Lithofacies and biostratigraphical correlation of marine Carboniferous ...

Facies (2014) 60:941–962 DOI 10.1007/s10347-014-0409-1

ORIGINAL ARTICLE

Lithofacies and biostratigraphical correlation of marine Carboniferous rocks in the Tindouf Basin, NW Africa P. Co´zar • A. Garcı´a-Frank • I. D. Somerville • D. Vachard • S. Rodrı´guez • P. Medina-Varea • I. Said

Received: 26 February 2014 / Accepted: 5 July 2014 / Published online: 25 July 2014 Ó Springer-Verlag Berlin Heidelberg 2014

Abstract Spatial and temporal variations of Carboniferous sediment accumulation within the northwestern part of the northern flank of the Tindouf Syncline in Saharan Morocco allowed to distinguish 16 lithofacies types. The predominant sedimentation pattern is cyclic, with the overall succession recording a major regressive trend. Outer platform siliciclastics in the lower part (Tournaisian and Vise´an) pass up to middle and inner platform mixed siliciclastic and carbonate sediments (late Vise´an–Serpukhovian) and finally to continental sandstones in the Bashkirian capping the marine carbonate sedimentation. The lack of similarities in a correlation with southern outcrops in the Tindouf Syncline suggests tectonically controlled sedimentation. The upper Tournaisian to lower Bashkirian succession records the incipient uplift of the

Anti-Atlas Mountains, changing the paleogeography and, therefore, affecting the paleoecologic conditions, as well as the sedimentary environments in the Tindouf Basin. It is suggested that from the Serpukhovian onwards, much of the Anti-Atlas was uplifted, leading to subaerial conditions, while during the late Vise´an, only a few small inliers had emerged. Although the number of Proterozoic emergent inliers of the Anti-Atlas is unknown, during the late Vise´an, the Anti-Atlas Mountain belt is regarded as an emerging structure, with a distinct influence on the paleobiogeography of the region.

P. Co´zar (&) Instituto de Geociencias (CSIC, UCM), c/Jose´ Antonio Novais 2, 28040 Madrid, Spain e-mail: [email protected]; [email protected]

Introduction

A. Garcı´a-Frank  S. Rodrı´guez  P. Medina-Varea Departamento de Paleontologı´a, Universidad Complutense de Madrid, c/Jose´ Antonio Novais, 12, 28040 Madrid, Spain I. D. Somerville UCD School of Geological Sciences, University College Dublin, Belfield, Dublin 4, Ireland D. Vachard Universite´ de Lille 1, UMR 8217 du CNRS: Ge´osyste`mes, Baˆtiment SN 5, 59655 Villeneuve d’Ascq, France I. Said Division du Patrimoine Ge´ologique (DPG), Direction du De´veloppement Minier (DDM), Ministe`re de l’E´nergie, des Mines, de l’Eau et de l’Environnement, Rue Abou Marouane Essadi BP: Rabat Instituts, Haut Agdal, 6208 Rabat, Morocco

Keywords Carboniferous  Morocco  Microfacies  Depositional environments  Tectonics  Biostratigraphic correlation

Carboniferous sedimentary rocks are well known in northern Africa, particularly those in the Sahara, which crop out in basins from Mauritania to Libya, south of the Anti-Atlas Mountains and the Atlas Transform Fault (Fig. 1a). They are part of the so-called Saharan or North African Platform (e.g., Guiraud et al. 2005). Details of the sedimentary facies in these basins is generally scarce and commonly restricted to unpublished reports of oil companies (see references in Peterson 1985, 1986; Boote et al. 1998; Guiraud et al. 2005; Villeneuve 2005). In fact, most of the published data are included in the mapping by national research projects during the 1950s–1970s. The largest Carboniferous outcrops include those of Taoudenni (Mauritania–Mali), Tindouf (Morocco–Algeria), Be´char (Algeria), Reggan (Algeria), Ahnet-Mouydir (Algeria), and Illizi-Ghamades (Algeria–Libya) (Fig. 1a).

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Fig. 1 a Location of the main Carboniferous basins in Morocco, Algeria, Libya, and Tunisia. 1 Central Massif (Azrou-Khenifra), 2 Rehamna, 3 Jebilet, 4 Tafilalt, 5 Be´char, 6 Jerada, 7 Tindouf, 8 Reggan, 9 Ahnet, 10 Mouydir, 11 subsurface Carboniferous, 12 IlliziGhamades, 13 Serdeles, 14 Taoudenni. AA Anti-Atlas, ATF Atlas

Transform Fault, EM Eastern Meseta, F Foum Defili, HA High Atlas, MA Middle Atlas, WM Western Meseta. b Location of the three stratigraphic sections in the Tinguiz Remz region. Dashed lines mark the approximate position of the formation boundaries

Fig. 2 a General view of the base of Section 2, starting in the Betaina Fm. (including limestone units 1 and 2, A–D), and the lower limestone units of the Ouarkziz Fm. (E–F); the thick sandstone ridge (base of sandstones) is highlighted, which is near the base of Fig. 3.

b General view of the middle part of Section 2 showing the cycles from limestone unit G (base and left) up to limestone unit L of the Ouarkziz Formation

The Mississippian marine succession of the Tindouf Basin is 1,800 m thick and consists of three formations: the Tazout Sandstone Fm. (oldest), Betaina Fm., and Ouarkziz Fm. (Figs. 1b, 2a) (Conrad 1985). The Tazout Sandstone Fm. comprises three sandstone units with rare interbedded shales. Its thickness is estimated to be 400 m (Conrad 1985). The Betaina Fm. is mostly composed of shales, with several intercalated thick sandstones in the upper part. Its thickness ranges from 700 to 1,100 m (Conrad 1985). Lenses with brachiopods and ammonoids have been

recorded throughout the formation (Hollard and Jacquemont 1956), which allowed the formation to be dated as late Tournaisian in the basal part to late Vise´an at the top. Limestones are virtually unknown from this formation in the northern flank of the Tindouf Syncline, although Sebbar et al. (2000) documented numerous limestones and mixed siliciclastic–carbonate deposits within the southern flank. According to Villeneuve (2005), those beds were considered mostly as dolomitic limestones and sandstones. However, this study is mostly focused on the

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sedimentology of the limestones of the Ouarkziz Fm. in the northwestern sector of the northern flank of the syncline (Fig. 1). The main feature of the Ouarkziz Fm. in the Tindouf succession is the distinct cyclic pattern (Fig. 2), with a predominance of shales and marly shales intercalated with thin limestone units, labeled 2, 1 and A–V (Fig. 3). A similar cyclic succession has been observed in three logged sections in the northern flank of the syncline (Fig. 1b). Cyclic successions have been documented in Be´char, NW Algeria (e.g., Pareyn 1961), although their characteristics and origin are poorly known. Exceptions include the work of Bourque et al. (1995), who analyzed fourth-order sequences containing Vise´an mud-mound bioherms in the eastern part of Be´char (Grand Erg Occidental), and of Malti et al. (2008), who analyzed the Vise´an third-order cycles of the Ben-Zireg Anticline (northern Be´char). In the Zrigat Fm. of the Tafilalt Basin, Morocco, Klug et al. (2006) described high-frequency cycles in a predominantly shale succession, interrupted by occasional limestone beds. Despite their interpretation, it is difficult to assess whether those cycles correspond to fourth- or third-order cycles, or a mixture of both. In Ahnet-Mouydir (Central Algeria), the lack of contemporaneous rocks makes any detailed comparison difficult (Wendt et al. 2009). The origin for this cyclicity is controversial, because this Saharan region was involved, from the late Vise´an to the Pennsylvanian, in the collision between Gondwana and Laurussia, and the closure of the Rheic Ocean (e.g., Nance et al. 2010; Murphy et al. 2010). Tectonic activity related to the closure of the Rheic Ocean started earlier (Devonian) in the western basins, and progressed eastwards. The marine succession was terminated during the early Bashkirian in Tindouf and possibly also in Taoudenni, Mali (Co´zar et al. 2014a), and much later, during the Moscovian, in the eastern Saharan basins, such as Be´char, Reggan, and Illizi-Ghamades (Conrad 1985; Massa 1985). Several Proterozoic inliers separated those basins, which are composed of Pan-African and Variscan crystalline massifs, and in some cases, rejuvenated Pan-African structures during Variscan movements. However, the precise timing of tectonic activity is not constrained yet. This debatable issue can be exemplified by the unknown duration when the Moroccan Meseta to the north was juxtaposed against the cratonic Saharan platform to the south (e.g., Hoepffner et al. 2005; Michard et al. 2008). The present study includes refined stratigraphic correlations based on additional biochronological data, illustrating the spatial and temporal variations of sediment accumulation within the Tindouf Basin. This analysis suggests the presence of active tectonic pulses as early as in the late Vise´an. Moreover, the scarce biostratigraphic and lithofacies correspondence between the northern and

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southern flanks of the Tindouf Basin also suggests a marked facies asymmetry between both flanks that was most likely influenced by tectonics.

