Tectonometamorphic evolution of an intracontinental ... · deformation, metamorphism and magmatism devel-oped in the interior of continental blocks away from active convergent plate

Tectonometamorphic evolution of an intracontinental orogenyinferred from P–T–t–d paths of the metapelites from theRehamna massif (Morocco)

P. WERNERT,1 K. SCHULMANN,1 ,2 F . CHOPIN,2 , 3 P . �ST�IPSK �A,1 , 2 D. BOSCH4 AND M. EL HOUICHA5

1Centre for Lithospheric Research, Czech Geological Survey, Kl�arov 3, 11000, Prague 1, Czech Republic2�Ecole et Observatoire des Sciences de la Terre, Institut de Physique du Globe – CNRS, UMR7516, Universit�e deStrasbourg, 1 rue Blessig, 67084, Strasbourg Cedex, France ([email protected])3Department of Geosciences and Geography, University of Helsinki, PO Box 68, 00014, Helsinki, Finland4G�eosciences Montpellier, UMR-CNRS 5243, Universit�e Montpellier, Place Eug�ene Bataillon, 34095, Montpellier, France5Laboratoire de G�eodynamique et G�eomatique, Facult�e des Sciences, Universit�e Choua€ıb Doukkali, 24000, El Jadida,Morocco

ABSTRACT New petrographic and microstructural observations, mineral equilibria modelling and U/Pb (mon-azite) geochronological studies were carried out to investigate the relationships between deformationand metamorphism across the Rehamna massif (Moroccan Variscan belt). In this area, typical Barro-vian (muscovite to staurolite) zones developed in Cambrian to Carboniferous metasedimentary rocksthat are distributed around a dome-like structure. First assemblages are characterized by the presenceof locally preserved andalusite, followed by prograde evolution culminating at 6 kbar and 620 °C inthe structurally deepest staurolite zone rocks. This Barrovian sequence was subsequently uplifted tosupracrustal levels, heterogeneously reworked at greenschist facies conditions, which was followedlocally by static growth of andalusite, indicating heating to 2.5–4 kbar and 530–570 °C. The206Pb/238U monazite age of 298.3 � 4.1 Ma is interpreted as minimum age of peak metamorphic con-ditions, whereas the ages of 275.8 � 1.7 Ma and 277.0 � 1.1 Ma date decompression and heating atlow pressure, in agreement with previous dating of Permian granitoids intruding the Rehamna massif.The prograde metamorphism occurred during thickening and associated horizontal flow in the deepercrust (S1 horizontal schistosity). The horizontally disposed metamorphic zones were subsequentlyuplifted by a regional scale antiform during ongoing N–S compression. The re-heating of the massiffollows late massive E–W shortening, refolding and retrograde shearing of all previous fabrics coev-ally with regionally important intrusions of Permian granitoids. We argue that metamorphic evolu-tion of the Rehamna massif occurred several hundred kilometres from the convergent plateboundaries in the interior of continental Gondwanan plate. The tectonometamorphic history of theRehamna massif is put into Palaeozoic plate tectonic perspective and Late Carboniferous reactivationof (Devonian)–Early Carboniferous basins formed during stretching of the north Gondwana marginand formation of the Palaeotethys Ocean. The inherited heat budget of these magma-rich basins playsa role in the preferential location of this intracontinental orogen. It is shown that rapid transitionfrom lithospheric stretching to compression is characterized by specific HT type of Barrovian meta-morphism, which markedly differs from similar Barrovian sequences along Palaeozoic plate bound-aries reported from Variscan Europe.

Key words: intracontinental orogeny; metapelite pseudosection; monazite dating; P–T–t–d paths;Variscan Morocco.

INTRODUCTION

Intracontinental orogens are represented by zones ofdeformation, metamorphism and magmatism devel-oped in the interior of continental blocks away fromactive convergent plate boundaries (Raimondo et al.,2010). Consequently, the tectonometamorphic evolu-tion of such orogens cannot be explained using con-ventional models of plate tectonic convergent

boundaries (e.g. Dewey & Bird, 1970) and a specificgeodynamic frame is required. The most prominentintracontinental orogenies are typified by the present-day Alpine Pyrenees (Tugend et al., 2014), Asian oro-gens (Tien Shan and Altai; e.g. Molnar & Tappon-nier, 1975; Cunningham et al., 2009) and ancientProterozoic to Phanerozoic central Australian oro-genic belts (Peterman and Alice Springs orogens, e.g.Hand & Sandiford, 1999). It has been suggested that

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J. metamorphic Geol., 2016, 34, 917–940 doi:10.1111/jmg.12214

location of these orogens, where no oceanic and/orcontinental subduction is observed, is predestinedmainly by structural inheritance, thermal effect andfluid–rock interaction (Raimondo et al., 2014). Thearchitecture of modern intracontinental orogens ischaracterized by positive flower structure geometry,which is relatively well constrained thanks to theirexcellent exposure (Cunningham, 2005). In contrast,the ancient intracontinental orogens (e.g. Peterman-type orogeny in Australia) show that the weak meta-morphic core was extruded in a form of an asymmet-rical wedge, similar to the tip of the Tibetan-Himalayan channel flow region (Grujic et al., 2002;Raimondo et al., 2010).

Intracontinental orogenic events have not beenreported so far in the European Variscan belt. Thisemblematic Palaeozoic orogenic system is character-ized by successive amalgamation of the Gondwana-derived continental blocks to Laurussia ranging inage from Lower Devonian to Upper Carboniferous(Stampfli et al., 2013). The multiple orogenic eventsare localized along convergent margins of successivelyadded blocks, and are characterized by Upper Sil-urian–Upper Devonian oceanic subduction of distaland proximal passive margins, followed by LowerCarboniferous collisional deformation and metamor-phism (Schulmann et al., 2009; Lardeaux et al.,2014). In summary, the European Variscides areknown as a locus of typical convergent boundarymetamorphism and deformation. However, thenorthern margin of the Gondwana supercontinentexposed in central Morocco reveals Upper Carbonif-erous metamorphism and deformation, which islocated away from convergent plate boundaries(Piqu�e & Michard, 1989; Hoepffner et al., 2005,2006; Michard et al., 2010) and may also represent atypical intracontinental orogen (Hoepffner et al.,2005). This orogenic system exemplified by theRehamna massif was recently interpreted as reflectingthe shortening of peri-Gondwanan Devonian toLower Carboniferous basins during Upper Carbonif-erous to Lower Permian (Chopin et al., 2014).

The main shortcoming of the reactivated continen-tal rift/basin orogenic models is a paucity of detailedand quantitative metamorphic petrology andmicrostructural studies, which allow the P–T–t–devolution of these systems to be assessed. To fill thegap, microstructural and petrological analysis cou-pled with in situ monazite U/Pb geochronology hasbeen undertaken. This allows us to evaluate thetectonometamorphic history of the Rehamna massifand to provide a thermal and mechanical model ofthis spectacular region. We discuss causes of litho-spheric weakening and thermomechanical implica-tions of the Rehamna orogenic event for reactivatedcontinental rift-basin orogens in general. It will beshown that young intra-continental rifts with inher-ited thermal budgets are loci of continental

lithosphere failure, leading to moderate thickening ofcrust and its extrusion as predicted by Thompson(1989) and Thompson et al. (2001). Finally, a plate-scale geodynamic perspective is provided for thisregion in the frame of the Variscan orogen.

GEOLOGICAL SETTING

The Alleghenian-Variscan orogeny in Morocco

In Morocco, the Alleghenian-Variscan orogenyresults from Devonian Rheic ocean closure and Car-boniferous to Permian collision of the Laurussia andGondwana supercontinents and associated terranes.This orogeny was ~8000 km long at the end of thePalaeozoic (Matte, 2001), before the Mesozoic Pan-gea continental breakup. The current witnesses arethe Palaeozoic massifs in Europe (Variscan orogeny;Edel et al., 2013), in North America (Alleghenianorogeny in the Appalachians; Hatcher, 2010) andNorth-West Africa (Mauritanides and MoroccanMeseta; Caby & Kienast, 2009; Michard et al., 2010).Away from the Cenozoic Atlas, the Moroccan Varis-can massifs are almost free of post-Variscan deforma-tion. Without considering the internal zone of theMagrebides, because of its controversial palaeogeo-graphic origin (Michard et al., 2010), almost all thePrecambrian and Palaeozoic north-west African mas-sifs are located in Morocco (Fig. 1). There, they canbe divided into two distinct structural domains(Hoepffner et al., 2005): the Meseta domain in thenorth and the Anti-Atlas domain in the south, at theWest African craton margin (Li�egeois et al., 2013).The Meseta domain is subdivided into the Westernand Eastern Meseta, and is considered to be thesouthwestern prolongation of the European Variscansegment (Michard et al., 2010). From west to east,the Western Meseta is divided into the Coastal Block,the Central zone, the nappe zone and the Sub Mesetazone (Fig. 1). Several granitoids associated with theAlleghenian-Variscan orogeny occur within thesePalaeozoic massifs (Mrini et al., 1992; Gasquet et al.,1996; El Hadi et al., 2006).

The Rehamna massif

Located in the Western Meseta domain (Fig. 1), theRehamna massif is divided into three lithotectonicunits (Michard, 1982): the Western Rehamna, whichis part of the Meseta Coastal block, and the Centraland Eastern Rehamna belonging to the Central Zoneof the Western Meseta domain (Fig. 2a). These threestructural units are separated by two major deforma-tion zones: the Median fault, which is part of theWestern Meseta Shear Zone, and the Ouled Zednesfault (Fig. 2a).Detailed lithology of the Rehamna massif is

described in Michard et al. (2010) and references

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918 P . WERNERT ET AL .

therein. The Western Rehamna consists of Cam-brian–Ordovician limestones and siliciclastic sedimen-tary rocks on Palaeoproterozoic basement (Pereiraet al., 2015). The Central Rehamna consists of aNeoproterozoic metarhyolitic basement covered byCambrian–Ordovician marble and metapelite toDevonian metaconglomerate (the Kef el Mouneb for-mation). The Eastern Rehamna is formed by Devo-nian metapelites (the Ouled Hassine fm.) and Viseanmetapelites with lenses of metabasites (the Lalla Tit-taf fm.). These formations are intruded by synoro-genic leucogranites and a post-tectonic graniticbatholith (Fig. 2a). A tectonic nappe (upper meta-morphic unit in Fig. 2a) consisting of Ordoviciansandstones and shales is located north to the Centraland Eastern Rehamna (e.g. Michard, 1982).

