatism. We propose that a significantproportion of the melting during the Jurassic–Cretaceous magma-tism took probably place in the lowermost enriched portions of thelithospheric mantle, as suggested by the enriched compositions ofthe studied rocks, which are different both from the Triassic–Juras-sic CAMP tholeiites (e.g. Marzoli et al., 2004; Martins et al., 2008)as well as basalts issued from the depleted asthenospheric mantle.

The Moroccan Jurassic–Cretaceous magmatism, here studied, ismoderately alkaline to transitional, similarly to some broadly coe-val magmatic rocks from the ‘‘Peri-Atlantic Alkaline Province(PAAP; Matton and Jébrak, 2009). The interpretation of the Moroc-can Jurassic–Cretaceous magmatism as being mainly due to exten-sional features related to the progressive ongoing extension ofcircum-Atlantic basins is consistent with the general picture pro-posed by Matton and Jébrak (2009) for PAAP, does not requiringmantle-plumes. However, these models differ in what concernsthe nature of mantle source. While Matton and Jébrak (2009) in-voke the asthenosphere, we propose the lowermost enriched por-tions of the lithospheric mantle as the most probable source.

7. Concluding remarks

The petrography, mineralogy, geochemistry and geochronologyof some of the High Atlas Jurassic–Cretaceous intrusions are pres-ently well known (e.g. Beraâouz and Bonin, 1993; Beraâouz et al.,1994; Armando, 1999; Lhachmi et al., 2001; Zayane et al., 2002).However, before this paper, little was known about the magma-tism associated with the Jurassic–Cretaceous ‘‘red beds’’ of theOued El-Abid syncline. We consider this zone a key area for thecomprehension of the magmatism of the High-Atlas because it isone of the rare locations where, over a relatively narrow area, allthe types of the known Jurassic–Cretaceous modes of emplace-ment (composite intrusion massifs, sills, dykes and lava flow se-quences) crop out side by side and in close relation (Bensalahet al., 2006). Furthermore, no attention was given before to thephysical volcanology of the volcanic successions.

The main conclusions of the present study can be grouped asfollow:

(i) The Jurassic–Cretaceous volcanic pile was formed duringtwo pulses of volcanic activity, represented by the Middleto Upper Jurassic B1 and the Lower Cretaceous B2 events.There is evidence for a variable number of eruptions in eachevent. The products of distinct eruptions can be separated byfine clastic sediments, indicating significant time intervalsseparating the emplacement of each package of lava flow-units. The B1 volcanic pulse was formed by 1–4 eruptions(Aït Attab and Aghzif Naour synclines, respectively), eachusually formed by flow fields composed of one to six flowunits. The B2 volcanic pulse is the result of three eruptionsmostly composed of 1–5 pahoehoe flows. The swarm ofdykes and inclined sheets and sills are petrographically iden-tical to the extrusive rocks, yet in the case of B1 pulse theycannot be considered strictly comagmatic.

(ii) The B1 eruptive pulse (lava flows and dykes of Aghzif-Naour/Tagleft) is moderately alkaline character, while the B2 one(lava flows and sills of Aït Attab/Ouzoud) is transitional,none of them presenting evidences for significant crustalcontamination. Chemical characteristics such as trace ele-ments ratios indicate that the Jurassic–Cretaceous igneousrocks associated with the ‘‘red beds’’ Group of the High Atlaspreserve the effects of their parental magmas equilibrium

with residual garnet. We propose that they were generatedby several events of magma genesis characterized by vari-able degrees of partial melting of subcontinental mantle lith-osphere, in opposition to the suggested elsewhere for coevalmagmatic occurrences in other continents to which anasthenospheric origin is assigned.

(iii) The location of the studied High-Atlas magmatism seems tobe controlled by pre-existing Hercynian fault system thatwas reactivated as normal faults during a Middle to UpperJurassic–Cretaceous rifting event. The associated lithospher-ic stretching induced melting by adiabatic decompression ofenriched low-solidus infra-lithospheric domains. This indi-cates that distinct sources (asthenosphere vs. lithosphere)are needed to explain the PAAP magmatism.

Acknowledgements

This work was carried out as part of a PhD thesis of MohamedKhalil Bensalah to be submitted to the Department of Geology ofthe Faculty of Sciences-Semlalia, Cadi Ayyad University of Marrak-ech). We thank Hamid Haddoumi for valuable bibliographic assis-tance. We thank also Chantal Douchet, Paul Capiez and PhilippeGrandjean (Laboratoire de Géologie de Lyon, France) and RaulCarampin (Dipartimento di Mineralogia e Petrologia, Universitàdegli Studi di Padova, Italy) for help and assistance in the acquisi-tion of chemical data on minerals and whole rocks. Financial sup-port for this work was provided by several research projects: (i)FCT (Portugal)–CNRST (Morocco) to José Munhá, Línia Martinsand Nasrrddine Youbi; (ii) CNRS (France)–CNRST (Morocco) to Her-vé Bertrand and Hassan Ibouh; (iii) PICS, CNRS (France)–CNRST(Morocco) to Hervé Bertrand and Nasrrddine Youbi; (iv). CNRi(Italy)–CNRST (Morocco) to Giuliano Bellieni and Nasrrddine You-bi, and (v) CSIS (Spain)–CNRST (Morocco) to Miguel Doblas andNasrrddine Youbi. We acknowledge thoughtful comments byJean-Paul Liégeois, an anonymous reviewer and editors-in-chief,P.G. Eriksson and Sospeter M. Muhongo.

Appendix A. Supplementary material

Supplementary data associated with this article can be found, inthe online version, at http://dx.doi.org/10.1016/j.jafrearsci.2013.01.004.

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