Geographic location of the stratigraphic sections Three stratigraphic sections (Fig. 1; see also Rodrı´guez et al. 2013a), up to 27 km apart, are the basis of the composite section (Fig. 3), where over 125 samples were collected and [1,000 thin-sections prepared for the microfacies analysis and biostratigraphy (foraminifers) (see Co´zar et al. 2014a for precise details on the position of the samples and thin-sections). The most representative section (Section 2 in Fig. 1; 28°240 2000 N, 9°240 0400 W) is situated along the Assa-Zag Road. There, limestone units 1–P (Fig. 3) were intensively studied and sampled, although the upper part of the section is considerably affected by pervasive dolomitization, and the youngest units are covered by Recent sands. Section 3 (28°260 0600 N, 9°130 4300 W) is located in the Tinguiz Remz valley, 17.5 km east of the road section, in the same area as that described by Mamet et al. (1966). In this section, exposure of limestone units O– T is better than in the road section, although units S1–U are difficult to access owing to military restrictions in this area. Approximately 9.5 km to the west of the road section, at Section 1 (28°210 2000 N, 9°290 0400 W) the youngest limestone units, V and U, as well as some of the oldest limestone lenses (labeled unit 2) near the base are exposed. The limestone units 1–O are recognized in all three sections, with remarkable uniformity in thickness, facies and fossil content throughout. Lateral changes are observed in limestone units B and C, poorly preserved to the east in Section 1, and limestone unit H2, only present in Section 1. In addition, sandstone beds are continuous but changing in thickness throughout the sections. The thickest sandstone intervals are observed in Section 2, whereas the basal sandstone levels (below limestone unit A) are thinner to the east (Section 3), and the higher sandstone levels (between limestone units 1–E) are thinner to the west (Section 1). Below the measured sections (not included in Fig. 3), the Betaina Fm. in the study area is composed of shales and siltstones, at least for a interval of 150 m, where bioclasts content and bioturbation are rare, and carbonate lenses or limestone beds were not recorded in the Tinguiz Remz area.

Lithofacies The dominant lithologies in the Ouarkziz Fm. are shales and limestones/dolostones, locally with some sandstones, mainly in the lower part (Fig. 3). The main variations are

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Facies (2014) 60:941–962 b Fig. 3 Composite stratigraphic section close to Tinguiz Remz

illustrating the main facies. In the central columns the percentage of the main bioclasts and other components in the limestone units are depicted; grey fills in the texture of the rocks represent totally or partly dolomitized carbonates; the two right-hand columns show environmental interpretation. sp supratidal, in intertidal, sb subtidal, I.R. inner ramp/shelf, M.R. middle ramp/shelf, O.R. outer ramp/shelf. Absolute ages based on radiometric data in Davydov et al. (2010)

observed in the carbonate facies (abbreviated as LstF), where up to nine facies types have been distinguished; the siliciclastic facies include five different types of shales (ShF) and two types of sandstones (SstF) (Table 1). Shales and marls Within the Ouarkziz Fm., shales are more abundant in the middle part, recording thicker intervals (up to 45 m thick) between the limestone horizons (H–M). The shales are also common in the upper part of the section, although those beds are poorly exposed or covered in the Tinguiz Remz area. They are described from northeastern outcrops around Foum Defili (Conrad 1972a). The younger part of the succession, from limestone unit M upward (Fig. 3), contains thinner shaly intervals, except for the covered interval (up to 50 m thick) of inferred shales below limestone unit S1 (Fig. 3). However, a general thinning-upwards trend of shale beds is observed in the entire Betaina Fm. (Fig. 2a). Shale facies ShF1 The shales are predominantly pale green to dark green in color, with intercalated 1–2-cm-thick siltstone levels. Intervals with shales and siltstones show a variable thickness, from 3 to 4.5 m, although they are predominantly thick intervals, (20–52 m; Fig. 2b). Thinner intervals contain a higher proportion of siltstones. These structureless shales are very fine-grained, lacking any fauna and contain rare plant remains. No sedimentary structures have been observed, except for the top part of the interval between beds I and J, where wave ripples were observed in the siltstones. This type occurs between limestone units D–L, Q–R and above V (Fig. 3). Shale facies ShF2 These fine-grained sandy shales are dark green in color and commonly interbedded with fine-grained sandstones in the lower part of the logged section between limestone units 2– D. Individual sandstone beds are up to 10 cm thick, generally 4–5 cm thick (Fig. 4a). Intervals containing this microfacies type are thinner than ShF1, and they vary between 2.8 and 5.3 m, and in the upper part, reach 13.5 m. Bioclasts, mainly brachiopods and crinoids are rare, as are

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plant fragments. Sedimentary structures such as horizontal lamination, cross-stratification, and hummocky crossstratification are recognized in the sandstones. Bioturbation is locally recorded. Shale facies ShF3 Shales and marls. In the upper part of the lower member of the Ouarkziz Fm. (between limestone units L–Q), finegrained, pale green shales are interbedded with reddish and yellowish marls (Fig. 4b). They occur in progressively thinner intervals, decreasing from 39 m in thickness below bed M, to 4.8 m below bed Q. They lack bioclasts and sedimentary structures. The only structure is a horizontal banding marked by differences in color and in the proportion of calcareous matrix in the marls (Fig. 4b). Shale facies ShF4 Dark bioturbated grey shales and siltstones interbedded with fine- to medium-grained sandstones beds (10–30 cm thick). They are common in the upper part of the Betaina Formation, below limestone unit 2 (Figs. 2a, 3, 4c). They contain a higher proportion of sandstone beds intercalated between the shales than in ShF2. Intense bioturbation with common surficial trace fossils (Fig. 4f) as well as burrows (Fig. 4d, e) are the most significant feature of this facies. The fauna (mostly brachiopods, crinoids, gastropods, and rare ammonoids) and plant remains are common. Sedimentary structures are restricted to rare wave ripples in sandstone beds (Fig. 4f, g). The total combined thickness of this facies is approximately 45 m. Shale facies ShF5 This shale facies consists of pyritic shales. Mostly covered, it occurs in the upper part of the sections (between limestone units R and U, Fig. 3). However, these shales are well exposed in Foum Defili, where they appear as prominent dark grey to green pyritic shales of the Middle Ouarkziz Member, with secondary gypsum lenses. In the same area, in Foum Ferhrech (Conrad 1972a), gypsum is considered to be a primary precipitate and constitutes thick beds (meter scale). Owing to the thickness and the biostratigraphy in beds above and below, this interval is considered to be equivalent to the covered interval between limestone units R and S1 (Fig. 3). Sandstone facies The thickest sandstone interval (8–20 m thick, from eastern to western sections) directly related to carbonates occurs near the base of the logged section, above the limestone

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Table 1 Main lithofacies recognized in the Tinguiz Remz area Facies

Name

Main components

Main features

Depositional environment

ShF1

Shales and siltstones

Rare to absent

Pale to dark green. Absence of bioturbation and sedimentary structures, virtual lack of fossils