The Rehamna massif is heterogeneously affectedby syntectonic metamorphism from anchimetamor-phism in the Western Rehamna to amphibolite faciesin the Central and Eastern Rehamna (Fig. 2)(Hoepffner et al., 1982; Aghzer & Arenas, 1998; Bau-din et al., 2003). According to the degree of meta-morphism and the intensity of deformation, the unitsare subdivided into an upper metamorphic unit (thenorthern Ordovician nappe sequences) and a lowermetamorphic unit (the central part of the Rehamna

massif, sometimes referred as the Rehamna dome)(Baudin et al., 2003; Razin et al., 2003). The OuledZednes fault does not significantly modify the trendof metamorphic zones, but the Median fault juxta-poses abruptly the anchimetamorphic Coastal blockrocks with medium to high grade metasedimentaryrocks of the Central Rehamna (Baudin et al., 2003).A geochronological 40Ar/39Ar study (Chopin et al.,

2014) showed that the southward nappe stacking andsyn-convergent extrusion is Upper Carboniferous(310–295 Ma, 40Ar/39Ar on metamorphic amphi-bole and mica). The subsequent orthogonal shorten-ing is Lower Permian (295–280 Ma, 40Ar/39Ar onmuscovite within a mylonite) and is associated withsyn-tectonic leuco-granites (285 Ma, 40Ar/39Ar onmuscovite). The intrusion of post-tectonic granite hasbeen dated at 275.8 � 5 Ma (40Ar/39Ar, Chopinet al., 2014) and 268.8 � 6 Ma (87Rb/86Sr, Mriniet al., 1992).

Previous metamorphic and structural studies

Two principal metamorphic episodes have beenrecognized in the Rehamna massif, a synkinematicBarrovian M1 metamorphism followed by a post-tec-tonic static event M2 (e.g. Hoepffner et al., 1982).

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20 40 600km

N

El Jadida

Marrakech

Casablanca

Saifi

Ouarzazate

Fès

Azrou

Khénifra

Midelt

Rabat

Anti AtlasAnti Atlas

C e n t r a l Zo n e

W.M

.S.Z

.

W.M

.S.Z

.Co a

s ta l

bl o

c k

W.M

.S.Z

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W.M

.S.Z

.

Western MesetaWestern Meseta

EasternMesetaEasternMeseta

R.T.F.Z.R.T.F.Z.

S.M.F.ZS.M.F.Z

nappe

zon e

M.M.F.

Z.

M.M.F.

Z.

Atla

ntic

ocea

n

8°W 7°W 6°W 5°W

34°N

34°N

32°N

31°N

Central Massif Central Massif

JebiletJebilet

RehamnaRehamna

S.A.F

0°10°W

50°N

40°N

30°NMA

DZ

FR

PT ES

Variscan granitoidLate Variscan basin

Mesozoic cover

Eo-variscan phase(Upper Dev.−Tournaisian)

Early/Late thrustStrike slip

Variscan phaseUndifferenciatedFoliation and fold strikeFaults

Eastern Meseta

Anti AtlasCarboniferousCambr.-Devon.

Cambr.-Devon.

Sub Meseta zoneCarboniferousCambr.-Devon.

Precambrian inlier

Sehoul block

Western MesetaAcadian block

Coastal block

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+

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+

CarboniferousCambr.-Devon.

Cambr.-Devon.

Cambr.-Devon.Central zone

Nappe zone

(a)

(b)

Abbreviations: Middle Meseta Fault Zone (M.M.F.Z.), Western Meseta Shear Zone (W.M.S.Z.), South Meseta Fault Zone (S.M.F.Z.), South Atlas Fault (S.A.F.), Rabat-Tiflet Fault Zone (R.T.F.Z.).

9°W

S u b M e s e t a Z o n e

Fig. 2

High-Atlas blockHigh-Atlas block

Fig. 1. The Meseta domain and the adjacent Anti-Atlas. (a) Location of the Moroccan Variscides. (b) Simplified geological andstructural map of the Moroccan Variscides (from Chopin et al., 2014), modified after Hoepffner et al. (2005), Michard et al. (2008,2010).


TECTONOMETAMORPH IC EVOLUT ION OF THE REHAMNA MASS I F 919

Related metamorphic isograds mapped over last50 years were summarized by Michard et al. (2010)(Fig. 2a). The prograde M1 metamorphism wasattributed to moderate tectonic burial while thegreenschist facies retrogression was considered to

reflect post-orogenic extension (Aghzer & Arenas,1995, 1998; Baudin et al., 2003; Razin et al., 2003).Late contact M2 metamorphism characterized by sta-tic growth of biotite and andalusite was interpretedas a result of thermal influence of late to post-

d

Sebt-Brikiine

LallaMoucha

Sidi AliRais El Abiod

Kef el Mouneb

Skhour massif

Sekhira es Slimane

Ouled Hassine

Lalla TittafDalaat

Sebt-Brikiine

LallaMoucha

Sidi Ali

Kef el Mouneb

Skhour massif

Sekhira es Slimane

Ouled Hassine

Lalla Tittaf

Kouiat el

Adam

Ko

t el A

dam

Jbel Kharrou

Jbel Kharrou

Dalaat

El Massiralake

Dam

Upper metamorphic unit

Coastal block (Western Rehamna)

Lower metamorphic unitCentral Rehamna Eastern Rehamna

Middle Ordovician (shale)Middle Ordovician (sandstone, shale)Upper Ordovician (sandstone, shale)

Lower Permian (volcanics, conglomerate)Upper Visean to Namurian (sandstone, limestone)

Lower/Middle Cambrian (greywacke)

Lower/Middle Cambrian (metapelite)

Lower to Middle Devonian(metaconglomerate)

Visean–Serpukhovian(metapelite ± metabasic rocks)Devonian (metapelite ± quartzite,metaconglomerate, marble)

Lower Cambrian (limestone)

Neoproterozoic gneiss

Lower Cambrian(metaconglomerate to marble)

Variscan granitoids Post variscan cover

Isograds in the metapelitesgst bi

mukyanite-quartz veins

contact aureole

Devonian (sandstone, conglomerate, limestone)

Wadi/River RoadFault D1, D3 thrust

N0 5km

WesternRehamna

EasternRehamnaCentral

Rehamna

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32°3

0'N

32°2

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8°00'W

chl

du

ai

Rais El Abiod

Draa el Kebir

A’

A A’NW SE

D3D3 D3

(a)

(c)

chloritoid garnet staurolitechloriteHighest metamorphic mineral in sample:

Median

fault

Median

fault

Median

fault

Ouled Ouggad fault

Ouled Ouggad fault

Ouled Ouggad fault

S3

S1

S3

1 km

1 km

12-09A

T-1 12-07II

SA-32

8°00'W 7°50'W 7°40'W

32°3

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32°2

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Structural symbolsPlanar fabrics Linear fabrics

50

Minerallineation

Foldaxis

(b)

N0 5km

Fig. 2. (a) Lithological map of theRehamna massif (modified after Chopinet al., 2014) showing the isograds inmetapelites and the location of samplesused for petrography and chemicalanalyses. (b) Simplified structural map ofthe Rehamna massif (modified after Chopinet al., 2014). (c) Cross-section showinginterpretation of the structures and locationof the samples used for petrology,modelling and geochronology.



kinematic granitoids (Michard et al., 2010). The M1peak P–T conditions were estimated to 530–560 °Cand 7–9 kbar using conventional thermobarometry(garnet–chlorite, garnet–muscovite, garnet–biotiteand garnet–rutile–ilmenite–plagioclase–quartz) byAghzer & Arenas (1998) and El Mahi et al. (1999).

Based on structural and geochronological investi-gations, Chopin et al. (2014) proposed a new tec-tonic model for the Rehamna massif. The firstmacroscopically recognized structure is a shallow-dipping foliation S1 present in the whole massif(Fig. 2b,c). It is interpreted as a result of horizontalflow of deep crustal infrastructure during supra-crus-tal thickening associated with thrusting of anOrdovician nappe over the Neoproterozoic base-ment, its sedimentary cover and the Devono–Carboniferous basins due to a SSW-directedcompression. Intensity of this deformation increaseswith metamorphic grade.

In the field, the D2 deformation is marked by tilt-ing of the S1 foliation in the whole massif, and bylarge to medium scale F2 folds locally connected toS2 axial planar crenulation cleavage recognized in theeasternmost part of the Lalla Tittaf formation(Fig. 2b,c). The D2 deformation, which is a conse-quence of a NNE-SSW compression, allowed extru-sion of the lower metamorphic unit in an E–Wtrending antiform.

The D3 deformation affects heterogeneously thewhole Rehamna massif. It is orthogonal to D1 andD2 events with structures striking from N to NNE(Fig. 2b,c). In the Eastern Rehamna, the intensity ofthe D3 deformation increases from east to west,starting with west-facing asymmetric F3 folds andassociated S3 axial planar crenulation cleavage S3,and ending with strong mylonitization in the OuledZednes fault. In the Central and Western Rehamna,the F3 folds are generally upright and connected toscarce S3 axial planar crenulation cleavage. The sub-circular domain with exhumed Neoproterozoic base-ment in the Central Rehamna results from superposi-tion of the D2 and D3 deformation events,interpreted as a large scale type-1 interference foldpattern. The latter results from massive shortening ofweak Devonian and Carboniferous sedimentary rocksdue to eastward indentation of rigid Coastal Block(Chopin et al., 2014).