Prodelta mudstones

ShF2

Fine-grained sandy shales

Rare brachiopods, crinoids, plants

Interbedded with thin sandstones. Sedimentary structures common. Weakly bioturbated

Subtidal marine shales in the middle ramp, rarely affected by storms

ShF3

Shales and marls

Rare to absent

Reddish to yellow in color

Prodelta mudstones in the middle ramp, with precipitation of carbonates

ShF4

Bioturbated shales

Brachiopods, crinoids, trace fossils

Intensely bioturbated medium-grained shales interbedded with thin sandstones

Marine shales in subtidal deposits in the outer-middle ramp

ShF5

Pyritic shales

Absent

Dark grey to green pyritic shales, secondary gypsum lenses

Supratidal sabkha

SstF1

Coarse-grained sandstones

Absent

Thick packages of multiple level, thickening-upward, slight erosional base and cross-stratified

Delta front and distributary channels

SstF2

Mediumgrained sandstones

Absent

Tabular sand bodies, composed of few individual levels, without sedimentary structures or granulometries trends

Deltaic sand bodies

LstF1

Sandy encrinitic limestones

Crinoids, quartz grains, brachiopods, bryozoans, ostracods

Fragmented, parallel and cross-laminated packstones (rare wackestone and grainstone). Bioturbated

Incipient bioclastic/sand bars in subtidal middle ramp

LstF2

Laminated sandy/ bioclastic limestones

Quartz grains, peloids

Bioturbated sandy micrites with rare bioclastic packstone–wackestone bands, cross-lamination and waves ripples

Deltaic sands/carbonates in inter- to subtidal setting, commonly affected by wave action

LstF3

Oolitic limestones

Ooids

Large radial ooids and rare small superficial ooids. Fining-upward sequences

Oolitic shoals in inter- to supratidal inner to middle ramp

LstF4

Bioclastic and lithoclastic limestones

Crinoids, brachiopods, intraclasts, foraminifers, molluscs

Amalgamated wackestones to packstone with fining-upward sequences, low/ moderate fragmented bioclasts

Bioclastic accumulations in the inner ramp, subtidal to intertidal locally. Local erosion and redeposition of limestone clasts

LstF5

Brachiopodrich limestones

Brachiopods, crinoids

Wackestone to grainstone, floatstone/ rudstone, in some levels with common foraminifers, intraclasts, ooids, peloids. Cross-laminated and fining-upward sequences. Gigantoproductoid bands mostly in life position, and in the top of the beds

Brachiopods and bioclastic concentrations mostly in subtidal middle ramp where most are in life position. More rarely, also in intertidal settings. Deposits have commonly been resedimented by storms

LstF6

Biostromal limestones

Fasciculate rugose corals

Colonies mostly in growth position, within a crinoidal, brachiopod wackestone sandy matrix. Unsorted, low fragmentation

Autobiostromes in subtidal in the outer to middle ramp setting

LstF7

Parabiostromal limestones

Fasciculate, massive and solitary rugose corals, crinoids, brachiopods, foraminifers

Transported and broken colonies within a wackestone/packstone matrix, highly bioclastic, poorly sorted and fragmented

Autoparabiostrome in subtidal (rarely intertidal) middle ramp setting, randomly affected by storms

LstF8

Laminated dolomicrites

crinoids, brachiopods, coated grains, ostracods

Peloidal mudstones and dolomicrites with rare thin bioclastic bands. Cyanobacterial structures, fenestral porosity and mud cracks are common

High intertidal to supratidal sabkha

LstF9

Dolostone

Absent

Ghosts of structures and biota

Destructive diagenetic dolomitization of other limestones

unit 1 of the Betaina Fm. (Figs. 2a, 5a). However, there is a thicker sandstone unit, nearly 50 m thick, which is the first thick sandstone logged in the Betaina Fm. (identified as ‘base of sandstones’ in Figs. 2a, 4c). Most

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sandstones are recorded in the lower part of the section, although some thin sandstone horizons are exposed at higher levels (above limestone units H and L, and above O2, Fig. 3).

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At the top of the logged sections, non-marine sandstones of Namurian age are recorded in the Re´ouina Formation, but are not part of this study. The Westphalian is represented by the Merkala Fm., with conglomerates being more abundant than in the Re´ouina Fm. Both formations together constitute the Betana Beds (Fig. 1). Two types of marine sandstones are distinguished, based on differences in grain size and sedimentary structures. Sandstone facies SstF1 Medium- to coarse-grained, unfossiliferous, brown sandstones. The sandstones are mostly quartzarenites, with subangular to sub-rounded quartz grains, and rare feldspar and mica. Packages are 6–45 m thick and composed of multiple sandstone beds, 0.2–1.2 m thick, thicker in the upper part of each package (Figs. 4c, 5a). Lateral variation of the sandstone interval is observed at the level located just above unit 1 (Fig. 5a), which is 6 m thick in Tinguiz Remz (Section 3), but 30 m thick in the road section (Section 2). The lower sandstone beds are fine-grained with medium- to coarse-grained sandstones in the higher individual horizons. Erosive bases are rarely observed between sandstone horizons, as well as in the base of the package (Fig. 2b), with a sharp wavy contact (Fig. 5d). Cross-stratification is also commonly observed throughout the packages (Fig. 5c), locally with herring-bone structures (Fig. 5b). Sandstone facies SstF2 Medium-grained brown sandstones composed of a few beds, approximately 0.5 m thick, with non-erosional base and even top (Fig. 4a) and without any recognizable trend in granulometry. Composition is similar to SstF1. These sandstones occur mainly in the upper part of the Betaina Fm. and in the lower member of the Ouarkziz Fm., especially between limestone units 1–E and below unit N and above O2. No fossils or sedimentary structures were recorded.

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all three stratigraphic sections. However, some of the limestone beds are more local, such as units 2 and H2, which form lenses, which laterally extend for a few hundred meters (Figs. 2b, 3). The interval composed of units B–D is continuous through the region (Fig. 4a), although individual units and thickness are not always clearly distinguishable. Nine carbonate facies types are distinguished (Table 1) based on a combination of field observations and petrographic data (textures and components of Dunham’s 1962 classification scheme, modified by Embry and Klovan 1971) determined by thin-section analysis and summarized below. Limestone facies LstF1 Sandy encrinitic limestone (present in units 2, 1, C, D, E and N) in packages up to 2 m thick, composed of individual beds 0.2–0.3 m thick. Encrinites are medium- to coarse-grained, and are interbedded with shale or thin sandstone beds. At higher levels (units C–D), surficial traces are more common, particularly at the top of unit C (Fig. 6a), which is also characterized by horizontal lamination and wave ripples. Variations are observed between individual limestone samples, but the most prominent feature in thin-section is the abundance of sub-rounded/ sub-angular quartz grains, together with a high percentage of crinoid ossicles. Sandy limestones of the Betaina Fm. (unit 2, and locally, 1; Fig. 3) are mostly encrinitic packstones with large fragments of brachiopods, bryozoans and ostracods (Fig. 7a). Textures range from wackestone to grainstone and rudstone (Fig. 7b). Locally there are coarsening-upward sequences, but without erosional surfaces between beds. The lower bioclastic bands are moderately sorted, with horizontal laminations, a high degree of fragmentation and densely packed grains, whilst the grainstones are well sorted, moderately packed, and highly fragmented, composed mostly of crinoid ossicles (Fig. 7b). Higher levels in beds contain lower amounts of brachiopods, bryozoans and ostracods (Fig. 3).

Carbonate facies (limestones and dolostones) Limestone facies LstF2 Carbonate sediments are restricted to the upper 100 m of the Betaina Fm. and occur throughout the Ouarkziz Fm. (Fig. 3). No carbonates have been recorded below the logged sections in the Tinguiz Remz area. Carbonate units are more abundant and thicker in the upper part of the lower member and middle member of the Ouarkziz Fm. (between units M–R; Fig. 3). Siliciclastic input in the limestones is mostly restricted to the lower part of the logged sections (up to unit G), in which sandstone beds dominate (Fig. 3). Most limestone units extend laterally over a distance of at least 27 km, since they are recorded in

This facies is composed of laminated sandy and bioclastic limestones (present in units F, G, N, S2, S4 and S5). The facies type is mostly composed of fine- to medium-grained sandy micrite, with scarce peloidal and bioclastic layers. In units F and G, wave ripples and cross-lamination are observed at the top of individual beds, typically only 10–15 cm thick. Thicker sandy horizons (20–25 cm) contain hummocky cross-lamination (Fig. 6b). In unit N, individual beds are exposed in packages, 0.5–1 m thick, separated by thin (\1 cm) shale partings. The predominant

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Facies (2014) 60:941–962 b Fig. 4 a Fine-grained sandy shales in between thick sandstone ridges

in Section 1. Limestone beds A–D are highlighted in between two sandstone beds corresponding to SstF2 (sharp flat parallel base of sandstone with black arrows). Note 10-cm-thick sandstone beds intercalated between the shales (white arrows). b Yellow to red marls (black arrows) above green shale levels (white arrows) below limestone unit M (top of photo) in Section 3. c General view of the bioturbated sandy shales in Section 2. Base of sandstone is the same as in Fig. 2a. Note the thickening-upward of the upper sandstone ridge, as well as the common sandstone beds intercalated between the shales (each 10–30 cm thick) (for scale, the sandstone ridge is 45 m thick). d Bioturbated fine-grained sandstone with Zoophycus (Section 2). e Bioturbated fine-grained sandstone with Rhizocorallium (Section 2). f Common horizontal trace fossils on top of a thin finegrained sandstone in Section 2, in the higher bed, wave ripples are observed on top of the bed. g Detail of wave ripples in vertical section included within the bioturbated shales facies in Section 1 (height of picture = 9 cm). Hammer is 40 cm long

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and intraclasts, showing fining-upward trends (Fig. 7f) or concentrated in thin layers. Sorting and packing is moderate to high, with a high fragmentation, and common orientation of bioclasts parallel to the bedding plane.

Limestone facies LstF3

microfacies is bioturbated sandy mudstone (Fig. 7e), with scattered bioclasts. Bioclastic intervals are packstone–wackestone with diverse bioclasts, as well as very small superficial ooids, abundant peloids, and rare coated grains

Oolitic limestones (present in units H2, J and T), composed of uniform, massive beds 0.4–0.5 m thick, commonly amalgamated in packages 2.5 m thick. Rarely, a nodular aspect is observed in some beds. Large radial ooids (0.5–1 mm in diameter) are common in bed J (Fig. 7h), although small superficial ooids are also recognized in units H2 and T (0.1–0.2 mm in diameter). Bed J, with the large ooids, also is markedly fining-upward, and ooids at the top of the level are much smaller. The bioclast content is restricted to sparse molluscs, except in the nuclei of ooids, where archaediscid foraminifers are recorded (Co´zar et al. 2014a). Ooids coarsen-upward in grain-size.

Fig. 5 a Distributary delta channel with sandstone fill in Section 3; base is slightly erosional (over limestone unit 1) and the sequence is coarsening- and thickening-upward (hammer for scale in the lower right center of the photo). b Herring-bone cross-stratification in Sst1

at the base of sandstones in Section 1 (Figs. 2a, 4a). c Planar crossstratification in Sst1 at the base of sandstones in Section 1 (Figs. 2a, 4a). d Sharp, wavy contact at the base of delta front sandstones and prodelta mudstones in Section 2. Hammer is 40 cm long

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Facies (2014) 60:941–962 b Fig. 6 a Strongly bioturbated top of limestone unit C in Section 2.

b Hummocky cross-stratification in sandy limestone of unit F in Section 2. c Massive bioclastic limestone with a gigantoproductoid band in Limestone bed H in Section 2. Note that the brachiopod shells are mostly in the hydrodynamically stable position (convex-up). d Concentrations of gigantoproductoids in life position on top of Limestone bed L in Section 1. e Thalassinoides-type burrows filled by iron mineral in the firmground at the top of limestone unit N in Section 3. f Bioturbated surface at the top of limestone unit O2 with silicified solitary rugose corals and brachiopods in Section 3. g Biostromal limestone with all the colonies of Siphonodendron in growth position (way-up to the head of the hammer) in Section 2. h Parabiostromal limestone with overturned colonies of Actinocyathus (way-up to the head of the hammer). Hammer is 40 cm long

Limestone facies LstF4 Bioclastic and lithoclastic limestones (present in units K and N). These comprise thick-bedded limestones (0.8–2 m thick) with an abundant and highly diversified fauna (corals, brachiopods and crinoids), along with common levels of limestone intraclasts (Fig. 3). In thin-section, individual beds are rich in crinoids, foraminifers, brachiopods and bivalves, with sparse calcareous green algae (Fig. 8d). Dolomitization is conspicuous, replacing individual beds within limestone units. Some beds fine-upward and consist of low to moderately fragmented and well-sorted crosslaminated packstone in the lower part, whereas towards the top, more uniform parallel-laminated, poorly sorted, moderately fragmented and densely packed wackestone–packstone are developed. At the higher levels, lime mudstone and coated grains are observed (Fig. 8b). At the top of unit N, Thalassinoides-like structures occur (Fig. 6e). Limestone facies LstF5 Brachiopod-rich limestones (present in units B, E, H, L, O1, O2 and V). They are composed of individual beds of massive limestone (0.4–1.6 m thick), amalgamated in packages of 1–4.5 m and separated by millimetric shale partings. The occurrence of gigantoproductoid brachiopod layers is characteristic. Brachiopod shells are mostly in the hydrodynamically stable convex-up position at the base and middle part of the beds (Figs. 3, 6c), but also in life position (concave-up) at the top of beds (Fig. 6d). The units are commonly composed of several amalgamated limestone beds some of which are dolomitized. Textures are predominantly packstone, although wackestone (Fig. 8c), grainstone and floatstone (Fig. 7g) are also observed. Micritization of the bioclasts is common in most units, with abundant peloid layers alternating with bioclastic layers (Fig. 7d). Foraminifers are present in most beds. Horizontal lamination is common, with local crosslamination, and with fining-upward sequences. Sorting

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varies from moderate in the lower part of individual beds to well sorted in their upper parts, with a low packing density and moderate fragmentation of bioclasts (Fig. 7g). At the top of unit O2, associated with common remobilization of iron minerals, there are concentrations of silicified solitary corals, brachiopods, and burrows (Figs. 6f, 8c). Limestone facies LstF6 Biostromal limestone (present in units 1 and A). Individual beds are 0.5–1 m thick. Fasciculate colonial corals are the predominant fossils, with colonies mostly in growth position, and rarely overturned (Fig. 6g; see also Rodrı´guez et al. 2013a). The matrix is predominantly wackestone, bioturbated, with rare brachiopods and crinoids, and common quartz grains (Fig. 7c). Limestone facies LstF7 Parabiostromal limestones (present in units I, M, P and U). The limestones contain numerous colonial and solitary rugose corals, generally transported and broken (Fig. 6h; see also fig. 7 in Rodrı´guez et al. 2013a; fig. 4 in Rodrı´guez et al. 2013b). The matrix in between the corals is predominantly bioclastic-rich packstone–wackestone. Poor sorting, low fragmentation and packing of components are characteristic (Fig. 8a, d). Additional skeletal components are mainly crinoids and brachiopods, along with common foraminifers. Bioturbation occurs in some beds. Limestone facies LstF8 Laminated dolomicrite and peloidal limestones (present in units Q, R, S1, S3 and T). They occur in packages up to 8 m thick, individual dolomite beds commonly being 20–30 cm thick. The fossil content is generally poor, occurring as rare bioclastic layers which contain crinoids and foraminifers, as well as common ostracods. The thickness of the units decreases upwards (Fig. 3). Thin lamination (cyanobacterial mats) and fenestral porosity are present in units Q, R and S3 (Fig. 8e), as well as mud cracks in unit Q (Fig. 8f). In some units (S3 and T), mudstone and dolomicrite alternate with layers of micropeloidal limestone and a grapestone-like texture, and show fenestral and shelter porosities in thin-section (Fig. 8g). The content in bioclasts is variable (Fig. 8h), including crinoids, brachiopods, bryozoans, ostracods and rare foraminifers (particularly with concentrations of Miliolata). In the uppermost part of unit T a bioturbated packstone–grainstone with sparry calcite cement is present, predominantly composed of superficial ooids, small intraclasts and peloids.