PETROGRAPHY AND MINERAL CHEMISTRY

In order to quantify the P–T–t–d evolution of theorogenic crust, crystallization–deformation relation-ships in metapelites were studied (sample locationsin Fig. 2a). In the description, we adapt the logicof relationship between multiple foliation develop-ment and porphyroblast growth, which establishesthe relative timing between structures and metamor-phism (Bell & Johnson, 1989; Skrzypek et al.,2011a,b). Succession of foliation development is

after Chopin et al. (2014). However, based onmicrostructural study, it is suggested here that theD1 deformation involves two nearly orthogonalstages, with the early D1 stage (introduced here asD1a) being preserved only in porphyroblasts andabsent as macroscopic foliation in the field, whilethe late D1 stage (introduced here as D1b) is pre-served in porphyroblasts and matrix and is also vis-ible macroscopically in the field (see descriptionsand discussion below). Metamorphic isogradsrelated to D1 event (Michard et al., 2010) havebeen slightly amended (Fig. 2a).First, a general description of metapelites is pre-

sented according to metamorphic zones, followed bydescription of selected samples used for modelling ofthe P–T path and geochronology. Representativephotomicrographs are shown in Figs 3 and 4. Theinterpretation of crystallization–deformation relation-ships is presented in Fig. 5, representative mineralanalyses are summarized in Tables 1–3, and garnet,chloritoid and staurolite compositional profiles areshown in Figs 6 and 7.The mineral abbreviations are after White et al.

(2007), alm = Fe/(Ca + Fe2+ + Mg + Mn), prp =Mg/(Ca + Fe2+ + Mg + Mn), grs = Ca/(Ca + Fe2+

+ Mg + Mn), sps = Mn/(Ca + Fe2+ + Mg + Mn),XFe = (Fe)/(Fe + Mg), XMn = Mn/(Mn+Fe+Mg).Trends in mineral composition from core to rim aremarked with ‘⟹’ and the sign ‘–’ is used for a rangeof mineral compositions, p.f.u. = per formula unit.

Crystallization–deformation relationships in metamorphiczones

Most of the samples are reworked by the D3 defor-mation and exhibit penetrative or spaced S3 cleavagewith more or less uniform orientation. As D2 isexpressed only as large-scale tilting of the S1 foliation(except for the easternmost part of the Lalla Tittafformation), there is no mineral crystallization duringthe D2 event. The pristine S1 foliation not affectedby the D3 deformation is only rarely preserved.Based on this macroscopic structural framework, thedeformation–crystallization relationships of the stud-ied samples therefore result from the D1 and D3deformation events only.

Garnet zone

The main fabric in the garnet zone is the S1 foliationreworked by the F3 folds leading in places to com-plete transposition into the S3 cleavage. Micaschistsin the garnet zone are fine-grained with the S1 folia-tion defined by alternation of quartz-rich and mica-rich layers composed of fine muscovite, chlorite andbiotite. The S3 cleavage is defined by fine muscoviteand chlorite with opaque minerals. Garnet is small(<2 mm), in places it includes numerous quartz andilmenite, rare muscovite, chloritized biotite and



Fig. 3. Microphotographs (plane-polarised light) of metapelites from the Rehamna massif illustrating crystallization–deformationrelationships. (a) Sigmoidal inclusion trails S1 in garnet at high angle to the external S3 fabric (garnet zone, sample 12-01). (b)Andalusite porphyroblasts with inclusion trails oriented at high angle to the external S3 foliation, interpreted as a result of growthin the S1 fabric (andalusite zone, sample 12-02F). (c) Folded inclusion trails of ilmenite in chloritoid. (d) Euhedral garnet grainsincluded in a staurolite porphyroblast. Inclusion trails S1a of ilmenite in garnet are systematically subparallel and at high angle toinclusion trails in staurolite and the S1b foliation. Inclusion trails of ilmenite and quartz in staurolite are parallel with the externalS1 foliation (staurolite zone, sample R061). (e) Lower end of a sigmoidal staurolite porphyroblast: inclusion trails in staurolite arecurved and continuous with the external S1 foliation. Small garnet occurs in the matrix (staurolite zone, sample SA-32). (f)Andalusite with folded inclusion trails of ilmenite, interpreted as growing after development of S3 foliation (staurolite–andalusitezone, sample 12-09A).



epidote. Inclusion trails in garnet are straight orcurved and oriented at high angle to the S3 foliation(Fig. 3a) suggesting that garnet grew during thedevelopment of the S1 fabric.

Andalusite zone M1

In sample 12-02F from the garnet zone, andalusite(up to 0.7 mm in size) was found with inclusion trailsof ilmenite and biotite that are straight and at highangle to the external S3 foliation (Fig. 3b). Therefore,andalusite in this area is attributed to the M1 meta-morphic event.

Garnet–chloritoid zone

A few samples in the garnet zone contain chloritoid.Small garnet (up to 2 mm) has no oriented inclusiontrails. Chloritoid (up to 1 cm) has commonly inclu-sion trails of ilmenite and quartz that are straight orcurved (Fig. 3c) and in places continuous with the S1

external foliation, interpreted as chloritoid growthduring the D1 event. Most samples are reworked bythe S3 cleavage defined mostly by fine-grained mus-covite, chlorite and ilmenite. The zones of the S3cleavage are commonly at high angle to the internalinclusion trails in chloritoid (Fig. 3c).

Staurolite zone

Micaschists in the staurolite zone are medium-grainedwith the S1 foliation defined by alternation of quartz-rich and mica-rich layers and by preferred orientationof muscovite, biotite and ilmenite. In some samples,chlorite occurs also in the S1 foliation and is inter-preted as replacing biotite. Most samples arereworked by the S3 crenulation and zones of S3cleavage are defined by fine-grained muscovite, bio-tite, chlorite and ilmenite. Garnet forms small por-phyroblasts with straight inclusion trails oriented athigh angle to the S1 foliation in staurolite (Fig. 3d).This led us to distinguish an early S1a foliation that

(a) (b)

(c) (d)

Fig. 4. Microphotographs (plane-polarised light) of metapelites from the Rehamna massif illustrating crystallization–deformationrelationships. (a) Straight and folded S1 inclusion trails of ilmenite in staurolite in samples affected by the S3 cleavage (staurolitezone, sample R079D). (b) Foliation S1 folded by F3 and transposed into S3 cleavage. On the right, staurolite shows foldedinclusion trails of ilmenite and quartz (staurolite–andalusite zone, sample 12-09A). (c) Hand-drawing showing relations ofinclusion trails in staurolite and F3 folds. The inclusions trails are perpendicular or parallel to the S3 cleavage. Staurolite growthprecedes the F3 folding and development of the S3 cleavage (staurolite–andalusite zone, sample 12-09B). (d) Hand-drawingshowing the relation between inclusions trails in staurolite, F3 folds, and post-kinematic andalusite (staurolite–andalusite zone,sample 12-09A).



was likely almost perpendicular to the S1b subhori-zontal foliation and therefore was subvertical in thefield. Staurolite (up to 1 cm in size) has straight orcurved inclusion trails, in places continuous with theS1 fabric (Fig. 3e), indicating its growth during thelate stages of the D1 (D1b) event.

Staurolite–andalusite zone

In the staurolite–andalusite zone, andalusite (0.5 mm insize) in occasionally found with curved inclusion trailsthat are continuous with the external S3 crenulation(Figs 3f & 4d). There are no S3 pressure shadowsaround andalusite and therefore its growth is consideredas post-tectonic with respect to the D3 event. Moreover,in both the staurolite and staurolite–andalusite zones,the inclusion trails in staurolite in samples affected bythe S3 cleavage show variable geometrical relations andare not always continuous with the S1 fabric (Fig. 4),which will be discussed below.

Petrology and mineral chemistry of selected samples

Mineral and whole-rock analyses were performed atthe Charles University in Prague, Czech Republic, on

a scanning electron microscope TESCAN VEGAequipped with an X-Max 50 electron dispersive spec-trometer detector (Oxford Instruments), with operat-ing conditions of 15 kV accelerating voltage and1.5 nA beam current. Whole rock analyses wereobtained by area scanning (see below).

Garnet–chloritoid micaschist sample 12-07II

Sample 12-07II is a garnet–chloritoid-bearing meta-pelite from the Visean-Serpukhovian Lalla Tittaf for-mation, located 3 km east from the Rais El Abiodgranite (Fig. 2a). It is taken from an outcrop withthe S3 foliation, which is microscopically defined byalternation of quartz- and muscovite–chlorite-richbands. The sample contains chloritoid, garnet, mus-covite, chloritized biotite, chlorite, quartz and ilme-nite. Garnet is small and rare (<1%, 250 lm in size),does not contain inclusions with preferred orienta-tion, and is surrounded by a fine-grained aggregateof chlorite and oxides, suggesting its partial resorp-tion (Fig. 6d). Garnet is not zoned (alm0.69 prp0.07sps0.17–0.18 grs0.08, XFe(0.91)) (Fig. 6a; Table 1). Chlori-toid (2 vol.%, up to 1 mm length) forms porphyrob-lasts with straight or curved inclusions trails that areorientated oblique to the external S3 foliation andare in places continuous with the S1 external fabric(Fig. 3c). In the absence of the D2 deformation inthis region, this suggests that chloritoid grew duringthe D1 deformation event. Curved inclusion trails inchloritoid may result from folding and transitionfrom the subvertical S1a foliation to the subhorizon-tal S1b fabric. Chloritoid is zoned from core to rimwith decreasing XFe = 0.86)0.83 and XMn =0.04)0.03 (Fig. 7; Table 1). Muscovite has Si =3.03–3.08 p.f.u. and Na = 0.19–0.24 p.f.u (Table 1).