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Facies (2014) 60:941–962 b Fig. 7 Carbonate microfacies of the lower half of the succession

(limestone units 2–J). Scale bar is 5 mm, except for H, which is 1 mm. Way-up to the top of the picture. a Crinoidal-quartz packstone, slightly shortened in flat parallel bands, unit 2, sample Pc4213. b Siliciclastic-rich packstone (base) passing up into crinoidal grainstone (upper part), unit 1, sample Pc4154. c Bioturbated wackestone, rich in quartz grains, rare crinoids and brachiopods in between the biostrome (see Siphonodendron corallites in the upper part of the picture), unit A, sample Pc4075. d Fine-grained bioclastic grainstone with microbacterial films (cortoids) and superficial ooids with coarse-grained brachiopod–bryozoan grainstone bands; erosional surfaces between bands are not marked and cross-laminations are highlighted by orientation of brachiopod shells, unit E, sample Pc4084. e Quartz-rich packstone, strongly bioturbated unit F, sample Pc4087. f Grainstone of micritized small bioclasts (bahamites), parallel laminated, with bands of well-preserved brachiopods (lower part), unit F, sample Pc4089. g Brachiopod–crinoidal wackestone to packstone (top), unit H, sample Pc4094. h Well-sorted ooidal grainstone, unit J, sample Pc4099

Limestone facies LstF9 Dolostones and dolomitized limestones (present in units H, H1, I, J, K, O2 and P). The bedded dolostones are more abundant and form thicker units (4–8 m) in the upper part of the section (unit O2; Fig. 3). However, individual dolomitized horizons are also present within the thinner limestone units (typically 1–4 m thick). Destructive dolomitization left only some ghosts of bioclasts. Laterally, in some of these beds textures are preserved, which correspond to bioclastic packstones and wackestones.

Interpretation of the depositional environments Mixed carbonate–siliciclastic successions are best developed during times of high-amplitude sea-level change, and are typical of icehouse conditions (Catuneanu et al. 2011). The succession in the Tinguiz Remz area depicts transgressive–regressive (T–R) cycles, which can be compared to typical Yoredale mixed siliciclastic–carbonate cycles (Tucker et al. 2009; Catuneanu et al. 2011), although oversimplified (lack of coal seams in Morocco and most of the cycles are sandstones and shales in the Carboniferous of northern England). Shaly facies ShF1 (shales and siltstones) and ShF3 (shales and marls) are considered to represent similar environments, recording so-called prodelta mudstones (shales and siltstones in the Tinguiz Remz area) associated with distal sediments in deltaic deposits. This continental origin as source for most of the shales and siltstones is suggested by the absence or virtual absence of marine fossils and of bioturbation. In contrast, facies types ShF2 (fine-grained sandy shales) and ShF4 (bioturbated shales) suggest an open marine environment, mostly in subtidal conditions within

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an outer-middle ramp setting. The pyritic shales with gypsum (ShF5) do not crop out in the sections studied. However, although most of the gypsum in Tinguiz Remz area seems to be diagenetic, there are published data of primary gypsum close to Foum Defili, and together with the interbedded laminated algal mats and mud cracks of LstF8, they are considered as sabkha deposits, as already suggested by Conrad (1972b) and Cavaroc et al. (1976). Sandstones are predominant in the lower part of the sections, where the SstF1 are interpreted as delta front sandstones in the lower part of each level, and as distributary channel fills in their upper parts (Figs. 4c, 5a). Individual channels can be traced for tens of meters, whereas the poorly developed sandstone beds of SstF2 are interpreted as proximal delta front deposits (Tucker 2003). Within the limestone facies, some units have been interpreted as bioclastic and oolitic bars and shell concentrations (LstF1, LstF2, LstF3, and LstF4) in subtidal to supratidal settings (Table 1). Brachiopod-rich limestones (LstF5) characterize the typical carbonate background sedimentation in the inner ramp, where gigantoproductoids in life position were not affected by storms, whereas shells in stable hydrodynamic positions were commonly resedimented by storms. The interpretation of biostromal and parabiostromal limestones (LstF6 and LstF7 respectively) is similar to that of the gigantoproductoid beds, with the biostromal limestone developed in subtidal conditions and unaffected by storms in the lower part of the sections, whereas the parabiostromes in the middle and upper part of the sections were affected by storms. The laminated dolomicrite (LstF8), with fenestral porosity, peloidal textures, cyanobacterial mats and mud cracks is considered to have been deposited in high intertidal to supratidal settings. Most of the siliciclastic-rich Betaina Fm. can be considered to represent outer-middle ramp/shelf deposits, characterized by common bioturbation, small brachiopods and ammonoids (ShF4), typical of prodelta deposits. The interbedded limestone units with high amounts of siliciclastic grains (LstF1, mostly quartz) mark this continental influence in the lower part of the succession and display the transition from the distal delta to the proximal delta front. The lower calcareous units were commonly reworked by storms, with muddy matrices, and thus a deposition within the middle shelf/ramp, above storm wave-base is inferred (Fig. 3). This environment also favored the concentration of brachiopod layers (LstF5) as well as the development of coral biostromes (LstF6). This middle shelf/ramp setting prevailed up to unit E, from incipient stages of the development of carbonate sand shoals within the cross-laminated brachiopod-rich packstone (LstF5), interpreted as subtidal deposits. The depositional environment shallowed to the top, with the moderately sorted ooid grainstone indicative of an inner shelf setting, above the fair-weather wave-base.

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Facies (2014) 60:941–962 b Fig. 8 Carbonate microfacies of the upper part of the succession

(limestone units M–V). Scale bar is 5 mm. Way-up to the top of the picture. a Bioclastic packstone–wackestone, rich in foraminifers, unit M, sample Pc4120. b Bioturbated wackestone rich in coated grains and ostracods, top of unit N, sample Pc4127. c Crinoidal–brachiopod wackestone, with burrows filled by micrite, unit O2, sample Pc4135. d Wackestone–packstone with recrystallized matrix, rich in brachiopods, micritized bioclasts, molluscs and foraminifers, unit P, sample Pc4141. e Dolomudstone with cyanobacterial mats and fenestral porosity, unit Q, sample Pc4146. f Dolomudstone with cyanobacterial mats and mudcracks, bed Q, sample Pc4147. g Micropeloidal wackestone rich in brachiopods and molluscs, unit S3, sample Pc4157. h Mudstone (lower part) passing up to crinoidal–brachiopodgrapestone grainstone (upper part); contact is sharp, unit S3, sample Pc4158

This regression culminated in unit F, where LstF2 occurs. There is a predominance of wave ripples and crossstratification in each bed, with moderate to high sorting and highly fragmented bioclasts pointing to deposition close to the intertidal zone, although locally hummocky crossstratification occurs. This facies likely developed in an inner shelf/ramp setting and generated a shoal relief, being affected by waves and tides. A similar depositional setting is inferred for units H2 and J (LstF3) owing to the common oolitic grainstone facies, although in those, clearer intertidal conditions can be inferred. This facies can be interpreted as the development of carbonate sand shoals in an inner shelf/ramp setting. The interval comprising limestone units K–P is interpreted to represent a more stable subtidal facies, with thicker carbonate deposits (Fig. 3). These carbonates have been generally affected by storms, which seem to have been locally important within levels of units N, P and mostly in O1 (LstF 4, 5 and 7). Sedimentary structures are mostly absent, but the high biodiversity suggests a quieter environment in a middle shelf/ramp setting, between storm wave-base and fair-weather wave-base, and with little or no siliciclastic input. Units N and O2 exhibiting a firmground at their tops with iron crusts covering the substrate and bioclasts. They are intensely bioturbated and seem to represent the deepest conditions (Fig. 6e, f). The interval between units Q–T represents a marked change in the depositional setting of the shelf/platform, with a predominance of unfossiliferous laminated and cross-laminated limestone and dolomicrite with fenestral cavities, and horizons with ooids (LstF8, LstF2 and LstF3). Intertidal to supratidal conditions in the inner shelf/ramp are inferred, with development of sabkhas at some carbonate levels, even in shales (gypsum described close to Foum Ferhrech, ShF5). However, oscillations with shallow subtidal environments are indicated by the presence of thin bioclastic beds (such as units S3 and S4), which contain a stenohaline fauna with crinoids, brachiopods, foraminifers, and bryozoans (Fig. 3).

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At the top of the Ouarkziz Fm., coral biostromes (unit U, facies Lst7) and brachiopod-rich beds (unit V—Titanaria bed, facies LstF5) are interpreted to represent a return to subtidal conditions in the mid shelf/ramp (Fig. 3), as they are commonly affected by storms with broken coral colonies, as well as in situ colonies (Rodrı´guez et al. 2013b). The limestones also contain a rich and diverse fauna of crinoids, brachiopods, foraminifers, and calcareous algae (Fig. 3). The sandstones of the Re´ouina Formation constitute the end of marine sedimentation and the commencement of continental conditions within the Tindouf Basin, interpreted as fluvio-deltaic facies, which has been studied in detail by Cavaroc et al. (1976). Most limestone units of the Ouarkziz Fm. are interpreted to form the base of T–R cycles. Owing to the duration of the Brigantian and Serpukhovian, they can be considered as high-frequency fourth-order cycles. However, this is not clear enough for the late Asbian with common interbedded deltaic sandstone beds and the early Bashkirian, and these rocks need to be further studied. Nevertheless, a generalized regression can be suggested which culminated in limestone units F and G (LstF2), interpreted to be one of the shallower carbonate facies in the middle part of the succession, close to the intertidal setting. From unit G to unit O2, it is possible to recognize a progressive deepening in the carbonate facies, gradually up to limestone unit K, and more marked from limestone unit L upwards. From the top of unit N to unit Q (laminated dolomicrite), it can be defined as a rapid regression event, which culminated in subaerial exposure on unit Q, marked by the development of mudcracks. The platform was relatively stable for a long period, when all the carbonate facies (and the inferred shales) were representative of shallow-water deposits. Limestone units U and V, however, represent a gradual transgression. From the top of unit V, a regression is observed in the succession, with deposition of non-marine fluvial–deltaic sandstones. In total, a basal regressive event and two large transgressive–regressive cycles are recognized in the Tinguiz Remz succession, with two maximum flooding surfaces on top of limestone units N and U, respectively. Considering that their duration spans the late Asbian to early Bashkirian (333–322 Ma ca. 11 Myr, according to data in Davydov et al. 2010; Fig. 3), they correspond to third-order sequences.

Correlation between the northern and southern flanks of the Tindouf Syncline The proposed correlation between the northern and southern flanks of the Tindouf Syncline is based on a

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combination of sedimentological and paleontological data (mostly brachiopods and foraminifers). Gigantoproductoid brachiopods were described in detail from the northeastern sector of the northern flank of the Tindouf Syncline at Foum Defili by Legrand-Blain (1973, 1987) and foraminifers by Lys (1979) (Fig. 9). Some levels were correlated by Conrad and Legrand-Blain (1971), Conrad (1972a, b), Legrand-Blain (1973, 1987) and Co´zar et al. (2014a). In Co´zar et al. (2014a), it was assumed that limestone units 1 and 2 are not present in Foum Defili, and thus the horizons of Gigantoproductus africanus (Stache, 1883) [which also corresponds to the ‘‘Premier niveau a` Productus no. 10’’ by Mamet et al. (1966), dated as V3b– V3c, based on brachiopods] were correlated with the limestone units A–E (see arrows in Fig. 9). Furthermore, the comparison between both sections shows that limestone units A–D do not occur in the Foum Defili section (Fig. 9). The proposed base of the Ouarkziz Fm. at Foum Defili is therefore somewhat imprecise, because it was defined by the first appearance of limestones in the succession, but it is recognized that in the northwestern part of the northern flank at Tinguiz Remz and in the southern flank limestone beds occur within the underlying Betaina Fm. A more reliable marker bed is the thick sandstone unit in the upper part of the Betaina Fm., which forms the main ridge of the Djebel Ouarkziz Mountain range in Tinguiz Remz, and appears to be also present at Foum Defili (Figs. 2a, 9). To allow a more accurate correlation of the Ouarkziz Fm. both at Foum Defili and Tinguiz Remz, the base of this formation should be placed at the top of the sandstone bed below limestone E. This carbonate unit contains similar foraminiferal assemblages in both regions (Co´zar et al. 2014a). Furthermore, the limestone units B and E, which represent the two lower Gigantoproductus beds in Tinguiz Remz coincide with the two limestone beds of G. africanus in Foum Defili (‘a’ and ‘b’ in Fig. 9). This modification implies that there are more numerous limestone beds at the top of the Betaina Fm. in the northwestern sector as previously supposed. In addition, the age of the Ouarkziz Fm. would range from the Brigantian to the Krasnopolyanian (Fig. 9), and the late Asbian is recorded only in the Betaina Fm. Some younger beds containing brachiopods can be correlated as well. For instance in the northeastern area of Djebel Ouarkziz range, levels with Gigantoproductus sp. 1 (V3c sup. = latest Vise´an) were interpreted as equivalent to unit H in Co´zar et al. (2014a), and levels with Gigantoproductus menchikoffi Legrand-Blain, 1973 (V3c sup.) were considered as probably equivalent to unit L in Co´zar et al. (2014a) (Fig. 9). However, gigantoproductoids seem to be less abundant at Tinguiz Remz than at Foum Defili (Fig. 9), so the level characterized by Gigantoproductus meridionalis Legrand-Blain, 1973 (dated as uppermost

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Vise´an or Namurian) is difficult to correlate, due to the absence of reliable foraminiferal markers. However, as described in Co´zar et al. (2014a), early Serpukhovian foraminifers occur from unit L at Tinguiz Remz, which permits a more precise correlation. In the absence of brachiopod markers, the upper beds of the lower member of the Ouarkziz Fm. described by Lys (1979) at Foum Defili were dated as E1 by foraminifers, and can be correlated with units N, O, and P in Tinguiz Remz. They include the upper part of the early Serpukhovian and lower part of the late Serpukhovian. Lys (1979) recognized the Namurian Zone H (currently lowermost Bashkirian; Fig. 9) from the coral biostrome with ‘‘Lithostrotion’’ at the top of the succession (bed U here) based on abundant Pseudoglomospira? subquadrata Potievskaya and Vakarchuk in Brazhnikova et al. (1967). These authors identified common Eosigmoilina and Neoarchaediscus postrugosus (Reitlinger, 1949) at the level above, containing the brachiopod Titanaria africana. These foraminifers were used as markers of the Namurian Zone H by Lys (1979) in most of the North African basins. However, Mamet et al. (1966), Conrad and Legrand-Blain (1971), Conrad (1972a, b), and Legrand-Blain (1987) assigned both horizons to the E2 Zone (Serpukhovian). Those levels correlate with units U and V, respectively (Fig. 9), although Pseudoglomospira? is very rare, but Hemigordiellina is abundant, and thus the two homeomorphous taxa might have been misidentified. Foraminifers described in Co´zar et al. (2014a) from units U and V contain rich assemblages (Fig. 10), of which the majority are typically Bashkirianin age, including several species of Globivalvulina, Bradyina, ‘‘Biseriella’’, Semistaffella, transitional forms to the genera Pseudostaffella and Varistaffella, Hemigordius, Seminovella, Millerella, and Iriclinella. In addition, Semistaffella variabilis (Reitlinger, 1961) is the foraminiferal marker for the Krasnopolyanian Substage (Fig. 9). Furthermore, the occurrence of the conodont Idiognathoides is also a marker of the Krasnopolyanian or younger substage (Co´zar et al. 2014a). The presence of forms transitional to the genera Pseudostaffella and Varistaffella would suggest an advanced Krasnopolyanian age (Fig. 9), because the typical representatives of both genera are markers of the Severokeltmian Substage (=Akavassian Substage of the Urals). The correlation between the northeastern and northwestern sectors of the northern flank of the syncline shows that deposition of limestones occurred earlier in the western sector, within the upper 100 m of the Betaina Formation (Fig. 9), but both successions can be perfectly correlated, and thus a parallel tectonic and sedimentological evolution, can be assumed for both areas. There are scarce data from the southern flank of the syncline, and only Sebbar (2000) and (Sebbar et al. 2000)