Staurolite–garnet micaschist sample SA-32

Sample SA-32 is a staurolite–garnet-bearing meta-pelite from the Lower–Middle Cambrian part of theCentral Rehamna region, located 3 km northeastfrom the centre of the Sidi Ali dome and taken fromthe outcrop with the shallow-dipping S1 foliation(Fig. 2a). The sample contains porphyroblasts ofstaurolite and garnet sitting in a S1 foliation withmuscovite, chlorite, biotite, plagioclase, ilmenite andquartz. Staurolite forms large porphyroblasts (up to5 mm in size) with numerous inclusion trails ofquartz and ilmenite that are straight or slightlycurved and continuous with the external S1 foliation(Fig. 3d,e). This is interpreted as staurolite growthwithin the S1 foliation, locally undergoing limitedsyntectonic rotation. Small garnet (up to 500 lm) ispresent in the matrix and is also included in stauro-lite (Fig. 3d,e). Inclusion trails of quartz and ilmenitein garnet are oriented oblique to the inclusion trailsin staurolite and in the matrix (Fig. 6e similar as inFig. 3d), and therefore interpreted as garnet growing

chlmu

bigst

Min.Struct. S1early S1 S2 S3 Post S3

ilm

SA

-32

bi

mu

ilm

and


12-0

2F

chlmu

gctdilm


12-0

7II

chlmu

bigst


12-0

9A

ilmand

Gar

net-c

hlor

itoid

zon

eS

taur

olite

-and

alus

ite z

one

Fig. 5. Crystallization–deformation relationships inmetamorphic zones. Quartz is always stable. See text fordetails.



during the early stages of the D1 deformation (D1a).Garnet is zoned from core to rim showing decreasein spessartine (0.23⟹0.20), increase in almandine(0.56⟹0.58) and pyrope (0.13⟹0.14), constantgrossular (0.08) and XFe (0.80–0.81), and at the outerrim increase in spessartine and decrease in almandineand pyrope to alm0.56 prp0.13 sps0.22 grs0.08, XFe(0.81)

(Fig. 6b,e; Table 2). Staurolite XFe from core to rimdecreases and increases at the outer rim(0.76⟹0.74⟹0.77), Ti = 0.04–0.05 p.f.u.,XMn = 0.04 and Zn = 0.03 p.f.u. (Fig. 7b,e; Table 2).Muscovite has Si = 3.03–3.07 p.f.u. and Na = 0.20–0.24 p.f.u., biotite XFe = 0.36–0.40 (Table 2).

Staurolite–garnet–andalusite micaschist sample 12-09A

Sample 12-09A is a staurolite–garnet–andalusite-bear-ing metapelite from the Visean-Serpukhovian LallaTittaf West region, located within the contact aureoleof the Rais El Abiod leucogranite, 2 km east from themagmatic rocks (Fig. 2a). It is sampled in an outcropwith dominant S3 foliation and contains staurolite,garnet, andalusite, muscovite, chlorite, chloritized bio-tite, ilmenite, rutile and quartz. Microscopically,domains of the S1 foliation, defined by muscovite andbiotite with subordinate chlorite (1–10 lm) are foldedand transposed into the S3 cleavage defined by finegrained muscovite and chlorite (0.5–5 lm). Stauroliteporphyroblasts (2 vol.%, up to 4 mm in size) present

straight or locally curved inclusions trails that are gen-erally oblique to the external S3 fabric (Fig. 4b,d).Garnet (<1 vol.%, up to 650 lm) occurs either in thematrix or included in staurolite, it contains inclusionsof ilmenite that are not oriented. Ilmenite and rutileoccur also in the matrix. These microstructures areinterpreted as ilmenite, garnet and staurolite growingin the S1 foliation. Andalusite porphyroblasts (0.6vol.%) contain folded inclusion trails of ilmenite, con-tinuous with the matrix crenulated foliation, indicatingthat andalusite growth occurred after the D3 deforma-tion (Fig. 3f). Garnet shows from core to rim decreas-ing spessartine and XFe, increasing almandine andpyrope and constant grossular (alm0.70⟹0.75

prp0.08⟹0.10 sps0.16⟹0.09 grs0.07, XFe 0.91⟹0.89)(Fig. 6c,f; Table 3). Staurolite XFe from core to rimfirst decreases, then at the outer rim increases(XFe = 0.86⟹0.84⟹0.88), XMn follows a similartrend (XMn = 0.03⟹0.02⟹0.04), Zn = 0.04 p.f.u.and Ti = 0.04 p.f.u. (Fig. 7c,f; Table 3). Muscovitehas Si = 3.04–3.10 p.f.u. and Na = 0.17–0.21 p.f.u(Table 3).

PSEUDOSECTION MODELLING

Calculation methods

The pseudosections were calculated using THERMO-

CALC v. 3.37 (Powell et al., 1998) and the data set 6.2

Table 1. Representative mineral analyses for sample 12-07II. The number of oxygen and cations for normalization is indicated.

Sample 12-07II (garnet–chloritoid zone – Eastern Rehamna)

Analysis number 303 312 306 316 113 23 16 101 106

Mineral mu-matrix mu (shear zone) chl-matrix chl (shear zone) chl-near g ctd-rim ctd-core g-rim g-core

wt%

SiO2 46.56 46.65 24.86 23.98 24.79 24.31 24.11 37.24 37.06

TiO2 0.39 0.28 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Al2O3 37.01 37.17 23.96 23.39 22.75 40.18 39.84 21.00 20.97

FeO 1.04 1.05 27.16 27.37 26.79 23.35 23.83 30.31 30.56

MnO 0.00 0.00 0.41 0.36 0.46 0.94 1.26 7.41 7.47

MgO 0.34 0.32 10.98 11.37 11.50 2.51 2.12 1.73 1.60

CaO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 2.44 2.31

Na2O 1.55 1.80 0.00 0.00 0.12 0.00 0.00 0.00 0.00

K2O 8.82 9.06 0.16 0.05 0.00 0.00 0.00 0.00 0.00

BaO (ZnO for st) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 95.71 96.34 87.52 86.51 86.41 91.29 91.17 100.13 99.98

Cation/Oxygen 7/11 7/11 20/28 20/28 20/28 8/12 8/12 8/12 8/12

Si 3.07 3.05 5.36 5.23 5.40 2.03 2.03 3.01 3.01

Ti 0.02 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Al 2.88 2.86 6.09 6.01 5.84 3.96 3.95 2.00 2.01

Fe3+ 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Fe2+ 0.06 0.06 4.90 4.99 4.88 1.63 1.67 2.05 2.08

Mn 0.00 0.00 0.07 0.07 0.09 0.07 0.09 0.51 0.51

Mg 0.03 0.03 3.53 3.69 3.73 0.31 0.27 0.21 0.19

Ca 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.21 0.20

Na 0.20 0.23 0.00 0.00 0.05 0.00 0.00 0.00 0.00

K 0.74 0.76 0.04 0.01 0.00 0.00 0.00 0.00 0.00

Ba (Zn for st) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 7.00 7.00 20.00 20.00 20.00 8.00 8.00 8.00 8.00

XFe (total) 0.84 0.86 0.91 0.91

Xalm 0.69 0.70

Xprp 0.07 0.06

Xsps 0.17 0.17

Xgrs 0.07 0.07



(Holland & Powell, 2011; February 2012 upgrade) inthe system MnO–Na2O–CaO–K2O–FeO–MgO–Al2O3–SiO2–H2O–TiO2–O2 (MnNCKFMASHTO).The following activity models are used: garnet, stau-rolite, chlorite, chloritoid, biotite and ilmenite arefrom White et al. (2014b), white mica from Whiteet al. (2014a), feldspar from Holland & Powell(2003), epidote from Holland & Powell (2011).Rutile, kyanite, sillimanite, andalusite and quartz areconsidered pure end-members. Molar and composi-tional isopleths were plotted for the phases of inter-est, especially to study garnet and staurolite growthand zoning to infer P–T trajectories. The isoplethnotations used are: x(g) = Fe/(Fe + Mg), z(g) = Ca/(Ca + Fe + Mg + Mn), m(g) = Mn/(Ca + Fe + Mg +Mn), x(st) = Fe/(Fe + Mg).

In order to approach as closely as possible the bulkcomposition effective at the scale of the thin section,whole-rock compositions were estimated by quantita-tive analysis of a representative area of the thin-sec-tions using a SEM equipped with an EDS detector(for more details see Appendix S1; Figs S1–S6;Table S2). The size of the selected areas for scanningwas 1.5 9 1 cm and particular attention was paid toinvolve garnet, staurolite, andalusite and matrix. Thewhole-rock compositions are presented as insets in thediagrams in molar per cent normalized to 100%(Figs 9 & 10). Muscovite, quartz and H2O are set inexcess and the proportion of Fe2O3 (O value) for

modelling is set to 0.01 to reflect the generally reducedcharacter of pelitic rocks. This value allows both ilme-nite and rutile to be stable and magnetite to be absentin the modelled P–T domains.

Pseudosection for garnet–chloritoid micaschist sample 12-07II

Several attempts were made to calculate a P–T pseu-dosection of sample 12-07II, including variation ofthe selected area for the whole-rock analysis, andvariation of the unknown proportion of Fe3+ andconsequently its effect on the whole rock XFe. How-ever, the calculated pseudosections did not satisfacto-rily reproduce the observed assemblage, especially thecoexistence of ilmenite and chloritoid, or the chlori-toid composition. The problems encountered aremost likely due to the chlorite activity-compositionrelations (R. Powell, pers. comm.). Thus, the pseudo-section for this sample is not presented.

Pseudosection for staurolite–garnet micaschist sample SA-32

Mineral assemblage and deformation relationships insample SA-32 indicate first growth of garnet followedby growth of staurolite in the stability field of ilme-nite in the S1a and S1b fabrics. This is compatiblewith a prograde path from g–mu–bi–chl–pl–ilm or g–

Table 2. Representative mineral analyses for sample SA-32.

Sample SA-32 (staurolite zone – Central Rehamna)