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Fig. 9 Correlation of the northern flank (Foum Defili, modified from c Conrad 1985) and southern flank (modified from Sebbar et al. 2000) of the Tindouf Syncline. Arrows 1–7 are levels of ammonoids and brachiopods in Foum Defili (modified from Legrand-Blain 1973): 1 Beyrichoceras micronotum, 2 Gigantoproductus africanus (level a), 3 Gigantoproductus africanus (level b), 4 Gigantoproductus sp. 1, 5 Gigantoproductus menchikoffi, 6 Gigantoproductus meridionalis, 7 Titanaria africana

studied the microfacies and biostratigraphy of some sections close to the margin of the Reguibat Shield (Fig. 9). The lithofacies and sedimentological model proposed here does not coincide precisely with that of the southern flank when biostratigraphic intervals are closely compared. The outer shelf/ramp deposits in Tinguiz Remz are represented in the southern flank by the Tazout and Betaina formations, locally passing into intertidal conditions in the lower part of the Ouarkziz Fm. A middle shelf/ramp environment was inferred for the upper part of the lower Ouarkziz Member. No data were presented for the middle and upper members of the Ouarkziz Fm. Based on a lithological correlation, the main difference between both limbs of the Tindouf Syncline is the earlier occurrence of limestones and more common sandy limestones in the south (Tazout Sandstone Fm. and from the base of the Betaina Fm.). This higher siliciclastic input in the southern flank can be explained by the proximity to the emergent landmass of the Pan-African Reguibat Shield to the south (Fig. 9), where the Paleozoic rocks rest on the crystalline basement (Villeneuve 2005). Lack of similarities between both flanks is also due to the absence of thick sandstone bodies with sedimentary structures (indicative of distributary channel fills or delta front deposits). In fact, the only interval that could be biostratigraphically correlated between both areas (Fig. 9) is the late Asbian; it is mostly composed of limestone beds in the southern flank, whereas there is a predominance of thick sandstone beds further to the north. The Serpukhovian in the southern flank is represented exclusively by shales and marls. This age—facies relationship is consistent with the original Namurian age assigned to the Middle Ouarkziz Member in Tinguiz Remz by (Mamet et al. 1966). However, no samples were studied from these upper levels by Sebbar (2000) and Sebbar et al. (2000). In fact, the biostratigraphic scheme proposed by the previous authors posed some problems. It includes an amalgamation of datings from Foum Defili (Lys 1979) and Tinguiz Remz (Mamet et al. 1966) from both the northern flank and their own investigations of the southern flank, as well as unpublished borehole data by Pelipenko (1990) close to the axial part of the Tindouf Syncline. Furthermore, the precise location of each sample or foraminiferal

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Fig. 10 Selected foraminiferal markers in the Tinguiz Remz area (scale bar 100 microns). a Endothyranopsis crassa (Brady) emend. Cummings, unit H. b Eosigmoilina explicata elongata Ganelina, unit S3. c Neoarchaediscus incertus (Grozdilova and Lebedeva), unit J. d Neoarchaediscus aff. postrugosus (Reitlinger), unit J. e Globivalvulina moderata Reitlinger, unit U. f Eostaffellina ‘‘paraprotvae’’ (Rauser-Chernousova), unit M. g Eostaffella pseudostruvei (Rauser-Chernousova and Beljaev), unit M.

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h Plectostaffella jakhensis Reitlinger, unit U. i Plectostaffella varvariensis (Brazhnikova and Potievskaya), unit U. j Globivalvulina bulloides (Brady), unit U. k ‘‘Biseriella’’ scaphoidea (Reitlinger), unit U. l Seminovella elegantula Rauser-Chernousova, unit U. m Semistaffella minuscularia Reitlinger, unit U. n Iriclinella spirilliniformis (Brazhnikova and Potievskaya), unit U. o Semistaffella variabilis (Reitlinger), unit U. p transitional to Varistaffella, unit U. q transitional to Pseudostaffella, unit U

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stratigraphic ranges were not provided, except for those of samples 21349–21389 in the Aouinet bel Legraˆa-Ain el Berka section (Fig. 9) in Sebbar (2000). It must be also noted that no foraminifers were illustrated from Tindouf in Mamet et al. (1966), Lys (1979, 1985), Sebbar (2000) or Sebbar et al. (2000). In the Aouinet bel Legraˆa-Ain el Berka section few limestones contain foraminifers, which generally are not very diagnostic. The late Vise´an age of the rocks can be confirmed from sample 21369 (Fig. 9) due to the occurrence of bilaminar paleotextulariids. This bed contains the diagnostic foraminifer Endothyranopsis crassa (Brady, 1870) emend. Cummings, 1955, which might indicate a late Asbian age (Conil et al. 1980, 1991; Co´zar and Somerville 2004; Somerville and Co´zar 2005; Co´zar et al. 2008). However, the listed E. crassa does not seem to be reliable enough. The single specimen illustrated in Sebbar (2000, pl. 8, fig. 9) from the Saoura Valley in Be´char (Algeria) should be transferred to Endothyranopsis sp. or could be a deformed E. compressa (Rauser-Chernousova and Reitlinger in Rauser-Chernousova et al. 1936). Certainly, there is a specimen of E. crassa illustrated in Sebbar (2000, pl. 8, fig. 10) from the late Serpukhovian Bent el Goumi Fm. in the Mezarif Massif, but it was misidentified as E. sphaerica (Rauser-Chernousova and Reitlinger in Rauser-Chernousova et al., 1936). In contrast, Endothyranopsis compressa (sensu Sebbar 2000, pl. 8, fig. 4) from the late Bashkirian of the Oued el Hamar Fm., is a Plectogyranopsis ampla (Conil and Lys, 1964). The identification of primitive Neoarchaediscus was the key taxon for the recognition of the base of Zone 16 inf., whereas N. incertus (Grozdilova and Lebedeva 1954) was used for defining the Zone 16 sup. As mentioned above, no illustration of these taxa exists, and in fact, the latter species was only illustrated in Sebbar (1998) from Reggan, but it should be transferred to Neoarchaediscus sp. The listed N. incertus in the Aouinet bel Legraˆa-Ain el Berka section was recorded in the topmost sample (21389; Fig. 6). Regarding the rest of the assemblage, also because of the presence of Bradyina sp., this upper part of the section might be no younger than the early Brigantian (Fig. 6), and could even be latest Asbian in age, where Bradyina and Neoarchaediscus first occur together (Conil et al. 1980, 1991; Co´zar and Somerville 2004; Somerville and Co´zar 2005; Somerville 2008). Sebbar (2000) and Sebbar et al. (2000) identified only two possible cycles in the time-equivalent lower Ouarkziz Member, interpreted as deposits on a middle ramp setting. Regarding their most probable age, of late Asbian to earliest Brigantian (?), these cycles could not be correlated with the lower levels of Tinguiz Remz, and thus correspond just to the Betaina Fm., where only a regressive event is recognized in the northern flank. However, apart from that

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age determination for a short interval, few data are shared by the southern and northern limbs, which could support this correlation (Fig. 9).