Analysis number 302 125 204 117 110 117 65 90 105 103

Mineral mu-matrix chl-in st chl-matrix bi-matrix g-core g-rim st-core st-rim pl ksp

wt%

SiO2 45.37 25.21 26.67 36.79 37.10 37.45 27.70 28.10 62.31 64.64

TiO2 0.41 0.00 0.00 1.30 0.26 0.00 0.48 0.45 0.00 0.00

Al2O3 34.68 22.65 19.14 18.38 20.46 20.65 53.73 54.42 23.33 18.23

FeO 2.43 17.27 23.66 14.60 25.48 26.36 13.44 12.25 0.00 0.00

MnO 0.00 0.34 0.38 0.14 10.45 8.95 0.71 0.71 0.00 0.00

MgO 0.54 19.31 15.93 13.06 3.30 3.63 2.41 2.42 0.00 0.00

CaO 0.00 0.00 0.00 0.00 2.96 3.02 0.00 0.00 4.66 0.07

Na2O 1.82 0.00 0.00 0.46 0.00 0.00 0.00 0.00 8.82 0.28

K2O 8.80 0.00 0.00 9.01 0.00 0.00 0.00 0.00 0.00 16.78

BaO (ZnO for st) 0.00 0.00 0.00 0.00 0.00 0.00 0.35 0.29 0.00 0.00

Total 94.05 84.77 85.78 93.75 100.01 100.06 98.85 98.64 99.12 100.00

Cation/Oxygen 7/11 20/28 20/28 8/11 8/12 8/12 15/23 15/23 5/8 5/8

Si 3.05 5.28 5.73 2.84 2.98 2.99 3.88 3.93 2.78 2.99

Ti 0.02 0.00 0.00 0.08 0.02 0.00 0.05 0.05 0.00 0.00

Al 2.75 5.60 4.85 1.67 1.94 1.95 8.87 8.97 1.23 0.99

Fe3+ 0.10 0.00 0.00 0.00 0.08 0.07 0.00 0.00 0.00 0.00

Fe2+ 0.03 3.03 4.25 0.94 1.63 1.70 1.57 1.43 0.00 0.00

Mn 0.00 0.06 0.07 0.01 0.71 0.61 0.08 0.08 0.00 0.00

Mg 0.05 6.03 5.10 1.50 0.39 0.43 0.50 0.50 0.00 0.00

Ca 0.00 0.00 0.00 0.00 0.25 0.26 0.00 0.00 0.22 0.00

Na 0.24 0.00 0.00 0.07 0.00 0.00 0.00 0.00 0.76 0.03

K 0.76 0.00 0.00 0.89 0.00 0.00 0.00 0.00 0.00 0.99

Ba (Zn for st) 0.00 0.00 0.00 0.00 0.00 0.00 0.04 0.03 0.00 0.00

Total 7.00 20.00 20.00 8.00 8.00 8.00 15.00 15.00 5.00 5.00

XFe (total) 0.81 0.81 0.76 0.74

Xalm (g), Ab (pl,ksp) 0.56 0.58 0.77 0.02

Xprp (g), An (pl,ksp) 0.13 0.14 0.23 0.01

Xsps (g), Or (pl,ksp) 0.23 0.20 0.00 0.97

Xgrs 0.08 0.08



mu–bi–chl–pa–pl–ilm fields to st–g–mu–bi–pl–ilmfield in the pseudosection (Fig. 8). Calculated molarisopleths of garnet predict growth until stauroliteappearance, then no or infinitesimal increase ofthe amount of garnet within the staurolite stabil-ity field. Staurolite molar isopleths predict itsgrowth during increase in pressure and temperature.Decreasing spessartine and grossular in garnet(sps0.23⟹0.20grs0.08) are interpreted as a result of gar-net growth along the prograde path at ~520–550 °Cand 4–5 kbar (Fig. 8). Garnet XFe (0.80) and stauro-lite XFe (0.76⟹0.74) are close to isopleths calculatedat higher pressure part of the st–g–mu–bi–pl–ilm fieldwhich may indicate reequilibration of the XFe at peakfor garnet, and reflect growth of staurolite. Parago-nite absence in the sample is compatible with inferredpeak conditions. Increasing spessartine in garnet andincreasing XFe in staurolite at the outer rim likelyrepresent partial diffusional reequilibration onexhumation, compatible with trend of the isoplethswith decreasing pressure. The P–T path during theD1 deformation inferred from this sample goes from4 to 5 kbar and 520–550 °C to peak estimated at6–7 kbar and 620–660 °C (Fig. 8).

Pseudosection for staurolite–garnet–andalusite micaschistsample 12-09A

Crystallization–deformation relationships involvegrowth of ilmenite, garnet and staurolite in the S1

foliation, presence of ilmenite and rutile in thematrix, and growth of andalusite after the F3 folding(Figs 3f & 4b,d). This succession of crystallization iscompatible first with the prograde path in the ilme-nite stability, from st–g–mu–chl–pl–ilm or st–g–mu–chl–pa–ilm field to st–g–mu–bi–pl–ilm–ru field(Fig. 9). Paragonite absence in the sample is compati-ble with inferred peak assemblage. Molar isopleths ofgarnet predict garnet growth up to the biotite-in line,then no garnet growth or its consumption, which isalso compatible with observed garnet zoning(sps0.16⟹0.09, XFe(0.91⟹0.89)) and calculated isopleths(m0.15⟹0.09, x0.91⟹0.89). Staurolite molar isoplethspredict staurolite growth up to the st–g–mu–bi–pl–ilm–ru stability field, also compatible with trend ofXFe in staurolite core (0.86⟹0.84). The XFe iso-pleths calculated in the st–g–mu–bi–pl–ilm–ru fieldare lower (x = 0.80) than the lowermost measuredvalues (XFe = 0.84), but as the rims of stauroliteshow increase in XFe (up to 0.88), this is interpretedas a result of diffusional reequilibration in the anda-lusite stability field, and therefore the staurolitechemistry from peak is not preserved. Andalusitegrowth after the F3 folding indicates the end of theP–T path in the andalusite stability field, likely in thest–mu–bi–pl–ilm–and field. The inferred P–T pathinvolves therefore first a prograde path from 4 to5 kbar and 550 °C to 6 kbar and 600 °C during theD1 deformation (Fig. 9). This is followed by regionalexhumation during the D2 which is not reflected by

Table 3. Representative mineral analyses for sample 12-09A.

Sample 12-09A (staurolite zone – Eastern Rehamna)

Analysis number 314 301 317 308 13 24 76 103

Mineral mu-matrix mu-matrix chl-matrix chl-matrix g-core g-rim st-core st-rim

wt%

SiO2 45.92 46.39 25.32 25.20 36.31 36.63 28.28 28.60

TiO2 0.21 0.28 0.00 0.00 0.00 0.00 0.42 0.51

Al2O3 36.74 36.23 23.36 23.07 20.80 21.05 55.28 55.47

FeO 0.84 0.74 23.76 24.08 31.50 33.70 12.71 12.24

MnO 0.00 0.00 0.28 0.26 6.25 3.89 0.42 0.49

MgO 0.37 0.42 12.76 12.92 1.77 2.22 1.29 1.10

CaO 0.00 0.00 0.06 0.00 2.56 2.32 0.00 0.00

Na2O 1.57 1.33 0.00 0.00 0.00 0.00 0.00 0.00

K2O 9.07 9.26 0.63 0.40 0.00 0.00 0.00 0.00

BaO (ZnO for st) 0.27 0.00 0.00 0.00 0.00 0.00 0.45 0.51

Total 94.98 94.64 86.16 85.94 99.17 99.81 98.86 98.91

Cation/Oxygen 7/11 7/11 20/28 20/28 8/12 8/12 15/23 15/23

Si 3.05 3.10 5.46 5.45 2.97 2.97 3.96 4.01

Ti 0.01 0.01 0.00 0.00 0.00 0.00 0.04 0.05

Al 2.88 2.85 5.93 5.88 2.00 2.01 9.14 9.16

Fe3+ 0.00 0.00 0.00 0.00 0.06 0.05 0.00 0.00

Fe2+ 0.05 0.04 4.28 4.35 2.10 2.23 1.49 1.43

Mn 0.00 0.00 0.05 0.05 0.43 0.27 0.06 0.06

Mg 0.04 0.04 4.10 4.16 0.22 0.27 0.27 0.23

Ca 0.00 0.00 0.01 0.00 0.22 0.20 0.00 0.00

Na 0.20 0.17 0.00 0.00 0.00 0.00 0.00 0.00

K 0.77 0.79 0.17 0.11 0.00 0.00 0.00 0.00

Ba (Zn for st) 0.01 0.00 0.00 0.00 0.00 0.00 0.05 0.05

Total 7.00 7.00 20.00 20.00 8.00 8.00 15.00 15.00

XFe (total) 0.91 0.90 0.85 0.86

Xalm 0.71 0.76

Xprp 0.07 0.09

Xsps 0.14 0.09

Xgrs 0.07 0.07



mineral growth in this area. The exhumation duringthe D3 deformation reaches the chlorite stability asthe chlorite and fine muscovite dominate the S3cleavage. It is therefore likely, that decompressionfirst involves decompression with cooling followed byheating to the andalusite field at ~2.5–4 kbar and530–570 °C rather than isothermal decompression(Fig. 9).

MONAZITE U–Pb IN SITU GEOCHRONOLOGY

Analytical procedure

Prior to analysis, monazite grains were examinedwith backscattered electron and secondary electron

imaging with a TESCAN VEGA scanning electronmicroscope equipped with an X-Max 50 electron dis-persive spectrometer (Oxford Instruments), at theInstitute of Petrology and Structural Geology of theCharles University (Prague, Czech Republic). Theoperating conditions were 15 kV accelerating voltageand 1.5 nA beam current. Monazite U/Th–Pb datawere acquired at the Geosciences department of theUniversity of Montpellier (Montpellier, France), withthe in situ technique ‘LASER-ICP-MS’ using aLambda Physik CompEx 102 Excimer laser coupledto an Element XR sector field ICP-MS (inductivelycoupled plasma-mass spectrometry). Laser spot sizeswere fixed to 26 lm. The laser was first fired to cleanthe sample surface (10 pulses) and then material was

0.5

0.6

0.7

0.8

0.9

0.1

0.2

0.3

Mol

e %

in the matrix

423 µm

included in staurolite

+chl +q

alm prp sps grs XFe

Central Rehamna Eastern Rehamna

SA-32(Staurolite zone)

12-09A(Staurolite zone)

618 µm+st +st

(b) (c)

825 µm+q +chl

(a)

Eastern Rehamnain the matrix

12-07II(Garnet chloritoid zone)

S1b

S1a

200 µm S1S3

(d) (e) (f)

100 µm200 µm

g

g

g

SA-3212-07II 12-09A

st

ilmilm

q

q

Fig. 6. Garnet profiles from samples 12-07II. (a), SA-32 (b) and 12-09A (c) and corresponding back-scattered electron images (d–f). Location of profiles indicated by arrows, dashed lines indicate foliation.



ablated during 45 s under ultrapure helium in a15 cm3, circular shaped cell with an energy densityof 10 J cm�2 at a frequency of 3 Hz. Each analysisconsisted of a 15 s period devoted to measure thebackground, followed by measurement of the sampleand one minute of washout in order to purge thecell. Unknowns were bracketed by analyses ofManangotry monazite standard (Montel et al., 1996)used for inter-element fractionation and for massbias factor corrections. The ages and the 238U/206Pband 207Pb/206Pb isotopic ratios for each analysesplotted on Tera-Wasserburg diagrams (Fig. 10;Table S2) were calculated using Isoplot/Ex (Ludwig,2003).

U–Pb results

The only samples with monazite grains without numer-ous inclusions and with diameter >15 lm were fromthe Central Rehamna (samples SA-32) and the EasternRehamna (sample T-1) (sample locations in Fig. 2a).Other rare monazite-bearing samples contain grains

<20 lm or larger xenoblastic monazite with numerousinclusions. All dated grains occur in the matrix; nomonazite included in garnet or staurolite was found.In staurolite–garnet micaschist sample SA-32

described above, monazite is elongated parallel withthe S1 fabric. Seventy-five spots were performed on25 grains (Fig. 10a). Three spots are significantlyolder than the remaining population and yield a206Pb/238U weighted mean age of 298.3 � 4.1 Ma(MSWD = 0.53, n = 3). The other data yield a scat-tering of 206Pb/238U ages ranging from 268 � 3 Mato 292 � 6 Ma. This spread is attributed to variousproportions of common lead. In a Tera-WasserburgConcordia plot, (Fig. 10a) all data points fall on aline (MSWD = 1.7) defining an age of277.0 � 1.1 Ma (n = 72). The latter is attributed tothe metamorphic crystallization of monazite.The micaschist sample T-1 is composed of st–g–

mu–chl–and–ilm–q; the S1 foliation is folded andreworked by the F3 folding leading to developmentof the S3 cleavage. Garnet and staurolite containinclusion trails parallel with the S1 fabric, andalusite

0.83

0.84

0.85

0.86

0.72

0.74

0.76

0.78

0.80

0.03

0.04

0.05 0.05

0.02

0.03

0.04

0.05

0.020.02

0.03

0.04

0.85

0.86

0.87

0.88

0.890.87

Mole %

(a) (b) (c)

Mole % Mole %

XMnXFe XMnXFe XMnXFe

268 µm+q +q

5118 µm+q +mu

12-09A (Staurolite zone)12-07II (Garnet chloritoid zone) SA-32 (Staurolite zone)Central Rehmana

stauroliteEastern Rehamna

chloritoidEastern Rehamna

staurolite

+mu +ilm4242 µm

1 mm

SA-3212-07II 12-09A

1 mm

ilm

ilm

ilm

mug

ctd

stst

S3

S1

S1aS3 S1

(d) (e) (f)

Fig. 7. Chemical profiles: (a) chloritoid in sample 12-07II, (b) staurolite in sample SA-32 and (c) staurolite in sample 12-09A; (d–f)corresponding back-scattered electron images, location of profiles indicated by arrows, dashed lines indicate foliation.



contains folded inclusion trails and is interpretedas postdating the S3 fabric. Fifty-five spots were per-formed on 20 monazite grains. In the Tera-Wasser-burg Concordia diagram, the data yield a continuumof 206Pb/238U ages ranging from 269 � 4 Ma to292 � 6 Ma (Fig. 10b). This spread is identical to

that observed for sample SA-32 and is interpreted asparticipation of a common lead component in theanalyses. The analysed monazite grains form a singlepopulation and yield a discordia age of 275.8 �1.7 Ma (MSWD = 1.2, n = 53). The latter is withinerror of the age of sample SA-32.

Metapelite SA-32SiO2 Al2O3 CaO MgO FeO

69.00 14.01 1.05 3.73 6.31K2O Na2O TiO2 MnO O2.35 2.60 0.77 0.17 0.01

7

123456

8910111213

141516

st g bi pa pl rust g bi pl ru sillst g bi pa pl ilm rug bi pa pl ilm rug bi pa ilm rug bi chl pa ilm rug bi chl pa pl ilm rubi chl pa pl ilmbi chl pa pl ilm rug bi chl pl ilm rug bi chl pl rubi chl pl rubi chl pl ilm ru

st g bi chl pa pl ilmst g bi chl pl ilmst g bi pl ilm sill

7

123456

89

10111213

141516

20

1819

212223242526

17 st g bi pl ilm andst g bi chl pl ilm andg bi chl pl ilm andg bi pl ilm ksp sillg bi pl ilm ksp sill -mug bi pl ilm ksp cd sill -mug bi pl ilm ksp andg bi pl ilm ksp and -mug bi pl ilm ksp cd and -mug bi pl ilm ksp cd -mu

500 550 600 6502

4

6

8

T (°C)

P (k

bar)

P (k

bar)

2

4

6

8

500 550 600 650T (°C)

500 550 600 650T (°C)

(a)

(b) (c)

st g bi pl ilm

g bi pl ilm sill

g bi pl ilm and

g chl bi pa ru

g chl bi pa pl ru

g bi pa ru

g bi pa pl ru

st g bi pl ru

st g bi pl rustg

bi pa

pl ilm

g bipa pl ilm

g bi chl pa pl ilm

g bi chl pl ilm

bi chl pl ilm

7

1

23

456

891011

12

13

14

15

16

7

1

23

456

891011

12

13

14

15

16

20

1819

21

2223

24 2526

17

g

stst

0.040.080.

12

0.24

0.16

0.20

0.01

0.02

0.030.04

0.050.06

0.04

0.05

0.03

0.02

0.8

0.81 0.76

0.720.74

0.78

Garnet mol. %

m(g)*100z(g)*100

x(g)*100Staurolite mol. %x(st)*100

S1aS1b

g

chl pl

ru

ru

stst

ksp

ksp

st

st

pl

pa

and

+ q, mu, H20

Fig. 8. (a) P–T pseudosection for the metapelite sample SA-32 (whole-rock composition in moles adjusted to 100%), withmodelled compositional isopleths of garnet (b) and staurolite (c). The arrow indicates P–T–d evolution deduced from thecrystallization-deformation relationships and garnet and staurolite chemistry compared with the calculated mineral assemblagesand isopleths. Outcrop geographical coordinates (WGS84): 32�23’18.0“N, 7�57‘09.1”W. See text for further details.



DISCUSSION

Metamorphic evolution

Combining crystallization–deformation relationshipsinferred from textures, mineral chemistry and pseudo-section modelling allows us to constrain the prograde

and retrograde paths for samples from various meta-morphic zones (Fig. 11). The P–T paths from thestaurolite and staurolite–andalusite zones are basedon pseudosection modelling (Figs 8 & 9), the P–Tpaths from chloritoid and andalusite zones areinferred qualitatively, based on relative stability ofchloritoid and andalusite (e.g. Ko�suli�cov�a & �St�ıpsk�a,

2

4

6

8

P (k

bar)

500 550 600T (°C)

500 550 600T (°C)

500 550 600T (°C)

2

4

6

8

P (k

bar)

Metapelite 12-09A(a)

(b) (c)

SiO2 Al2O3 CaO MgO FeO63.06 21.55 0.40 2.84 6.22

K2O Na2O TiO2 MnO O3.69 1.00 1.07 0.16 0.01

chl ctd pa ru

st chl ctd pa ru

st gchl

pa ru

st g chl ru

st g chlilm ru

st gbi ru

st g bipl ilm ru

st g bipl ilm

st bi pl ilm

g bi pl ilm and

st g chlpl ilm

st chlpl ilm

st g chlpa ru

st g chl pa

ilm ru

st chlpl ilm

st chl pl ilm and

stg

bi chl

pl ilm

stbi

chl pl

ilm st chlpa ru

st g chl ilm

st g chl

pa ilm

7

1

2

34

5 6

8

91011121314

15 16

17

2018

19

2122

2324

2526

2728

29

303132

20

1819

212223242526272829303132

7

123456

89

10111213141516

17st ctd chl rust g chl bi rust g bi chl ilm rust g bi pl rust g chl bi ilmst g bi ilm rust g bi chl pl ilm rust g chl pa pl ilmst g chl pl ilm rust g chl pa pl ilm rust chl pa ilm rust chl pa pl ilm rust g chl pa pl rust g chl pa pl ilm rust g chl pa ilm rust g chl pl ru

st chl pa pl rust chl pa pl ru andst chl pa ru andst chl pl ru andchl pl ilm ru andst chl pl ruchl pl ilm andst chl pl ilm ru andst g bi pl ilm sillst bi pl ilm sillst g bi pl ilm andst bi pl ilm andst g chl bi pl ilm andst chl bi mu pl ilm andst g chl mu pl ilm andg bi mu pl ilm and

ilm

ru

pa

plpa

ilmru

ctd

g

st

ctd

stg

and silland

g

ru

g

sillg

ru

bi

bich

l

chl

and

pl

+ q, mu, H20

+and

g

g

0.1

0.05

0.01

0.02

0.03

0.06

0.04

0.82 0.

80

0.90

0.88

0.86

0.80

0.84

0.82st st

S1

S3

post S3

S1

S3

post S3

S1

S3

post S3

0.050.100.150.200.25

0.30

0.35

0.15

0.09

0.89

0.91

Garnet mol. %m(g)*100x(g)*100

Staurolite mol. %x(st)*100

Fig. 9. (a) P–T pseudosection for the metapelite sample 12-09A (whole-rock composition in moles adjusted to 100%) withmodelled compositional isopleths of garnet (b) and staurolite (c). The arrows indicate P–T–d evolution deduced from thecrystallization–deformation relationships and garnet and staurolite chemistry compared with the calculated mineral assemblagesand isopleths. Outcrop geographical coordinates (WGS84): 32�21‘05.9“N, 7�51‘46.0”W See text for further details.