Discussion The poor similarities between the Vise´an of the southern and northern flanks of the Tindouf Syncline is difficult to interpret. The Serpukhovian succession in both flanks seems to be even more distinct, because Sebbar et al. (2000) did not study Serpukhovian rocks. They assumed a thick interval exclusively composed of shales (Fig. 9) based on unpublished data by the Bureau des Recherches du Pe´trole. In order to reconstruct the basin, this discrepancy might be attributed to strong lateral facies changes across the basin, or to more significant synsedimentary tectonic activity than has been assumed previously (Sebbar et al. 2000). Transgressions and regressions during this period (prograding to the east/northeast) cannot justify the asymmetry observed between the northern and southern flanks. To consider a lateral facies change across the basin, the outer ramp sandy limestone of Sebbar et al. (2000) has to be re-interpreted as shallower middle ramp deposits. Thus, the environments will be more similar in both flanks, and the presence of more sandy limestones could reflect the proximity of the emergent land of the Reguibat Shield, and a shallower-water facies with a higher siliciclastic input. This model would imply a nearly west–east-oriented shoreline close to the southern flank, and an epicontinental platform slightly deepening to the north, where it should be connected with the Paleo-Tethys Ocean. This model would agree with that proposed by Villeneuve (2005) and Guiraud et al. (2005), with the deepest water facies far from the Reguibat Shield. Michard et al. (2008) also considered the source of the siliciclastics to be the Reguibat Shield. However, our study and that of others (Hollard 1970; Cavaroc et al. 1976) do not support that model. Although according to Cavaroc et al. (1976), the Reguibat and Ougarta structures to the south and east, respectively (Fig. 9) were considered as slightly submerged, they acted as barriers to confine the sand input from the north and northwest. Hollard (1970) envisaged a source of siliciclastics to the northeast of the syncline, close to the Ougarta Mountains (Fig. 9), and Cavaroc et al. (1976) considered as an additional source of siliciclastic sediments, a partially exposed and eroded sandstone complex to the northwest of the basin, in the region of the Anti-Atlas Mountains. The lithological features, maturity and degree of metamorphism of the upper Devonian–lower Carboniferous sandstones in the Tindouf Basin were similar to those observed in the Anti-Atlas, and different to the Precambrian Reguibat

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Shield (Cavaroc et al. 1976). They also observed a migration basinward (from north to south) of the Devonian and Carboniferous depocenters, which was interpreted as a response to a progressive uplift of the Anti-Atlas region. However, as those authors admitted, there is no evidence of subaerial exposure of the crystalline core in the Anti-Atlas. There is a general consensus that the Anti-Atlas was uplifted, and thus eroded, from Serpukhovian time upward (e.g., Cavaroc et al. 1976; Fabre 2005; Michard et al. 2008). However, it is not clear when this process started (it is generally dated imprecisely as ‘late Vise´an–Namurian’), and whether this uplift event involved some parts of the Anti-Atlas and the main crystalline core. Choubert (1952) documented the discontinuity between the upper Vise´an and older Devonian and Lower Carboniferous rocks. This event was later included within an intra-Vise´an tectonic pulse (e.g., Pique´ 2001; Hoepffner et al. 2005). This pulse was the cause for the exhumation of a Precambrian inlier, the Ouzellarh area (Fig. 9), which was considered as exposed and eroded during the late Vise´an by Choubert (1952). This region is part of the so-called Sagro Block (Fig. 9) by Michard et al. (1982), an uplifted area during the late Vise´an. Hoepffner et al. (2005) also considered an uplifted Proterozoic Anti-Atlas as the source of the Devonian–Carboniferous siliciclastics, including areas close to Tafilalt and Maider (Fig. 9), although the source of the siliciclastics in those basins was considered to be located in the south by Wendt (1985). Michard et al. (2008), although considering the Reguibat Shield as the source of the siliciclastics, assumed synsedimentary fault tectonics along the Anti-Atlas during the ‘late Vise´an– Namurian’, with the uplift of a ridge resulting in bioclastic limestones, chaotic breccias and flysch-like deposits. Recently, Sebti et al. (2009) demonstrated using zircon fission track dating a heating episode due to tectonic events and a subsequent rapid cooling (interpreted as exhumation and erosion) in the Ifni and Kerdous Proterozoic inliers (Fig. 9). This area corresponds to that interpreted as the source of the siliciclastics northwest of the Tindouf Basin by Cavaroc et al. (1976). The age of the heating and subsequent cooling was constrained to 330–320 Ma, although some of the peaks were dated as early as 337 Ma (mean age). This process would involve the late (latest) Vise´an, Serpukhovian, and base of the Bashkirian (Fig. 3). In summary, it could be considered that from the Serpukhovian onward, the Anti-Atlas was uplifted, in vast tracts, whereas during the late Vise´an, only some smaller inliers seem to be uplifted. Analysis of the early Carboniferous paleocurrents suggests a transport direction from west to east with the river valleys of the continental Re´ouina and Merkala formations draining southwards (Cavaroc et al. 1976), which does not solve per se the problem of an uplifted late Vise´an Anti-Atlas. However,

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during the late Vise´an, distributary channels and mouth bars (delta front deposits) are only recorded in the northern flank of the syncline. The presence of those proximal deltaic facies suggest an uplifted and eroding source to the north, in the Anti-Atlas, which supports the interpretation by Cavaroc et al. (1976). It is difficult to establish precisely how many of the Proterozoic inliers of the Anti-Atlas were emergent, but during the late Vise´an, the Anti-Atlas can be readily considered as an emerging structure, with a certain influence on the paleobiogeography of the region (Somerville et al. 2013; Co´zar et al. 2014b). Thus, the classical model with a single epeiric platform with the shallower facies forming an onlap on the Reguibat Shield and deeper facies to the north (e.g., Foum Defili, Tinguiz Remz) is not sustained by the data. Instead, a basin with parallel platforms across the Reguibat Shield and the Anti-Atlas, and a central trough or at least deeper water facies, seem to be a more probable paleogeographic model. Despite this earlier tectonic activity for the Tindouf region and at least local uplifted inliers, a relationship with the Ahnet–Reggan–Mouydir basins is not recognized. There, Wendt et al. (2009, 2010) considered a vast uplifted area during the middle to late Vise´an and the entire Serpukhovian, with Bashkirian carbonate rocks directly overlying the early Vise´an siliciclastics. This model has been already questioned (Legrand-Blain et al. 2010). The stratigraphical record and context seem to be different to those in the Tindouf Basin, which prevents any further comparison.

Conclusions The northwestern part of the northern flank of the Tindouf Syncline contains a mostly siliciclastic succession, with local intercalated carbonate units in the late Vise´an, Serpukhovian and lowermost Bashkirian. The succession is notably cyclic, and represents a major regressive event from outer platform sediments in the lower part (mostly Tournaisian and Vise´an), with middle and inner platform facies higher up, and continental sandstones from the Bashkirian capping the marine sedimentation. Internally, high frequency cyclicity is also recognized. The depositional environments in the Tinguiz Remz area are shallower than those of the counterpart in the southern flank of the syncline. The northern flank is noteworthy for the occurrence of delta front and distributary channel fills deposits during the late Vise´an. The environments are representative of a proximal facies, which suggest the presence of an uplifted source area close to the Anti-Atlas region, earlier than was previously supposed, from the late Vise´an onwards. The size of this uplifted area is not constrained, which could only imply small inliers within the

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Anti-Atlas axis, but an emerging Anti-Atlas from the late Vise´an onward is supported by the spatial facies distribution in the northern flank of the Tindouf Basin. This uplift was probably a geographic barrier for faunal migration between the Moroccan Meseta to the north and the Sahara to the south, which partially isolated both areas, leading to some differences in their faunal assemblages. Acknowledgments We would like to thank Dr Elias Samankassou and an anonymous reviewer for their constructive comments. We are also grateful to Franz Fu¨rsich for his editorial expertise. Fieldwork funded by the Spanish Ministerio de Economı´a y Competitividad (research project CGL2012-30922BTE) was carried out thanks to the collaboration of the Department of Mine Development of the Ministe`re de l’Energie et des Mines of Morocco.

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