2007). The beginning of the P–T paths in the stauro-lite zone is recorded by garnet core compositions at500–550 °C and ~4 kbar (Fig. 11a,b), garnet molarisopleths and type of zoning indicate a prograde P–Tpath into the staurolite stability field. Entering thisfield, the garnet molar isopleths are subparallel withthe prograde P–T path suggesting cessation of garnetgrowth. This is compatible with the different charac-ter of the inclusion trails in garnet, which are perpen-dicular to the S1 inclusion trails in staurolite. Basedon this observation, we suggest that garnet growthoccurred during early stages of burial, in a subverti-cal fabric (S1a). The S1 inclusion trails in stauroliteare parallel with the external subhorizontal S1 fabric,or in some crystals they are curved. Therefore, wesuggest that the beginning of staurolite growth coin-cided with transition of the subvertical fabric to thesubhorizontal dominant foliation (D1b phase). Thepeak conditions are determined by mineral assem-blage and staurolite rim (or near rim) composition to6–6.5 kbar and 600–650 °C, and are in accordancewith previous thermobarometric studies (Aghzer &Arenas, 1995). Some rocks in the staurolite zone havedeveloped S3 cleavage with chlorite and muscovite,indicating decompression accompanied by cooling(Fig. 11a,b). These samples show staurolite inclusiontrails sometimes parallel, oblique or even orthogonalto the external S1 foliation (Fig. 4). Such amicrostructural pattern has in many places beeninterpreted as a result of polyphase pre-D3 deforma-tion-porphyroblast growth (Bell & Sanislav, 2011;e.g. Bell et al., 2013). These geometrical relationshipsare interpreted as a result of rigid body rotation of astaurolite porphyroblast within an incompetent weakmatrix during the F3 folding. Samples which are not

affected by the S3 cleavage have staurolite internalinclusion trails parallel to S1, and we suggest thatthey represent passively exhumed low strain domainsthat are surrounded by zones of D3 deformation(Fig. 11a). There are also areas within the staurolitezone, where the S3 cleavage is marked by fine-grainedmuscovite and chlorite, and is statically overgrownby andalusite. This microstructural succession of thechlorite–muscovite S3 cleavage overgrown byandalusite indicates that rocks underwent firstdecompression with cooling to the stability field ofchlorite, followed by heating to the andalusite stabil-ity field (Fig. 11b). In the chloritoid zone, the inclu-sion trails in chloritoid are in place both curved andstraight and at high angle to the external S3 fabric(Fig. 11c). We suggest that they reflect chloritoidgrowth during D1 burial, and possibly reflect thechange from subvertical shortening to subhorizontalflow, as in the case of transition from garnet to stau-rolite in the staurolite zone. The presence of the S3cleavage with chlorite and muscovite is interpreted asan indication of syn-D3 exhumation and cooling. Inone area andalusite is found with inclusion trails athigh angle to the external S3 fabric (Fig. 11d), whichis interpreted as andalusite growth within the subhor-izontal S1 fabric. This may imply that some areashad localized higher thermal gradient during orbefore main D1 event. A schematic P–T path basedon andalusite stability shows cooling associated withD3 deformation (Fig. 11d).Those results are similar to the metamorphic evolu-

tion of the southern Jebilet massif, where a pseudo-section for a staurolite–garnet micaschist gave peakconditions of 4–5 kbar and 560–585 °C for the regio-nal Barrovian metamorphism and up to 650 °C at

0.06

0.08

0.10

0.12

0.14

20 21 22 23 24 20 21 22 23 24 24

310

320 290

2700.

060.

080.

10

320 300 260

280

(a) (b)

275.8 ± 1.7 Ma(MSWD = 1.2, n = 53)

298.3 ± 4.1 Ma (MSDW = 0.53, n = 3)

238U/206Pb 238U/206Pb

207 P

b/20

6 Pb

207 P

b/20

6 Pb

277.0 ± 1.1 Ma(MSWD = 1.7, n = 72)

SA-32 T-1

Fig. 10. Tera–Wasserburg concordia plots for laser-ablation ICP-MS analyses of monazites from metapelites samples SA-32 (a)and T-1 (b). Error ellipses are 1r. Outcrop geographical coordinates (WGS84) are 32�23‘18.0“N, 7�57‘09.1”W and 32�21‘03.3“N,7�51’45.1”W for the samples SA-32 and T-1 respectively.



1.5–3.5 kbar (combined Raman Spectroscopy of Car-boniferous Material and pseudo-section) for the lateto post tectonic contact metamorphism affectingheterogeneously this massif (Delchini et al., 2016). Inthe Rehamna massif, rare occurrence of kyanite inthe matrix of the micaschists is reported by

Hoepffner et al. (1975, 1982), El Mahi et al. (1999)and Baudin et al. (2003). During this study, kyanitewas found only in quartz-kyanite veins that precipi-tated from circulating fluid (El Mahi et al., 2000). Insuch a case, P–T conditions for this kyanite forma-tion cannot be determined from the metapelite

Fig. 11. Sketches show interpretative petrological successions and inferred P–T–d paths based on observations in individualsamples. The orientation of structures shows an attempt to interpret structural development of mineral assemblages during D1, D2and D3 events. (a) For documentation of P–T see Fig. 8, for relation with D1 see Figs 3d,e and 7e, for relation with D3 seeFig. 3a, and for absence of S3 in low strain domains see Fig. 3d,e. (b) For documentation of P–T see Fig. 9 and for relation withstructural development see Figs 3f, 4b,d and 6f. (c) The character of inclusion trails at high angle with external S3 in chloritoidsuggests its growth during burial connected with D1 deformation (Fig. 4c). P–T path is only qualitatively inferred based onchloritoid stability (e.g. Ko�suli�cov�a & �St�ıpsk�a, 2007). (d) The character of inclusion trails in andalusite at high angle to theexternal S3 foliation (Fig. 3b) indicates its growth in S1 foliation. P–T path is only qualitatively inferred based on stability ofandalusite.



pseudosections. In the absence of other P–T determi-nations, as for example, study of fluid inclusions, P–T conditions of qtz–ky formation can be anywhere inthe kyanite stability field, including for example,greenschist facies conditions. It remains thereforeunclear whether all kyanite crystallization in thematrix of the micaschists is related to peak metamor-phic conditions and thus could indicate a local occur-rence of a classical kyanite zone, or whether itscrystallization is related to fluid circulation throughthe matrix of the micaschists contemporaneous withformation of quartz–kyanite veins, or even whetherthere occur both types of kyanite with different P–Tsignificance.

Interpretation of the monazite dates

An attempt is now made to correlate the monazitedates (Fig. 10) with the inferred P–T paths. Thereare numerous studies that show empirical correlationof monazite presence or absence, abundance andshape with metamorphic grade (Kingsbury et al.,1993; Rubatto et al., 2001; Wing et al., 2003; Gasseret al., 2012; �St�ıpsk�a et al., 2015). The first metamor-phic monazite appears at the garnet isograd, and israre, small and irregular in shape; it is slightly moreabundant, but still small at the staurolite isograd,and it is abundant, large and may be euhedral at thekyanite, sillimanite or andalusite isograd. It has beenshown in natural examples (Ayers et al., 1999; Poi-trasson et al., 2000; Townsend et al., 2001; Martinet al., 2007; Kelly et al., 2012; Seydoux-Guillaumeet al., 2012; Didier et al., 2013) and in experiments(Seydoux-Guillaume et al., 2002; Williams et al.,2011) that monazite is easily susceptible to recrystal-lization, especially in the presence of fluids and thatthe isotopic ratios may be partially or completelyreset at temperature far below 800 °C. According tothe above described empirical correlation of monaziteabundance and size with metamorphic grade in meta-pelites, it is proposed that in the studied samplessome monazite grew on the prograde path in the gar-net and staurolite stability fields (e.g. Smith & Bar-reiro, 1990), but it was not abundant, while afterentering the andalusite stability field, most monazitehas grown (sample T-1). The presence of abundantchlorite in sample SA-32 suggests that monazite mayhave recrystallized along the retrograde P–T path togreenschist facies conditions. Consequently, the mon-azite ages in both samples reflect the end of the P–Tpath, for sample T-1, it represents the entering theandalusite stability field, and for sample SA-32 enter-ing in the greenschist facies. Therefore, we interpretthe mean date of 277 Ma obtained for both samplesas approximating the exhumation to low pressure. Insuch a context we interpret the oldest dates of c.298 Ma to reflect the peak of the burial, when somemonazite was likely to grow contemporaneously withthe staurolite.

Tectonic setting of Morocco Palaeozoic metamorphism

The Barrovian metamorphic sequences are classicallydescribed and interpreted in the European Variscanbelt to result from Lower Carboniferous (350–340 Ma) continental thickening (e.g. Rubio Pascualet al., 2013 in the Iberian Variscides; Oliot et al.,2015 in the Maures massif) related to the main colli-sion of the Gondwana derived blocks with Laurussia.The metamorphic isograds are commonly tectonicallyinverted due to continental underthrusting and imbri-cation of deeply buried continental crust (Burg et al.,1989; �St�ıpsk�a et al., 2015) or buried during crustalthickening (Rubio Pascual et al., 2013).In contrast in the Rehamna the metamorphic zones

are not inverted, but are disposed around a dome likestructure (Baudin et al., 2003). In addition, the ageof metamorphism is not Lower Carboniferous butUpper Carboniferous to Lower Permian (310–280 Ma; monazite ages of this work; Huon et al.,1987; Chopin et al., 2014).In Morocco, the deformation and metamorphism

spatially coincide with lower Carboniferous NE-SWstriking intracontinental basins (Piqu�e & Michard,1989; Beauchamp et al., 1991; Hoepffner et al., 2005;Michard et al., 2010) that are present in the WesternMeseta (Fig. 12), and for which many tectono-sedi-mentary reconstructions have been proposed (Bouab-delli & Piqu�e, 1996; Chakiri & Tahiri, 2000; ElKamel & El Hassani, 2006; Piqu�e et al., 2007). In theCentral Massif, two basins (the Sidi Bettache-Tilouine and Fourhal-Azrou-Kh�enifra basins) areseparated by a topographic high (the Za€er rise). Theymerge to the south with the Mechra ben Abbou basinand Lalla Tittaf and Dalaat units in the Rehamna,and terminate into the Sahrlef and Kharouba basinsin the Jebilet for which Vis�ean-early Serpukhovianage was proposed by Playford et al. (2008) based onpalynostratigraphy of black shales. Basin formationis associated with (hydro-) thermal alteration and/oreruption of mafic lavas massifs and felsic to maficintrusions in the Moroccan Central (e.g. Roddazet al., 2002; Bamoumen et al., 2008), Rehamna (e.g.Remmal et al., 1997; Baudin et al., 2003) and Jebilet(e.g. Essaifi & Hibti, 2008; Lotfi et al., 2010). Theyare interpreted as a result of intraplate magmaticactivity during tectonic extension (Kharbouch, 1994;Bamoumen et al., 2008; Lotfi et al., 2010) (Fig. 12).Indeed, a thermal event dated at 330–320 Ma in theCentral massif (K–Ar isotopic dating of a clay frac-tion from slates, Huon et al., 1987), the emplace-ments of magmatic bodies such as cordierite-bearinggranodiorites at c. 327 Ma (Rb–Sr isochrons, Mriniet al., 1992), a granitic sill dated at c. 330 Ma (U–Pbzircon, Essaifi et al., 2003) and formation of massivesulphide deposits at c. 332 Ma (40Ar/39Ar method onhydrothermal sericite, Marcoux et al., 2008) are coe-val with the deposition of late Visean (Asbian) sedi-ments, including the Lalla Tittaf formation in the



D3

D2

D1

Isograds in metapelitesgst bi

D0

Lalla Tittaf basin

Saxonian(295−280 Ma)

Westward convergence

600km

N

D3

Westphalian to Autunian(310−295 Ma)

Southward convergence

600km

N

D1

Upper Visean to Lower Namurian(335−320 Ma)

Intracontinental basins opening

Intracont. basins, magmatismPre-Carboniferous basement

600km

N

ZRZRZR

SBBSBBSBB

JJRR

CMCMCM

SBSBSB

HABHABHAB D0

FBFBFB AKAKAK

Fig. 12. Maps and three-dimensional idealized model for the evolution of the Moroccan Meseta and the Rehamna massifrespectively (see text for discussion). SB, Sehoul block; CM, Central Massif; SBB, Sidi Bettache-Tilouine basin; ZR, Za€er rise; FB,Fourhal-Azrou-Kh�enifra basin; R, Rehamna; J, Jebilet; HAB, High-Atlas block. Open circles: location of main cities.



Rehamna massif. High-grade metamorphism reachinggranulite facies at 328 Ma below the Upper Viseanto Lower Namurian intracontinental basins is alsoinferred from xenoliths hosted by Triassic lampro-phyres from the Jebilet massif (Dostal et al., 2005).All of this supports a rift-related magmatic and meta-morphic activity together with a high geothermal gra-dient at c. 335–320 Ma. These intracontinental basinswere recently interpreted by Frizon de Lamotte et al.(2013) to extend along the whole north Gondwanamargin and testify for the opening of the Palaeo-Tethys ocean (Stampfli et al., 2013).

It can be concluded that when the internal part ofthe main European Variscan belt was tectonicallythickened and experienced peak metamorphismrelated to continental thickening, the Moroccan Var-iscides suffered an extensional event and intraconti-nental rifting related to stretching of the northGondwana margin (see also Frizon de Lamotte et al.,2013). The convergent deformation of the northGondwana Palaeo-Tethysian passive margin is char-acterized by important N–S shortening event (Chopinet al., 2014; Baidder et al., 2016) at 310–295 Ma(Fig. 12), which resulted in formation of theRehamna metamorphic dome (D1 and D2 events ofChopin et al., 2014). Subsequent change in plate con-figuration during Lower Permian resulted in E–Wsubsequent shortening of the intracontinental Mesetadomains within the Gondwana margin (Engelder &Whitaker, 2006; Hatcher, 2010; Chopin et al., 2014;Edel et al., 2015) (Fig. 12). In summary, the tectonicsetting of Moroccan metamorphism exploits inheritedthermal and tectonic anomalies at the margin of theGondwana continent, which is in contrast to classicalsyn-collisional European Variscan metamorphismthat is emblematic for Himalayan type collisionalprocesses.

Thermal and mechanical aspects of the Moroccanintracontinental Variscan event

Metamorphic evolution of the Rehamna massif isconsistent with the theoretical scenario of thickeningof thermally softened rifted continental region. Theearly syntectonic andalusite (Sample 12-02F) couldhave developed during an extensional event, which isconsistent with observations of an early hot regimeby Essaifi et al. (2014) from the nearby Jebilet massif.Subsequent a Barrovian metamorphic gradient devel-oped along a field geotherm, which is slightly hotter(26–30 °C km–1) compared to classical Variscan colli-sional geotherms (14–20 °C km–1) as exemplified byKo�suli�cov�a & �St�ıpsk�a (2007). The Barrovian meta-morphic isograds today disposed along the marginsof the Rehamna massif indicate that they have beensub-horizontal and reflect a normal, that is, non-inverted metamorphic gradient. The subsequent ret-rograde path is related to doming phase and reflectsprogressive cooling of the whole domain as the heat

budget of the crust is consumed by the orogenicevent (Fig. 12).A characteristic feature of the Rehamna massif is

the telescoping of metamorphic isograds (Fig. 1).This feature is classically interpreted in this area asan indication of extensional thinning associated withgravitational collapse, subsequent to orogenic thick-ening (Aghzer & Arenas, 1995, 1998). However, thepost-thickening extensional fabrics were not reportedin the studied region, neither close to the main faultzones where the distance of isograds is particularlyreduced. It was shown in other regions of the Varis-can belt that the metamorphic isograds can be juxta-posed and thinned by upright folding (e.g. �St�ıpsk�aet al., 2012) and by thrust related shearing (e.g.�St�ıpsk�a & Schulmann, 1995). Therefore, we suggestthat in the Rehamna, the isograds were telescopedfirst by doming related to N–S shortening and subse-quent E–W shortening, which juxtaposed isogradsalong shear zones that reactivated margins of theinverted Devonian–Carboniferous basin.Even if the Rehamna metamorphic dome does not

occur in the middle of the continent, it is located out-side the continental boundary and surrounded bynon-metamorphosed continental blocks. The LateCarboniferous collisional boundary of north-Gond-wana with the Variscan belt is probably hidden inthe Rif Mountains in northern Morocco, 350–400 km north of the Rehamna Massif. Here, HPgranulites show metamorphic conditions typical forthe Himalaya-Tibetan type thickening (Bouybaou�eneet al., 1998) and were dated at 284 � 27 Ma (Montelet al., 2000), reflecting well the metamorphic age ofthe Rehamna massif presented in this work. Theimplication thus, is that the Rehamna and Jebiletmassifs were thickened far from any collisionalboundary in the Late Carboniferous and can serve asa proxy for intracontinental orogenic systems as pro-posed by Raimondo et al. (2014).In the case of the Rehamna massif, the cause of

the orogenic event should be correlated with aninherited structural and thermal anomaly developedduring (Devonian-) Lower-Carboniferous intraconti-nental rifting (Piqu�e et al., 1980; Piqu�e, 1981; Cor-sini, 1991; Michard et al., 2010) (Fig. 12). It wasproposed that the thermal softening of crust relatedto rifting can localize subsequent moderate thickeningof crust, when the previously rifted region is short-ened (Thompson, 1989; Michard et al., 2010). Theeffect of thermal weakening of the lithosphere on thethickening processes was examined by Thompsonet al. (2001) and Schulmann et al. (2002) whoshowed that the thermal budget of the rifted regionpersisted for c. 30–40 Ma. In Morocco, the time lagof thermal relaxation between the rifting at 335–320 Ma and the onset of collision at 310 Ma (thiswork and Chopin et al., 2014) is a maximum of25 Ma. This implies that the crust beneath theRehamna, Jebilet and other rifted regions in



Morocco (Fig. 1) was sufficiently hot to localizedeformation during the Upper Carboniferous conver-gence (Fig. 12). In contrast, other regions surround-ing thermally weakened domains like the Coastalblocks or the Anti-Atlas behaved like rigid domains,where limited deformation and metamorphismoccurred. Thompson et al. (2001) quantified pro-cesses of reactivation of rifted domains and showedthat their thickening can be only moderate, and isinevitably followed by extrusion-doming phase as thethickened domains cool and rigidify.

The Permian re-configuration of plates is a giantprocess that influenced deep fluxes in the non-relaxedlithospheric mantle and asthenosphere and resultedin Permian mafic magmatism and volcanism in south-ern Europe (e.g. Arthaud & Matte, 1977; Carreras &Druguet, 2014; Den�ele et al., 2014; Casini et al.,2015). This event is reflected by last thermal pulseresulting in syn- to post-D3 magmatism at c. 280–270 Ma and associated contact metamorphism inMorocco (Diot & Bouchez, 1989; Mrini et al., 1992;El Hadi et al., 2006; Chopin et al., 2014; Delchiniet al., 2016) (Fig. 12).

CONCLUSIONS

The Rehamna case study shows that the inheritedpost-rift thermal anomaly is a first order cause forformation of orogenic zone in the interior of previ-ously extended continent. Our study shows that thetime lag of 10–25 Ma between the rifting and theonset of shortening (from 335–320 to 310 Ma) keptthe crust hot enough to generate a moderately thick-ened domain characterized by a classical Barrovianmetamorphic sequence. The crustal thermal budgetwas sufficient to keep the orogenic domain tectoni-cally active for a further 30 Ma (from 310 to 280 Ma)even during superimposed polyphase tectonic events.

ACKNOWLEDGEMENTS

O. Bruguier from the University of Montpellier isacknowledged for supervising the U/Pb dating and M.Racek is thanked for providing valuable assistance atthe Scanning Electron Microscope Laboratory at theCharles University of Prague. J.-F. Ghienne is thankedfor helpful discussions. We gratefully acknowledge thefinancial support of the Czech National Grant Agency(13-16315S to P.S.) and of the Ministry of Educationof the Czech Republic (grant LK11202 to K.S.). Weare grateful to R. Arenas and G. Cruciani for detailedconstructive reviews. D. Robinson is thanked for hiscomments and editorial work.

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SUPPORTING INFORMATION

Additional Supporting Information may be found inthe online version of this article at the publisher’sweb site:Appendix S1. Whole rock calculations.Figure S1. BSE image of the thin section of sample

12-09A.Figure S2. BSE images of scanned areas from sam-

ple 12-09A.Figure S3. BSE image of the thin section of sample

SA-32.Figure S4. BSE images of scanned areas from sam-

ple SA-32.Figure S5. BSE image of the thin section of sample

12-07II.Figure S6. BSE images of scanned areas from sam-

ple 12-07II.Table S1. Compositions of scanned areas from

sample 12-09A (a) shown in Figs S1 and S2, fromsample SA-32 (b) shown in Figs S3 and S4 and fromsample 12-07II (c) shown in Figs S5 and S6 and cal-culated means.Table S2. Laser-ablation ICP-MS analyses of mon-

azites from metapelite samples SA-32 (a) and T-1 (b).

Received 2 March 2016; revision accepted 26 July 2016.



Tectonometamorphic evolution of an intracontinental ... · deformation, metamorphism and magmatism devel-oped in the interior of continental blocks away from active convergent plate

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